Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup (Geological Society Special Publication No. 282)

Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup

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) PHIL LEAT (UK) NICK ROBINS (UK) JONATHAN TURNER (UK) Society Books Advisors

MIKE BROWN (USA) ERIC BUFFETAUT (FRANCE )

RANDELL

RETO GIERE´ (GERMANY ) JON GLUYAS (UK) DOUG STEAD (CANADA ) STEPHENSON (THE NETHERLANDS )

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: KARNER , G. D., MANATSCHAL , G. & PINHEIRO , L. M. (eds) 2007. Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282. HOPPER , J. R., FUNCK , T. & TUCHOLKE , B. E. 2007. Structure of the Flemish Cap margin, Newfoundland: insights into mantle and crustal processes during continental breakup. In: KARNER , G. D., MANATSCHAL , G. & PINHEIRO , L. M. (eds) Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 47– 61.

GEOLOGICAL SOCIETY SPECIAL PUBLICATION NO. 282

Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup

EDITED BY

G. D. KARNER, ExxonMobil Upstream Research Company, Houston, USA

G. MANATSCHAL Universite´ Louis Pasteur, Strasbourg, France and

L. M. PINHEIRO Universidade de Aveiro, Aveiro, Portugal

2007 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 2007. All rights reserved. No reproduction, copy or transmission of this publication may be made without written permission. No paragraph of this publication may be reproduced, copied or transmitted save with the provisions of the Copyright Licensing Agency, 90 Tottenham Court Road, London W1P 9HE. Users registered with the Copyright Clearance Center, 27 Congress Street, Salem, MA 01970, USA: the item-fee code for this publication is 0305-8719/07/$15.00. British Library Cataloguing in Publication Data A catalogue record for this book is available from the British Library. ISBN: 978-1-86239-228-1 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 þ 1 800-972-9892 Fax þ 1 802-864-7626 E-mail [email protected] For individual and corporate orders: AAPG Bookstore, PO Box 979, Tulsa, OK 74101-0979, USA Orders: Tel þ 1 918-584-2555 Fax þ 1 918-560-2652 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 KARNER , G. D., MANATSCHAL , G. & PINHEIRO , L. M. Imaging, mapping and modelling continental lithosphere extension and breakup: an introduction The Iberia– Newfoundland continental extensional system (geological and geophysical constraints) TUCHOLKE , B. E., SAWYER , D. S. & SIBUET , J.-C. Breakup of the Newfoundland–Iberia rift HOPPER , J. R., FUNCK , T. & TUCHOLKE , B. E. Structure of the Flemish Cap margin, Newfoundland: insights into mantle and crustal processes during continental breakup SIBUET , J.-C., SRIVASTAVA , S. P., ENACHESCU , M. & KARNER , G. D. Early Cretaceous motion of Flemish Cap with respect to North America: implications on the formation of Orphan Basin and SE Flemish Cap –Galicia Bank conjugate margins RESTON , T. J. The formation of non-volcanic rifted margins by the progressive extension of the lithosphere: the example of the West Iberian margin The Iberia– Newfoundland continental extensional system (dynamic modelling) HUISMANS , R. S. & BEAUMONT , C. Roles of lithospheric strain softening and heterogeneity in determining the geometry of rifts and continental margins BUROV , E. The role of gravitational instabilities, density structure and extension rate in the evolution of continental margins HARRY , D. L. & GRANDELL , S. A dynamic model of rifting between Galicia Bank and Flemish Cap during the opening of the North Atlantic Ocean The Iberia– Newfoundland continental extensional system (kinematic modelling) EGAN , S. S. & MEREDITH , D. J. A kinematic modelling approach to lithosphere deformation and basin formation: application to the Black Sea HEALY , D. & KUSZNIR , N. J. Early kinematic history of the Goban Spur rifted margin derived from a new model of continental breakup and sea-floor spreading initiation Observational characteristics of non-Atlantic extensional systems (offshore) GOODLIFFE , A. M. & TAYLOR , B. The boundary between continental rifting and sea-floor spreading in the Woodlark Basin, Papua New Guinea DIREEN , N. G., BORISSOVA , I., STAGG , H. M. J., COLWELL , J. B. & SYMONDS , P. A. Nature of the continent– ocean transition zone along the southern Australian continental margin: a comparison of the Naturaliste Plateau, SW Australia, and the central Great Australian Bight sectors Observational characteristics of non-Atlantic extensional systems (onshore) COCHRAN , J. R. & KARNER , G. D. Constraints on the deformation and rupturing of continental lithosphere of the Red Sea: the transition from rifting to drifting MANATSCHAL , G., MU¨ NTENER , O., LAVIER , L. L., MINSHULL , T. A. & PE´ RON -PINVIDIC , G. Observations from the Alpine Tethys and Iberia–Newfoundland margins pertinent to the interpretation of continental breakup ROBERTSON , A. H. F. Overview of tectonic settings related to the rifting and opening of Mesozoic ocean basins in the Eastern Tethys: Oman, Himalayas and Eastern Mediterranean regions Revisiting fundamental concepts of continental extension KUSZNIR , N. J. & KARNER , G. D. Continental lithospheric thinning and breakup in response to upwelling divergent mantle flow: application to the Woodlark, Newfoundland and Iberia margins

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9 47 63

77

111 139 157

173 199

217 239

265 291

325

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CONTENTS

CHRISTIE -B LICK , N., ANDERS , M. H., WILLS , S., WALKER , C. D. & RENIK , B. Observations from the Basin and Range Province (western United States) pertinent to the interpretation of regional detachment faults DYKSTERHUIS , S., REY , P., MU¨ LLER , R. D. & MORESI , L. Effects of initial weakness on rift architecture MORESI , L., MU¨ HLHAUS , H.-B., LEMIALE , V. & MAY , D. Incompressible viscous formulations for deformation and yielding of the lithosphere Index

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Imaging, mapping and modelling continental lithosphere extension and breakup: an introduction G. D. KARNER1, G. MANATSCHAL2 & L. M. PINHEIRO3 1

ExxonMobil Upstream Research Company, P.O. Box 2189, Houston, TX 77252-2189, USA 2

Ecole et Observatoire des Sciences de la Terre, Universite´ Louis Pasteur, 1 Rue Blessig, 67084 Strasbourg, France

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Departamento Geocieˆncias and CESAM, Universidade de Aveiro, Campus de Santiago, 3810-193 Aveiro, Portugal Abstact: This Special Publication is a direct outcome of a small but dedicated group of researchers who met in Pontresina, Switzerland, to review and define the fundamental observations characterizing extensional systems and their application in guiding and constraining modelling efforts and results. The various summaries of the keynote addresses give an objective overview of the state of the art in modelling lithospheric extensional systems, both from the regional scale using dynamic models to individual basins using kinematic models with an emphasis on capturing the extensional history of the Iberia and Newfoundland margins. At the heart of all of these efforts is a simple question: Exactly what mechanisms allow the continental lithosphere to be thinned to the point of rupture? Related questions are: (1) Do crustal and mantle faults play a major role in this thinning process? If so, what is their geometry and does their importance and geometry change with time? (2) Are there other mechanisms of lithospheric and crustal thinning that cannot be imaged on seismic sections? (3) How is deformation accommodated in space and time? (4) What role do inherited mechanical, thermal and/or chemical heterogeneities play in controlling strain distribution and localization? (5) When, how and to what degree does magma production affect the distribution and localization of extension? And (6) what is the stratigraphic record of continental extension and how does it document the extension of the crust and thinning of the lithospheric mantle? The aim of this Special Publication is to address many of these fundamental questions concerning the extreme extension and thinning of continental lithosphere.

The Pontresina workshop – the source of the book During the week of 11–16 July 2004, in the Swiss Alps, 46 researchers working on various problems related to lithospheric extensional systems met in Pontresina, Switzerland. The aim of the workshop was to bring together observationalists and numerical modellers working on rifted continental margins. Its main objectives were to: (1) summarize the cutting-edge research related to extensional basin and lithospheric deformation studies (both from field observations and modelling), and recent advances and applications in basin modelling codes; (2) benchmark lithospheric deformation codes in terms of temperature history, brittle and ductile strain history, strain rates, and subsidence/uplift patterns in space and time; and (3) identify a set of field constraints (both onshore and offshore) of extensional systems (geological, geophysical, petrophysical and petrological data) that fundamentally constrain theoretical and conceptual models for the extensional deformation of the lithosphere. The workshop attracted

participants from Australia, Brazil, Canada, Egypt, France, Germany, India, Ireland, Norway, Portugal, Switzerland, the UK and the USA. The workshop was made possible with funding from the European Science Foundation, the UK National Environment Research Council, InterMARGINS and the US National Science Foundation. Key questions that need to be answered by academia relate to defining and testing the mechanisms responsible for thinning continental lithosphere. However, the key questions that need to be answered by the hydrocarbon industry for the exploration of deep-water rifted margins are: (1) Where is the location of the ocean –continent transition? (2) What are the structures, compositions and modes of deformation within the ocean–continent transition? And (3) what are the subsidence and heat-flow history of the ocean –continent transition and the zone of lithospheric thinning? There is a complete parallelism in the themes being asked and investigated by both groups. Consequently, a new period of industry–academic collaboration has begun, centred around these themes, and

From: KARNER , G. D., MANATSCHAL , G. & PINHEIRO , L. M. (eds) Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 1– 8. DOI: 10.1144/SP282.1 0305-8719/07/$15.00 # The Geological Society of London 2007.

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in the spirit of this collaboration BP, Conoco – Phillips, ExxonMobil and Statoil have generously supported the colour publication of this book. We are very grateful to their managements and also for publication support from InterMARGINS.

The Pontresina experience – the purpose of the book With the advent of sophisticated finite-element schemes for modelling the brittle and plastic extensional deformation of the lithosphere, fully integrated thermal and mechanical geodynamic models simulating the large-scale lithospheric systems have been possible. To perform such modelling, a myriad of geological assumptions need to be made concerning the rheological zonation, composition, strain rate and strain-rate dependency, initial temperature structure, radiogenic element distribution and crustal thickness of the lithosphere. Likewise, kinematic models of extensional basin formation, while computationally simpler than dynamic models, are evolving to include the wide range of geological processes responsible for basin development. Comparing the results of kinematic and dynamic models is difficult, primarily because the model philosophies address very different spatial scales of the real-world problem. For the present modelling schemes, the prediction of syn- and post-rift subsidence and crustal structure allow a comparison to be made between predicted dynamic and kinematic heat-flow histories. New field observations and syntheses concerning the detailed development of the ultra-deep regions of passive margins and the ocean–continent transition (OCT) have produced new and exciting constraints for structural reconstructions of latestage extension, implications for magma production and intrusion, infiltration of synextensional magmas into the mantle, thermal constraints for the proximal and distal margin using radiometric techniques on exposed mantle rocks, and relationships between sediments and magmatic rocks within various extensional systems. This book focuses primarily on mapping and modelling extensional processes in deep magma-poor extensional systems using the Iberia –Newfoundland and other comparable examples for benchmarking of modelling results and modelling codes. Each chapter provides a reflective overview of a specific modelling technique or study area, highlighting the present state of knowledge and the first-order problems and challenges facing tomorrow’s researchers.

Chapter arrangement The book has been organized into six main sections and 17 Chapters. The first section – ‘The

Iberia –Newfoundland continental extensional system (geological and geophysical constraints)’ – presents new and exciting geological and geophysical data and observational constraints that are necessary to model the extensional history of the Iberian and Grand Banks/Newfoundland margins. This is followed immediately by the simulation or benchmarking experiments in the sections. ‘The Iberia– Newfoundland continental extensional system (dynamic modelling)’ and ‘The Iberia–Newfoundland continental extensional system (kinematic modelling)’. However, cutting-edge research on mapping and modelling extensional systems is happening across a wide front, not only in Iberia and Newfoundland. Many other passive continental margins are also being extensively surveyed and intensively investigated, such as the western and southern Australian margins, the Antarctic margin, the MARGINS Gulf of California focus site, the margins of the South Atlantic and back-arc basins such as the Woodlark Basin. Attempts to re-energize studies in interesting scientific (but logistically and politically difficult) regions, such as the northern Red Sea, have also been initiated. These research themes are highlighted in the next section – ‘Observational characteristics of non-Atlantic extensional systems (offshore)’. Perhaps one of the most dramatic realizations over the last decade has been the applicability of mapped Alpine extensional systems in supplying a fundamental insight into the boundary and initial conditions prior to Iberian margin extension and the implications for the changing role and range of geological processes in space and time responsible for the regional and local extensional deformation of the European and Adriatic continental lithospheres. Field mapping of exposed extensional systems, such as the collisional systems of the Mediterranean, offer equally important insights into the role, complexity and range of processes responsible for the thinning of continental lithosphere. These issues are the subject of the section – ‘Observational characteristics of non-Atlantic extensional systems (onshore)’. The final section – ‘Revisiting fundamental concepts of continental extension’ – contains a number of eclectic chapters that herald future directions and/ or a need to reflect on the current paradigm related to extensional systems. For example, do active lowangle normal faults really exist? Arguably an old question and, to some, a question that has already been answered by field observations and relationships, predicting theoretically low-angle faulting nevertheless remains elusive. Further, it is now clear that lithospheric extension and thinning cannot be uniform with depth given the exhumation of continental lithospheric mantle or the paucity of brittle deformation observed at many rifted continental margins. But how does this occur? These fundamental questions are addressed in this section.

CONTINENTAL LITHOSPHERE EXTENSION AND BREAKUP

Research essentials Despite numerous seismic reflection and refraction and ODP drilling cruises of the Iberian, Galician and Newfoundland margins (Leg 47B, Leg 103, Leg 149, Leg 173 and Leg 210), a physical understanding of the extension processes responsible for margin development remains rather incomplete. For example, we do not yet understand the driving forces of rift initiation, segmentation and continuation, nor what controls the switch from rifting to drifting. We do not know the controls on the loci of rifting or how extension partitions itself laterally or vertically within the lithosphere. We do not know the scale and role of low-angle normal faults and whether these faults are early or late extensional features. We do not know how synextensional sagging, i.e. subsidence in the absence of controlling normal faults, can occur early in the extensional system, or the feedback between sedimentation and the style of deformation. We do not understand how the crust can thin dramatically from 30–40 km onshore to about 10 km within 100 km of the coast. We do not understand the role of fluids, volatiles, and magma in facilitating rifting, or the mechanisms by which huge volumes of magma are generated very quickly over wide areas at many rifted margins. We do not know the heat budget associated with rifting or how it changes as a function of space and time. In short, we still do not know a lot about the fundamentals of extensional systems, in general, and the conjugate Iberian, Galician and Newfoundland margins, in particular. The conjugate aspect of continental margin studies is crucial as it represents the entire extensional system. Defining the first-order observations that necessarily constrain any modelling effort of extensional systems, dynamic or kinematic, are sometimes straightforward, such as: – What is the pre-extension crustal and lithosphere thicknesses? – What is the pre-event relief and environments of deposition? – What are the spatial and temporal scales of deformation? – What are the strain rates, regions of strain localization and modes of deformation? – What is the age of the onset of rifting and how can it be recognized? – What is the composition, facies and palaeobathymetry of synrift/post-rift sediments, subsidence history and the nature and age of unconformities/discontinuities? Defining other equally important constraints, however, is not so easy and their definition is either controversial or even contentious, as for example: – What is the distribution of mechanical, thermal and chemical heterogeneities?

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– What is the age of breakup (i.e. when is normal ocean crust generated in a well-defined oceanic spreading centre)? – Then again, what exactly is breakup? – How many rift events are there and what is their timing? – Did these rift events pulse or was extension continuous? – What is the age, composition and volume of synextensional magmas? – How has magmatic underplating modified the crustal thickness of the region and when did it occur? Does it affect the rifting process? – Do volcanic (i.e. totally non-magmatic) margins really exist or is magma simply emplaced within the lower crust or lithospheric mantle? – What is the compositional variation of the lithospheric mantle, and how does it control the rheology, magma production history and flexural strength of the extending lithosphere? – What is the distribution of crustal radioactivity?

The Iberia – Newfoundland extension system: research controversies, contentions and unanticipated results Tucholke et al., Hopper et al., Sibuet et al. and Reston address many of these objectives while summarizing the geological history of the Iberian –Newfoundland conjugate margin. The key data sets are multichannel reflection and crustal refraction seismic data and drilling results, principally from the Ocean Drilling Program (ODP) and industry wells within the Jeanne d’Arc and Orphan Basins. As summarized by Reston, the Iberian –Newfoundland conjugate margins are characterized as non-volcanic margins exhibiting a deficit of synrift magmas, a zone of exhumed subcontinental mantle in the continent–ocean transition zone and an apparent extension discrepancy. The existence of basement ridges of serpentinized peridotite of continental affinity in the transition zones is extremely important. First, mantle exhumation is a direct indicator of depth-dependent extension. Second, these continental peridotites are both faulted and associated with magnetic anomalies presumably of M5 –M3 age. Apparently, the serpentinization process can create sea-floor spreading magnetic anomalies during continental mantle exhumation. This observation challenges the long-held view that the existence of correlatable magnetic anomalies unambiguously defines oceanic crust. The lack of synrift magmas is somewhat of a problem given the extreme extensional deformation that characterizes the Iberian and Newfoundland continental lithospheres. Reston suggests that the lack of melt may be explained by a combination

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of heterogeneous extension of the lower lithosphere (i.e. depth-dependent extension) and a cool subcontinental geotherm. An infertile continental mantle due to earlier magma extraction events (e.g. during Hercynian orogenesis) may be another factor. So what now is the real meaning of breakup or even the ‘breakup unconformity’? As a part of the depth-dependent extension process, there are various geological forms of ‘breakup’: (1) the cessation of brittle deformation within the upper crust (although lower crustal and continental mantle thinning may continue to occur); (2) breaching of the lower crust, which is concomitant with mantle exhumation; and (3) the breaching of the continental mantle and thus the inception of oceanic crust generation (or the exposure of oceanic mantle). For the Jeanne d’Arc Basin, the crustal blocks last moved in the late Barremian–early Aptian. But, is this the time of oceanic crust generation? Tucholke et al. argue that the final breakup of the Newfoundland– Iberia rift and the beginning of ‘true’ sea-floor spreading occurs at the Aptian–Albian boundary (c. 112 Ma), which is some 15 Ma younger than the oldest proposed oceanic crust (anomaly M3). The Aptian–Albian boundary is represented by a strong, rift-wide reflection that is onlapped by post-rift sediments deposited as a stable sea-floor-spreading regime was established. Tucholke et al. have termed this unconformity the ‘Aptian event’. Tucholke et al. suggest that this unconformity was produced by a change of in-plane stress induced by the final breaching of the subcontinental mantle lithosphere. Although depth-dependent extension has been surmised by a number of authors based on the paucity of synrift brittle deformation on some margins, Reston questions the significance of this observation. In particular, Reston suggests that ‘extension discrepancies’, as he terms them, may simply be the result of incomplete recognition of the entire polyphase faulting history. In the absence of mappable preand synrift stratigraphic packages, it just is not possible to define the extensional history of the Iberian margin, polyphase or otherwise (because the stratigraphic ‘tape recorder’ was not particularly effective). For margins exemplified by well developed pre-, syn- and post-rift stratigraphies, the polyphase faulting history is known and is minor (e.g. Karner & Driscoll 1999). Hopper et al. concentrated on interpretations of seismic reflection and refraction data from the Flemish Cap margin off Newfoundland. The seismic data highlighted important asymmetries at a variety of scales that developed during the final stages of continental breakup and the onset of oceanic sea-floor spreading. In strong contrast to the conjugate Galicia Bank margin, Flemish Cap shows an abrupt necking of continental crust,

thinning from 30 km to 3 km thick over a distance of 80 km with a distinct absence of controlling normal faults. At the ocean–continent transition, oceanic crust averages only 3 km to 4 km thick and in places is less than 1.3 km thick, suggesting that there were large spatial and temporal variations in the available melt supply following continental breakup as ocean crust was being generated. There is little evidence for significant extensional deformation of the Flemish Cap, consistent with the hypothesis that it behaved as a microplate throughout the Mesozoic. This fact was used by Sibuet et al., who recognized that rifting between Iberia and North America involved two larger plates (Iberia and North America) and two smaller microplates (Flemish Cap and Galicia Bank), to infer that Flemish Cap rotated approximately 438 relative to Galicia Bank and Iberia and moved 200–300 km SE with respect to North America. The seismic data for the Newfoundland Basin highlight important asymmetries at a variety of scales that developed during the final stages of continental breakup and the onset of oceanic sea-floor spreading. In strong contrast to the conjugate Galicia Bank margin, Flemish Cap shows abrupt necking, thinning the continental crust from 30 to 3–5 km over a distance of 80 km with no clear evidence for horizontal detachments or normal faults within the continental crust. This remains a fundamental question– exactly by what process does the crust undergo dramatic thinning without obvious brittle deformation?

The modelling effort: dynamic models The Pontresina benchmarking exercise consisted of modelling a combined Iberian –Newfoundland transect, using ISE-1 and SCREECH-1 multichannel seismic reflection data, refraction inversions and tomography for crustal velocity structure and Moho topography, and ODP drilling constraints. In terms of results, some models dealt with the details but offered no insight into the larger picture, while other modelling efforts provided no details but gave important insights to understanding general crustal architecture. The chapters by Huismans & Beaumont, Burov and Harry & Grandell represent the dynamic modelling of the Iberia –Newfoundland transect, while Egan & Meredith and Healy & Kusznir are the custodians of the kinematic approach to basin modelling. A dynamic model is designed to model the extension of the continental lithosphere by allowing for the spontaneous location of the deformation in shear zones approximating faults as well as the accommodation of the deformation by ductile behaviour. This is in contrast to kinematic models where the structural deformation of the crust and mantle lithosphere deformation are prescribed, and the effects

CONTINENTAL LITHOSPHERE EXTENSION AND BREAKUP

of basement rebound, sediment loading, erosion, compaction, eustasy, and lateral and vertical heat flow are easily included in the basin models both during, and after, extension. In general, it is very difficult to compare kinematic and dynamic models because these models address different scales of the extensional system. Huismans & Beaumont used plane-strain thermo-mechanical finite-element modelling experiments to investigate the effects of frictional-plastic strain softening and inherited weakness on the style of lithospheric extension, comparing the model results to the Newfoundland –Iberia conjugate margin. The fundamental result is that coupling between the plastic –viscous layering, acting together with frictional-plastic strain softening localized on inherited weak heterogeneities, can explain the initially wide rifting and distributed rift basins that are later abandoned in favour of a narrow rift in which mantle lithosphere is exhumed to the surface. Lithosphere-scale models with a single weak seed exhibit a range of asymmetric and symmetric rifting modes. That is, pure shear-type extension can be transformed into simple shear-type extension as functions of rheological layering and/ or rifting velocity, thus disqualifying the concept of simple pure v. simple shear end-member models. Burov takes a very different approach by investigating Rayleigh –Taylor instabilities in the lithosphere triggered by abrupt density and viscosity contrasts and steep boundaries between cold deformed lithosphere and hot upwelling asthenosphere induced by lithospheric extension and thinning. Burov suggests that continental lithosphere necessarily must extend at relatively high strain rates compared with the average extension rates usually assumed. A problem? Perhaps, but then again, extensional strain rates of almost all rift systems are poorly known. Because the influence of Rayleigh – Taylor instabilities is strongly controlled by extension rate, density, rheology and thermal structure of the lithosphere, there is a clear need for an increased attention to this problem to better constrain strain rates. In a more focused modelling effort, Harry & Grandell used a finite-element model to simulate the Late Jurassic –Early Cretaceous rifting between the Flemish Cap and Galicia Bank continental margins. In particular, the style and timing of rifting can be explained as a consequence of the interaction of two pre-existing lithospheric weaknesses: (1) a crustal weakness attributed to structural fabrics in the Variscan front; and (2) a deep-seated weakness attributed to the thick crust beneath the central Variscan Orogen, which controls the location of mantle necking. Extension is initially diffuse, but accelerates and becomes more focused with time. Most interestingly, some

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time after rifting begins, the locus of crustal extension shifts from the region of preweakened crust into the region of preweakened mantle, marking the end of subsidence in the Galicia Interior Basin and the onset of subsidence in the Flemish Cap and Galicia Bank marginal basins. The modelling results are interesting but the serendipitous assigning of the all-important crustal and mantle weaknesses preordains the results – more regional geological data are required to constrain and/or validate these assumptions.

The modelling effort: kinematic models While basin kinematic models may not have the pizzazz of dynamic models, they nevertheless are significantly more flexible when applied to real basins. Egan & Meredith, after providing the fundamental rationale for kinematic modelling of extension, apply such a model to the Black Sea and the Iberia –Newfoundland margin. Healy & Kusznir, working with a conceptually new kinematic approach to lithospheric thinning (explained fully in Kusznir & Karner later in this Special Publication), model the development of the Goban Spur margin in the eastern North Atlantic. While Healy & Kusznir model satisfactorily the observed bathymetry and the free air gravity anomaly, they do not predict the exhumation of continental lithospheric mantle, which has only been inferred at Goban Spur from seismic refraction and reflection studies. If serpentinized mantle does exist at this margin, then the possibility exists that the mantle exhumation at non-volcanic margins is melt suppression during sea-floor spreading initiation (rather than continental lithospheric mantle exhumation) analogous with mantle exhumation during ultra-slow sea-floor spreading at the Gakkel and SW Indian Ridges.

The knowledge base of other onshore and offshore extensional systems Five papers are dedicated to the study of other, nonAtlantic extensional systems. Perhaps one of the most important rift basins studied over the last decade is the Woodlark Basin of Papua New Guinea, which included ODP drilling, field mapping, seismic tomography and underway geophysical studies. Goodliffe & Taylor summarize the fundamental characteristics of this extensional system: west of the spreading tip, the southern margin is characterized by large fault blocks and a northern margin devoid of large offset normal faults but extreme subsidence, indicative of preferential lower crustal thinning. The margin asymmetry primarily reflects across-strike inherited differences in the prerift geology, morphology and rheology. ODP drilling

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results identify an initial phase of synrift sag subsidence interrupted by periods of relatively minor normal faulting. The brittle deformation (low-angle fault) that characterizes the Moresby seamount is very young, less than 0.5 Ma. The 258–308 fault dip is explainable in terms of a low coefficient of friction, being formed or reactivated in serpentinized ophilitic basement. The progression from extension to spreading is characterized by a decrease in sedimentation as the conjugate margins are progressively thinned, subside below sea level, trap sediments in proximal basins and develop significant palaeo-water depths. There is no evidence for the outcrop of lower crust or continental mantle as the lateral balance to depth-dependent extension, which is a problem. While Goodliffe & Taylor’s preferred mode of crustal thinning is via lower crustal ductile flow, crustal flow nevertheless tends to minimize Moho topography, which is not the conclusion of the seismic tomography studies of Abers et al. (2002) and Ferris et al. (2006). A problem still exists as to the mechanisms that produce synrift sag subsidence. Direen et al. document the geological development of the ocean– continent transition of the Naturaliste Plateau and the Great Australian Bight of southern Australia using multichannel seismic reflection, gravity and magnetic data. The ocean – continent transition zone of these margins is geologically and geophysically complex, but interpretation of all the geophysical data and dredge hauls (principally in the Diamantina Zone) is consistent with the presence of a mixture of modified continental lower crust, volcanics related to breakup and exhumed continental mantle. The decision by the US National Science Foundation to ‘postpone indefinitely’ MARGINS research in the Red Sea region is a major setback in gaining access to sites crucial in mapping, modelling and understanding extensional processes. The Red Sea remains a fundamentally important region in studying cratonic rifting. As stated in the MARGINS Science Plans for Rupturing Continental Lithosphere, ‘The Red Sea and Gulf of Aden are the closest modern analogs to the rifting and rupturing of continental lithosphere that formed the vast majority of “Atlantictype” continental margins, and are the location where the processes that shaped the early development of rifted continental margins can be studied with the fewest tectonic complications. Nearly all of the passive continental margins of the Atlantic, Indian, Southern and Arctic oceans were formed by the nucleation of an oceanic spreading centre within a continental rift following an extended period of cratonic lithosphere extension’. Cochran & Karner revisit the geological setting of the northern Red Sea and underscore its absolute importance in understanding rift systems and the rift–drift transition.

As summarized by Cochran & Karner, the Red Sea basins formed by rifting of Precambrian continental lithosphere in the late Oligocene, leading to breakup and sea-floor spreading by approximately 5 Ma in the southern Red Sea and significant continental block rotation and the instalment of volcanic centres in the northern Red Sea. Comparing and contrasting the extension of the southern and northern Red Sea remains the definitive experiment to understand the development of Atlantic-type passive continental margins. For decades it has been simply assumed that the northern Red Sea had experienced insufficient extension to breach the continental lithosphere, but, in time, would develop into a spreading centre. However, as shown by Kusznir & Karner later in this Special Publication, rifts characterized by large offset fault systems, i.e. faults that generate synrift accommodation, tend not to proceed to breakup. Thus, Cochran & Karner explore the tantalizing possibility that the northern Red Sea region, with its interpreted rotated fault blocks, is rheologically stronger than the lithosphere of the southern Red Sea and will never evolve to sea-floor spreading. In marked contrast to the Woodlark Basin, Eastern Tethyan margins and even the Gulf of California margins (e.g. Persaud et al. 2003), the Alpine Tethys margin is characterized by a continent–ocean transition zone and exposed continental mantle in a style very similar to that of the Iberian margin (Manatschal et al.). The Tethys margin, now exposed in various nappes in the Alps, shows a change from initially regionally distributed and decoupled extension to later localized, coupled and asymmetric extension that results in thinning of the crust and exhumation of subcontinental mantle. The change in the mode of extension, together with the localization of deformation, reflects an evolution of the bulk rheology of the extending lithosphere. This is therefore a fundamental natural laboratory to study how the crust thins during extension. As underscored by the dynamic rifting model results, initial rifting of the Tethyan margin appears to be controlled by inherited heterogeneities (cf. chapters by Dysterhuis et al. and Harry & Grandell in this Special Publication) and mechanical localization processes, whereas the site of final extension and lithospheric rupture is controlled by serpentinization, magmatic and thermal weakening of the continental mantle. In the final paper of this section, Robertson compares and contrasts the geology and tectonics of the rifted margins of the Eastern Tethys with the Iberia –Newfoundland margin and the East Greenland margin. None of the Eastern Tethyan rifted margins show evidence of features characteristic of the non-volcanic rifted margins of the Atlantic (e.g. exhumation of continental mantle).

CONTINENTAL LITHOSPHERE EXTENSION AND BREAKUP

Combinations of geophysical studies and deep-sea drilling have in the past suggested that rifted margins fall into two end-members: volcanic-rifted margins (e.g. East Greenland) and non-volcanic rifted margins (e.g. Iberia–Newfoundland conjugate). Most of the Eastern Tethyan rifted margins are neither, and appear to be ‘intermediate’ between these end-member margin types, being characterized by pulsed rifting (extending over more than 50 Ma), limited rift volcanism and a narrow but well-developed continent– ocean transition zone. In terms of tectonic setting, there is little evidence to support previous suggestions that the Eastern Tethyan rifts are back-arc basins above either northward-dipping or southwarddipping subduction zones. All of the rifted margins of the Eastern Tethys are associated with rift-related volcanic rocks. The dominant controls of rifting are seen as the traction of rising asthenosphere on the base of the lithosphere, related deviatoric tensional stresses, inherited and thermally induced weaknesses in the crust, and slab pull.

Revisiting fundamental concepts of continental extension As stated earlier, this final section contains a number of eclectic chapters that may herald future directions and/or a need to revisit or reflect on supposedly well-established paradigms. Four papers comprise the last section of the book, Kusznir & Karner, Christie-Blick et al., Dyksterhuis et al. and Moresi et al. Although depth-dependent lithosphere thinning has been implicitly discussed in earlier chapters, Kusznir & Karner review the prime evidence for this process. Kusznir & Karner demonstrate that depth-dependent thinning characterizes the margins of the North and South Atlantic, Norway, South China Sea, Papua New Guinea, NW Australia and Vietnam. In short, depth-dependent lithosphere thinning appears to be the norm, rather than the exception, for both non-volcanic and volcanic extensional systems. However, depth-dependent thinning, if it is an extension process, may create a ‘space problem’ depending on how the extension is distributed with depth, i.e. the extension needs to be laterally balanced. For regions characterized by regionally significant synrift sags, such as for the Congo–Angola and NW Australian margins, an adjacent region of exposed lower continental crust and continental mantle needs to exist – clear for the Iberian margin but a problem, for example, for the Woodlark Basin. Recognizing or reconciling this extensional balance, in general, has been problematic. Consequently, Kusznir & Karner describe a fundamentally different form of kinematic extension model in which deformation and thinning of

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continental lithosphere occurs in response to an upwelling divergent flow field. Such a model successfully predicts depth-dependent stretching of continental margin lithosphere and mantle exhumation for both non-volcanic and volcanic margins while maintaining a balance of extensional strain. This strain balance is achieved by redistributing lower crust and continental mantle away from the zone of active extension into the hinterland region, implying a broad region of crustal and continental thickening and broad hinterland uplift during and for some time following extension. Christie-Blick et al. report recent fieldwork that casts doubt on the interpretation of some well-studied regional detachment faults in the US Basin and Range Province. What has this to do with the Iberian – Newfoundland margin? Nothing, but it directly addresses the question of how crust thins during extension. Given that the Basin and Range examples were influential in the development of the active low-angle normal fault concept, it is somewhat disturbing to learn that the Mormon Peak detachment, for example, a structure that was one of the most influential in arguing for the importance of lowangle normal faults and for high extensional strain in regions not associated with metamorphic core complexes, can be reinterpreted as a series of rootless slide blocks on the basis of detachment characteristics and spatially variable kinematic indicators. Among issues addressed by Christie-Blick et al. are the manner in which brittle deformation relates to mylonitization, the degree to which faults are folded or tilted to lower dip and the role of unusual materials such as talc in reducing frictional coefficients. Christie-Blick et al. explicitly state that their work does not falsify the active low-angle normal fault paradigm nor do they deny the reality of extreme crustal extension – they ask for a critical re-examination of interpreted detachments elsewhere in the Basin and Range Province and in other extensional and passive margin settings. Dyksterhuis et al. investigate the primary factors controlling rift architecture. Key results are that assumed weaknesses fundamentally control rift mode rather than bulk rheology, initial lithospheric temperature structure or strain softening. This result is in complete agreement with the modelling results of Harry & Grandell. A secondary conclusion of Dyksterhuis et al. is that low-angle faults can indeed form as the result of rotation of initially highangle faults, i.e. the rolling hinge model of low-angle normal fault development. It is gratifying to know that the principles of isostasy in extensional systems remain very much intact. The final chapter by Moresi et al. presents an extremely important theoretical development that allows viscous formulations to include 2D and 3D approximations to elastic effects and plastic

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failure – that is, a way to represent Mohr –Coulomb failure within the mathematical framework used by mantle-convection fluid dynamics codes. In the words of the authors ‘The algorithms described in this paper address a prominent problem in geodynamics: how to model the brittle deformation of the uppermost lithosphere that occurs when mantle convection deforms the continents at the same time as modelling the underlying fluid convection which drives the surface deformation’. In many ways, formulations of this type represent the future of complex, large strain approximations of lithospheric extensional deformation. As with any professional publication, the quality of the final product is a strong function of the dedication of reviewers within the earth science community. The volume editors would like to acknowledge the following people for their unselfish donation of time to critically review the various chapters of this book: G. Bertotti, R. Buck, E. Burov, N. Christie-Blick, J. Cochran, P. Cowie, N. Dawers, S. Deemer, N. Direen, S. Egan, M. Enachescu, T. Funck, A. Goodliffe, R. Govers, M. Pe´rez-Gussinye´, D. Harry, J. Hopper, G. Houseman, R. Huismans, N. Kusznir, L. Lavier, G. Lister, K. Louden, F. Martinez, T. Minshull, L. Montesi, L. Moresi, D. Mu¨ller, J. Nunn, W. Powell, D. Roberts, A. Robertson, D. Sawyer, D. Shillington, T. Reston,

M. Tischer, F. Tsikalas, B. Tucholke, H. Whitechurch, R. Whitmarsh, C. Wilson, P. Ziegler and last but certainly not least, our production editor, S. Oberst.

References A BERS , G. A., F ERRIS , A., C RAIG , M. ET AL . 2002. Mantle compensation of active metamorphic core complexes at Woodlark Rift in Papua-New Guinea. Nature, 418, 862– 865. F ERRIS , A., A BERS , G. A., Z ELT , B., T AYLOR , B. & R OECKER , S. 2006. Crustal structure across the transition from rifting to spreading: the Woodlark rift system of Papua New Guinea. Geophysical Journal International, 166, 622– 634. K ARNER , G. D. & D RISCOLL , N. W. 1999. Style, timing and distribution of tectonic deformation across the Exmouth Plateau, northwest Australia, determined from stratal architecture and quantitative basin modelling. In: MAC NIOCAILL , C. & RYAN , P. D. (eds) Continental Tectonics. Geological Society, London, Special Publications, 164, 271–311. P ERSAUD , P., S TOCK , J. M., S TECKLER , M. S. ET AL . 2003. Active deformation and shallow structure of the Wagner, Consag, and Delfin Basins, northern Gulf of California, Mexico. Journal of Geophysical Research, 108(B7), 2355, doi:10.1029/2002 JB001937.

Breakup of the Newfoundland– Iberia rift B. E. TUCHOLKE1, D. S. SAWYER2 & J.-C. SIBUET3 1

Department of Geology and Geophysics, Woods Hole Oceanographic Institution, Woods Hole, MA 02543, USA (e-mail: [email protected]) 2

Department of Earth Science, Rice University, Houston, TX 77005, USA 3

Ifremer Centre de Brest, B.P. 70, 29280 Plouzane´, France

Abstract: The Newfoundland–Iberia rift is considered to be a type example of a non-volcanic rift. Key features of the conjugate margins are transition zones (TZs) that lie between clearly continental crust and presumed normal (Penrose-type) oceanic crust that appears up to 150– 180 km farther seaward. Basement ridges drilled in the Iberia TZ consist of exhumed, serpentinized peridotite of continental affinity, consistent with seismic refraction studies. Although the boundaries between continental crust and the TZs can be defined with relative confidence, there are major questions about the position and nature of the change from rifting to normal sea-floor spreading at the seaward edges of the TZs. Notably, drilling of presumed oceanic crust in the young M-series anomalies (,M5) has recovered serpentinized peridotite, and this basement experienced major extension up to approximately 15 million years after it was emplaced. In addition, existing interpretations place the ‘breakup unconformity’ (normally associated with the separation of continental crust and simultaneous formation of oceanic crust) near the Aptian–Albian boundary, which is also some 15 million years younger than the oldest proposed oceanic crust (anomaly M5–M3) in the rift. To investigate and potentially resolve these conflicts, we analysed the tectonic history and deep (pre-Cenomanian) stratigraphy of the rift using seismic reflection profiles and drilling results. Rifting occurred in two main phases (Late Triassic–earliest Jurassic and Late Jurassic– Early Cretaceous). The first phase formed continental rift basins without significant thinning of continental crust. The second phase led to continental breakup, with extension concentrated in three episodes that culminated near the end of Berriasian, Hauterivian and Aptian time. The first two episodes appear to correlate with separation of continental crust in the southern and northern parts of the rift, respectively, suggesting that the rift opened from south to north in a two-step process. The third episode persisted through Barremian and Aptian time. We suggest that during this period there was continued exhumation of subcontinental mantle lithosphere at the plate boundary, and that elevated in-plane tensile stress throughout the rift caused intraplate extension, primarily within the exhumed mantle. This rifting may have been interrupted for a time during the Barremian when melt was introduced from the southern edge of the rift by plume magmatism that formed the Southeast Newfoundland Ridge and J Anomaly Ridge, and the conjugate Madeira– Tore Rise. We propose that the rising asthenosphere breached the subcontinental mantle lithosphere in latest Aptian– earliest Albian time, initiating sea-floor spreading. This resulted in relaxation of in-plane tensile stress (i.e. a pulse of relative compression) that caused internal plate deformation and enhanced mass wasting. This ‘Aptian event’ produced a strong, rift-wide reflection that is unconformably onlapped by post-rift sediments that were deposited as a stable sea-floor-spreading regime was established. Although previously considered to be a breakup unconformity associated with separation of continental crust, the event instead marks the final separation of the subcontinental mantle lithosphere. Our analysis indicates that interpretation of tectonic events in a non-volcanic rift must consider the rheology of the full thickness of the continental lithosphere, in addition to spatial and temporal changes in extension that may occur from segment to segment along the rift.

The Newfoundland–Iberia rift (Fig. 1) is considered to be a type example of non-volcanic rifted margins, and because of its accessibility it has been extensively studied by geophysical means and by drilling over the past two decades. Most of these investigations have concentrated on the Iberia margin, but recent geophysical surveys and deep-sea drilling off Newfoundland also provide critical information

that contributes to understanding the evolution of the rift. A key feature observed on each margin is a transition zone (TZ) that lies between clearly continental crust and assumed normal oceanic crust that appears up to 150–180 km farther seaward in the vicinity of magnetic anomaly M3 (Figs 2 & 3). (Normal ocean crust is here considered to be crust that approximates the Penrose ophiolite

From: KARNER , G. D., MANATSCHAL , G. & PINHEIRO , L. M. (eds) Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 9– 46. DOI: 10.1144/SP282.2 0305-8719/07/$15.00 # The Geological Society of London 2007.

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Fig. 1. North Atlantic bathymetric map (Smith & Sandwell 1997) showing features of the conjugate margins of the Newfoundland–Iberia rift. Boxes locate Figures 2 & 3. White dashed lines show Palaeozoic sutures and terrane boundaries in continental crust (after Keen et al. 1990; Silva et al. 2000) and black lines delineate major rift basins. Yellow line in the Tagus Abyssal Plain locates refraction profile of Pinheiro et al. (1992) (Fig. 4g, section I1). AM, Armorican Massif; BB, Bay of Biscay; CA, Collector magnetic anomaly; CIZ, Central Iberian Zone; CZ, Cantabrian Zone; F, Flemish Pass Basin; FC, Flemish Cap; GB, Galicia Bank; GI, Galicia Interior Basin; GRB, Gorringe Bank; GZ, Galicia-Tra´s-os-Montes Zone; H, Horseshoe Basin; IAA, Ibero-Armorican Arc that presumably connected to the northern European sutures prior to opening of the Bay of Biscay; IAP, Iberia Abyssal Plain; J, Jeanne d’Arc Basin; JAR, J Anomaly Ridge; L, Lusitanian Basin; MAR, Mid-Atlantic Ridge; MEG, Meguma Terrane; MTR, Madeira– Tore Rise; NNB, northern Newfoundland Basin; NS, Newfoundland Seamounts; O, Orphan Basin; OMZ, Ossa Morena Zone; P, Porto Basin; SB, Salar– Bonnition Basin; SENR, Southeast Newfoundland Ridge; SNB, southern Newfoundland Basin; SPZ, South Portuguese Zone; TAP, Tagus Abyssal Plain; TS, Tore Seamounts; W, Whale Basin; WZ, Western Asturian–Leonese Zone.

model (Penrose Conference Participants 1972) and consists of gabbro, sheeted-dyke and volcanic layers that accreted at a magmatic spreading centre). Basement ridges drilled in the TZ on the Iberia margin consist of serpentinized peridotite of apparent continental affinity (e.g. Whitmarsh & Wallace 2001), and seismic refraction studies show velocity structure that is consistent with the presence of serpentinized mantle (e.g. Chian et al. 1999; Dean et al. 2000). Thus, the TZ been referred to as the Zone of Exhumed Continental Mantle (ZECM) (Whitmarsh et al. 2001). Although the boundaries between continental crust and the TZ can be defined with relative confidence on the basis of basement morphology, fault patterns and velocity structure (Figs 2 & 3), major questions remain about the nature and position of

the seaward limit of the TZ and the first occurrence of normal oceanic crust. Notably, in basement identified as oceanic crust from seismic refraction studies, drilling at Site 1070 on the Iberia margin and at Site 1277 on the Newfoundland margin has recovered serpentinized peridotite (Whitmarsh et al. 1998; Tucholke et al. 2004). This puts into question the validity of identifying oceanic crust only on the basis of velocity structure. In addition to this uncertainty, several existing interpretations place the final breakup of the Newfoundland– Iberia rift near the Aptian –Albian boundary, about 112 Ma. This event has been interpreted to be marked by a breakup unconformity (e.g. Mauffret & Montadert 1987; Boillot et al. 1989c), which is conventionally considered to mark the separation of continental crust and

BREAKUP OF THE NEWFOUNDLAND– IBERIA RIFT

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Fig. 2. Bathymetry of the Newfoundland Basin (Smith & Sandwell 1997) with basin features and profile locations. Identified magnetic anomalies are shown as black solid lines, from Srivastava et al. (2000) and Shillington et al. (2004). The approximate distribution of the high-amplitude magnetic J anomaly, which contains anomaly M0, is indicated by a heavier line. The dashed black line at the boundary between Aptian and Albian crust is based on interpolation between anomalies M0 and 34. Black dots locate ODP drill sites. The approximate seaward limit of certain continental crust is the white short-dot line (CL), and the white wide-dot line (TCL) shows the seaward limit of highly thinned continental crust interpreted from refraction data (Lau et al. 2006b; Van Avendonk et al. 2006). Multichannel seismic reflection lines illustrated in other figures are solid yellow lines, dotted yellow lines show the extent of crustal-thickness profiles (,10 km thickness where over continental crust) derived from seismic refraction data and illustrated in Figure 4g, and black/yellow lines show locations of SCREECH data not illustrated in this paper. Refraction line R94 is from Reid (1994), SCREECH2 is from Van Avendonk et al. (2006) and SCREECH1 is from Funck et al. (2003).

simultaneous formation of the first oceanic crust (Falvey 1974). The age of this postulated breakup is some 15 million years younger than the oldest magnetic anomaly (M5) that has been proposed as possible oceanic crust off Iberia (Russell & Whitmarsh 2003), and it is 9 million years younger than anomaly M0 which most investigators would argue is certain oceanic crust. There are three

possible ways to explain this paradox: (1) the breakup unconformity has been misidentified; (2) the identification has no meaning because continental breakup does not generate a consistent marker in the stratigraphic record; or (3) the breakup unconformity can represent an event other than simultaneous separation of continental crust and generation of the first normal oceanic crust.

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Fig. 3. Bathymetry of the Iberia margin (Sibuet et al. 2004a) with basin features and locations of illustrated profiles. Explanation as in Figure 2. Magnetic anomalies are from Miles et al. (1996) and Russell & Whitmarsh (2003). Peridotite ridges identified from various sources are indicated by white lines and ‘R’ labels. Seismic refraction profile W96 is from Whitmarsh et al. (1996) and CAM127-IAM9 is from Dean et al. (2000).

In this paper, we review the rift history of the Newfoundland–Iberia margins with particular attention to the deep (pre-Cenomanian) stratigraphy of the rift. We focus on the issue of Late Jurassic – Early Cretaceous rift history, and in particular on development of the previously proposed breakup unconformity and the time that the first normal oceanic crust formed. From this analysis, we propose that the third explanation above is likely to be true in the Newfoundland–Iberia rift. We suggest that the first normal oceanic crust was not formed until the Aptian –Albian boundary, approximately 112 Ma, coincident with the formation of the previously proposed breakup unconformity. Up to that time new sea floor was emplaced largely by exhumation of subcontinental mantle lithosphere,

and the entire rift was under elevated in-plane tensile stress. This caused local extension throughout the rift, even though the last principal episode of extension in continental crust probably ended in Hauterivian –Barremian time. ‘Breakup’ is represented by the final separation of subcontinental mantle lithosphere, rise of the asthenosphere to shallow levels and release of large volumes of melt, and commencement of normal sea-floor spreading. Relaxation of plate-wide tensile stress at this time led to a pulse of relative compression within the plates, resulting in internal plate deformation (e.g. topographic amplification), enhanced mass wasting of serpentinite ridges in the deep basins, and proximal-margin erosion and basin flooding by turbidites. We suggest that this

BREAKUP OF THE NEWFOUNDLAND– IBERIA RIFT

deformation, the accumulation of coarse debris in the deep basins and the subsequent change to sedimentation in a low-stress environment combined to create a unique stratigraphic marker in seismic reflection records. According to our interpretation, this marker is not a breakup unconformity in the classical sense (i.e. simultaneous separation of continental crust and commencement of sea-floor spreading) but is the structural and stratigraphic expression of the final separation of the full thickness of continental lithosphere. Thus, we will refer to this feature as the ‘Aptian event’.

Present structural framework Continental crust The Newfoundland margin is dominated by a broad, shallow-water (,100 m) platform that forms the Grand Banks (Fig. 1). This platform is formed by thick continental crust (27 –35 km: Reid 1988; Reid & Keen 1990) that is deeply incised by large, deep rift basins. The seaward edge of the platform is a seaward-dipping basement hinge zone that runs mostly beneath the continental slope from the southern transform margin of the Grand Banks northwards to Flemish Cap. Along the southern Grand Banks, rift basins seaward of the hinge zone (Salar –Bonnition basins) contain an apparently thick sequence of evaporites (Austin et al. 1989) and they are marked at their eastern edge by a landward-dipping hinge in basement (Keen & de Voogd 1988; Austin et al. 1989). Flemish Cap, at the NE end of the Grand Banks, has thick (c. 35 km) continental crust (Funck et al. 2003) and it is separated physiographically from the Grand Banks by Flemish Pass (1100 m water depth), which overlies a deep rift basin in continental crust. A prominent east-dipping hinge zone in basement on the seaward side of Flemish Cap, together with the landward-dipping hinge zone along the margin to the south, is generally considered to be the seaward edge of clearly continental crust (CL, Fig. 2); these interpretations are based on crustal thickness and velocity structure, basement morphology, and the occurrence of evaporites (where present). Highly thinned (2–6 km) continental crust has been interpreted to extend up to about 60 km farther seaward (TCL, Fig. 2) based on refraction data (Lau et al. 2006b; Van Avendonk et al. 2006). The Iberia margin contrasts with the Newfoundland margin in that it has a much narrower continental shelf and limited development of rift basins in thick continental crust (Fig. 1). The principal rift basin in the shallow Iberia margin is the Lusitanian Basin, which is roughly

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co-linear with the Porto Basin beneath the continental shelf and slope and with the deep-water Galicia Interior Basin to the north. Thinned continental crust underlies both the Galicia Interior Basin and Galicia Bank to the west (Boillot et al. 1989b; Murillas et al. 1990). Continental crust under Galicia Bank thins westward to zero thickness on the landward side of peridotite basement ridges (R1 –R3, Fig. 3). Based on geophysical and drilling results, the seaward edge of clearly continental crust runs SE along the southern margin of Galicia Bank and then southward in a poorly defined position to the southern limit of the rift (CL, Fig. 3). The seaward limit of continental crust is significantly younger in the northern part of the rift between Flemish Cap and Galicia Bank than in the southern part (Figs 2, 3 & 4g). Some interpretations suggest that this reflects relatively uniform northward propagation of rifting and crustal separation (e.g. Whitmarsh & Miles 1995). However, it is equally possible that rifting was a stepwise process. Although data are limited for the southern part of the rift, extension there appears to have caused separation of continental crust during the latest Jurassic–earliest Cretaceous (Fig. 4g, sections N1 and I2). This extension probably also affected the northern rift (Boillot et al. 1989c; Murillas et al. 1990), but Flemish Cap and Galicia Bank did not separate until near the Hauterivian–Barremian boundary approximately 16–19 million years later (Fig. 4g, sections N2, N3 and I3). Thus, breakup of the rift may have occurred in two successive and relatively discrete episodes.

Transitional crust On both sides of the rift south of the Flemish Cap– Galicia Bank conjugates, there is a wide (c. 150–180 km) TZ where basement has unusual geophysical properties and a complex origin. This zone extends from the outer limits of crust of certain continental origin seawards to about anomaly M3. The most extensive studies of the TZ are on the Iberia side of the rift, largely in the region of the southern Iberia Abyssal Plain (Fig. 3). In that area, refraction and multichannel seismic reflection studies document thin ‘crust’ (2–4 km) defined by low seismic velocities of approximately 4–4.5 km s21 (line IAM9, section I2 in Fig. 4g) (Whitmarsh et al. 1990; Discovery 215 Working Group 1998; Dean et al. 2000). Seismic velocities, velocity gradients and a general absence of a Moho reflection have been interpreted as representing intense serpentinization of mantle peridotites at the top of the basement and significantly reduced serpentinization with depth (c. 7.2 km s21 increasing to c. 7.9 km s21) (Discovery 215 Working Group 1998; Chian et al. 1999). Along the IAM9 line the upper

Fig. 4. Summary of geological features of the Newfoundland–Iberia rift v. time from the latest Jurassic through Early Cretaceous. The data are compiled from numerous sources, with principal sources discussed in the text. The light green bar marks the position of the ‘Aptian event’ at the Aptian– Albian boundary. (a) Timescale from Channell et al. (1995) and Gradstein et al. (1995). (b) Principal seismic and lithological sequences on the conjugate margins. Grey bars indicate biostratigraphic age constraints on the Aptian event (¼U reflection and orange reflection) at DSDP/ODP drill sites. (c) Principal rift events: red, magmatic, with significant magmatism as heavier lines; blue, tectonic. (d) Interpretation of in-plane tensile stress in the rift system; interpreted high stress in the Jurassic followed by decreasing stress up to Aptian time is from Tankard & Welsink (1987). (e) Rift phases defined from studies of extensional basins in the shallow margins of the rift and in Galicia Interior Basin. Albian extension in the northern Jeanne d’Arc Basin (grey shade) reflects detachment of basin fill associated with early extension of the Labrador Sea to the north of the Grand Banks. (f) Major unconformities;

14 B. E. TUCHOLKE ET AL.

Fig. 4. (Continued) all except the Galicia ‘breakup unconformity’ are based on data from extensional basins in the shallow margins of the rift and in Galicia Interior Basin (i.e. well and onshore outcrop data, and related seismic interpretations). (g) Representative thicknesses of velocity layers in basement v. age, derived from refraction profiles in the southern, central and northern parts of the rift. Note that these sections depict only layer thickness and not topography. Three sections are arranged from south to north (left to right) for each of the Newfoundland (N) and Iberia (I) margins. Sections N3 and I3 are conjugates, I2 is offset about 25 km south of N2 at anomaly M0, and N1 and I1 are approximately conjugate. Locations of sections are shown in Figures 2 and 3 (I1 is located in Fig. 1). Pale blue and pale green colours indicate thickness of interpreted oceanic and transitional ‘crust’, respectively. These have compressional wave velocities in the range of 3.5–7.2 km s21 and commonly exhibit steep velocity gradients. The oceanic layer is anomalously thick at the young end of section I2 because the profile intersects a seamount (see Fig. 3). Broken patterns at the young ends of sections N3 and I3 indicate where velocity structure is not well constrained. Dark blue is interpreted continental crust; continental crust thicknesses are shown only up to 10 km, although all sections except I1 extend onto thicker crust. Red indicates lithosphere with velocities from 7.2 up to 8.0 km s21 that typically has been interpreted as serpentinized mantle because velocities are considered to be too high to represent magmatic underplating; the velocity model in N2 did not detect 8 km s21 velocities where shown by white dotted lines, so in those locations the red area represents minimum thickness of the serpentinized layer. The transitional ‘crust’ above the red zones can be serpentinite and has specifically been interpreted as such on sections N2 and I2, and at the peridotite ridge on I3. Black dots show peridotite ridges sampled by drilling at ODP Sites 1277 (N2) and 637 (I3); open circles locate peridotite ridges interpreted from refraction data and also drilled at ODP Sites 1070, 897 and 899 (seaward to landward) approximately 50 km to the north. Basement ages are pinned to magnetic anomaly M0 (or in the case of I2, to both M0 and M5; in the case of N2, to both M0 and M3). Away from these anomalies the ages are calculated assuming that basement was emplaced at a rate of 9 km per million years (the average rate seaward of M0 to anomaly 34); ages landward of the seaward limits of continental crust have no meaning in this context. Data sources: N1, Reid (1994); N2, Van Avendonk et al. (2006); N3, Funck et al. (2003); I1, Pinheiro et al. (1992); I2, Dean et al. (2000); I3, Whitmarsh et al. (1996).

BREAKUP OF THE NEWFOUNDLAND– IBERIA RIFT 15

basement has a markedly non-reflective character that also has been interpreted as a layer of highly serpentinized mantle (Pickup et al. 1996). The Ocean Drilling Program (ODP) Leg 149–173 drilling transect, about 50 km to the north, drilled basement ridges that can be traced northward from the IAM9 line. Serpentinites were recovered from these ridges at Sites 897, 899 and 1070 (Figs 3 & 4g, open circles), consistent with seismic-velocity interpretations. The boundary between the serpentinized mantle and continental crust of Galicia Bank appears to run between Sites 899 and 1069, with some additional complexity just to the east at Sites 900, 1067 and 1068 where gabbros, amphibolites and serpentinites were recovered from basement. On the opposing margin the 565 km-long SCREECH3 (Studies of Continental Rifting and Extension off the Eastern Canadian sHelf Line 3) profile crosses the Grand Banks and reaches seaward to beyond anomaly M0 (Fig. 2). Based on combined seismic reflection and refraction data, Lau et al. (2006a, b) interpreted the inner part of the TZ there to be thinned continental crust (,3 km) over a thick (.5 km) layer of serpentinized mantle (7.6–7.9 km s21). Seaward to near anomaly M3, they proposed that an approximately 80 km wide zone with a thin (2 km) low-velocity basement and underlying 7.6–8.0 km s21 layer is probably exhumed and serpentinized mantle. The TZ in the southern part of the Newfoundland–Iberia rift has been much less studied. On the Iberia margin beneath the Tagus Abyssal Plain, Mauffret et al. (1989) inferred from multichannel seismic reflection and magneticanomaly studies that Tithonian–Hauterivian seafloor spreading formed the TZ, prior to a westward ridge jump at chron M10. However, a refraction profile shot over this area (Pinheiro et al. 1992) identified exceptionally thin ‘crust’ (2 km) that overlies mantle with low velocities of 7.6–7.9 km s21 (Fig. 4g, section I1), very much like the IAM9 basement farther north under the Iberia Abyssal Plain. Pinheiro et al. (1992) interpreted the Tagus basement to be formed by sea-floor spreading beginning at chron M11, but they noted that the low mantle velocities may indicate widespread serpentinization. In the conjugate Newfoundland TZ, Reid (1994) found mantle velocities beginning about 6 km below the basement surface just east of the Salar –Bonnition hinge. His results (Fig. 4g, section N1) show thin, low-velocity upper basement (4.5 km s21) and 7.2–7.7 km s21 velocities in the deeper basement, also similar to serpentinized mantle on the IAM9 profile. In the northern part of the rift, the TZ narrows abruptly between the Flemish Cap and Galicia Bank conjugates (Figs 2 & 3). The SCREECH2

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reflection–refraction profile off Newfoundland is located where this narrowing occurs, conjugate to the ODP Leg 149 –173 drilling transect at the southern margin of Galicia Bank. Van Avendonk et al. (2006) interpreted approximately 60 km of the landward part of the TZ along the SCREECH2 profile as thinned, 2 –6 km thick continental crust (velocity 5.0–6.5 km s21); this overlies a 7.2– 8.0 km s21 layer that has laterally variable thickness (1–.4 km) and may be serpentinized mantle (Fig. 4g, section N2). An approximately 25 km wide zone at the outer edge of the TZ has basement velocities of 6.3–7.7 km s21 and was interpreted as exhumed, serpentinized mantle. Within the narrow TZ farther north, Todd & Reid (1989) observed a thin (2–3 km) 4.0–5.0 km s21 layer overlying probably serpentinized mantle (7.3 km s21) in one profile and overlying 8.0 km s21 mantle in a second profile. There is little, if any, TZ present still farther north at the position of the SCREECH1 and W96 conjugate profiles (Figs 2 & 3). Continental crust occupies the position where anomaly M3 would be expected on the Newfoundland side (Fig. 4g, section N3), but the anomaly could be present off Galicia Bank (Fig. 4g, section I3). In this part of the western Galicia margin, a layer of serpentinized peridotite underlies thin continental crust, a peridotite ridge, and a narrow zone of thin crust interpreted as oceanic (Whitmarsh et al. 1996). Two features are nearly universal among the seismic-refraction observations of the TZ. One is the presence of low-velocity (7.2–7.9 km s21), apparently serpentinized, mantle that is up to approximately 5 km thick. The second is the occurrence of a thin (c. 2–4 km) low-velocity layer that forms the overlying basement. Where thin continental crust is interpreted in the landward part of the TZ, it is likely that serpentinization of the underlying mantle occurred when the crust thinned sufficiently for fluids to penetrate there (Pe´rez-Gussinye´ & Reston 2001). The outer TZ appears to be exhumed mantle, based primarily on the combined drilling and seismic results around Galicia Bank (Fig. 4g, sections I2 and I3). The highly variable velocity structure of the thin upper ‘crustal’ layer in the outer TZ probably is caused by strong but heterogeneous serpentinization. The magnetic field over the TZ on both margins is characterized by low-amplitude magnetic anomalies. Srivastava et al. (2000) interpreted these as sea-floor-spreading anomalies and suggested that oceanic crust extends all the way from anomaly M3 to the edges of continental crust (to as old as anomaly M20, Tithonian, in the southern rift). However, considering the fact that the TZ appears to be serpentinized peridotite (with overlying thin continental crust in some parts of

the inner TZ) there may be two possible explanations for the source of the anomalies. One is that there are igneous intrusions in the peridotites and thinned continental crust. Detailed magnetic analyses on the Iberia margin (Whitmarsh & Miles 1995; Russell & Whitmarsh 2003) suggest that enough melt was emplaced at deep levels about the time of chron M5–M3 to form recognizable magnetic anomalies, and small amounts of melt emplaced at earlier times might help account for older anomalies. A second possibility is that the anomalies are generated by magnetization of the serpentinites. Magnetic studies of serpentinites cored on both margins show NRM intensities of 1–9 A m21, comparable to values in basalts, suggesting that they could generate significant magnetic anomalies (Zhao 1996; Shipboard Scientific Party 2004). In this case, the magnetization contrast between peridotite ridges and intervening sedimentary basins might explain at least part of the weak magnetic anomalies. Observed co-linearity of magnetic anomalies and peridotite ridges in the TZ is consistent with this idea (Fig. 3).

Oceanic crust Normal oceanic crust is considered by most investigators to have formed in the Newfoundland–Iberia rift beginning by about chron M3, but certainly no later than chron M0 (c. 121 Ma, earliest Aptian). There is less information on the velocity structure of this basement than that of the TZ landward of chron M3 and unfortunately there are no refraction profiles that extend seaward onto Albian crust, i.e. beyond the time when the previously defined ‘breakup unconformity’ is thought to have formed (Mauffret & Montadert 1987). The IAM9 profile seaward of anomaly M3 on the Iberia margin (Fig. 4g, section I2) shows crustal thickness and velocity that appear to be normal for oceanic crust, except at the western end where crust thickens significantly as the profile intersects a seamount. On the opposing margin (Fig. 2), the SCREECH3 profile seaward of anomaly M3 also has crustal thickness and velocity expected for normal ocean crust (Lau et al. 2006b). It is noteworthy, however, that the seaward ends of both the IAM9 and SCREECH3 profiles lie near seamounts and the northern edge of a zone that may have been influenced by plume magmatism (see below). In contrast to these profiles, the SCREECH2 profile (Fig. 4g, section N2) shows apparent oceanic crust that is only 2 –4 km thick above a 7.2–8.0 km s21 layer that has highly variable thickness and may be serpentinized mantle. Within approximately 80 km to either side of this line, refraction profiles over supposed oceanic crust at anomaly M0 (Srivastava et al. 2000) and

BREAKUP OF THE NEWFOUNDLAND– IBERIA RIFT

up to 50 km seaward of M0 (Todd & Reid 1989) also show a thin, low-velocity layer (2–2.5 km, 4.5–5.3 km s21) over probably serpentinized mantle (6.9–7.5 km s21) that is at least 4–5 km thick. To the north, the SCREECH1 profile (Figs 2 & 4 g, section N3) shows a large thickness (up to 5 km) of apparently serpentinized mantle with velocities of 7.2–7.9 km s21 (Funck et al. 2003). The overlying basement layer is very thin (1–3 km), and is interpreted as oceanic on the basis of velocity structure and reflection characteristics (Hopper et al. 2004). Normal-thickness oceanic crust is interpreted at the seaward end of the profile beyond anomaly M0, although this is not well constrained by the refraction data. The conjugate Galicia Bank profile (Fig. 4g, section I3) also shows a short section of apparent oceanic crust near its western end (Whitmarsh et al. 1996). A separate feature of the young M-series basement that suggests a magmatic component during its emplacement is the presence of the highamplitude J anomaly. This magnetic anomaly is thought to have formed in association with anomalous melt production and northward intrusion from the southern margin of the rift from about the time of chron M4 to a time younger than M0 (Tucholke & Ludwig 1982) (Fig. 4c; see the subsection on ‘Rift magmatism’ later). The anomaly amplitude is highest at the southern edge of the rift in the area of the Southeast Newfoundland Ridge (SENR), J Anomaly Ridge (JAR) and Madeira–Tore Rise (MTR) (Fig. 1); away from this area the high amplitude is confined to the zone around chrons M1 and M0 (Rabinowitz et al. 1978). Within the Newfoundland–Iberia rift the anomaly gradually loses its high amplitude north of the Newfoundland Seamounts and Tore Seamount on the Newfoundland and Iberia margins, respectively (bold lines, Figs 2 & 3). Farther north, anomaly M0 is identified with normal amplitudes between Flemish Cap and Galicia Bank, and it is also interpreted to be present in the Bay of Biscay (Sibuet et al. 2004b). If we assume that the amplitude of the J anomaly reflects the distribution of magma introduced from the southern edge of the rift, the effect reached to near the latitude of the ODP Leg 149– 173 drilling transect. Thus, the thickened crust in the young part of the IAM9 and SCREECH3 profiles might reflect this influence (e.g. Fig. 4g, section I2). The above data show that the structure of basement younger than approximately M3 is highly variable, ranging from thin crust with apparently limited serpentinization of underlying mantle, to thin with seemingly extensive serpentinization, to normal or thicker than normal with no clear serpentinization. Thus, in some places there appears to be a magmatic component while in others none is

17

evident. It is noteworthy that in the two locations where the presumed oceanic crust has been drilled, it is exhumed mantle. Basement at Site 1070, west of anomaly M3 on the Iberia margin (Figs 3 & 4g, section I2), is serpentinized peridotite with minor gabbro intrusions (Whitmarsh et al. 1998). Basement at Site 1277 near anomaly M1 off Newfoundland is also serpentinized peridotite beneath a thin interval of allochthonous basaltic, gabbroic and serpentinite debris (Figs 2 & 4g, section N2; see also Fig. 10 later) (Tucholke et al. 2004). It is well known that serpentinized peridotites can have compressional-wave velocities ranging from less than 5 km s21 up to normal mantle velocities of 8þ km s21, depending on the degree of serpentinization (e.g. Christensen 2004). Thus, in the absence of other constraints such as Poisson’s ratio or velocity anisotropy, and considering that serpentinites were recovered at Sites 1070 and 1277, it is possible that much of the basement layer that is interpreted from seismic velocity data to be normal oceanic crust in the area of the young M-series anomalies is variably serpentinized peridotite.

Development of the rift Pre-rift crust Continental crust in which the Newfoundland– Iberia rift developed was an assemblage of Precambrian– Palaeozoic rocks that accreted during the closing of the Palaeozoic Iapetus and Rheic oceans. Along the present Atlantic Canadian margin these rocks form the Appalachian Orogen, and from NW to SE they appear in three terranes comprising the Dunnage, Gander and Avalon zones (Fig. 1) (Williams & Hatcher 1982, 1983). The Avalon Zone occupies easternmost Newfoundland and most of the Grand Banks platform. Another zone (Meguma) lies south of the Avalon terrane at the southern edge of the Grand Banks, and a prominent magnetic anomaly (Collector Anomaly) marks its northern boundary. The first three terranes were accreted during the Taconian, Salinic and Acadian orogenies (Early–Middle Ordovician, Early Silurian and Devonian respectively) (van Staal et al. 1998; Waldron & van Staal 2001; Percival et al. 2004). The Meguma probably was emplaced during the Acadian orogeny, but it may have been reactivated during the Alleghanian orogeny (Carboniferous –Permian). Large granitic plutons were emplaced in the then-existing terranes during the Acadian orogeny. The Gander–Avalon and Avalon–Meguma boundaries are steep ductile shears and brittle faults that imply assembly by transcurrent motion rather than subduction/obduction (Keen et al. 1990). Original structural fabrics within the Avalon and Meguma terranes are well

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defined by gravity and magnetic data, and in many places they appear to have controlled subsequent development of rift basins within the Grand Banks (Welsink et al. 1989). Continental basement of the Iberia margin was accreted by closure of Palaeozoic ocean basins and suturing during the Hercynian (Variscan) orogeny in approximately Middle Devonian –Carboniferous time, i.e. at the same time as the late Acadian and Alleghenian orogenies in North America. In Iberia, the accreted terranes have NW–SE orientations (Fig. 1); offshore, they are thought to have curved back to the NE and east to form the Ibero-Armorican Arc, continuous with approximately east– west sutures in NW France (Armorican Massif ), SW Britain and Germany (Ziegler 1982; Capdevila & Mougenot 1988; Silva et al. 2000). These terranes were intruded by large granitoid batholiths both during and following the Hercynian orogeny (Pinheiro et al. 1996). Tonalites also were intruded in the original continental crust of Galicia Bank (ODP Sites 1067) during the Hercynian and earlier (Rubenach 1999). Late Hercynian deformation included complex brittle faulting in generally NE–SW sinistral and conjugate NW–SE dextral shears (e.g. Ribeiro et al. 1990), and these features exerted significant control on the subsequent structure of the Mesozoic rifted margin (Wilson et al. 1989).

Mesozoic rifting Rifting between Newfoundland and Iberia occurred primarily in two phases. The initial phase occurred in the Late Triassic –earliest Jurassic, at which time rift basins formed over a broad region within the Grand Banks (e.g. Jeanne d’Arc, Whale and Horseshoe basins; Fig. 1) and the Iberia margin (Lusitanian Basin, possibly Porto and Galicia Interior basins: Murillas et al. 1990). The Grand Banks basins accumulated siliciclastic ‘red-bed’ sediments during Carnian– Norian time, and these were succeeded by evaporite deposits that reached into the earliest Jurassic (Hettangian– Sinemurian) in both the Grand Banks and Lusitanian basins (Jansa & Wade 1975; Wilson 1988; Rasmussen et al. 1998; Alves et al. 2003). During the Early and Middle Jurassic, the Grand Banks experienced epeirogenic subsidence without further rifting (Tankard & Welsink 1987) while an episode of faulting and basin subsidence occurred in the Lusitanian Basin south of the Nazare Fault (Rasmussen et al. 1998). A second phase of extension began in the Late Jurassic. During this phase, extension shifted from a wide-rift mode that encompassed future proximal margins, and it focused at distal margins where continental crust eventually separated (e.g. Manatschal & Bernoulli 1999). On the Grand Banks, extension

appears to have persisted with variable intensity from Oxfordian to Aptian time, with the most intense rifting occurring in the Kimmeridgian – Tithonian (Tankard & Welsink 1987). Unconformities that were developed in the Jeanne d’Arc Basin in late Barremian–early Aptian and late Aptian time (Fig. 4f ) have been interpreted by Driscoll et al. (1995) to have formed during the final rifting events on the main Grand Banks, immediately preceding the onset of sea-floor spreading. After this time, extension shifted to the northern margin of the Grand Banks (e.g. Orphan Basin) where it was associated with opening of the Labrador Sea (Enachescu 1988; Grant et al. 1988). On the Iberia margin, rifting in the Lusitanian Basin was limited largely to the Oxfordian–early Kimmeridgian, although there may have been later minor extension in Tithonian–Barremian time (Wilson 1988; Wilson et al. 1989; Rasmussen et al. 1998). In the Galicia Interior Basin to the north, Murillas et al. (1990) analysed seismic stratigraphy and ties to wells on the shallower Iberia margin. Their interpretations show that probable Tithonian– Berriasian shallow-water carbonates (seismic sequence 6, Fig. 4b) are highly faulted, while the overlying Valanginian–Hauterivian sequence 5 is only locally faulted, primarily towards the centre of the basin (see also Pe´rez-Gussinye´ et al. 2003). Thus, rifting in the Interior Basin appears to have occurred largely in Berriasian time, with some extension continuing into the Valanginian. On thinned continental crust at the deep western and southern margins of Galicia Bank, drilling and dredge/submersible sampling have recovered shallow-water Tithonian– (?)Berriasian carbonates and claystones of sequence 6, succeeded by deeper-water carbonates and clastics (Boillot et al. 1989c; Whitmarsh & Sawyer 1996; Whitmarsh & Wallace 2001). From these data the first rifting of Galicia Bank has been interpreted to date to the Berriasian (Fig. 4e). Rifting continued through the Valanginian and Hauterivian on the western margin of Galicia Bank, disrupting sequences 6 and 5 (Figs 5–7) (Mauffret & Montadert 1987). Some parts of the southern margin also show signs of extension during this period (Figs 8b & 9) but other parts were inactive (Fig. 8c). The last stage of rifting on the Iberia margin is generally interpreted to have occurred in the latest Aptian –earliest Albian (Mauffret & Montadert 1987; Boillot & Winterer 1988). Wilson (1988) discussed an Aptian unconformity in the Lusitanian Basin (Fig. 4f ) and suggested that it developed at the onset of sea-floor spreading. In contrast, Rasmussen et al. (1998) did not identify this event in the basin, although their figures show a local unconformity between Lower and Upper Cretaceous beds that may be equivalent. On the deep Iberia

Fig. 5. (a) Section of ISE 14 multichannel reflection profile across the SW margin of Galicia Bank. Location in Figure 3 (ISE 14E). The basement highs are continental fault blocks. Closed arrows mark the Aptian event between seismic sequences 3 and 4, and open arrows mark the unconformity between sequences 4 and 5. Note the level attitude of sequence 4. See Figure 4b for sequence ages and lithology. (b) Enlarged segment of the profile with interpreted structure and stratigraphic relations of the deep sedimentary sequences. Sequence 6 capping basement is probably Tithonian– Berriasian limestone and dolomite. Sequence 5 (Valanginian–Hauterivian) shows little divergence of sedimentary reflections and was faulted rapidly together with sequence 6 near the end of the Hauterivian. Sequence 4 (Barremian– Aptian) is horizontal and was not faulted or rotated. Note the unconformable onlap of sediments onto sequence 4.

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Fig. 6. (a) Section of ISE 2 multichannel reflection profile across the western margin of Galicia Bank just south of the ODP Leg 103 drilling transect. Location in Figure 3. The basement ridges at the right are continental fault blocks above the S reflection, which is interpreted to be a detachment contact with underlying mantle (e.g. Reston et al. 1995); some of the distortion of S is probably an artifact of velocity pull-up in this time section. Closed arrows at top mark the Aptian event between seismic sequences 3 and 4, and open arrows mark the unconformity between sequences 4 and 5. Note the strong reflectivity of sequence 4; reflectivity is particularly strong near the peridotite ridge at left, which suggests that the sequence contains abundant serpentinite debris shed from the ridge. See Figure 4b for sequence ages and general lithology. (b) Enlarged segment of the profile with interpreted structure and stratigraphic relations of the deep sedimentary sequences. Rotated and diverging reflections in sequence 5 document extension in the Valanginian– Hauterivian (note that divergence of reflections would be enhanced in a depth section because of velocity increase with depth). Sequence 4 (Barremian–Aptian) was not faulted. Note the unconformable onlap of sediments onto sequence 4. ODP Site 640, 7 km to the north, penetrated about 65 m of deformed Hauterivian –Barremian sandy turbidites and marlstones that correspond to the top of sequence 5 and the base of sequence 4 (Boillot et al. 1987).

BREAKUP OF THE NEWFOUNDLAND– IBERIA RIFT

21

margin, a prominent, widespread reflection between Barremian–Aptian seismic sequence 4 and overlying Albian sediments has been interpreted as the ‘breakup unconformity’ (Groupe Galice 1979; Mauffret & Montadert 1987, 1988). This Aptian event (Figs 5–8) is the same as the ‘orange reflection’ at DSDP Site 398 on the southern part of Galicia Bank (Figs 4b & 9). The end-of-Aptian breakup that has been interpreted off Iberia is nearly synchronous with the development of the late Aptian ‘rift-onset’ unconformity in the Jeanne d’Arc Basin (Fig. 4f ) (Driscoll et al. 1995). In the offshore Newfoundland TZ, recent drilling at ODP Site 1276 dated a strong, basin-wide reflection as the same age (U reflection, Figs 4b & 10) (Tucholke et al. 2004). The seismic- and time-stratigraphic similarity of this horizon and the Aptian event off Iberia indicates that both originated from the same process. Together, these seismic markers document a major event during the evolution of the rift. The above interpretations that rifting continued throughout the entire Early Cretaceous into Aptian time are consistent with radiometric age dates on seafloor submersible/dredge samples and drill cores in the deep basin (Fig. 4c). Metagabbros that were emplaced in continental basement at ODP Sites 900 and 1067 during Hercynian time (Rubenach 1999; Manatschal et al. 2001) passed through the hornblende blocking temperature (c. 500 8C) at 161 + 1 Ma (Callovian) and through the plagioclase blocking temperature (c. 150 8C) at approximately 136–137 Ma (Berriasian) (Fe´raud et al. 1996); the latter age is interpreted to mark the time of exhumation (Manatschal et al. 2001). It is not known how much longer extension continued at this location, but it may have persisted through the Barremian or later (Whitmarsh & Wallace 2001). On peridotite ridges R1 and R2 bounding extended continental crust at the western margin of Galicia Bank (Fig. 3), granulites and sheared intrusive diorites and gabbros in the peridotites were cooled, either before or during exhumation, at ages ranging from 129.3 + 13.4 Ma to 117.7 + 1.8 Ma (approximately Hauterivian –middle Aptian) (Fe´raud et al. 1988; Boillot et al. 1989a; Fuegenschuh et al. 1998; Scha¨rer et al. 1995, 2000). To the south at ODP Site 1070, gabbro pegmatites in peridotites passed through the hornblende blocking temperature at 119 + 0.7 Ma (Aptian) and are interpreted to have been exhumed during the earliest Albian (110.3 + 1.1 Ma) as they passed through the plagioclase blocking temperature (Whitmarsh & Wallace 2001). Extension of basement at the SW margin of Galicia Bank during or following late Aptian time has also been interpreted at ODP Sites 897 and 899, both of which were drilled on the tops of

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Fig. 7. (a) Section of GP 101 multichannel reflection profile over continental crust near ODP Sites 638, 639 and 641 on the western margin of Galicia Bank (from Winterer et al. 1988). Location in Figure 3. Site 639 is about 0.5 km south of the profile, and Sites 638 and 641 are about 1 km to the north. Closed and open arrows mark the tops of seismic sequences 4 and 5, respectively. See Figure 4b for sequence ages and lithology. (b) Enlarged section from the western part of the profile with interpreted structural and stratigraphic relations. In its upper part, sequence 5 (Valanginian–Hauterivian) shows divergence and rotation of sedimentary reflections, indicating that extension here occurred largely in the Hauterivian (note that divergence of reflections would be enhanced in a depth section because of velocity increase with depth). Sequence 4 (Barremian– Aptian) lacks these features and shows subhorizontal onlap at the basin edges; a chaotic mass at the base of the sequence may be a slump from the basement ridge to the east. Indicated ages were determined at the drill sites (Boillot & Winterer 1988). Hole 639D recovered Tithonian limestones, and Holes 639E and 639F recovered rhyolites that are interpreted to come from continental, possibly Palaeozoic basement (BMT). Hole 639A penetrated lower Valanginian marlstone and sandstone and bottomed in dolomite, which was also cored at the bottoms of Holes 639B and 639C. The dolomite is barren of microfossils but must date to the latest Jurassic –earliest Cretaceous. The S reflection (see Fig. 6) underlies the continental block on which these sites were drilled. (c) Enlarged section from the eastern part of profile in a with interpreted structural and stratigraphic relations (note that vertical exaggeration is slightly enhanced compared to profile in a). Reflections diverging towards the fault footwall in the upper part of sequence 5 suggest basin extension in late Valanginian– early Hauterivian time. Barremian–Aptian sequence 4 exhibits horizontal reflections that lap unconformably onto the basin perimeter, and it in turn is unconformably overlain by sequence 3.

BREAKUP OF THE NEWFOUNDLAND– IBERIA RIFT

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peridotite ridges (Fig. 3) (Comas et al. 1996). At these sites, serpentinized peridotite olistostromes and debris flows were emplaced on the basement during early Aptian and early late Aptian time, respectively, before faulting and uplift formed the ridges.

Rift magmatism During most of the Late Jurassic–Early Cretaceous phase of extension, there was only minor magmatism in the rift (Fig. 4c). Tithonian–early Valanginian lamprophyre dykes and a small ultramafic intrusion are documented in Newfoundland, and recoveries of basalt and volcaniclastics both of this age and of late Barremian–Aptian age have been reported from wells in the South Whale Basin and southern Grand Banks (Jansa & Pe-Piper 1988; Sinclair 1988 and references therein). Dykes and sills of primarily Berriasian–Valanginian age are present in the southern Lusitanian Basin of Iberia (Montenat et al. 1988; Pinheiro et al. 1996). All these occurrences are volumetrically minor in the context of the larger rift and all are restricted to the southern part of the rift. However, the age of their emplacement is consistent with the timing of the most intense rifting in the latest Jurassic–Early Cretaceous (Fig. 4). By far the most extensive magmatism within the rift dates to the Barremian–early Aptian (approximately chron M4 to younger than M0; Fig. 4c) (Tucholke & Ludwig 1982). At this time the SENR and JAR formed on the North American side of the plate boundary, and the conjugate MTR and perhaps parts of Gorringe Bank (GRB; later deformed during the Cenozoic) were emplaced on the Iberia side (Fig. 1). Magmatism was centred at the southern edge of the rift in the position of the SENR –GRB, and at approximately chron M2– M0 (Barremian) it was channelled both southwards along the Mid-Atlantic Ridge axis to form the JAR –MTR (Tucholke & Ludwig 1982) and northwards into the Newfoundland –Iberia rift to form comparable basement edifices that reach toward the present positions of the Newfoundland Seamount and Tore Seamounts (Fig. 1) (Tucholke et al. 1989). At least at the southern margin of the rift, some of this magmatic crust was at and above sea level. Deep-sea drilling at DSDP Site 384 near anomaly M2 on the JAR recovered upper Barremian/lower Aptian –lower Albian reef deposits overlying tholeiitic basalt that most probably was subaerially weathered (Tucholke et al. 1979); the reef appears subsequently to have been exposed to meteoric waters, probably during Albian time, before it subsided to its present depth of 4100 m. Overall, the broad morphological expression of the SENR – JAR –MTR magmatic complex was very much like that of modern Iceland and the associated

Fig. 8. (a) Section of ISE 10 multichannel reflection profile along the southern margin of Galicia Bank. Location in Figure 3 (ISE 10E). The basement ridges are continental fault blocks, based on their continuity with interpreted continental crust at the ODP drill sites 30–40 km to the south. Closed arrows mark the Aptian event between seismic sequences 3 and 4, and open arrows mark the unconformity between sequences 4 and 5. Note the strong reflectivity and drape-and-fill character of sequence 4. See Figure 4b for sequence ages and lithology. (b) Enlarged segment of the profile with interpreted stratigraphic relations of the deepest sedimentary sequences. Sequence 5 (Valanginian – Hauterivian) in its lower part shows rotation and divergence of sedimentary reflections towards the fault footwall, indicating basin extension in the Valanginian. Note that sequence 4 (Barremian –Aptian) shows horizontal attitude and onlap at the basin edges and that sequence 3 laps unconformably onto sequence 4. (c) Enlarged segment of the profile in a with interpreted stratigraphic relations of the deepest sedimentary sequences. There is no apparent rotation or divergence of reflections in either sequence 4 or 5, indicating that extension in this basin ceased by about late Berriasian time. Note the chaotic debris immediately below the Aptian event on the right, indicating mass wasting from the adjacent basement ridge.

24 B. E. TUCHOLKE ET AL.

BREAKUP OF THE NEWFOUNDLAND– IBERIA RIFT

25

Fig. 9. (a) Section of GP 19 multichannel reflection profile across DSDP Site 398 on the southern margin of Galicia Bank. Location in Figure 3. (b) Interpretation of the profile, with seismic sequences as in Figure 4b. Sequence 5 is faulted but sequence 4 is basin-filling and relatively flat-lying. Note the strong reflectivity at the top of sequence 4. Ages determined in the lower part of the borehole are indicated.

Reykjanes and Kolbeinsey ridges (although at a somewhat smaller scale), and it appears likely that the magmatism was associated with a mantle plume. Either a thicker or more magnetized basaltic crust emanating from this plume probably accounts for the development of the magnetic J anomaly between M1 and M0, which has a high amplitude near the SENR–GRB but decays with distance both to the north (Figs 2 & 3) and south (Rabinowitz et al. 1978; Tucholke & Ludwig 1982). The younger age limit of this magmatic complex is not well constrained, but it appears that magmatism retreated rapidly towards the southern margin of the rift following chron M0 and died out by early Albian time. Post-rift magmatism in the Newfoundland Basin is observed in the form of diabase sills that were injected into sediments at and below the U reflection at ODP Site 1276 during late Albian –early Cenomanian time (105 –98 Ma; Hart & Blusztajn 2006). These sills are not clearly observed in seismic reflection records around the drill site

(Fig. 10), but probable sills at this stratigraphic level appear in profiles from the southern part of the basin (Fig. 11). The source of the magma is uncertain, but it may have been associated with formation of the Newfoundland Seamounts near the centre of the basin (Karner & Shillington 2005); a single age date indicates that one of these seamounts formed at 97.7 + 1.5 Ma (Sullivan & Keen 1977). Only minor magmatism of this age is documented on the Iberia margin (Fig. 4c), and the event seems primarily to have affected the Newfoundland margin.

Questions related to breakup The features of the Newfoundland–Iberia rift described above, and prevailing interpretations that the ‘breakup unconformity’ dates to the end of the Aptian, lead to several fundamental problems in understanding final continental separation and the initiation of sea-floor spreading. First, existing

Fig. 10. (a) Section of SCREECH2 multichannel seismic reflection profile (Shillington et al. 2004) from the seaward part of the TZ in the north-central Newfoundland Basin. Location in Figure 2 (SCREECH 2NW). Locations of magnetic anomalies M3 and M1 are indicated. ODP Site 1276 did not reach basement but penetrated 105 –98 Ma base sills that intrude uppermost Aptian –lowermost Albian sediments at the level of the U reflection, which is the time-stratigraphic equivalent of the ‘orange reflection’ and Aptian event on the Iberia margin (see Fig. 4). Note the basin-levelling character of U. The basement surface is poorly imaged beneath the U reflection because of strong reflectivity of U, because of low impedance contrast with the U-basement interval, or both. Drilling at Site 1277 recovered massive serpentinized peridotite, interpreted as basement, beneath approximately 57 m of mass-wasting debris consisting of basalt, gabbro and serpentinite (Tucholke et al. 2004). (b) Enlarged segment of the profile with interpreted structural and stratigraphic relations. If the identification of U is correct, the sediments of seismic sequence A and the underlying basement block were rotated during approximately latest Aptian time, prior to deposition of overlying, onlapping Albian sediments of sequence B. See Figure 4 for sequence ages and lithology.

26 B. E. TUCHOLKE ET AL.

BREAKUP OF THE NEWFOUNDLAND– IBERIA RIFT

27

Fig. 11. Section of Conrad 2510 multichannel reflection profile from the southern Newfoundland Basin (Tucholke et al. 1989). Location in Figure 2 (NB26). See Figure 4b for sequence ages and lithology. Sedimentary and basement structure below the U reflection (Aptian event) is poorly imaged because of the strong reflectivity of U, because of low impedance contrasts, or both. Short sections of high-amplitude reflections near the level of U probably are diabase sills, similar to those drilled at this stratigraphic level at ODP Site 1276 in the northern part of the basin (see Fig. 10).

magnetic analyses (e.g. Russell & Whitmarsh 2003) and refraction studies (Dean et al. 2000; Lau et al. 2006b; Van Avendonk et al. 2006) suggest that some form of oceanic crust was emplaced in the rift at least by the time of chron M3 (late Hauterivian –Barremian) (Figs 2 & 3). Thus, the Aptian event would appear not to be a ‘breakup unconformity’ in the conventional sense that it correlates to the final separation of continental crust and coincident onset of sea-floor spreading (Falvey 1974). Second, there is a conflict between previously described evidence for extension during Barremian–Aptian time and the corresponding structural and stratigraphic record as imaged in seismic reflection records both over continental crust of Galicia Bank and over the Newfoundland and Iberia TZs. Sediments of this age in seismic sequence 4 off Iberia and sequence A off Newfoundland (Fig. 4b) are mostly horizontal, and they rarely are faulted or show splayed reflections that we would expect if they were deposited over rotating fault blocks during extension (Figs 5–15).

Finally, in at least three cases, there was a significant lag between the time when upper mantle was first emplaced in the rift and a time when it was later extended. One instance is at peridotite ridges R1 and R2 at the western margin of Galicia Bank. Recent magnetic anomaly compilations (Miles et al. 1996) and identifications (Srivastava et al. 2000; Sibuet et al. 2004b) indicate that anomaly M0 lies approximately 50 km west of these ridges (Fig. 3). If we extrapolate basement age east of anomaly M0, assuming an accretion rate of 9 mm year21 (the average rate between anomaly M0 and anomaly 34 to the west), the age of the ridges is about 126–127 Ma (see section I3, Fig. 4g). Gabbro protoliths of chlorite rocks in these ridges were emplaced at 122.1 + 0.3 to 121.7 + 0.4 Ma (Scha¨rer et al. 1995, 2000) and subsequently sheared during exhumation. In the area of ODP Site 637, tectonized diorite dykelets in peridotite mylonites cooled below approximately 500 8C at about 122 Ma and below approximately 150 8C at 117.7 + 1.8 Ma as they were exhumed (Fe´raud et al. 1988; Boillot et al. 1989a). Thus, exhumation of the peridotites appears to post-date

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initial emplacement of the basement by at least 4–8 million years. A similar interval between the time of basement emplacement and later extension is indicated by the study of Comas et al. (1996) at ODP sites 897 and 899 farther to the south. According to magnetic anomaly identifications (Fig. 3), basement ages there are approximately 125– 126 Ma, but faulting and uplift of the peridotite ridges did not occur until late Aptian time (c. 115 Ma) or later, a delay of at least 10 million years. An even larger time lag is observed at ODP Site 1070. Basement there was emplaced at about 124 Ma according to the identification of anomaly M1 (Fig. 3), and small gabbro pegmatites intruded the peridotites at essentially the same time (127 + 4 Ma; Beard et al. 2002). The gabbros cooled through approximately 600 8C at 119 + 0.7 Ma and through about 150 8C at 110.3 + 1.1 Ma (earliest Albian) (Whitmarsh & Wallace 2001). These small intrusions would have equilibrated quickly with the temperature of surrounding peridotites, so their long cooling history indicates that they were emplaced in hot rock at significant depth before they were later exhumed. Assuming a geothermal gradient of 25–30 8C km21 (continental, Whitmarsh & Wallace 2001) to approximately 90 8C km21 (oceanic, Expedition 309 Scientists 2005), the rocks were still at a depth of about 2–6 km at c. 110 Ma. The exhumed peridotite is covered by upper Aptian sediments (Whitmarsh et al. 1998), potentially as young as 112.2 Ma but still at least 0.8 million years older than the earliest possible exhumation suggested by the radiometric date (111.4 Ma). This inconsistency probably can be attributed to uncertainties in the absolute ages. However, unless the radiometric date is grossly in error, it appears that exhumation was rapid and that it post-dated the age of original basement emplacement by approximately 14 million years. We suggest that these discrepancies can best be explained by elevated plate-wide tensile stress and rifting that persisted throughout the Early Cretaceous until the end of Aptian time. Below, we examine this hypothesis by evaluating the deep structure and stratigraphy of the conjugate Newfoundland and Iberia margins.

Deep stratigraphy and structure of the rift Seismic sequences To examine the timing and character of rift events we analysed the deep stratigraphy and structure of Newfoundland and Iberia margins using existing seismic reflection profiles and correlations to DSDP and ODP drill sites (Figs 2 & 3). The Aptian event can be traced continuously throughout

the seismic grid except at the top of Galicia Bank where jump correlations have to be made on the basis of seismic character (e.g. Fig. 7a, right-hand side). Deeper seismic sequences in the Galicia Bank area commonly are in graben separated by basement ridges. Where they could not be traced continuously between basins these sequences were correlated on the basis of reflection character, relations to adjacent sequences (including development of bounding unconformities) and the limited control provided at drill sites (Figs 7 & 9). On the Iberia margin we used profiles from the Iberia Seismic Experiment (ISE) lines (Henning et al. 2004), CAM lines (Discovery 215 Working Group 1998; Chian et al. 1999), GP lines (e.g. Mauffret & Montadert 1987; Winterer et al. 1988), and the Lusigal 12 (Beslier 1996) and IAM9 lines (Pickup et al. 1996). On the Newfoundland margin we analysed profiles from R/V Conrad cruise 2510 (Tucholke et al. 1989), Lithoprobe (de Voogd & Keen 1987), ERABLE (Srivastava & Sibuet 1992) and R/V Ewing cruise 2007 (SCREECH program: Hopper et al. 2004; Shillington et al. 2004). The principal seismic sequences and their lithological correlations on the Iberia margin (Fig. 4b) were defined by Groupe Galice (1979) and later extended to include sequence 5 (Boillot et al. 1979, 1987). Mauffret & Montadert (1987) subdivided sequence 5 into 5a and 5b, and subsequently renamed 5a as 5 and 5b as 6 (Mauffret & Montadert 1988); this last classification is adopted in Figure 4. The age of the sequence 6–5 boundary is not well constrained; drilling at ODP Site 639 (Fig. 7) recovered lower Valanginian sediments over undated (possibly Berriasian) dolomites near the top of sequence 6 (Boillot & Winterer 1988), and these in turn overlie Tithonian shallow-water limestones. In Figure 4 we tentatively place the sequence boundary at the Berriasian–Valanginian boundary. The sequence 5 –4 boundary at nearby Site 638 dates to late Valanginian–Hauterivian, depending on the fossil group considered (Boillot et al. 1988). At Deep Sea Drilling Project (DSDP) Site 398 on the southern Galicia Bank (Fig. 9), the boundary is probably upper Hauterivian –lower Barremian (Sibuet & Ryan 1979). Although we place the 5–4 boundary at this latter level in Figure 4, it is possible that the boundary is somewhat diachronous and reflects locally different timing of extension events around Galicia Bank. The sequence 4– 3 boundary (Aptian event) is manifested as the ‘orange reflection’ in the area of Galicia Bank (Fig. 9) (Groupe Galice 1979), and the correlative sequence A –B boundary in the Newfoundland Basin is represented by the highamplitude U reflection (e.g. Figs 10 & 15) (Tucholke et al. 1989). Dates on the Aptian event consistently place it in the uppermost

BREAKUP OF THE NEWFOUNDLAND– IBERIA RIFT

Aptian –lowermost Albian where it has been drilled at Sites 398 and 641 on Galicia Bank (Sibuet & Ryan 1979; Boillot et al. 1987) and at Site 1276 on the Newfoundland margin (Tucholke et al. 2004).

Tithonian – Berriasian (sequence 6) Seismic sequence 6 (Figs 5–8) contains carbonates and mudstones deposited in water less than a few hundred metres deep on continental crust of the Iberia margin during the Tithonian and possibly Berriasian (Boillot & Winterer 1988; Whitmarsh & Wallace 2001). These sediments often are difficult to distinguish seismically from the underlying basement, probably because they include coarse, high-velocity debris from the underlying Palaeozoic basement as well as high-velocity dolomites (Boillot & Winterer 1988). Extensional faulting of this sequence occurred at two different times, as illustrated by contact relations with the overlying sequence 5. Over much of Galicia Interior Basin and the southern margin of Galicia Bank (Fig. 8), sequence 5 is not conformable to sequence 6 or to basement (Murillas et al. 1990), indicating that an episode of rifting faulted sequence 6 shortly before the Valanginian. Because the internal structure of sequence 6 is hard to detect, however, it is difficult to determine (e.g. by divergence of reflections) whether the extension began much earlier, for example in the Tithonian. In a few cases where internal reflections are observed (Figs 5b & 6b), they are mostly parallel. This suggests that faulting was confined to a short period in the Berriasian, and it is consistent with the development of unconformities in the Lusitanian and Porto basins (Fig. 4f ) that are ascribed to rifting. Age dates of 136–137 Ma on metagabbros at ODP Sites 900 and 1067 on the southern edge of the Galicia Bank crust show that basement there was being cooled and presumably exhumed near the end of Berriasian time (Fig. 4c) (Manatschal et al. 2001). In contrast to the above observations, sequence 5 is largely conformable to sequence 6 over the western, distal margin of Galicia Bank (e.g. Figs 5–7) and both sequences were faulted together. Numerous published reflection profiles show this relationship (e.g. Mauffret & Montadert 1987, 1988), indicating that most of the extension at the distal, western Galicia margin post-dates extension in the Interior Basin (see the next section). On the Newfoundland margin, time-stratigraphic equivalents of sequence 6 appear in the deep sedimentary basins of the Grand Banks (Grant et al. 1988; Tankard & Welsink 1988; Balkwill & Legall 1989) and extend seawards into basins beneath the continental slope. These rocks are mostly shales– sandstones with interbedded limestones, and they reflect rapid fill that kept pace with fault-controlled

29

basin subsidence. Interpretation of well and seismic data indicates that the most intense extension was in Late Jurassic time, decreasing thereafter into the Early Cretaceous (Fig. 4e). A widespread unconformity in the Jeanne d’Arc Basin (Tankard & Welsink 1988) suggests enhanced extension there at the Jurassic–Cretaceous boundary. The above data illustrate two important aspects of the rifting process. First, although there was probably rift-wide extension throughout this period, normal faulting and basin development were often amplified during relatively short pulses. Second, these pulses were not necessarily synchronous or co-located within the rift. However, they do tend to group at the end of the Jurassic into the Berriasian, suggesting a general extensional episode that culminated approximately in Berriasian time.

Valanginian– Hauterivian (sequence 5) Seismic sequence 5, where drilled off Iberia, consists of limestones, marlstones, mudstones and turbidites that generally reflect deeper-water deposition than in the underlying sequence (Sibuet & Ryan 1979; Boillot et al. 1987). The sequence displays three kinds of configuration with respect to the underlying, faulted basement and/or sequence 6. By far the most common is a conformable relation in which internal reflections are parallel both to one another and to the underlying formation in rotated fault blocks (Fig. 5b). This configuration is widespread across the western margin of Galicia Bank. The general lack of reflections that diverge towards fault footwalls in sequence 5 indicates that normal faulting occurred during a geologically brief period of time, focused near the Hauterivian–Barremian boundary. This extensional episode may correlate with interpreted exhumation and cooling of granulites associated with peridotite ridges R1 and R2 at approximately 129–126 Ma (Fig. 4c) (Fuegenschuh et al. 1998). Less frequently, beds in sequence 5 thicken towards the footwall (Figs 6b, 7b & 8b). In the case of the profile in Figure 8b along the southern margin of Galicia Bank, weakly divergent reflections are present in the lower part of sequence 5 and relatively parallel reflections appear in the upper part. This suggests that the footwall comprising the Hobby High ridge, drilled at ODP Sites 900, 1067 and 1068 about 40 km to the south, was being exhumed during the Valanginian, but that extension slowed or ceased by Hauterivian time. This is consistent with the ages at which metagabbros drilled on Hobby High are interpreted to have been exhumed (Fig. 4c) (Manatschal et al. 2001). On parts of the western Galicia margin (Figs 6b & 7b, c) reflections diverging towards fault footwalls in

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sequence 5 indicate that extension continued throughout the Valanginian–Hauterivian. Finally, in some places sequence 5 is represented by tectonically undeformed basin fill which demonstrates that extensional faulting ended no later than about Berriasian time (Fig. 8c). This is commonly observed in profiles across the Galicia Interior Basin (Murillas et al. 1990), although faulting of sequence 5 is apparent locally along the western margin of the basin (eastern margin of Galicia Bank) (e.g. Fig. 9). Where sequence 5 is faulted, it shows both diverging and parallel reflections, indicative of both prolonged and geologically brief extension at various locations. In the basins of the Grand Banks off Newfoundland, time-stratigraphic equivalents of sequence 5 have facies similar to sequence 6, noted above. In the Jeanne d’Arc Basin the sediments fine upwards, which Grant et al. (1988) interpreted as indicating peneplanation of the source area and, by inference, continued decline of crustal extension that would lead to elevation of the sources.

Significance of Hauterivian extension As noted above, most normal faulting that disrupts sequence 5 appears to have occurred during a relatively brief period of time, late during deposition of the sequence (i.e. in the late(?) Hauterivian and possibly into the Barremian, c. 127 –126 Ma). This was an interval associated with several significant events in the rift (Fig. 4). First, the voluminous magmatism that formed the SENR –JAR and the conjugate MTR began in the southern part of the rift about this time (Tucholke & Ludwig 1982). Perhaps not coincidentally, small volumes of magmatic products as well as other indications of magmatism also appeared farther north in the rift. Gabbro pegmatites were intruded into mantle at ODP Site 1070 (Beard et al. 2002), and within a few million years gabbros and diorites were intruded into the peridotite ridges farther north along the western margin of the Galicia Bank (Fe´raud et al. 1988; Boillot et al. 1989a; Scha¨rer et al. 1995, 2000). In addition, according to magnetic anomaly studies by Russell & Whitmarsh (2003), intrusive magnetic source bodies became more organized and linked at this time, beginning at about chron M5–M4. Results from refraction studies also suggest that ‘normal thickness oceanic crust’ developed by about 125 Ma on the Iberia margin (Fig. 4g, sections I2 and I3) (Whitmarsh et al. 1996; Dean et al. 2000) and on the Newfoundland margin (SCREECH3: Lau et al. 2006b), so this period is commonly interpreted to be the start of sea-floor spreading in the rift (e.g. Whitmarsh & Miles 1995; Russell & Whitmarsh 2003). However, accretion of

exceptionally thin oceanic crust up until about 117 Ma has been interpreted seaward of Flemish Cap, with possible normal thicknesses in younger crust (Fig. 4g, section N3) (Funck et al. 2003; Hopper et al. 2004). In the central Newfoundland Basin Van Avendonk et al. (2006) found no significant change in crustal thickness from the outer limit of continental crust (Fig. 2) seaward to the end of the refraction profile on approximately 114 Ma crust (Fig. 4g, section N2). Although Van Avendonk et al. (2006) interpreted most of this as thin oceanic crust, the recovery of peridotite basement at ODP Site 1277 seaward of anomaly M1 (Fig. 10) suggests that the basement may be dominantly exhumed mantle. A similar conclusion might be drawn for the Iberia side, where peridotite basement was drilled at ODP Site 1070 near anomaly M2 (open circle at youngest peridotite ridge: Fig. 4g, section I2) (Whitmarsh & Wallace 2001). From these observations, it seems that while a component of melt was introduced into the rift beginning at about chron M5 –M3, emplacement of basement at the plate boundary was still being accomplished primarily by exhumation of mantle. An increase in basement roughness is observed on the Newfoundland side of the rift at this time (Tucholke et al. 1989), and a similar change may have occurred on the Iberia side although it presently is not well documented by existing reflection data. It may be that the introduction of increased melt and heat facilitated longer-lived extension on normal faults, thus increasing the amplitude of the basement topography. This also could have temporarily reduced rift-wide tensile stress. However, the subsequent extension of mantle in existing ocean floor at ODP Sites 897, 899 and 1070 and at peridotite ridges R1 and R2 (Fig. 4c) shows that tensile stress was elevated in the Aptian during deposition of seismic sequence 4/A, discussed below.

Barremian– Aptian (sequence 4 and A) Where it has been cored at ODP drill sites off Iberia, seismic sequence 4 is dominated by muddy turbidites and debris flows derived largely from the adjacent continental margin (Sibuet & Ryan 1979; Boillot et al. 1987). These sediments accumulated at rates ranging from ,5 to c. 65 m per million years, with the highest rates in the Aptian at Site 398. Only the uppermost part of correlative sequence A has been cored in the Newfoundland Basin (Site 1276: Tucholke et al. 2004), but it has a similar facies consisting of sandy turbidites and mudstones. These sequences are characterized by several notable features in reflection profiles. Normally they are relatively flat-lying and their internal reflections are parallel to subparallel

BREAKUP OF THE NEWFOUNDLAND– IBERIA RIFT

(Figs 5, 6, 8 & 9). Where reflections diverge, their spacing usually increases toward basin centres rather than towards footwall fault blocks. Faulting and rotation of the sequences (e.g. Fig. 10) are not common but are observed in scattered locations, including the eastern margin of Galicia Bank (Pe´rezGussinye´ et al. 2003). Thus, most basins accumulated sediments in the absence of obvious extension. Where sequence 4 fills valleys between peridotite ridges on the Iberia margin, its reflectivity appears to be related to the basement composition. The rounded shape of peridotite ridge crests, as well as possible slumps and debris slides (Figs 12 & 14), suggests that the serpentinite was weak and mobile. In fact, drilling of peridotite ridges within the rift has consistently recovered serpentinite breccias and olistostromes that reflect mass wasting of the serpentinized basement (Whitmarsh et al. 1998; Tucholke et al. 2004). In many places mass wasting is manifested by exceptionally strong reflectivity of sequence 4 near the peridotite ridges and by weakening reflectivity with distance from the ridges (Figs 6 & 14). These deposits were probably emplaced by serpentine-rich turbidity currents and highly mobile debris flows. The decrease of reflectivity with distance from the ridges is most probably an effect of decreasing grain size. In other cases, mass wasting from both peridotite ridges and continental fault blocks is demonstrated by chaotic fill (Figs 7b & 12) and by instances where sequence 4 laps high onto the bounding fault blocks (Figs 6b & 8c). Where cored at ODP Sites 897 and 899 (see Fig. 12a), these materials are serpentine-rich olistostromes and breccias that were emplaced before faulting and uplift of the underlying peridotite ridges (Comas et al. 1996; Gibson et al. 1996). Thus, some ridge-flank deposits identified in reflection profiles may predate extension (e.g. Fig. 12b, right-hand side), although thickening of the deposits in valleys shows that mass wasting also occurred during or after the faulting. Elsewhere, the ridge-flank deposits clearly were emplaced following extension, and their configuration shows that they were derived largely from the local topography (Figs 6b & 8c). The above patterns are not well defined in sequence A in the Newfoundland Basin, although rounding of presumed serpentinite ridges and possible displacement of allochthonous blocks and debris are observed (Figs 10a & 13). Tucholke et al. (1989) observed rare instances where the Aptian event (U reflection) at the top of the sequence appears to truncate basement topography. A similar possible occurrence is shown on the left of Figure 13b. If these basement highs consisted of exceptionally weak serpentinite, it is conceivable that their exposed tops were eroded away by mass wasting or even by passing turbidity currents.

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On a basin-wide scale, sequences 4 and A are widespread, laterally continuous and thick in locations away from protruding fault blocks that occur at the seaward limits of the sequences and at the edges of Galicia Bank. Sequence 4 is observed all along the northern part of the Iberia margin (and presumably along the southern margin as well, although it is not well mapped there), and sequence A extends some 600 km along the Newfoundland margin. The sequence thicknesses reach more than 1.5 km near the continental margins (Tucholke et al. 2004; Tucholke unpublished data). Thus, at a basin-wide scale the sequences appear to represent rapid deposition of sediments eroded from proximal-margin sources. A particularly noteworthy characteristic of the Aptian event (U reflection) at the top of sequence A in the Newfoundland Basin is its strong reflectivity (Tucholke et al. 1989). In the central and southern Newfoundland Basin the reflection is so strong that deeper structure is masked (Figs 11 & 15). Either the impedance of sequence A is so high that signal penetration is limited, the impedance contrast with basement is small (e.g. the shallow basement is strongly serpentinized), or both. In the northern part of the basin the masking is reduced, although deeper structure often is still difficult to detect (Fig. 10a). A major unconformity was eroded across the southern Grand Banks during the Early Cretaceous in response to deformation and doming associated with the Avalon uplift (e.g. Grant et al. 1988), and it is likely that coarse clastics shed from this region help to account for the strong reflectivity in at least the central to southern Newfoundland Basin. Off Iberia the Aptian event (orange reflection) is locally strong, but it generally has lower amplitude than off Newfoundland and it seldom masks underlying structure except where debris has been shed from peridotite ridges (Figs 6a & 14).

Albian – Cenomanian (sequence 3 and B) Seismic sequence 3 off Iberia and correlative sequence B off Newfoundland are Albian to Cenomanian in age (Fig. 4b). At Sites 398 and 641 off Iberia sequence 3 is largely laminated to massive claystone and mudstone (Sibuet & Ryan 1979). Sequence B off Newfoundland contains muddy to sandy turbidites at it base, changing upwards to mudstones, and then back to sandy turbidites in the Cenomanian at Site 1276 (Tucholke et al. 2004). Sedimentation rates at Sites 398 and 1276 varied from approximately 10 up to 100 m per million years and generally tapered off in the late Albian and Cenomanian, while rates at Site 641, although poorly determined, were at the low end of this range. A general trend towards

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Fig. 12. (a) Section of ISE 10 multichannel reflection profile SW of Galicia Bank, just north of the ODP Leg 149–173 drilling transect. Location in Figure 3 (ISE 10W). Magnetic anomalies M1 and M3 are identified. The basement ridges are serpentinized peridotite where they were drilled to the south. Arrows mark the Aptian event between seismic sequences 3 and 4. Seismic sequence 5 predates emplacement of this basement and is not observed here. Note the very strong reflectivity of sequence 4, which probably has abundant serpentinite debris shed from the adjacent ridges. The reverse fault at the centre was caused by later Cenozoic compression within the plate (e.g. Boillot et al. 1979; Masson et al. 1994) and illustrates the weakness of fault zones in the serpentinized peridotite. (b) Enlarged segment of the profile with interpreted structure and stratigraphic relations. Note the chaotic character of sequence 4, which is interpreted to be mostly mass-wasting debris from the adjacent ridges, and the unconformable onlap of sediments above this sequence. Faults and an interpreted peridotite allochthon (shaded) are speculative, but they are consistent with the geological history of the adjacent ridge as indicated at Site 1070 (i.e. emplacement of the mantle in the rift and gabbro impregnation at about 124 Ma, followed by exhumation some 14 million years later; see text). The interval labelled ‘SP’ has somewhat lower reflectivity than the underlying basement and may be strongly serpentinized, much as has been postulated for poorly reflective upper basement observed on the IAM9 line to the south (Fig. 3) (Pickup et al. 1996).

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finer-grained deposits in sequence 3/B compared to underlying sequence 4/A suggests reduced sediment supply from margin source areas. Seismic sequences 3 and B characteristically lap unconformably onto underlying sequences 4 and A, respectively (Figs 5–8 & 12 –15). This implies either a tectonic event or a sharp change in sedimentation pattern, or both, at the Aptian –Albian boundary.

Breakup of the rift Four puzzling features are observed in the Newfoundland –Iberia rift that require explanation: (1) the continued (or late) extension of peridotite basement at times up to 14 million years or more after it was emplaced at the plate boundary; (2) the lack of significant coeval extension manifested in the structural and stratigraphic relations of seismic sequence 4 off Iberia and sequence A off Newfoundland; (3) the sharp change in reflectivity marked by the Aptian event at the top of sequence 4/A; and (4) the seismic unconformity at this level. We propose that these can be explained by continued exhumation and rifting of subcontinental mantle lithosphere through Aptian time, followed by initiation of normal sea-floor spreading at the Aptian –Albian boundary, as follows. Stage A, late Aptian (Fig. 16a). Up to this time, sea floor was being emplaced primarily by exhumation of subcontinental mantle, and in-plane tensile stresses were elevated throughout the rift. The plates extended internally in local weak zones, probably where existing faults were highly serpentinized. Mass wasting and slumps contributed coarse, reflective, sometimes chaotic debris to sequence 4/A, particularly in the vicinity of peridotite ridges. These deposits were interbedded with turbidites derived from the proximal continental margins. Stage B, Aptian– Albian boundary (Fig. 16b). Normal sea-floor spreading began in latest Aptian –earliest Albian time. This introduction of heat and melt at the plate boundary caused in-plane tensile stress to drop sharply, resulting in a pulse of relative compression in the adjoining plates. The compression may have amplified topography in the deep basins and thus briefly enhanced mass wasting from basement highs. Although the common concave-upwards shape of the Aptianevent reflection between basement highs primarily reflects input of debris from adjacent ridges (e.g. Figs 6, 8 & 12), its curvature may have been augmented by the compression. The change in stress regime probably also caused differential vertical motions on the proximal continental margins, stimulating erosion and sediment transport to the

Fig. 13. (a) Section of SCREECH2 multichannel seismic reflection profile (Shillington et al. 2004) seaward of anomaly M0 in the northern Newfoundland Basin. Location in Figure 2 (SCREECH 2SE). The U reflection (Aptian event) marks the boundary between seismic sequences A and B, equivalent to sequences 4 and 3 on the Iberia margin (see Fig. 4b). Note the unusual, rounded form of the basement highs which suggests that they may be serpentinized peridotite that was subject to degradation by mass wasting. (b) Interpretation of part of the profile. Note the unconformable onlap of sediments onto the Aptian event (U reflection). The faults and possible allochthons (grey shade) are speculative, but are consistent with the idea that the basement is weak, serpentinized peridotite that failed easily; note the possible reverse fault in sequence A on the right. The low-reflectivity intervals marked ‘SP’ may be more completely serpentinized, as has been suggested for the IAM9 line off Iberia (Pickup et al. 1996).

34 B. E. TUCHOLKE ET AL.

Fig. 14. (a) Section of ISE 14 multichannel reflection profile across the SW margin of Galicia Bank. Location in Figure 3 (ISE 14W). The basement ridges at the right are probably continental fault blocks and the ridge on the left is peridotite ridge R3. Arrows mark the Aptian event between seismic sequences 3 and 4. See Figure 4b for sequence ages. Note the exceptionally strong reflectivity of sequence 4 and the Aptian event near the peridotite ridge, which we interpret to be caused by mass wasting of serpentinite debris from the ridge. (b) Enlarged segment of the profile with interpretation. Note the flat-lying nature of sequence 4, the chaotic bedding at the top of the sequence that indicates mass wasting and the unconformable onlap of sediments above the Aptian event.

BREAKUP OF THE NEWFOUNDLAND– IBERIA RIFT 35

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Fig. 15. Section of ERABLE multichannel reflection profile from the centre of the Newfoundland Basin (Srivastava & Sibuet 1992). Location in Figure 2. See Figure 4b for sequence ages and lithology. The strong reflectivity of the Aptian event (U reflection) is a feature observed throughout most of the Newfoundland Basin (Tucholke et al. 1989). Deeper sedimentary and basement structure is poorly imaged because of the strong reflectivity of U, because of low impedance contrasts, or both. White arrows below U show apparent sedimentary reflections and grey arrows mark dipping, possibly sedimentary reflections that may indicate rotation of fault blocks. Note the unconformable onlap of sediments onto the U reflection.

deep basins. The Aptian event is represented by these deposits combined with debris eroded from basement highs. Stage C, Albian (Fig. 16c). Following the Aptian event, relaxation of stresses reduced longwavelength flexural topography, and sediment flux to the basins was attenuated. Slopes on basement highs also equilibrated, and there no longer was significant mass wasting from these sources. Finegrained turbidity currents spread widely across the basins and lapped unconformably onto the curved Aptian-event reflection in troughs between basement ridges. Deposition was largely conformable on the broader basin plains away from these ridges.

Discussion The relatively prolonged phase of extension that persisted from the Late Jurassic through to the Early Cretaceous appears to have been punctuated by three rifting episodes: one that initiated in the Tithonian and culminated near the Berriasian– Valanginian boundary; a second that peaked near the Hauterivian–Barremian boundary; and a third that is manifested largely in the Aptian. It is noteworthy that the timing of the first two episodes correlates relatively well with the times at which continental crust appears to have separated in the southern and northern parts of the rift. South of Galicia Bank and Flemish Cap (Fig. 4g, sections N1 and I2), extrapolated ages of the edges of

Fig. 16. Schematic representation (not to scale) of interpreted events through the time of the Aptian event. See text for discussion. Grey is continental crust, white is subcontinental mantle lithosphere and black is normal oceanic crust. (a) Latest Aptian emplacement of basement at the plate boundary, primarily by exhumation of subcontinental mantle lithosphere. Elevated in-plane tensile stress affects the entire rift system, and crests of fault blocks are locally uplifted by normal faulting. (b) At the end of Aptian time, introduction of heat and melt from the rising asthenosphere (As) relaxes in-plane tensile stress and results in relative compression of the adjacent plates, causing differential vertical motion (broad arrows) as discussed in the text. Mass wasting from basement highs together with turbidites and debris flows from the uplifted lower continental slope and continental rise creates a highly reflective stratigraphic marker (Aptian event). Relative compression may also cause reverse faulting, locally amplify short-wavelength topography

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(a) LATE APTIAN: MANTLE EXHUMATION AND INTRA-PLATE EXTENSION Sediment input from shelf

Sea level Intra-plate normal faulting and uplift, mass wasting of serpentinites

Br

Rift-wide in-plane tensile stress

le i t t ile ct Du

(b) APTIAN-ALBIAN BOUNDARY: PLATE-BOUNDARY RELEASE, ONSET OF NORMAL SEA-FLOOR SPREADING Probably reduced sediment input from shelf

Sea level Possible topographic amplification and warping of inter-block strata, local reverse faulting, continued mass wasting Aptian event

Sediment input from uplifted lower slope and rise

Plate boundary release Pulse of relative compression

Melt As

(c) ALBIAN: SEA-FLOOR SPREADING Limited sediment input from shelf

Sea level Normal turbidite and pelagic sedimentation Aptian event

Normal sea-floor spreading Melt As

Fig. 16. (Continued) and warp the Aptian-event reflection. (c) Normal sea-floor spreading follows in Albian time. Sea-floor distortions caused by relative compression at the end of the Aptian are levelled out, major turbidity currents and mass wasting cease to be important, and hemipelagic sediments and flat-lying, fine-grained turbidites from distal sources cover the Aptian-event reflection. Rising eustatic sea level probably limited sediment input from the continental shelf.

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continental crust fall near the Jurassic–Cretaceous boundary and thus within the envelope of the Tithonian–Berriasian rift episode. To the north, the edges of continental crust all are close to the Hauterivian–Barremian boundary (Fig. 4g, sections N2, N3 and I3), correlating with the peak of the second rift episode. If this correlation is valid, it suggests that the rift opened stepwise, south to north, during two principal extensional episodes rather than by continuous northward rift propagation. It is interesting to consider the reasons that these rift episodes may have been relatively short-lived. It has been proposed that once continental crust is thinned to approximately 6–10 km thick, fluids can penetrate to the mantle and facilitate serpentinization, thus weakening the lithosphere and localizing extension on large-scale detachments (Pe´rezGussinye´ & Reston 2001). If this process engenders further fluid penetration, serpentinization and fault weakening, extensional collapse make take place quite rapidly once the initial conditions for fluid access to the mantle are met. We suggest that this process may account for the rift episodes that mark the boundaries between seismic sequences 6, 5 and 4. The second rift episode correlates in time with the beginning of magmatism that formed the SENR–JAR and the conjugate MTR at the southern edge of the rift. This event began near the end of Hauterivian time and extended into the early Aptian (Fig. 4c). The high-amplitude J anomaly in the southern part of the rift and its northward decay (Figs 2 & 3) suggest that magma was introduced northward into the rift to near the latitude of southern Flemish Cap and Galicia Bank. This might explain the presence of the young M-series magnetic anomalies as well as apparently normal oceanic crust at the seaward ends of refraction lines IAM9 off Iberia (Fig. 4g) and SCREECH3 off Newfoundland; it might also account for thin oceanic crust overlying serpentinized mantle between Flemish Cap and Galicia Bank (Fig. 4g, sections N3 and I3). We have suggested, however, that elevated in-plane tensile stress continued throughout the rift (rift episode 3) until normal seafloor spreading began near the Aptian– Albian boundary. If this is correct, then it is likely that magma entering the rift was confined to the shallow lithosphere and that the deeper lithosphere retained its strength. Alternatively, the magmatic event may have temporarily relieved in-plane tensile stress during the Barremian, with high stress levels and mantle exhumation resuming in Aptian time. The clustering of observed extensional events in the later part of the Aptian (Fig. 4c) suggests that this may be the case. Refraction profiles that extend seaward across Aptian and Albian basement are needed in order to investigate these alternatives.

The fact that it is rare to see faulting and rotation in sequence 4/A suggests that Barremian–Aptian extension within the plates was limited and/or localized. However, Wilson et al. (2001) suggested several alternate explanations for why rotated and splayed reflections might not develop above fault blocks during extension; these include: (a) sediments too thin to be resolved seismically; (b) deposition on level, non-rotating fault surfaces; and (c) chaotic re-sedimentation that destroys layering. The first explanation is unlikely because it would require a drastic reduction in sediment supply during deposition of sequence 4/A, for which there is no evidence. Sedimentation rates at ODP Sites 638 and 641 (Fig. 7) were fairly constant or decreased from Valanginian through to Aptian time (Boillot et al. 1987), but rates at DSDP Site 398 (Fig. 9) increased from the Barremian into the Aptian (Sibuet & Ryan 1979). More importantly, sequence 4 clearly has thicknesses comparable to sequence 5 (e.g. Figs 5– 8). Considering the time spans of these sequences (Fig. 4), average sedimentation rates decreased in sequence 4, but only by about 40%. Deposition on a relatively level fault surface, such as might be produced by a rolling-hinge detachment, is a possibility for certain areas such as the part of Hobby High drilled at ODP Sites 900, 1067 and 1068 at the southern margin of Galicia Bank (Manatschal et al. 2001). However, this explanation is not applicable over most of the rest of Galicia Bank where faults have significant dip (e.g. Figs 5, 7 & 8). Chaotic redeposition appears to be a realistic possibility. The clearest evidence for extension during the Barremian–Aptian is confined to areas of exhumed mantle (Fig. 4c). Faulting and uplift of structurally weak serpentinite blocks could produce massive slumps, slides and debris flows that would obscure most structural information about the original faults. This kind of remobilization is suggested by the complex intrabasement structure in many parts of the exposed mantle (Figs 12 & 13), and it is also attested to, perhaps at a smaller scale, by the occurrence of serpentinite olistostromes and breccias on the peridotite ridges drilled at ODP Sites 397 and 399 (Comas et al. 1996; Gibson et al. 1996). As slopes stabilized following faulting and major mass failure at a given location, deposition would shift to more coherent patterns of relatively flat-lying, semi-layered to layered bedding (Figs 6 & 12 –14). This explanation, however, does not apply over most areas of continental crust, where sequence 4 overlies sequence 5 with little or no evidence of major mass wasting, faulting or rotation (Figs 5–9). Although some faults in continental crust probably soled out in serpentinized mantle, the continental

BREAKUP OF THE NEWFOUNDLAND– IBERIA RIFT

faults overall were probably stronger than those in the exhumed mantle. Thus, we suggest that during Barremian–Aptian time there was very limited extension in continental crust and that most intraplate extension occurred within areas of exhumed, serpentinized mantle (Fig. 16a). Why the proposed large amounts of heat and melt were introduced at the plate boundary to initiate normal sea-floor spreading at the end of Aptian time is uncertain, but there are several possibilities. Finite-element models of continental extension by Bowling & Harry (2001) suggest that asthenospheric upwelling and production of melt can accelerate rapidly very late in the history of rifting. They also note that the period of amagmatic rifting leading to this event can be protracted if the extension rate is low and if the lithosphere is homogeneous, or if the lithosphere is heterogeneous and zones of crustal and mantle weakness are offset from one another. In the Newfoundland –Iberia rift, both the heterogeneity of the prerift Palaeozoic crust (Fig. 1) and the observed spatial and temporal shifts in the locus of rifting suggest that the latter may be the case. Reston & Phipps Morgan (2004) proposed that the onset of sea-floor spreading might be triggered by invasion of melt and heat from an adjacent rift segment or mantle plume. In the present instance, this seems unlikely because no increased magmatism is known to have been present in adjacent spreading or rift segments at the end of Aptian time. In addition, the large magmatic event that created the SENR, JAR and MTR in Barremian–early Aptian time seems not to have resulted in normal sea-floor spreading, at least on a permanent basis. A further possibility is that the mantle lithosphere was chemically or thermally heterogeneous, and that sea-floor spreading began with the rise and decompression melting of melt-prone mantle. Such heterogeneity is well documented along the MidAtlantic Ridge (e.g. Bonatti et al. 1992) and the ultra-slow spreading Gakkel Ridge in the Arctic (Michael et al. 2003). However, the observed scale of heterogeneity is generally smaller (,c. 100– c. 200 km) than the length of the Newfoundland– Iberia rift (c. 800 km). Considering these observations, it seems likely that the proposed late-stage onset of normal seafloor spreading probably is an inherent feature of the rift evolution of non-volcanic margins (e.g. Bowling & Harry 2001). We suggest that the arrival of large volumes of melt at the rift axis, and thus the commencement of normal sea-floor spreading, reflects shoaling and decompression melting of the asthenosphere. Rigid, largely subcontinental mantle lithosphere would have been exhumed up until this time (with the possible exception of the aforementioned period of Barremian

39

magmatism), and this would have caused elevated in-plane tensile stress throughout the rift. Sharp reduction or release of the tensile stress (and thus relative compression) associated with introduction of heat and melt from the rising asthenosphere would be expected to produce a well defined event in the geological record. The character of the Aptian event is consistent with this postulated stress change. Relative compression may have amplified local basement topography and stimulated mass wasting, as suggested by chaotic deposits at the top of sequence 4 off Iberia (Figs 8c & 14b). It may also have enhanced the curvature of the Aptian-event reflection and even caused reverse faulting (Fig. 13b). Unfortunately, any reverse faulting that may have developed on the Iberia margin has been obscured by the effects of Cenozoic compression (e.g. Fig. 12) (Boillot et al. 1979; Masson et al. 1994). It is likely that the stress change was accompanied by long-wavelength differential vertical motions along the proximal margins (Fig. 16b). Numerical models indicate that release of in-plane tensile stress would cause uplift of the proximal rift basin (lower continental slope and continental rise), subsidence of the distal rift basin and the outer shelf, and uplift of the inner shelf and coastal plain (e.g. Braun & Beaumont 1989; Kooi & Cloetingh 1992). Details of positions and amplitudes of such vertical motions vary depending on flexural state and rheology of the lithosphere, but in general the models show that the lower slope and rise can be uplifted by several hundred metres. Thus, although most sediments eroded from the hinterland might not pass beyond the depressed outer shelf, the deeper proximal margin could provide a significant source of sediment to more distal basins. Eustacy may also have affected sediment input into the rift. Eustatic sea level was low during the Aptian (Haq et al. 1988), and it could have facilitated shallow-water erosion and sediment transport to the deep basins. By itself, such a eustatic effect is not a likely explanation for the Aptian-event reflection because numerous other eustatic lowstands had no comparable effect on the sedimentary record in the deep basins. Moreover, a eustatic effect would not explain the curvature of the reflection between basement ridges or the unconformable onlap of overlying sediments. The numerical models (e.g. Kooi & Cloetingh 1992) suggest that the long-wavelength vertical distortion of the lithosphere caused by relative compression would have decayed rapidly (Fig. 16c). With the introduction of a new and relatively stable stress regime, regional flux of coarse sediments in debris flows and turbidity currents would have been sharply curtailed. Slope stabilization

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probably also resulted in a rapid reduction of mass wasting in local basins between basement ridges on the distal margins. Thus, finer-grained sediments were carried in turbidity currents following the Aptian event, and these dilute flows distributed sediments more widely and uniformly. This is consistent with available data from drill sites, which generally show reduced average grain size of sediments deposited during the Albian. These changes account for the sharp drop in reflectivity above sequence 4/A, and for onlap of beds onto the Aptian-event reflection (Figs 5 –8 & 12 –15).

Conclusions The Newfoundland–Iberia rift developed in structurally and lithologically heterogeneous continental crust that had previously been shortened and thickened in Palaeozoic suture zones during the closure of the Iapetus and Rheic oceans. These structures influenced, but did not dominate, the geometry of Mesozoic rifting. Initial rifting in the latest Triassic– earliest Jurassic created large, deep rift basins across a wide region, predominantly in the Grand Banks of Newfoundland and to a lesser extent on the Iberia margin, but this rifting failed to thin the continental crust substantially. A second, prolonged rift phase in the Late Jurassic – Early Cretaceous first thinned and separated the continental crust, and then exhumed subcontinental lithospheric mantle in wide transition zones on the conjugate margins. Emplacement of basement primarily by mantle exhumation appears to have continued until normal sea-floor spreading commenced near the Aptian –Albian boundary. The rift experienced elevated in-plane tensile stress throughout the Late Jurassic –Early Cretaceous rift phase, but most extension of continental crust was concentrated in two episodes: one culminating near the Berriasian –Valanginian boundary and one focused near the Hauterivian –Barremian boundary. In most places these rifting episodes occurred over geologically short intervals, as indicated by faulted sedimentary sequences that are rotated but commonly show only weakly divergent reflections towards fault footwalls. We suggest that this is a manifestation of crustal thinning to the point where fluids could penetrate to and serpentinize the mantle (Pe´rez-Gussinye´ & Reston 2001), which led to rapid collapse of the crustal blocks. The first rift episode primarily affected the Lusitanian and Galicia Interior basins, and it was probably associated with separation of continental crust and initial exhumation of lithospheric mantle in the transition zone south of the Flemish Cap – Galicia Bank conjugate margins. The second rift episode focused in continental crust between Flemish Cap and Galicia Bank, but

to a small extent it also affected the eastern Galicia margin. It also culminated in exhumation of lithospheric mantle and appears to have led to emplacement of thin and/or spatially variable occurrences of magmatic oceanic crust between Flemish Cap and Galicia Bank. Magnetic anomaly analyses and refraction studies along and south of the southern Flemish Cap–Galicia Bank margins suggest that some form of magmatic oceanic crust may have formed there during this rift episode, beginning by about the time of chron M3. However, drilling in this crust near anomaly M1 on both margins has yielded basement consisting of serpentinized peridotite with only minor occurrences of igneous rocks. This suggests that the interpreted oceanic crust seaward of anomaly M3 is not normal oceanic crust but may be largely exhumed mantle that is variably intruded by igneous rocks. Such basement could be difficult to distinguish from normal magmatic oceanic crust in magnetic and refraction data. A possible source of magma during this period is a mantle plume that constructed the SENR, JAR and MTR at the southern edge of the Newfoundland – Iberia rift from about chron M4 (Barremian) to a time somewhat younger than chron M0 (earliest Aptian). The high but northward-decaying amplitude of the J anomaly (c. M1 –M0) suggests that the influence of this plume temporarily extended northwards to near the southern margins of Flemish Cap and Galicia Bank. The ages of the two rift episodes appear to correlate with ages at the edges of continental crust in the southern and northern parts of the rift. This suggests that the rift opened in a stepwise, twophase process rather than by continuous northward propagation of rifting. Existing peridotite basement off Iberia (ODP Sites 897, 899 and 1070, and peridotite ridges R1 and R2) experienced strong normal faulting during a third rift episode in late Aptian time, and possibly throughout both the Barremian and Aptian. We suggest that this was caused by elevated, rift-wide in-plane tensile stress that persisted while subcontinental mantle lithosphere was being exhumed at the plate boundary. The normal faulting probably was focused at the weakest zones in the lithosphere and thus was concentrated in serpentinized, exhumed mantle rather than in adjacent continental crust. Although the high level of in-plane tensile stress may have been temporarily interrupted by the introduction of magma along the plate boundary during Barremian time, we do not consider this to mark the initiation of normal sea-floor spreading. We propose that sea-floor spreading began in the latest Aptian–earliest Albian. At this time the subcontinental mantle lithosphere was finally breached by the rising asthenosphere, and large amounts of

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heat and melt were introduced at the rift axis. This resulted in relaxation of the in-plane tensile stress and thus relative compression throughout the rift. This event is recorded by a prominent, rift-wide sedimentary reflection (Aptian event) that is unconformably onlapped by post-rift sediments. We suggest that the above stress history explains the following characteristics of the Barremian –Aptian sedimentary sequence and the Aptian event.

produced widespread changes in stress state of the rift and is marked by prominent structural and stratigraphic features in the geological record. Thus, interpretation of continental breakup in a nonvolcanic rift must consider the rheology not only of the crust but also that of the subcontinental mantle lithosphere, in addition to spatial and temporal changes in extension that may occur from segment to segment along the rift.

† High reflectivity: flanks of basement ridges (particularly those consisting of serpentinized peridotite) were comparatively unstable up to the end of the Aptian, first because of footwall uplift associated with normal faulting, and then because of the pulse of relative compression associated with stress release at the plate boundary. This resulted in mass wasting, deposition of coarse, thickly bedded debris in basins, and high reflectivity of the Barremian–Aptian sequence. At the time of stress release, there may also have been enhanced input of coarse sediments from the proximal margins because of differential vertical motions there and because of a coincidental eustatic lowstand. † Change to low reflectivity above the Aptian event: sea-floor slopes stabilized in the new, lower-stress environment, thus curtailing mass wasting and turbidity currents. This probably reduced both average grain size and occurrence of thick graded beds, thus leading to lower reflectivity. † Concave-upward configuration of Barremian– Aptian sediments in distal, inter-ridge basins: this was created primarily by mass wasting from adjacent ridge sources, but it may have been enhanced by warping in response to relative plate compression at the end of Aptian time. † Unconformable onlap onto the Aptian-event reflection: following slope stabilization, turbidity currents that entered the basins were finergrained and thus spread more uniformly across basins, depositing flat-lying beds that lapped unconformably onto the Aptian-event reflection.

This research was supported by National Science Foundation grants OCE98-19053 and OCE03-26714 to Woods Hole Oceanographic Institution, grants OCE95-21517 and OCE99-11725 to Rice University, and the Ocean Drilling Program. The Ocean Drilling Program is sponsored by the US National Science Foundation and participating countries under management of Joint Oceanographic Institutions (JOI), Inc. B. Tucholke also acknowledges support from JOI/USSSP grant F001846 and the Henry Bryant Bigelow Chair in Oceanography at Woods Hole Oceanographic Institution. Our research has benefited from discussions with R. B. Whitmarsh and G. Manatschal, and with SCREECH program colleagues W. S. Holbrook, J. R. Hopper, H. Lau, K. E. Louden, D. J. Shillington and H. van Avendonk. We thank S. P. Srivastava for permission to use a segment of the ERABLE multichannel seismic data in our figures. We also thank K. E. Louden, R. C. L. Wilson, D. G. Roberts, G. Karner and L. Pinheiro for careful reviews that helped to improve the manuscript; in some cases their views differ from ours, and the interpretations made here remain our responsibility. This paper was submitted 10 May 2005 and revised 20 April 2006. Contribution No. 11,371 of Woods Hole Oceanographic Institution.

According to our interpretations, the historically identified ‘breakup unconformity’ (here, the Aptian event) in the Newfoundland–Iberia rift does not mark the separation of continental crust and simultaneous initiation of sea-floor spreading. Instead, this event records the final separation of subcontinental mantle lithosphere, the rise of the asthenosphere to shallow levels and emplacement of the first normal oceanic crust. This occurred some 14 million years after the late Hauterivian– early Barremian rupture of continental crust in the northern part of the rift, and even longer after the Tithonian– Berriasian separation of continental crust in the southern part of the rift. Each of these three events

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the Ocean Drilling Program, Scientific Results, 149. Ocean Drilling Program, College Station, TX, 713– 733. W HITMARSH , R. B. & W ALLACE , P. J. 2001. The rift-to-drift development of the west Iberia nonvolcanic continental margin: A summary and review of the contribution of Ocean Drilling Program Leg 173. In: B ESLIER , M.-O., W HITMARSH , R. B., W ALLACE , P. J. & G IRARDEAU , J. (eds) Proceedings of the Ocean Drilling Program, Scientific Results, 173. Ocean Drilling Program, College Station, TX, 1–36. W HITMARSH , R. B., B ESLIER , M.-O. ET AL . 1998. Proceedings of the Ocean Drilling Program, Initial Reports, 173. Ocean Drilling Program, College Station, TX. W HITMARSH , R. B., M ANATSCHAL , G. & M INSHULL , T. A. 2001. Evolution of magma-poor continental margins from rifting to sea-floor spreading. Nature, 413, 150–154. W HITMARSH , R. B., M ILES , P. R. & M AUFFRET , A. 1990. The ocean–continent boundary off the western continental margin of Iberia – I. Crustal structure at 408300 N. Geophysical Journal International, 103, 509– 531. W HITMARSH , R. B., W HITE , R. S., H ORSEFIELD , S. J., S IBUET , J.-C., R ECQ , M. & L OUVEL , V. 1996. The ocean– continent boundary off the western continental margin of Iberia: Crustal structure west of Galicia Bank. Journal of Geophysical Research, 101, 28,291– 28,314. W ILLIAMS , H. & H ATCHER , R. D. J R . 1982. Suspect terranes and accretionary history of the Appalachian Orogen. Geology, 10, 530– 536. W ILLIAMS , H. & H ATCHER , R. D. J R . 1983. Appalachian suspect terranes. In: H ATCHER , R. D. JR , W ILLIAMS , H. & Z IETZ , I. (eds) Contributions to the Tectonics and Geophysics of Mountain Chains. Geological

Society of America Memoir, 159. Geological Society of America, Boulder, CO, 33– 53. W ILSON , R. C. L. 1988. Mesozoic development of the Lusitanian Basin, Portugal. Revista de la Sociedad Ge´olo´gica de Espana, 1, 393–407. W ILSON , R. C. L., H ISCOTT , R. N., W ILLIS , M. G. & G RADSTEIN , F. M. 1989. The Lusitanian Basin of west-central Portugal: Mesozoic and Tertiary tectonic, stratigraphic, and subsidence history. In: T ANKARD , A. J. & B ALKWILL , H. R. (eds) Extensional Tectonics and Stratigraphy of the North Atlantic Margins. AAPG Memoirs, 46, 265– 282. W ILSON , R. C. L., M ANATSCHAL , G. & W ISE , S. 2001. Rifting along non-volcanic passive margins: Stratigraphic and seismic evidence from the Mesozoic successions of the Alps and western Iberia. In: W ILSON , R. C. L., W HITMARSH , R. B., T AYLOR , B. & F ROITZHEIM , N. (eds) Non-volcanic Rifting of Continental Margins: A Comparison of Evidence From Land and Sea. Geological Society, London, Special Publications, 187, 429– 452. W INTERER , E. L., G EE , J. S. & W AASBERGEN , R. J. V. 1988. The source area for Lower Cretaceous clastic sediments of the Galicia margin: Geology and tectonic and erosional history. In: B OILLOT , G., W INTERER , E. L. ET AL . Proceedings of the Ocean Drilling Program, Scientific Results, 103. Ocean Drilling Program, College Station, TX, 697–732. Z HAO , X. 1996. Magnetic signatures of peridotite rocks from Sites 897 and 899 and their implications. In: W HITMARSH , R. B., S AWYER , D. S., K LAUS , A. & M ASSON , D. G. (eds) Proceedings of the Ocean Drilling Program, Scientific Results, 149. Ocean Drilling Program, College Station, TX, 431–446. Z IEGLER , P. A. 1982. Geological Atlas of Western and Central Europe. Elsevier, Amsterdam.

Structure of the Flemish Cap margin, Newfoundland: insights into mantle and crustal processes during continental breakup J. R. HOPPER1, T. FUNCK2 & B. E. TUCHOLKE3 1

Department of Geology and Geophysics, Texas A&M University, College Station, TX 77843, USA (e-mail: [email protected]) 2

Geological Survey of Denmark and Greenland, Øster Voldgade 10, DK-1350 Copenhagen K, Denmark 3

Department of Geology and Geophysics, Woods Hole Oceanographic Institution, Woods Hole, MA 02543, USA

Abstract: Seismic reflection and refraction data from the Flemish Cap margin off Newfoundland reveal the large-scale structure of a magma-starved rifted margin. There is little evidence for significant extensional deformation of the Flemish Cap, consistent with the hypothesis that it behaved as a microplate throughout the Mesozoic. The seismic data highlight important asymmetries at a variety of scales that developed during the final stages of continental breakup and the onset of oceanic sea-floor spreading. In strong contrast to the conjugate Galicia Bank margin, Flemish Cap shows: (1) an abrupt necking profile in continental crust, thinning from 30 km thick to 3 km thick over a distance of 80 km, and a narrow, less than 20 km-wide, zone of extremely thin continental crust; (2) no clear evidence for horizontal detachment structures beneath continental crust similar to the ‘S’ reflection; and (3) evidence for at least a 60 km-wide zone of anomalously thin oceanic crust that began accreting to the margin shortly after continental crustal separation. The oceanic crust averages only 3 –4 km thick and in places is as thin as 1.3 km thick, although seismic layer 3 is missing where this occurs. The data suggest that there are large spatial and temporal variations in the available melt supply following continental breakup as oceanic sea-floor spreading becomes established. In addition, wide-angle data show that anomalously slow mantle P-wave velocities appear approximately where continental crust has thinned to 6– 8 km thick, indicating that low-degree serpentinization begins where the entire crust has become embrittled.

Continental extension leading to lithospheric rupture and the creation of new oceanic sea-floor spreading systems has long been a cornerstone of global plate tectonics. Because rifted margins contain the geological record of key processes involved with extension and breakup, investigations of these margins are central to advancing our understanding of lithospheric rupture. Thinning of the crust and lithospheric mantle is the most obvious manifestation of continental extension. Two key consequences of thinning are brittle failure of the upper crust and decompression melting of asthenospheric mantle as it nears the surface. Despite the seeming simplicity of these two effects, considerable controversy over basic aspects of how they work has resulted from enigmatic observations from a number of areas. For the brittle behaviour of the crust, the possible role of low-angle normal faulting and subhorizontal detachments during extension is still debated (e.g. Lavier et al. 1999; Lavier & Manatschal 2006). In the case of mantle melting, there is no firm consensus on the primary causes of the extreme range of magmatic productivity during the initial stages of

sea-floor spreading (e.g. Boutilier & Keen 1999; Nielsen & Hopper 2004). The rifted passive margins around the North Atlantic are a unique natural laboratory where many of these questions can be studied. Magnetic anomalies permit the reconstruction of conjugate margin pairs with a reasonable degree of certainty (Srivastava & Tapscott 1986). This is critical as many rift systems exhibit asymmetry that can only be quantified by characterizing the entire rift system. In addition, the North Atlantic margins are themselves extremely diverse. Figure 1 shows the regional setting with measured oceanic crustal thicknesses marked along several of the margins. At one end are the margins along Greenland, Norway and the British Isles where 15 –30 kmthick oceanic crust accreted to the rifted continental margin (Holbrook et al. 2001). At the other extreme are the margins of SW Greenland and Iberia, where 0 –5 km-thick oceanic crust is observed (Chian & Louden 1994; Minshull et al. 2001). The lack of volcanism along these latter margins has proven particularly useful for studying

From: KARNER , G. D., MANATSCHAL , G. & PINHEIRO , L. M. (eds) Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 47–61. DOI: 10.1144/SP282.3 0305-8719/07/$15.00 # The Geological Society of London 2007.

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Fig. 1. Regional overview of the North Atlantic Ocean and rifted margins. The map is shaded relief of topography and bathymetry with shallow-water areas highlighted in darker shades. Bold numbers indicate the thickness of igneous crust produced immediately after continental crustal separation (values taken from Fowler et al. 1989; Chian & Louden 1994; Chian et al. 1995; Smallwood et al. 1999; Holbrook et al. 2001; Whitmarsh et al. 2001; Funck et al. 2003, 2004; Hopper et al. 2004). CGFZ, Charlie Gibbs Fracture Zone; FC, Flemish Cap; GIB, Galicia Interior Basin; GM, western Galicia margin; IAP, Iberia Abyssal Plain; MAR, Mid-Atlantic Ridge; NF, Newfoundland; RP, Rockall Plateau; UK, United Kingdom.

mechanical deformation of the upper crust because the extensional structures are accessible to seismic imaging techniques. On the Galicia Bank margin off Iberia, the strong regional ‘S’ reflection is observed that is now believed to be a detachment surface at the crust-mantle boundary (Reston et al. 1996). The detachment is thus a major rheological boundary that probably formed as a heavily serpentinized ductile shear zone during the final stages of extension (e.g. Reston et al. 1996; Pe´rez-Gussinye´ & Reston 2001). The uniform sense of shear and rotation indicated by fault blocks above ‘S’ has led to various hypotheses about possible asymmetry during extension and breakup. Farther south along Iberia, a broad, up to 100 km-wide, zone of exhumed serpentinized mantle was exposed at the sea floor (e.g. Whitmarsh

et al. 2001), reminiscent of Hess’ (1962) original hypothesis for the formation of oceanic basins. In this so-called transition zone, the mantle is mechanically unroofed by brittle extension and interacts with sea water to form a thin layer of serpentinite with crust-like seismic velocities (Dean et al. 2000). Initially, this exhumation is thought to be largely amagmatic, but as extension progresses more melt is eventually produced. In the Alps, where these melt products can be sampled and analysed, there is a progression from transitional mid-ocean ridge basalt (T-MORB) towards the continent to normal mid-ocean ridge basalt (N-MORB) towards the ocean (Desmurs et al. 2002; Schaltegger et al. 2002). Desmurs et al. (2002) suggest that this indicates progressive thinning of continental lithospheric mantle, which is eventually

STRUCTURE OF THE FLEMISH CAP MARGIN

replaced by asthenospheric ‘oceanic’ mantle as a magmatic sea-floor spreading system becomes established. It is becoming increasingly clear that resolving the nature and origin of ‘transition zone’ crust is central to understanding final lithospheric rupture along magma starved margins. In this contribution, results from a recent seismic reflection and refraction survey conducted off the coast of Newfoundland are summarized. The data shown are published in Funck et al. (2003) and Hopper et al. (2004, 2006). We do not attempt to replicate all of those results here, but focus on key summary figures and portions of the data set that are particularly relevant for the numerical modelling exercises described in this volume.

Regional tectonic setting and survey description The extensional history of the region between Iberia and Newfoundland spans nearly 100 million years, beginning with an early phase of rifting in the Late Triassic –Early Jurassic that formed the major basins within the Grand Banks and off Portugal (e.g. Tankard & Welsink 1987; Wilson 1988;

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Tucholke et al. 1989; Rasmussen et al. 1998; Wilson et al. 2001). A 40– 50 million year period of quiescence was followed by a second major phase of extension that ultimately led to complete lithospheric rupture and the establishment of magmatic sea-floor spreading in the North Atlantic some time between Barremian and Aptian time (127 –112 Ma using the timescale of Gradstein et al. 1994). Details regarding the full history and development of the region are summarized in Tucholke et al. (2007) and Tucholke & Whitmarsh (2007). The SCREECH survey (Studies of Continental Rifting and Extension off the Eastern Canadian sHelf) was conducted in 2000 to better understand the large-scale structure of the Newfoundland margin at positions conjugate to the major geophysical surveys and drilling transects off the Iberian margin (Funck et al. 2003; Hopper et al. 2004, 2006; Lau et al. 2006a, b; Shillington et al. 2004, 2006; Van Avendonk et al. 2006). Multichannel seismic reflection and wide-angle seismic refraction data were acquired along three primary transects. Each transect began on the shelf where continental crustal thinning should be minimal and continued seaward well past magnetic anomaly M0 identified over oceanic crust (Figs 2 & 3). Crustal thickness,

Fig. 2. Expanded view of the region offshore Newfoundland showing the SCREECH survey. Open circles are OBS locations along the three major transects, labelled Lines 1, 2 & 3. The thick black line is magnetic anomaly M0 as picked by Srivastava et al. (2000). FC, Flemish Cap. Bathymetric contours are spaced every 500 m.

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Fig. 3. Regional magnetic anomaly map based on the data of Verhoef et al. (1992). Seismic transects and the anomaly M0 pick from Figure 2 are also shown. Bathymetric contours are at 1000 m intervals, and the bathymetry is shaded with illumination from the NW.

seismic P-wave velocity and reflection imaging provide critical constraints on the large-scale structure of the full rift system between North America and Europe. Here, the primary focus is on the northernmost transect across the Flemish Cap (Line 1 in Fig. 2), a block of Appalachian continental crust that is conjugate to the northernmost Iberian peninsula where the Galicia Interior Basin (GIB) and Galicia Bank margins are located. The profile is also conjugate to the Ocean Drilling Program (ODP) Leg 103 drilling transect as well as a number of seismic reflection and refraction profiles described in Reston et al. (1996), Zelt et al. (2003) and Henning et al. (2004). Data from SCREECH Line 1 are summarized in Figures 4 & 5. The oldest rift-related sediments found by ODP Leg 103 are Tithonian in age (c. 142 Ma Boillot et al. 1987; Enachescu 1987) and, within the GIB, the oldest sediments that can be identified with confidence are also Tithonian (Murillas et al. 1990; Pe´rez-Gussinye´ et al. 2003). Murillas et al. (1990)

described a seismic reflection sequence beneath the Tithonian unit that they proposed is either Late Triassic–Middle Jurassic or older Palaeozoic sediments. In addition, the Flemish Cap graben along the SW part of the cap contains evaporites that are presumably Triassic (Enachescu 1987). However, available data suggest that the Late Triassic phase of extension may have only minimally affected the region this far north. In considering the timing of rifting and the duration of extension along the conjugate profiles, however, it should be noted that drilling and sampling along the Flemish Cap side of the margin is essentially non-existent. An additional point regarding the conjugate margin is that continental crust terminates just seaward of the ‘S’ reflection at a single peridotite ridge beyond which magmatically produced oceanic crust is thought to exist. This is unlike the area farther south along Iberia where a broad, approximately 100 km-wide, zone of exhumed continental mantle and multiple peridotite ridges were exposed at the sea floor prior to magmatic sea-floor

Fig. 4. Seismic velocity model along the SCREECH 1 transect derived from wide-angle data. See text and Funck et al. (2003) for details. Top panel is the model with no vertical exaggeration. Bottom panel is exaggerated 3:1 to show detail. Labelled contours are P-wave velocity.

STRUCTURE OF THE FLEMISH CAP MARGIN 51

Fig. 5. (a) Pre-stack depth migrated data from Hopper et al. (2004) shown from 200 to 320 km along the SCREECH 1 transect. No vertical exaggeration. (b) same as (a) but with seismic velocity model from Figure 4 superimposed. White lines are velocity contours at 0.2 km s21 interval. Greyed area towards the end of the profile indicates no ray coverage in the wide-angle data.

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spreading (Pickup et al. 1996; Dean et al. 2000; Whitmarsh et al. 2001). However, seismic refraction surveys along Galicia Bank terminate less than 40 km seaward of the peridotite ridge (Whitmarsh et al. 1996), so precisely where magmatically produced oceanic crust begins is not well constrained.

Initial condition and crustal thinning Flemish Cap itself is 30 km thick and appears to be unaffected by regional extension (Fig. 4). Large, sediment-filled rift basins are absent in the reflection images and the wide-angle data show no evidence for sediment cover (see Funck et al. 2003; Hopper et al. 2006). The basement has uniform P-wave velocities between 6.0 and 6.7 km s21, consistent with that typical of Appalachian continental crust. King et al. (1985) correlate Flemish Cap to the Avalon Terrane onshore Newfoundland, which has a seismic velocity structure very similar to that measured along the central part of the cap (Fig. 4) (Hall et al. 1998; Funck et al. 2003). Rock samples from the cap are primarily granodiorites and granites (King et al. 1985) and the crust onshore has whole-crust bulk Poisson’s ratios of 0.23–0.24, indicating that quartzo-feldspathic rocks dominate the crustal composition (Hughes et al. 1994; Hall et al. 1998). One important constraint for modelling exercises is to establish probable compositions for the lower continental crust. In the velocity model of Funck et al. (2003), there are no refracted arrivals from the lower-most layer in the crust. The velocity is thus constrained primarily by PmP phases (Moho reflections), which are clear on most of the Ocean Bottom Seismometers (OBSs) across the Flemish Cap. Despite the inherent ambiguity in determining velocity– depth profiles based only on reflections, lower crustal velocities in access of 6.75 km s21 are not supported by the data. This effectively rules out the possibility of significant gabbroic intrusions in the lower crust. At pressures appropriate for 25–30 km depth, gabbro has a P-wave velocity of more than 6.85 km s21 (Christensen & Mooney 1995). The Flemish Cap data are most consistent with measured velocities for intermediate composition rocks such as diabase (see the compilation in Christensen & Mooney 1995). More mafic rocks such as amphibolites, anorthositic granulites and mafic granulites are also possible, but require temperatures in access of 650 8C. No heatflow data exist on Flemish Cap, but such high temperatures seem unreasonable in this setting. Given the similarity of the velocity structure with onshore observations, where Poisson’s ratio provides additional constraints, and the difficulty with reconciling mafic compositions with the available data, it

is very likely that Flemish Cap is dominated by relatively weak crustal mineralogies. Towards the Flemish Pass basin at the NW end of the transect, reduced seismic velocities are observed in the upper crust. Although the shape of this unit is suggestive of a sedimentary basin, the lack of stratigraphic layering in the reflection section argues against its origin as a rift basin associated with Triassic or later rifting, which should be well preserved. In addition, the seismic velocities are much higher than observed in other Mesozoic basins around Newfoundland. Funck et al. (2003) suggested that this unit most probably corresponds to Precambrian and Palaeozoic sediments and volcanics associated with the Avalon Terrane of the Appalachians. The lack of significant basin formation indicates that Flemish Cap is an intact block of continental lithosphere and thus can be considered as a microplate. This is consistent with previous work on reconstructions of the region (Srivastava et al. 2000; Sibuet et al. 2007). Behaviour as a microplate implies that Flemish Cap must be a strong block of lithosphere that did not deform internally in response to regional extensional stresses responsible for lithospheric rupture. Given the likelihood that the crust is dominated by weaker lithologies, it is inferred that the strong part of the lithosphere must rest primarily in the mantle. This places a qualitative constraint on the initial thermal condition of the lithosphere at this location. In particular, a strong upper mantle indicates reduced temperatures and thermal gradients and implies a cool initial condition (Brace & Kohlstedt 1980). This probably precludes the possibility of significant lower crustal flow in response to crustal buoyancy forces as this requires relatively warm lithosphere even for very weak crustal mineralogies (e.g. Hopper & Buck 1998). Lower crustal flow, it should be noted, is distinct from ductile stretching of the lower and middle crust during extension. The seaward transition from relatively undeformed, 30 km-thick, continental crust to extremely thinned, less than 5 km-thick crust occurs over a narrow interval of 80 km. Moderately thinned continental crust is marked by a single sedimentary basin at 180 km (Figs 4 & 6). Although the basin is clearly observed in the both the wide-angle and normal incidence data, the reflection image quality is poor due to the incomplete removal of multiple energy. However, weak sedimentary reflections appear to fan seaward, indicating a halfgraben bounded by a landward-dipping fault. The crust beneath this basin is 20 km thick, but it thins abruptly seaward to a thickness of less than 5 km over only 40 –50 km. The abruptly necking zone of crust from 150 to 240 km is particularly interesting. Figure 7 shows

Fig. 6. Seismic reflection section in two-way travel time, where continental crust thins abruptly from depths 170 to 215 km. Top panel is uninterpreted data, bottom panel includes an interpretation as discussed in the text. Seaward of 215 km, the data are time-converted from the pre-stack depth migration from Hopper et al. (2004). The full depth migration is also shown in Figure 5.

STRUCTURE OF THE FLEMISH CAP MARGIN 53

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Thinning factor

5.00

4.00

3.00

UC

MC

LC

2.00

1.00 150

160

170

180

190

200

210

Distance along profile (km)

Fig. 7. Thinning factors of different seismically defined layers as a function of distance along the profile. The thinning factor is simply a reference thickness divided by the actual crustal thickness at a location. Reference thickness is taken from 110 km, which is approximately where each layer is thickest across the Flemish Cap on the SCREECH profile. UC, upper crust (Vp ¼ 6.0– 6.2 km s21); MC, middle crust (Vp ¼ 6.3–6.4 km s21); LC, lower crust (Vp ¼ 6.6– 6.7 km s21). See text for further details.

the thinning factors of different crustal layers as defined by the seismic refraction data. The upper crust has velocities from 6.0 to 6.2 km s21, middle crust has velocities from 6.3 to 6.4 km s21, and lower crust has velocities from 6.6 to 6.7 km s21 (Funck et al. 2003). From 150 to 210 km, the upper crust shows extensional thinning and strain concentrated in a local area defined by the rift basin at 160 –180 km. Little thinning of the upper crust is observed elsewhere over this interval. In contrast, the middle and lower crust thin uniformly as is expected for layers that have undergone ductile stretching by bulk pure shear. At least locally, these observations seem to require that deformation of the upper crust was detached from deformation deeper in the crust. The overall amount of extension in the upper crust seems to balance with the amount of extension indicated in the middle and lower crust over this interval despite a small lateral offset in the location where strain was focused. The amount of stretching in the middle and lower crust increases rapidly from 200 km seaward and one or both these layers may pinch out. The details of this pinch-out, however, are difficult to resolve. In addition, the use of seismic velocities to define original upper, middle or lower crust under such extreme thinning is questionable because decompression, porosity changes and mineral alteration from circulating fluids will alter the seismic velocity as deeper material is brought closer to the surface. Thus, we cannot attach significance to layer thickness changes defined solely by seismic velocities beyond 210 km. In the zone from 210 to 240 km, the middle and lower crust appear to be very thin, and the upper crust thins significantly. The presence of small rift

basins indicates brittle failure of the upper crust, but specific interpretation of the fault pattern is difficult because of the lack of reflections from fault surfaces. One possible interpretation of this section is shown in Figure 6. An important feature of the data is an interval of east-dipping reflections primarily in the lower crust that is interpreted to be Palaeozoic structure formed during the assembly of the Avalon terrane (Hopper et al. 2006). These reflections continue through and below reflection and refraction Moho, although the reflection Moho is discontinuous through the cross-over. It is possible that the Palaeozoic structure was reactivated by normal faulting and may have acted as a detachment surface at which overlying high-angle normal faults terminate. The small basins centred at 216 and 227 km appear to enclose an interval with high velocities of 6.0–6.1 km s21. On the basis of velocity alone, this might be considered to represent upper continental crust. However, the sections show horizontal layering that is unconformable to the basin margins, suggesting a possible sedimentary origin. It is possible that they are limestones and dolomites, which can have velocities of 6 km s21 or more (e.g. Press 1966). In addition, it is known from drilling at ODP site 639 that Upper Jurassic–Lower Cretaceous dolomite is present on the conjugate Galicia margin (Boillot et al. 1987). If the interpretation that much of the seismic upper crust is in fact sedimentary, then faulting probably formed the basins early during the second Newfoundland–Iberia rift phase. Faulting during the earlier, Late Triassic – Early Jurassic, rifting is unlikely because there is no indication of evaporites, which probably would be mobile and would distort the sedimentary fill. Deposition of evaporites was widespread in the basins formed on the Grand Banks and Iberia margins during the first rift phase (Tankard & Welsink 1987; Wilson et al. 1989). The origin of the basement highs at 230–240 km is uncertain. Hopper et al. (2006) present arguments for and against various hypotheses of the section as either continental blocks or as exhumed continental mantle. Figure 6 shows one possible interpretation where the blocks are continental crust bounded by east-dipping normal faults that formed during the early phase of extension that led to the abrupt thinning along the margin. Erosion of the tops of the blocks would explain their rounded tops and apparent coarse debris shed into the adjacent basins. The fault labelled 1.1 would post-date the main faulting associated with these blocks, thus forming the basins centred at 216 and 222 km probably during the Jurassic (see above). The orange faults within the sedimentary fill are related to extension that post-dates the main crustal thinning phase of extension.

STRUCTURE OF THE FLEMISH CAP MARGIN

Although it is not possible to rule out the possibility that the blocks from 230 to 240 km could be exhumed continental mantle, major large-scale asymmetry between Flemish Cap and Galicia is clearly supported by the data. The abrupt necking profile and narrow region of basin formation is in strong contrast to the conjugate, which shows a broad, 300 km-wide region of attenuated continental crust. Figure 8 shows the profile described here reconstructed to a composite profile of the Galicia Bank and Galicia Interior Basin based on data from Pe´rez –Gussinye´ et al. (2003), Reston et al. (1996), Gonza´lez et al. (1999) and Whitmarsh et al. (1996). Importantly, the asymmetry observed is independent of the precise placement of the seaward termination of continental crust along Flemish Cap profile. Moving the boundary 10 km landward and interpreting the blocks as exhumed upper mantle does not fundamentally alter the conclusion of asymmetric margin formation. The question of symmetry v. asymmetry on conjugate rifted margins has received a great deal of attention (e.g. Lister et al. 1986; Dunbar & Sawyer 1989b; Sibuet 1992; Louden & Chian 1999). Early models suggested that asymmetry requires strong crust– mantle decoupling, whereby crustal strain is displaced from mantle lithospheric strain, eventually resulting in asymmetric conjugate rifted margin pairs (e.g. Braun & Beaumont 1989; Dunbar & Sawyer 1989a). Such a mode of decoupling, however, may be difficult to achieve for realistic geological conditions (Hopper & Buck 1998). Hopper & Buck (1996) proposed instead that asymmetric margins form when wide-rift mode extension breaks along one edge. Both these possibilities, decoupling and wide-rift mode extension, probably require warm lithospheric conditions (Buck 1991), which seems at odds with the arguments presented above for cool, stronger lithosphere. Recently, a new class of dynamical models of extension have shown that symmetry and asymmetry may instead be controlled by strain-dependent

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rheologies (Huismans & Beaumont 2002, 2003; Lavier & Manatschal 2006). In particular, the results of Huismans & Beaumont (2003) suggest that strongly asymmetric rifting occurs when: (1) extension is slow; (2) there is significant strain softening of the dominant rheology; and (3) crust is coupled to the lithospheric mantle. Their conclusions are broadly consistent with data available from the Flemish Cap.

Mantle serpentinization It has become increasing clear that penetration of sea water into the upper mantle can result in large-scale serpentinization along divergent plate boundaries. This results in fundamental but still poorly understood changes in the rheology of the upper mantle that probably have a significant impact on mantle and crustal deformation during extension. In addition to having rheological effects, the serpentinization process is strongly exothermic and results in significant volume expansion and density reduction (e.g. Schroeder et al. 2002). Thus, it may play an important role in the thermal structure and evolution of divergent plates boundaries as well as having an impact on the dynamic evolution through buoyancy changes. One of the most important results of research on the Iberia margin is the recognition of the important role that mantle serpentinization plays in the final stages of non-volcanic margin development (e.g. Manatschal 2004). In particular, crustal separation is not always coincident with final lithospheric rupture. Instead, a broad zone of exhumed continental mantle forms and there is an absence of extrusive magmatic rocks and oceanic crust. A key result of the SCREECH Line 1 is that, with the possible exception of the crustal block at 230–240 km, there is no convincing evidence of mantle exhumation occurring in a similar environment. In previous work, Funck et al. (2003) and Hopper et al. (2004,

Fig. 8. Plot of Flemish Cap data summary from this work reconstructed to conjugate data sets. Data were spliced together at the seaward-most point where continental crust can be identified in each case and the data shifted so that the sea floor is at the same depth. Conjugate data are based on profiles published in Reston et al. (1996), Whitmarsh et al. (1996), Gonza´lez et al. (1999) and Pe´rez-Gussinye´ et al. (2003).

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2006) presented a number of arguments demonstrating that some type of oceanic crust exists seaward of 240 km. OBS 19, just on the eastern side of the basement blocks at 230 –240 km, shows clear evidence for a rapid seaward velocity change that is best explained by anomalously thin oceanic crust overlying serpentinized mantle beginning at this location (Figs 4 & 5) (Funck et al. 2003). Thus, it appears that a combination of extrusive magmatism and mantle exhumation characterized extension shortly after continental crustal separation along this transect. Nevertheless, mantle serpentinization is clearly indicated along the profile, and it must have played an important role in the late stages of margin development that led to full lithospheric rupture. Well-defined and easily identifiable upper mantle Pn phases from beneath the thin magmatic crust along the profile consistently show phase velocities of 7.6 km s21. This is significantly less than the 8 km s21 measured below Flemish Cap (Funck et al. 2003), and it suggests a maximum of approximately 15– 20% serpentinization of olivine peridotite immediately beneath the thin crust (100 –150 Mpa: e.g. Christensen 2004). Amplitude modelling further constrains the velocity gradient in the mantle and shows that velocities of 8 km s21 are reached by about 5 km below the base of the crust. The velocity data provide convincing evidence that serpentinization becomes important when the crust thins to less than 6–8 km thick. Figure 9 shows the mantle P-wave velocity just below Moho as a function of crustal thickness along the SCREECH Line 1. Assuming that reduced mantle velocities indicate the onset of serpentinization, this process appears to initiate at a crustal thickness of approximately 8 km. Simple thermal models of lithospheric extension show that this is about the point at which the entire crust becomes brittle (Pe´rez-Gussinye´ & Reston 2001). In the absence of significant melt production, extension can then only be accommodated by brittle failure, which

Vp (km/s)

8.25

8.00

7.75 onset of serpentinization

7.50 0

3

6

9

12

15

Crustal Thickness (km)

Fig. 9. Mantle P-wave velocity just below the Moho as a function of thickness of the overlying crust. See the text for explanation.

provides pathways for water to penetrate into the mantle and to facilitate serpentinization. This idea is generally consistent with the interpretation of the seismic reflection section of SCREECH transect, given that major faults bounding the basins at 210–230 km most probably cut deep into the crust and possibly into the upper mantle. These faults could provide the conduits necessary for fluids to circulate deep into the lithosphere.

Magma production and sea-floor spreading The last important observation the SCREECH data provide is information on the onset of melting and the generation of magmatically produced oceanic crust, which marks final lithospheric rupture. Historically, amagmatic margins have been studied to understand brittle and mechanical deformation of the crust. The absence of melt in the system has important consequences. From an observational point of view, extensional structures that result from thinning and stretching are easier to image and sample as they are not covered up by volcanic rocks. From a modelling point of view, extensional stress cannot be relieved by magmatic intrusion and dyke injection so only mechanical deformation of the lithosphere needs to be considered up to the time that volumetrically significant melt is produced. The SCREECH Line 1 results are fully consistent with an absence of significant decompression melting prior to breakup of the continental crust, despite extreme stretching and thinning. Across Flemish Cap and along the abruptly thinning crust farther seaward, seismic velocities measured in the lower crust are low, consistent with an absence of underplating or gabbroic intrusion into lower crust. Crustal velocities of more than 6.8 km s21 are not observed anywhere in the section until 240 km, where thin magmatic crust is interpreted. In addition, the reflection data show little evidence for significant sill intrusion into the few sedimentary basins that are observed. To the extent that melts may have formed and infiltrated the upper mantle, there was insufficient production to generate extrusive flows or geophysically resolvable intrusive layers within the crust. Melting and magmatism, when it did appear, produced anomalously thin, 3–4 km thick, oceanic crust, indicating a magma-starved crustal accretion system. Such crust is observed in modern settings along ultra-slow spreading ridges like the Gakkel Ridge where crust is only 4 km thick (Coakley & Cochran 1998; Jokat et al. 2003). This is further consistent with Srivastava’s estimate that extension at the time of opening was approximately 14 mm

STRUCTURE OF THE FLEMISH CAP MARGIN

year21 (full rate), which is also comparable to Gakkel Ridge spreading rates. Crust with closer to average oceanic thickness did not appear until about 295 km on SCREECH Line 1, some 20–25 km seaward of anomaly M0 as picked by Srivastava et al. (2000) (Fig. 5). Unfortunately, velocities and crustal thickness on this part of the refraction line are not well constrained (Funck et al. 2003), and the line does not extend far enough to confirm that the crust continues seawards with normal oceanic thicknesses. While it is tempting to conclude that extension rate was the controlling factor in suppressing melting in the Iberia –Newfoundland rift, it is not clear that this is the only important factor. Reston & Phipps-Morgan (2004) proposed that anomalously cool upper mantle may be important in limiting or suppressing melt production. A compilation of continental geotherms by Ro¨hm et al. (2000) shows that the potential temperature of the upper mantle may have a natural variation of several hundred degrees, and the possibility that the Newfoundland Iberia rift was underlain by an asthenospheric cool spot is one that must be considered. It is clear that large variations in melt supply are indicated once magmatism was established in the rift. At 273– 290 km, oceanic layer 3 is absent, and the crust thins to as little as 1.5 km thick in part of this zone. Hopper et al. (2004) attributed this to a cessation of magmatism that resulted in the formation of an oceanic core complex in which the lower crust was mechanically removed. They further proposed that magmatism was episodically more robust, in order to explain apparent volcanic stratigraphy that is observed in the upper crust. Such volcanic stratigraphy is also observed on data 10 km to the north of Line 1 (Hopper et al. 2006). This kind of variation is not unique to the Flemish Cap –Galicia transect. There are significant disparities in the amount of melt introduced throughout the rift following breakup of the continental crust, although a common denominator appears to be that all areas exhibit significant exhumation and serpentinization of mantle (see Tucholke et al. 2007). Chemical heterogeneities may have also played a key role in the temporal and spatial variability of magmatism. In a magma-starved rift, small pockets of anomalously fertile mantle could have a large effect on local areas. For example, passively imbedded mantle inhomogeneities have been proposed to explain isolated magmatic centres along the ultra-slow Gakkel Ridge (Cochran et al. 2003). The available data indicate that large spatial and temporal variations in melt supply may be important for understanding the detailed structure of non-volcanic margins during the extensional phase following complete separation of continental crust

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and prior to the establishment of a robust sea-floor spreading system. A complex interplay of brittle faulting and mantle exhumation, melting, and intrusion and extrusion of melt controls the detailed structure of this crust. Geophysical data alone cannot constrain these processes and further progress in this area will require basement sampling and geochemical analyses.

Conclusions The seismic reflection and refraction data from the Flemish Cap margin, together with seismic surveys over the conjugate Galicia Bank margin and Galicia Interior Basin, provide first-order constraints for numerical modelling of lithospheric rupture in a nonvolcanic rift. In particular, the data show important differences in margin development and document asymmetries on a variety of scales. Well-defined crustal and mantle arrivals provide detailed constraints on the crustal thickness across the entire transect. Flemish Cap itself has crustal seismic velocities consistent with velocities measured in Appalachian crust from the Avalon terrane throughout the Newfoundland region. Such crust is probably dominated by quartzo-feldspathic lithologies (Hughes et al. 1994). There is a conspicuous absence of significant rift basins in the continental block comprising Flemish Cap. Rift basins appear on the eastern margin of the block, but they are small compared to other rift basins on the western side of the block and within the Grand Banks. The continental crust thins rapidly on the eastern margin of Flemish Cap, from 30 to approximately 3 km thick, over a distance of about 80 km. Mantle velocities beneath Flemish Cap are 8.0 km21 and thus do not indicate anything unusual about the composition of the upper mantle. Reduced seismic velocities are observed under thinned continental crust farther seaward, beginning where the crust is approximately 8 km thick. Reflection images suggest that the embrittlement of the entire crust may have occurred at this point, providing pathways for fluids to penetrate into the mantle and facilitate serpentinization. However, with the possible exception of one 10 km-wide basement block at the edge of the thinned continental crust, there is no evidence that mantle was extensively exhumed as is observed off the Iberia margin. Nevertheless, serpentinization of the upper 5 km of mantle is indicated by anomalously low seismic velocities. Serpentinization is an exothermic reaction that results in a significant volume increase and density reduction. In addition, the rheology of serpentinite is considerably different than peridotite. Thus, there clearly are important, but as yet unexplored, implications of this process

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for understanding the dynamic evolution of rifted margins during the final stages of lithospheric rupture. We do not observe high seismic velocities in the lower crust that would be expected from underplating or gabbroic intrusion as a result of decompression melting during rifting of the continental crust. There is no evidence for melt production until after the breakup of the continental crust, at which point thin (3–4 km) magmatic crust with oceanic velocity structure was produced. This crust shows a strong variation in thickness and it consistently overlies serpentinized mantle several kilometres thick, indicating a complex phase of mantle extension with variable magmatism before a more magmatically robust sea-floor spreading system was established. The timing of rifting events on the Flemish Cap margin is uncertain because of a lack of drilling and sampling. It seems likely that the first phase of Newfoundland–Iberia extension (Late Triassic – Early Jurassic) did not significantly affect the crust between Flemish Cap and Galicia Bank. The oldest sediments drilled on Galicia Bank are shallow-water Tithonian carbonates and mudstones, which suggests that rifting of the crust there began during the second, Late Jurassic –Early Cretaceous, phase of extension. Possible high-velocity sediments in the small rift basins at the seaward edge of thinned Flemish Cap crust may be the counterparts of these Tithonian carbonates at the western edge of the rift. If it is assumed that rifting of the eastern Flemish Cap margin began in the Tithonian, then continental crustal extension followed by the formation of thin magmatic crust combined with mantle extension and exhumation occupied an interval of at least approximately 25 million years until apparently normal ocean crust was formed seaward of anomaly M0. There are several parameters critical for numerical modelling that the presently available data do not constrain very well. There is no information on the lithospheric thickness of the region, and we can only speculate on the thermal structure of the lithosphere prior to rupture. However, the data are consistent with Flemish Cap behaving as a microplate, implying that is a relatively strong block of lithosphere. Thus, a relatively cool geotherm and thick lithospheric mantle is indicated. The SCREECH survey was a joint project sponsored by the US National Science Foundation (NSF), the Danish National Research Foundation, and the Natural Science and Engineering Research Council of Canada. We thank the captains and crews of the R/V Maurice Ewing and R/V Oceanus for a successful experiment. Depth imaging was supported by European Union grant HPRI-CT1999-00037 to GEOMAR (now Leibniz Insititute for

Marine Science) in Kiel, Germany. This work has benefited from discussions with our SCREECH colleagues K. Louden, D. Shillington, H. van Avendonk, W. S. Holbrook, H. C. Larsen, H. Lau & S. Deemer. Reviews from S. Deemer, R. Huismans, G. Manatschal & M. Steckler helped to improve the manuscript. B. Tucholke acknowledges support by NSF grants 9819053 and 0326714, and by the Bigelow Chair in Oceanography at Woods Hole Oceanographic Institution.

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Early Cretaceous motion of Flemish Cap with respect to North America: implications on the formation of Orphan Basin and SE Flemish Cap –Galicia Bank conjugate margins J.-C. SIBUET1, S. P. SRIVASTAVA2, M. ENACHESCU3 & G. D. KARNER4,5 1

Ifremer Centre de Brest, B.P. 70, 29280 Plouzane´ Cedex, France (e-mail: [email protected])

2

Natural Resources Canada, Geological Survey of Canada, Bedford Institute of Oceanography, P.O. Box 1006, Dartmouth, N.S., Canada B2Y 4A2 3

Memorial University of Newfoundland, St John’s, Newfoundland, Canada 4

Lamont-Doherty Earth Observatory, Palisades, NY 10964, USA

5

Present address: ExxonMobil Upstream Research Company, Mail Stop URC-URC-S169A, P.O. Box 2189, Houston, TX 77252-2189, USA Abstract: Bouguer gravity anomalies together with deep seismic reflection and magnetic data on both sides of the North Atlantic are used to locate the hinge zones of the Flemish Cap and Galicia Bank within the Iberian and North American plates, regions across which there were abrupt changes in lithospheric extension. The characteristic shape and alignment of these hinge zones suggest that they were conjugate features generated during chrons M25– M0 (Late Jurassic– Early Aptian) around a distally located Euler pole of rotation. Rifting between Iberia and North America involved these two larger plates and the two smaller microplates – the Flemish Cap and Galicia Bank microplates. The motion of the microplates, which were adjacent to Eurasia, was much more complex than those of the larger plates. The motion between the microplates from chron M25 or older to chron M0 was complicated by the fact that they remained attached to each other for most of the time when regions to the south were rifting apart. As a result, continental regions landward of these segments were subjected to extension that created the Orphan and Flemish Pass basins on the North American side and the Galicia Interior Basin on the Iberian side. By comparing the hinge zones delineated off Galicia Bank and Flemish Cap using the Bouguer anomalies, we were able to infer that Flemish Cap rotated approximately 438 relative to Galicia Bank and Iberia, and moved 200 –300 km SE with respect to North America. Such motions of Flemish Cap and Galicia Bank agree remarkably well with extensional episodes deduced from industry multichannel seismic reflection data acquired in the Orphan Basin. Normal fault orientations identified in the West Orphan Basin trend N0208 and are approximately perpendicular to the flow lines of our proposed Flemish Cap–North American motion during the M25–M0 period, which provides an independent constraint on our proposed kinematic model. Therefore, extensional events affected not only the Galicia Bank–Flemish Cap conjugate margins but also the Galicia Interior and Orphan Basins, and need to be taken into account in any assessment of the geological development of the Iberian and North American continental margins.

During the InterMARGINS Workshop meeting ‘Modelling the Extensional Deformation of the Lithosphere’ (IMEDL) held in Pontresina in 2004, a benchmark exercise was proposed to model the crustal and mantle rifting processes between the western Galicia Bank and the SE Flemish Cap conjugate margins where high-quality seismic reflection and refraction data exist (SCREECH Line 1: Funck et al. 2003; Hopper et al. 2004) and ISE Line 17 (Pe´rez-Gussinye´ et al. 2003; Henning et al. 2004). During the meeting it became apparent that additional boundary conditions and constraints were needed to correctly characterize the known extension across

the Iberian–Newfoundland margin. The aim of this paper is to provide some of these conditions and constraints, namely a knowledge of the initial positions of plates and microplates before their separation by sea-floor spreading processes. The positions of Flemish Cap and Galicia Bank are well constrained at chron M0 (Fig. 1a) because of the identification of correlatable sea-floor magnetic anomalies in both the Bay of Biscay (Sibuet et al. 2004b) and in the North Atlantic (Srivastava et al. 2000). The resulting Euler pole position was constrained by the geometry of the triple junction that had remained at the mouth of the Bay of

From: KARNER , G. D., MANATSCHAL , G. & PINHEIRO , L. M. (eds) Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 63–76. DOI: 10.1144/SP282.4 0305-8719/07/$15.00 # The Geological Society of London 2007.

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Fig. 1. (a) Reconstruction of the North Atlantic between Africa (AF), Iberia (IB) and North America (NA) at chron M0 with respect to Eurasia (EU) as given by Srivastava et al. (2000) (IB– EU 43.858N, 5.838W, 244.768; AF– EU 43.618N, 6.938W, 240.898; NA –EU 69.678N, 154.268E, 23.178). M0 picks west of IB (crosses) and east of

EARLY CRETACEOUS ROTATION OF FLEMISH CAP

Biscay during this period and by palaeomagnetic declination data on the stable Iberia plate. However, similar constraints are not available for earlier time periods. Based on palaeomagnetic data, Srivastava & Verhoef (1992) derived the positions of the Eurasian (EU) and Iberian (IB) plates relative to North America (NA) for chron M25. For consistency with previous workers in this region, we have adopted the Kent & Gradstein (1986) timescale in this paper. The time of chron M25 (156 Ma, Late Jurassic) thus corresponds to the beginning of the Late Jurassic–Early Cretaceous rifting phase. Although M25 magnetic lineations have not been identified between the EU, IB and NA plates, for convenience we will use ‘M25’ either as the time of chron M25 or as the beginning of the Late Jurassic– Early Cretaceous phase of extension. When the positions of the three plates (EU, IB and NA) are reconstructed there is significant overlap. To remove the overlap the plates need to be translated, rotated and unstretched by an amount equal to the amount of overlap between them. Thus, the reconfigured plates at chron M25 time suggest that the Flemish Cap (FC) would lie NW of its present position. Similarly, the Galicia Bank (GB) would lie eastward of its present position at this same time. In deriving the relative positions of the Flemish Cap and Galicia Bank, Srivastava & Verhoef (1992) assumed that an equal amount of extension took place within all the basins lying significantly landward of the hinge zones. The detailed seismic reflection and refraction data recently acquired in this region suggest that such an assumption is not valid. To derive more accurate restored positions of these basins and plates, we have combined an interpretation of Bouguer gravity imaging with interpretations of commercial seismic reflection data from the various adjacent basins.

Quantifying the amount of extension between the Eurasian, North American and Iberian plates at M25 – M0 time Figure 1b shows the amount of crustal extension calculated during the M25–M0 period (156 – 118 Ma, Late Jurassic –Early Aptian) using parameters

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from Srivastava & Verhoef (1992) for chron M25 and from Srivastava et al. (2000) for chron M0. As the M25–M0 stage poles summarizing the rotation for EU –NA, IB –EU and IB –NA are distant from our study area, the amount of extension between each plate pair will not vary significantly along strike, at least within the region shown in Figure 1b. This gives an estimated extension of 350 km between the IB and EU plates, 130 km between the EU and NA plates, and 450 km between the IB and NA plates. In this paper we use the expression ‘transitional crust’ in the traditional sense of the expression; namely, as the transition between distal thinned continental crust and the first clearly recognized oceanic crust, even though we appreciate that the transitional crust between the margins of the Iberia Abyssal Plain and the Newfoundland Basin consists mostly of exhumed serpentinized continental mantle with some decompression melting products (e.g. Tucholke et al. 2007). Inversion of magnetic data (Sibuet et al. 2007) shows that magnetic sources are N–S-trending horizontal bodies located within the highly serpentinized upper crust rather than in the lower crust (cf. Russell & Whitmarsh 2003) and that the serpentinization process is able to produce magnetic lineations in an analogous way as that for the oceanic crust. Thus, within transition zones, sequences of magnetic anomalies can provide constraints concerning the timing of emplacement of crust and continental mantle within an extensional regime but not on the nature of that material. During the M25–M0 period, extension between the EU and IB plates resulted in the formation of the northern and southern Bay of Biscay continental margins, the Armorican Basin transitional crust and the subsequent formation of oceanic crust (or possibly transitional crust) in the Bay of Biscay during M3–M0 period (Montadert et al. 1979; Sibuet et al. 2004b). During this same period, extension also resulted in fault reactivation in the Paris Basin, the Celtic Sea and the English Channel (Ziegler 1988), south of the Ebro Variscan massif (Verge´s & Garcı´a-Senz 2001) in the Iberia, Catalan and Basque–Cantabrian basins (Casas 1993; Salas et al. 2001), in the Organya` Basin (southern Pyrenees)

Fig. 1. (Continued ) NA (white circles) are from Srivastava et al. (2000); north of IB (crosses) and south of EU (white circles) from Sibuet et al. (2004b). A, B and C are conjugate points of the fossil triple junction trajectories at this time. FC, Flemish Cap; FPB, Flemish Pass Basin; GB, Grand Banks; GS, Goban Spur; GB, Galicia Bank; JAB, Jeanne d’Arc Basin; PSB, Porcupine–Seabight Basin; WB, Whale Basin. (b) The calculated amount of extension between M25 and M0 (Late Jurassic–Early Aptian) is shown in grey on the M0 reconstruction. The parameters used are those from Srivastava & Verhoef (1992) for M25 and from Srivastava et al. (2000) for M0. M0 picks west and north of IB (crosses and open dots, respectively) are from Srivastava et al. (2000) and Sibuet et al. (2004b). Large transparent arrows show the M25–M0 directions of plate motions. Hachured lines show continental shelf basins and continental margins where extension occurred during that period. Small black arrows indicate rotation of Galicia Bank and Flemish Cap during this period. BCb, Basque–Cantabrian basins; Cb, Catalan basins; Ib, Iberian Basins; Mb, Maule´on Arzacq basins; NPF, North Pyrenean Fault; Ob, Organya` Basin; Pb, Parentis Basin.

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(Bera`stegui et al. 1990; Verge´s & Garcı´a-Senz 2001), and north of the Pyrenees in the Parentis, Maule´on and Arzacq basins (Bois et al. 1997; Masse 1997; Winnock 1971) (Fig. 1b). Between the EU and NA plates, continental rifting was still ongoing at chron M0. Transitional or oceanic crusts were not formed during this period but extension occurred on their continental margins, in the Porcupine and Seabight basins, and in the Rockall Trough and on the adjacent shelf areas. Between the IB and NA plates, extension gave rise to the development of conjugate margins, to the formation of a transitional crust, and then to oceanic crust since chrons M5–M3 (Russell & Whitmarsh 2003) or more probably since the Aptian–Albian time as implied from drilling results at sites 1067 (Whitmarsh & Sawyer 1996) and 1277 (Shipboard Scientific Party 2004) at time of chrons M2 and M1, respectively (Sibuet et al. 2007; Tucholke et al. 2007). For example, extension was recorded not only in the Salar and Carson basins located on the continental slope of the Grand Banks but also across the Grand Banks, in the Flemish Pass and in the Jeanne d’Arc, Horseshoe and Whale basins (Fig. 1) from Late Jurassic to Early Cretaceous. In general, this extension was significantly smaller than the earlier Late Triassic–Early Jurassic extensional phase in the region. Between the northern Newfoundland Basin (NNFB) and Iberia Abyssal Plain (IAP) conjugate margins, a transitional crust about 300 km wide was formed as evidenced from wide-angle seismic refraction and drilling data both in the IAP (Chian et al. 1999; Dean et al. 2000; Whitmarsh et al. 2001) and in the NNFB (Lau et al. 2003; Shipboard Scientific Party 2004). Perhaps even a wider transitional domain exists along the conjugate Tagus Abyssal Plain and southern Newfoundland Basin margins, but not to such a large extent between Galicia Bank and the SE Flemish Cap (Whitmarsh et al. 1996). This is owing to the simultaneous relative motion of Flemish Cap and Galicia Bank with respect to their adjacent NA and IB plates (Fig. 1b), as previously suggested by Le Pichon et al. (1977), Srivastava & Verhoef (1992), Srivastava et al. (2000) and Sibuet et al. (2004b). In this paper, we use new geophysical constraints in order to quantify the FC– NA motion, which has occurred since Late Jurassic, and compare the resulting motions with the direction of structural trends and the amount of extension compiled by Enachescu et al. (2004a– c) from industry seismic reflection data collected in the Orphan Basin located NW of FC.

North American and Iberian hinge zones Le Pichon & Sibuet (1981) demonstrated that if the thinning mechanism of continental margins is

basically pure shear, then the location of the plate boundary before stretching approximately follows a present-day isobath at a distance equal to the width of unstretched crust; for example, the 3000 m-isobath on the northern Biscay margin. In other words, if a is the width of the unstretched continental crust on a plate, which has been thinned during extension, then the isobath defining the prestretched position of the plate will be located at a distance a from the edge of the undeformed continental crust and will correspond to the maximum gradient in the Bouguer gravity anomaly (Fig. 2). We call these lines or boundaries ‘hinge zones’ in Figure 2, across which abrupt changes in lithospheric extension occur. Thus, if the extension mechanism is pure shear, the reconstruction of the hinge zones would correspond approximately to the restoration in space of the two portions of initial continental crusts before extension. The locations of these boundaries across the margins of NA and IB plates were deciphered from the Bouguer gravity anomalies determined from the 11 minute global free-air gravity grid of the world’s ocean generated by Sandwell & Smith (1994) and the 11 minute GEBCO topographic grid (IOC, IHO and BODC 2003) using a sediment –water contrast of 2200–1030 kg m23 (Karner et al. 1997). To decide what density contrast to use, for example 2200 v. 2800 kg m23, is not straightforward. This is because the edge-effect free-air gravity anomaly of a margin is partitioned between the gravity effects across the hinge zones and across the shelf-break, as shown by Karner & Watts (1982). These effects can play important roles in determining the location of hinge zones, depending on whether the margin is sediment rich or starved. Using 2200, 2400 and 2800 kg m23 as the reduction density made little difference in the location of the maximum gradients in the Bouguer anomaly maps across the NA and IB margins. That is, the gravity effect of the hinge zone is the dominant contribution to the gravity. Thus, we use a sediment –water reduction density of 2200– 1030 kg m23 to show our reconstruction of the NA and IB margins at chron M0. On each side of the IB–NA M0 boundary, the Bouguer values are significantly different, implying that the two crusts are of different thicknesses. This is confirmed by the depths of unloaded basement, which is 500–800 m deeper west of Galicia Bank than east of Flemish Cap, and by different mean values of heat flow data collected on both sides of the ocean (Louden et al. 1997; Louden & Lau 2004). Beneath the NA and IB continental slopes, the hinge zones roughly follow the 280 mGal isolines (white continuous lines in Fig. 3). The existence of such a crustal gravity gradient is the consequence of the oceanward thinning of the continental crust.

EARLY CRETACEOUS ROTATION OF FLEMISH CAP

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location of the initial rupture

continental crust before extension a

HZ

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undeformed continental crust

mantle a

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thinned continental crust

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Fig. 2. Pure shear model of continental crust extension. HZ, hinge zone; a, width of the deformed continental crust. The maximum Bouguer gravity anomaly gradient occurs at the vertical of the hinge zone located at distance a from the undeformed continental crust. Fitting the hinge zones identified on the Bouguer gravity map brings the extended plates to their prerift positions.

However, the trends of the hinge zones do not necessarily follow the bathymetric trends and gradients because bathymetric contours are influenced by local post-rift sedimentary processes. For example, along the IAP –NNFB conjugate margin transect, deep bathymetric trends (within the green colour of Fig. 3 at 448N; 138W) are 458 oblique to the gravity trends. Concerning the segment of margin located immediately to the south, which corresponds to the Estramadura Spur and the bathymetric ridge located between the IAP and the Tagus Abyssal Plain, the hinge zone is clearly defined by the Bouguer gravity anomalies. It is for this reason that plate tectonic reconstructions based solely on the fit of bathymetric trends between IB and NA are not convincing. Interestingly, south of the Galicia Bank and Flemish Cap, the hinge zones on the NA and IB plates have similar geometries. We suggest that they are conjugate features whose positions can be described by a distally located pole of rotation,

primarily because the distance between these zones is the same from south of Galicia Bank and Flemish Cap to the Newfoundland –Gibraltar fracture zone. Based on the Bouguer gravity and bathymetric data, flow lines can be inferred between the NA and IB conjugate segments (thick black lines in Fig. 3). Consequently, if we assume that the beginning of the rifting phase between IB and NA plates started in the Late Jurassic (close to M25 time), then the distance between the two hinge zones in the M0 reconstruction would correspond approximately to the M25–M0 displacement between the NA and IB plates. South of Galicia Bank and Flemish Cap, the M0 trend, as defined by the M0 picks (Fig. 3), is approximately linear, while the trends of the hinge zones are not. Such a modification of geometry occurred some time close to the onset of the emplacement of transitional crust. As commonly observed, the hinge-zone boundary evolves from an irregular continental rupture to a rather linear mid-oceanic rift system.

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Fig. 3. Plate reconstruction of the North Atlantic and Bay of Biscay (BB) at chron M0 (118 Ma) keeping Eurasia (EU) fixed. Bouguer gravity anomalies are in colour. The 200 m-interval bathymetric contours shown in black are extracted from the bathymetric map of the NE Atlantic (Sibuet et al. 2004a) and from Sandwell & Smith (1994) for the rest of the map. Magnetic anomaly picks of the M sequence between Iberia and North America are from Srivastava et al. (1988, 2000) and Sibuet et al. (2004b). Black dots are IPOD and ODP holes. FC, Flemish Cap; FPB, Flemish Pass Basin; GB, Galicia Bank; GS, Goban Spur; IAP, Iberian Abyssal Plain; JAB, Jeanne d’Arc Basin;

EARLY CRETACEOUS ROTATION OF FLEMISH CAP

Thus, for the same conjugate segment, the distance between chron M0 and the two hinge zones could vary along strike. For example, this distance is constant south of the IAP–NNFB segment but is different to the north of this segment. In conclusion, south of Galicia Bank and Flemish Cap, our proposed IB–NA M25–M0 kinematic pattern, based primarily on Bouguer gravity anomalies, seems to be coherent.

Flemish Cap and Galicia Bank motions The relative motion between the GB and FC conjugate segments is a more complex problem as it also involves additional interactions with the neighbouring larger plates. While the NA and IB plates were being rifted apart in the south, the FC and GB plates remained attached to each other. This required the crust landward of these features to undergo stretching, forming extensive and deep sedimentary basins in these regions, namely the Orphan and Galicia Interior basins. Ultimately, the NA and IB blocks were rifted apart, resulting in the formation of transitional and oceanic crusts between them. From the Bouguer gravity anomaly (Fig. 3) and using the same criteria as before, we have delineated the locations of the main hinge zones associated with the Galicia Interior and Orphan basins. Additional hinge zones have been delineated along the continental margins west of Galicia Bank and east of Flemish Cap. However, the hinge zone located on the eastern side of Galicia Bank cannot be clearly defined because the Moho gradient is too small (Pe´rezGussinye´ et al. 2003). In general, the motions between various plates, which for simplicity can be regarded as microplates, can only be inferred from geological constraints because there are no oceanic magnetic anomalies between these plates to constrain their motion. However, by using four independent geophysical constraints, we can uniquely define the FC– NA motion during the M25–M0 period. These constraints are as follows. (1) Because the Galicia Interior Basin seems to have approximately the same width along its whole length, we assume that the Galicia Bank is acting as a microplate that did not rotate with respect to IB (or that the GB– IB pole of rotation was distally located) during Late Jurassic –Valanginian extension (Murillas et al. 1990). Assuming that the Galicia Interior Basin was stretched between the Iberia continental

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shelf and Galicia Bank sensu stricto, we have calculated 105 km of crustal extension (yellow arrows in Fig. 3) from the geometry of the crust as seen in seismic reflection and refraction data (Murillas et al. 1990; Pe´rez-Gussinye´ et al. 2003). (2) Because the hinge zones east of the SE Flemish Cap and west of Galicia Bank (white lines in Fig. 3) have approximately the same strike length, but make a 438 angle between them, it suggests that the FC plate has rotated by this amount relative to GB (or IB). Prior to Late Jurassic–Early Cretaceous rifting, the hinge zones were parallel, similar to the hinge zones between NA and IB to the south. (3) Such a motion of the FC and GB plates would require that the total extent of the crust formed or thinned between them should be about the same as between the hinge zones to the south (440 km). This would suggest the FC rotated by 438 relative to the GB (or IB) and moved by 335 km (440 –105 km) with respect to the GB. (4) The position of Flemish Cap at chron M25 is constrained in the north–south direction by the geomorphologies and geometries of the Orphan Basin and the Orphan Knoll. Thus, the resulting motion of FC with respect to NA during the M25–M0 period can be described by the following parameters of rotation (pole at 46.178N, 49.098W, 2438). The location of FC with respect to NA can thus be uniquely defined at the time of chrons M0 and M25. The consequences of such a motion concerning the timing, amount and direction of extension in the Orphan Basin can be determined since chron M25 and compared with independent geological constraints.

Reconstruction of Flemish Cap with respect to North America at chron M25 time Figure 3 shows the initial position of Flemish Cap (dashed black lines) with respect to North America at chron M25 and its M25– M0 motion (yellow arrows). At chron M25, Flemish Cap was located SW of Orphan Knoll and between two large, east –west negative magnetic anomalies (Fig. 4), which are not connected to the well-known curved magnetic anomalies located farther west on the continental platform. In addition, the northern

Fig. 3. (Continued) NNFB and SNFB, north and south Newfoundland basins; GIB, Galicia Interior Basin; EOB and WOB, East and West Orphan basins; OK, Orphan Knoll; PB, Porcupine Bank; PSB, Porcupine –Seabight Basin; RT, Rockall Trough; TAP, Tagus Abyssal Plain. Explanation of poles of rotation and other symbols as in Figure 1. Solid pink lines are boundaries of basins. The large continuous white lines represent major gradients of the Bouguer gravity and the large continuous black lines the flow lines between IB and EU from Late Jurassic to Early Aptian (approximately M25–M0) deduced from the morphology and crustal Bouguer anomalies. Yellow arrows give the FC– NA and GB– IB motions.

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Fig. 4. Plate reconstruction of the North Atlantic and Bay of Biscay at chron M0 (118 Ma) with EU kept fixed. Shown in colour are the magnetic anomalies extracted from Verhoef et al. (1996). The 200 m-bathymetric contours are from Sandwell & Smith (1994) and Sibuet et al. (2004a). Explanation of poles of rotation and symbols as in Figures 1 & 2.

EARLY CRETACEOUS ROTATION OF FLEMISH CAP

M25–M0 flow line closely matches the EU– NA plate boundary (Figs 3 & 4). It is thus apparent that our interpreted position of Flemish Cap at chron M25 does not violate geophysical data presented in recent studies (e.g. Enachescu et al. 2004a–c; Skogseid et al. 2004). Furthermore, the position of FC at M25 time and its motion with respect to NA during the M25–M0 period compares favourably with geological data obtained by the oil industry in the Orphan Basin, as described below. Another major implication is that FC is bounded to the NE and SE by major shear zones during its motion with respect to NA and EU. Thus, for example, the Flemish Pass Basin and the Beothuk Ridge (Fig. 5) should have been formed in a transtensional regime.

Formation of the Orphan Basin by the SE motion of Flemish Cap The compilation of over 30 000 km of seismic data acquired by Geophysical Seismic Incorporated (GSI) in the Orphan Basin shows the presence of two central rift zones (Enachescu et al. 2004a–c; Skogseid et al. 2004) characterized by thin, stretched continental crust without obvious volcanic intrusions or high velocity layers (7.2 km s21) at the base of the crust (Chian et al. 2001) (Fig. 5a). Seismic and stratigraphic relationships indicate that the first episode of rifting in this basin could be as old as Late Triassic–Early Jurassic, based on the rifting age established from seismic correlations and well data in the Jeanne d’Arc Basin (e.g. Enachescu et al. 2004a, c), creating a deep trough (East Orphan Basin) immediately landward of Flemish Cap–Orphan Knoll ridge. A second welldated rifting phase occurred from Late Jurassic – Early Cretaceous creating the West Orphan Basin (Fig. 5a), located NW of the East Orphan Basin, but also reactivating the sedimentary features of the East Orphan Basin (Enachescu et al. 2004a–c; Skogseid et al. 2004). Further to the west, a third rifting phase occurred during Late Cretaceous. The Orphan Basin was thus created during three rifting phases sequentially from east to west. Figure 5b shows an interpreted seismic profile, which traverses the entire Orphan Basin from Flemish Cap to the Bonavista platform. From east to west, it shows that the East Orphan Basin is limited to the west by the White Sail bounding fault (Fig. 5b). Above the Palaeozoic basement, Early Jurassic, Triassic or older sediments were deformed during the first Late Triassic –Early Jurassic rifting phase. Here the structural trends are oriented north –south. The second Late Jurassic – Early Cretaceous rifting phase is also recorded in the East Orphan Basin, but to a lesser degree

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compared with rifting further to the west. This second rifting phase took place essentially in the eastern part of the West Orphan Basin and the oldest sediments are Early Cretaceous in age. The strike of tilted fault blocks shows a N0208 orientation. The third rifting phase occurred during the Late Cretaceous in the westernmost part of the West Orphan Basin. In the West Orphan Basin the basement corresponds to the top of prerift sedimentary units, which might be Palaeozoic, as indicated in Figure 5b. The trends of the main basement ridges and troughs (Enachescu et al. 2004a, c) are summarized in Figure 6, which shows that ridges correspond to N0208 linear tilted faults blocks (e.g. Chian et al. 2001). Relative positive magnetic anomalies are associated with the crest of these tilted blocks (Chian et al. 2001), the trend of which is perpendicular to the flow lines of our proposed FC–NA M25–M0 motion, suggesting that the West Orphan Basin was created by the FC –NA rotation and not by the NE motion of Orphan Knoll. In fact, Orphan Knoll is a portion of the Flemish Cap– Porcupine Bank arch feature existing before Late Jurassic, which was left behind at the beginning of FC rotation with respect to NA. One of the consequences of the Early Cretaceous FC microplate motion with respect to EU and NA is that shear zones would have developed on the NE and SW sides of Flemish Cap. A series of basement ridges, elongated and aligned in the NW–SE direction and located east of the Cumberland Ridge (Fig. 5), might be the topographic expression of this FC– NA motion. Similarly, a preliminary inspection of old industrial seismic reflection lines collected across the continental margin SE of Orphan Knoll (Fig. 5) shows elongated basement ridges parallel to the margins, which might be interpreted as the fossil trace of the FC–EU shear zone. For the M0 reconstruction, the East Orphan Basin is adjacent to the Porcupine–Seabight Basin and the West Orphan Basin is adjacent to Rockall Trough (Figs 3 & 6). Two phases of rifting are recorded in the Porcupine–Seabight Basin and are dated as Early Jurassic and Late Jurassic–Early Cretaceous (Dore´ et al. 1997; Tate et al. 1993), identical to that of the East Orphan Basin. For the Rockall Trough, the rifting phase is considered to be Late Jurassic–Early Cretaceous (Dore´ et al. 1997; Shannon et al. 1999), as for the eastern part of the West Orphan Basin. For the Labrador Sea Basin, located NW of Orphan Basin, the rifting phase is Late Cretaceous (Enachescu et al. 2004a), as is the rifting age for the westernmost part of the West Orphan Basin. The width of the Rockall Trough, being approximately the same width as the West Orphan Basin (perpendicularly to the M25–M0 flow lines) during the Late Jurassic–Early Cretaceous rifting

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SCHEMATIC GEOLOGICAL CROSS-SECTION ACROSS ORPHAN BASIN

Fig. 5. (a) Structural map of the Orphan Basin established from Geophysical Seismic Incorporated (GSI) seismic reflection data (modified from Enachescu et al. 2004a– c). Green-coloured features and thin black lines in the Orphan Basin are basement ridges and axes of depressions, respectively. Distribution and shapes of other basins (light–dark yellow) from Enachescu et al. (2004c). Basement platform in brown. CGFZ, Charlie-Gibbs Fracture Zone. Black dots are well locations with their names. Dashed lines are oceanic fracture zones lying in the prolongation of continental or thinned continental crust features. (b) Schematic geological cross-section of the Orphan Basin from the Flemish Cap to the Bonavista platform showing that basins become younger westward (Enachescu et al. 2004a– c). The location of the cross-section is defined in the upper panel.

EARLY CRETACEOUS ROTATION OF FLEMISH CAP

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Fig. 6. Plate reconstruction of the North Atlantic Ocean and Bay of Biscay at chron M0 (118 Ma) with EU kept fixed. Explanation of poles of rotations and symbols are as in Figures 1 and 2. Basement trends (thick black lines are lows and thin black lines are highs) in the East (EOB) and West (WOB) Orphan basins from Enachescu et al. (2004a –c). Note that the basement trends are approximately perpendicular to the FC–NA motion. Thick dashed line shows the position of Flemish Cap rotated back to chron M25 time.

phase, suggests that the Porcupine–Seabight and East Orphan basins opened as a single system and remained adjacent to each other during the M25– M0 rifting period. The FC displacement along the flow lines during the M25–M0 period varies from 200 km in the south to 300 km along the EU boundary. The distance between our M25 position of Flemish Cap and the Bonavista platform further to the NW is 200– 250 km. If the region located between the M25 position of Flemish Cap and the Bonavista platform corresponds to the initial non-deformed continental crust, the formation of Orphan and Flemish Pass

basins corresponds to a mean stretching factor of approximately 2, which is similar to the 2–2.5 mean stretching factor given by J. Skogseid (pers. comm., 2005) for the entire Orphan Basin. The Late Triassic –Early Jurassic plate boundary was located along the adjacent Porcupine–Seabight and East Orphan basins. This boundary extended northward through a system of basins connected to the North Sea basins (Dore´ et al. 1997; Shannon et al. 1999). To the south, it extended through the Flemish Pass, Jeanne d’Arc, Horseshoe and Whale basins. During the Late Jurassic –Early Cretaceous rifting phase, this plate boundary

74

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extended from the Vøring Basin, Møre Basin, Faeroe– Shetland Basin, Rockall Trough, the entire Orphan Basin, Flemish Pass, Jeanne d’Arc, Carson and Salar basins, and then mostly between the Iberia margin and the Grands Banks. Thus, the FC–NA motion was controlled by the propagation of rifting through the EU and NA plates, explaining the unexpected diachronous structure of the crust within the Orphan Basin.

Conclusions The main conclusions of this study are as follows. † Using Bouguer gravity anomalies, we have been able to define the hinge zones of the North American and Iberian plates at chron M25 time. From chron M25 to chron M0 time, a single far-field Euler pole of rotation can describe the relative motion between these plates, which includes regions north of the Newfoundland–Gibraltar Fracture Zone and south of Flemish Cap and Galicia Bank. Rifting between Iberia and North America involved not only these two large plates but also two small microplates (the Flemish Cap and Galicia Bank microplates). † The motion of these microplates, which were adjacent to Eurasia, was much more complex than for the larger plates. The motion between the Flemish Cap and Galicia Bank microplates from chron M25 time or older to chron M0 time was complicated owing to the fact that they remained attached to each other for most of the period when regions to the south were rifting apart, as suggested by Srivastava et al. (2000) and others (e.g. Le Pichon et al. 1977). As a result continental regions landward of these segments were subjected to extensional thinning that created the Orphan and Flemish Pass basins on the North American side and the Galicia Interior Basin on the Iberian side, as well as shear transtensional regimes NE and SW of Flemish Cap. † By delineating hinge zones using Bouguer gravity anomalies at the edges of the conjugate margins, we were able to infer that Flemish Cap rotated 438 relative to Galicia Bank and Iberia, and moved 200– 300km SE with respect to North America. † During the M25–M0 period, most of the Orphan and Flemish Pass basins were initiated. The flow lines describing the translational motion of Flemish Cap are orthogonal to the linear trends of the tilted fault blocks, which were mapped by the oil industry from a compilation of seismic reflection data in this region, giving independent support for the location of the FC–NA pole of rotation.

† New seismic reflection data collected by the oil industry in the Orphan Basin show two central rift zones characterized by highly stretched continental crust without volcanic intrusions at the base of crust. Seismic stratal and stratigraphic relationships indicate that the first episode of rifting could be as old as the Late Triassic – Early Jurassic in the East Orphan Basin. A second rifting phase occurred from Late Jurassic to Early Cretaceous giving rise to the West Orphan Basin but also reactivating earlier structures within the East Orphan Basin. A third phase of rifting occurred in the western part of the West Orphan Basin during the Late Cretaceous. † In our M0 reconstruction, the East and West Orphan basins are adjacent to the Porcupine – Seabight Basin and Rockall Trough, respectively. Two phases of rifting are suggested in the Porcupine–Seabight Basin: Late Triassic – Early Jurassic and Late Jurassic–Early Cretaceous, the same as for the East Orphan Basin. For the Rockall Trough, the rifting phase is Late Jurassic –Early Cretaceous, as it is for the West Orphan Basin. The width of the Rockall Trough, similar to the width of the West Orphan basin, and with structural trends perpendicular to the M25–M0 flow lines imply that the Porcupine–Seabight and East Orphan basins opened simultaneously as a single Basin and remained juxtaposed during the M25–M0 motion of FC– NA. Finally, the Late Cretaceous phase of extension recorded in the westernmost part of the West Orphan Basin is also observed in the south Labrador Basin, which is in the northwestward prolongation of the West Orphan Basin.

The GMT software package was used to make some of the figures (Wessel & Smith 1991). We acknowledge numerous fruitful discussions with W. Roest, J. Skogseid and P. Werner, as well as very helpful and constructive reviews provided by K. Louden, B. Tucholke and R. Whitmarsh. Geophysical Seismic Incorporated (GSI) of Calgary is acknowledged for the permission to use their Orphan Basin seismic data.

References B ERA` STEGUI , X., G ARCIA , J. M. & L OSANTOS , M. 1990. Structure and sedimentary evolution of the Organya` basin (Central South Pyrenean Unit) during the lower Cretaceous. Bulletin de la Socie´te´ Ge´ologique de France, 8, 251 –264. B OIS , C., P INET , B. & G ARIEL , O. 1997. The sedimentary cover along the ECORS Bay of Biscay deep seismic reflection profile. A comparison between the Parentis Basin and other European rifts. Bulletin de la Socie´te´ Ge´ologique de France, 171, 143– 165.

EARLY CRETACEOUS ROTATION OF FLEMISH CAP C ASAS , A. M. 1993. Tectonic inversion and basement thrusting in the Cameros Massif (Northern Spain). Geodinamica Acta, 6, 202–216. C HIAN , D., L OUDEN , K. E., M INSHULL , T. A. & W HITMARSH , R. B. 1999. Deep structure of the ocean–continent transition in the southern Iberia Abyssal Plain from seismic refraction profiles: 1. Ocean Drilling Program (Legs 149 and 173) transect. Journal of Geophysical Research, 104, B4, 7443–7462, doi:10.1029/1999JB900004. C HIAN , D., R EID , I. D. & J ACKSON , H. R. 2001. Crustal structure beneath Orphan Basin and implications for nonvolcanic continental rifting. Journal of Geophysical Research, 106, B6, 10,923–10,940. D EAN , S. M., M INSHULL , T. A., W HITMARSH , R. B. & L OUDEN , K. E. 2000. Deep structure of the ocean– continent transition in the southern Iberia Abyssal Plain from seismic refraction profiles: The IAM-9 transect at 408200 N. Journal of Geophysical Research, 105, B3, 5859– 5886, doi:10.1029/1999JB900301. D ORE´ , A. G., L UNDIN , E. R., F ICHLER , C. & O LESEN , O. 1997. Patterns of basement structure and reactivation along the NE Atlantic margin. Journal of the Geological Society, London, 154, 85– 92. E NACHESCU , M., H OGG , J. & M EYER , K. 2004a. East Orphan Basin, offshore Newfoundland and Labrador: A deep water super extended rift with potential petroleum system. In: CSPG Annual Convention, expanded abstract. E NACHESCU , M. ET AL . 2004b. Orphan Basin, offshore Newfoundland, Canada: Structural and tectonic framework, petroleum systems and exploration potential. In: 74th SEG Annual Meeting and Exposition, expanded abstract. E NACHESCU , M., M EYER , K. & H OGG , J. 2004c. East Orphan Basin: Structural setting and evolution with seismic and potential fields arguments. In: CSEG Annual Convention, expanded abstract. F UNCK , T., H OPPER , J. R., L ARSEN , H. C., L OUDEN , K. E., T UCHOLKE , B. E. & H OLBROOK , W. S. 2003. Crustal structure of the ocean–continent transition at Flemish Cap: Seismic refraction results. Journal of Geophysical Research, 108, B11, 2531, doi:10.1029/2003JB002434. H ENNING , A. T., S AWYER , D. S. & T EMPLETON , D. C. 2004. Exhumed upper mantle within the ocean–continent transition on the northern West Iberia margin: Evidence from prestack depth migration and total tectonic subsidence analyses. Journal of Geophysical Research, 109, B05103, doi:10.1029/2003JB002526. H OPPER , J. R., F UNCK , T. & T UCHOLKE , B. E. 2004. Continental breakup and the onset of ultraslow sea floor spreading off Flemish Cap on the Newfoundland rifted margin. Geology, 32, 93–96. IOC, IHO AND BODC. 2003. Centenary Edition of the GEBCO Digital Atlas. GEBCO Digital Atlas (GDA). British Oceanographic Data Centre, Liverpool. K ARNER , G. D. & W ATTS , A. B. 1982. On isostasy at Atlantic-type continental margins. Journal of Geophysical Research, 87, 2923– 2948. K ARNER , G. D., D RISCOLL , N. W., M C G INNIS , J. P., B RUMBAUGH , W. D. & C AMERON , N. R. 1997. Tectonic significance of syn-rift sediment packages across the Gabon-Cabinda continental margin. Marine and Petroleum Geology, 14, 973–1000.

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K ENT , D. V. & G RADSTEIN , F. M. 1986. A Jurassic to Recent chronology. In: V OGT , P. R. & T UCHOLKE , B. E. (eds) The Geology of North America, Volume M, The Western North Atlantic Region. Geological Society of America, Boulder, CO, 45– 50. L AU , K. W. H., L OUDEN , K. E., F UNCK , T. ET AL . 2003. Seismic velocity model across the eastern Newfoundland (Grand Banks) continental margin. In: 38th Annual Meeting of the Geological Society of America, Abstracts with Programs, Northeastern Section. L E P ICHON , X. & S IBUET , J.-C. 1981. Passive margins: a model of formation. Journal of Geophysical Research, 86, 3708–3710. L E P ICHON , X., S IBUET , J.-C. & F RANCHETEAU , J. 1977. The fit of continent around the North Atlantic Ocean. Tectonophysics, 38, 169–209. L OUDEN , K. E. & L AU , H. 2004. Anomalous heat flow and basement depth in the Newfoundland Basin ocean–continent transition compared with the Iberia Abyssal Plain conjugate. Eos transactions, AGU, 85, 461. Washington, DC. L OUDEN , K. E., S IBUET , J.-C. & H ARMEGNIES , F. 1997. Variations in heat flow across the ocean–continent transition in the Iberia Abyssal Plain. Earth and Planetary Science Letters, 151, 233– 254. M ASSE , P. 1997. The early Cretaceous Parentis basin (France), a basin associated with a wrench fault. Bulletin de la Socie´te´ Ge´ologique de France, 171, 177–185. M ONTADERT , L., DE C HARPAL , O., R OBERTS , D., G UENNOC , P. & S IBUET , J.-C. 1979. Northeast Atlantic passive continental margins: Rifting and subsidence processes. In: T ALWANI , M., H AY , W. & R YAN , W. B. F. (eds) Deep Drilling Results in the Atlantic Ocean: Continental Margins and Palaeoenvironments. American Geophysical Union, Washington, DC, 154– 186. M URILLAS , J., M OUGENOT , D., B OILLOT , G., C OMAS , M. C., B ANDA , E. & M AUFFRET , A. 1990. Structure and evolution of the Galicia Interior basin (Atlantic western Iberian continental margin). Tectonophysics, 184, 297–319. P E´ REZ -G USSINYE´ , M., R ANERO , C. R. & R ESTON , T. J. 2003. Mechanisms of extension at nonvolcanic margins: Evidence from the Galicia Interior basin, west of Iberia. Journal of Geophysical Research, 108, B5, 2245, doi:10.1029/2001JB000901. R USSELL , S. M. & W HITMARSH , R. B. 2003. Magmatism at the west Iberia non-volcanic rifted continental margin: Evidence from analyses of magnetic anomalies. Geophysical Journal International, 154, 706– 730. S ALAS , R., G UIMERA , J., M AS , R., M ARTIN -C LOSAS , C., M ELE´ NDEZ , A. & A LONSO , A. 2001. Evolution of the Mesozoic Central Iberian Rift system and its Cainozoic inversion. In: Z IEGLER , P. A., C AVAZZA , W., R OBERTSON , A. H. F. & C RASQUIN -S OLEAU , S. (eds) Peri-Tethyan Rift/Wrench Basins and Passive Margins. Me´moires du Muse´um National d’Histoire Naturelle, Paris, 145– 185. S ANDWELL , D. T. & S MITH , W. H. F. 1994. New global marine gravity map/grid based on stacked ERS1, Geosat and Topex altimetry. Eos, 75, 321. S HANNON , P. M., J ACOB , A. W. B., O’ REILLY , B. M., H AUSER , F., R EADMAN , P. W. & M AKRIS , J. 1999.

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Structural setting, geological development and basin modelling in the Rockall Trough. In: F LEET , A. J. & B OLDY , S. A. R. (eds) Petroleum Geology of NW Europe: Proceedings of the 5th Conference. Geological Society, London, 421–431. SHIPBOARD SCIENTIFIC PARTY . 2004. Leg 210 summary. In: T UCHOLKE , B. E., S IBUET , J.-C. & K LAUS , A. (eds) Proceedings of the Ocean Drilling Program, Initial Reports, 210. Ocean Drilling Program, College Station, TX, 1 –78. S IBUET , J.-C., M ONTI , S., L OUBRIEU , B., M AZE´ , J.-P. & S RIVASTAVA , S. 2004a. Carte bathyme´trique de l’Atlantique nord-est et du golfe de Gascogne: implications cine´matiques. Bulletin de la Socie´te´ Ge´ologique de France, 175, 429– 442. S IBUET , J.-C., S RIVASTAVA , S. & S PAKMAN , W. 2004b. Pyrenean orogeny and plate kinematics. Journal of Geophysical Research, B08104, doi:10.1029/2003 JB002514. S IBUET , J.-C., S RIVASTAVA , S. P. & M ANATSCHAL , G. 2007. Exhumed mantle forming transitional crust in the Newfoundland–Iberia rift and associated magnetic anomalies. Journal of Geophysical Research. (in press) S KOGSEID , J., B ARNWELL , A., A ARSETH , E. S., A LSGAARD , P. C., B RISEID , H. C. & Z WACH , C. 2004. Orphan basin: Multiple failed rifting during early opening of the north Atlantic. Eos, Transactions of the American Geophysical Union, 85, 17, Joint Assembly Supplement, Abstract T41B-03. S RIVASTAVA , S. & V ERHOEF , J. 1992. Evolution of Mesozoic sedimentary basins around the north Central Atlantic: a preliminary plate kinematic solution. In: P ARNELL , J. (ed.) Basins on the Atlantic Seaboard: Petroleum Geology, Sedimentology and Basin Evolution. Geological Society, London, Special Publications, 397–420. S RIVASTAVA , S., S IBUET , J.-C., C ANDE , S., R OEST , W. R. & R EID , I. R. 2000. Magnetic evidence for slow sea-floor spreading during the formation of the Newfoundland and Iberian margins. Earth and Planetary Science Letters, 182, 61– 76. S RIVASTAVA , S., V ERHOEF , J. & M ACNAB , R. 1988. Results from a detailed aeromagnetic survey across the northeast Newfoundland margin, Part 1: Spreading anomalies and relationship between magnetic anomalies and the ocean– continent boundary. Marine and Petroleum Geology, 5, 306–323.

T ATE , M., W HITE , N. & C ONROY , J. J. 1993. Lithospheric extension and magmatism in the Porcupine Basin west of Ireland. Journal of Geophysical Research, 98, B8, 13,905–13,924. T UCHOLKE , B. E., S AWYER , D. S. & S IBUET , J.-C. 2007. Breakup of the Newfoundland–Iberia rift. In: K ARNER , G. D., M ANATSCHAL , G. & P INHEIRO , L. M. (eds) Imaging Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 9–45. V ERGE´ S , J. & G ARCI´ A -S ENZ , J. 2001. Mesozoic evolution and Cainozoic inversion of the Pyrenean rift. In: Z IEGLER , P. A., C AVAZZA , W., R OBERTSON , A. H. F. & C RASQUIN -S OLEAU , S. (eds) Peri-Tethyan Rift/Wrench Basins and Passive Margins. Me´moires du Muse´um National d’Histoire Naturelle, Paris, 187–212. V ERHOEF , J., R OEST , W. R., M ACNAB , R. & A RKANI -A HMED , J. Members of the Project Team. 1996. Magnetic Anomalies of the Arctic and North Atlantic Oceans and Adjacent Land Areas. Geological Survey of Canada, Open File, 3125. W ESSEL , P. & S MITH , W. M. F. 1991. Free software helps map and display data. Eos, 72, 441– 446. W HITMARSH , R. B. & S AWYER , D. S. 1996. The ocean/ continent transition between the Iberia Abyssal Plain and continental-rifting to sea floor spreading processes. In: W HITMARSH , R. B., S AWYER , D. S. & K LAUS , A. (eds) Proceedings of the Ocean Drilling Program, Scientific Results. Ocean Drilling Program, College Station, TX, 713– 733. W HITMARSH , R. B., M ANATSCHAL , G. & M INSHULL , T. A. 2001. Evolution of magma-poor continental margins from rifting to sea floor spreading. Nature, 413, 150– 154. W HITMARSH , R. B., W HITE , R. S., H ORSEFIELD , S. J., S IBUET , J.-C., R ECQ , M. & L OUVEL , V. 1996. The ocean–continent boundary off the western continental margin of Iberia – III. Crustal structure west of Galicia Bank. Journal of Geophysical Research, 101, 28,291–28,313. W INNOCK , E. 1971. Ge´ologie succincte du bassin d’Aquitaine. In: D EBYSER , J., L E P ICHON , X. & M ONTADERT , L. (eds) Histoire Structurale du Golfe de Gascogne. Technip, Paris, 1 –30. Z IEGLER , P. A. 1988. Evolution of the Arctic– North Atlantic and the Western Tethys. AAPG Memoirs, 43. AAPG, Tulsa, OK.

The formation of non-volcanic rifted margins by the progressive extension of the lithosphere: the example of the West Iberian margin T. J. RESTON1,2 1

School of Geography, Earth and Environmental Sciences, University of Birmingham, Birmingham B15 2TT, UK 2

Previous address: IFM-GEOMAR, Wischhofstrasse 1-3, D24148 Kiel, Germany (e-mail: [email protected]) Abstract: Non-volcanic margins such as the West Iberian margin exhibit certain characteristics, such as a deficit of synrift igneous rock, a zone of exhumed subcontinental mantle in the continent –ocean transition and an apparent extension discrepancy. These observations can be explained as a consequence of the progressive extension of the lithosphere above relatively cool mantle. The evolving rheological stratification of the lithosphere controls the style of extension at different lithospheric levels at different times; extension is probably heterogeneous at all stages, with lower crustal and upper mantle boudinage controlling the patterns of thinning and mantle upwelling early in the rift history, and complete crustal embrittlement and mantle serpentinization controlling the formation of late-stage detachment faults. Extension in the brittle crust is via multiple phases of faulting, with a general focusing of extension towards the incipient ocean. The lack of melt is explained by a combination of heterogeneous extension of the lower lithosphere and a cool subcontinental geotherm. The extension discrepancy may in places be controlled by depth-dependent stretching of the crust through lower crustal boudinage, but may also simply be the result of incomplete recognition of the entire polyphase faulting history. The latter seems to be the case for West Iberia. Evidence for all these processes can be found at the West Iberian rifted margins as well as those preserved and partially exposed in the Alps.

Rifted margins include two main end-members: those termed ‘volcanic rifted margins – VRMs’ where magmatism is much more voluminous than predicted by passive asthenospheric upwelling (e.g. White & McKenzie 1989), and those where magmatism is consistent or even less than the same predictions. The latter are termed ‘non-volcanic rifted margins – NVRMs’ to emphasize the contrast with the VRMs: the name does not exclude the presence of minor amounts of magmatic activity. The NVRMs are typified by the North Biscay (Le Pichon & Barbier 1987), South Australian (Sayers et al. 2001) and the West Iberian margins (Boillot & Winterer 1988; Krawczyk et al. 1996; Whitmarsh et al. 1996), which share a number of common characteristics (Fig. 1): † low –moderate sediment accumulation rates; † extreme crustal thinning, increasing towards the ocean; † rotated fault blocks – however, at the featheredge of the continent the amount of extension that can be inferred from the geometry of these faults is less than that indicated by the crustal thinning and/or observed: this is the so-called extension discrepancy;

† presence, in places, of a detachment fault at the base of the fault blocks; † little evidence for voluminous synrift magmatism and in general less magmatism than predicted by standard extension models; † the presence of a broad zone of partially serpentinized mantle (Boillot & Winterer 1988; Krawczyk et al. 1996; Pickup et al. 1996; Whitmarsh et al. 1996), both occurring beneath the highly thinned and faulted continental crust of the continental rise region, and as a zone of exhumed continental mantle (ZECM: Whitmarsh et al. 2001), now largely buried by post-rift sediments. I suggest that such margins are the logical result of progressive extension of cool, sediment-starved continental lithosphere away from hotspots. I focus first on the expected evolution before considering the application of this to the West Iberian margin in particular. The key factors are the occurrence of multiple phases of faulting and the rheological evolution of the lithosphere and the way this controls the four-dimensional (4D) distribution of deformation within the lithosphere and the serpentinization of the uppermost mantle.

From: KARNER , G. D., MANATSCHAL , G. & PINHEIRO , L. M. (eds) Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 77–110. DOI: 10.1144/SP282.5 0305-8719/07/$15.00 # The Geological Society of London 2007.

897

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50 km

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6 TWT [s] 10

Fig. 1. Cross-sections at approximately true scale of typical non-volcanic margins, showing the presence of a wide zone of unroofed mantle rocks and the presence of local detachments. Fault block structure is invariably incompletely resolved.

sediments

oceanic crust

SW Greenland margin (Chian & Louden 1994) - no vertical exageration

Ligurian Margin (Alps) Froitzheim + Manatschal 1996

11

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S Australia margin (Sayers et al. 2001)

W. Iberia

78 T. J. RESTON

RIFTED MARGINS BY PROGRESSIVE EXTENSION

Terminology In this paper, I will discuss various aspects of margin evolution, including results from seismic imaging, from wide-angle studies of crustal structure, numerical, thermal and rheological modelling, and structural geology. As the terminology between the different fields is commonly different, it is worth defining some of the terms to be used: it is important that the terminology used is appropriate, not misleading and not too specific for general application. In particular, the description of the different rheologies of the lithosphere can be a source of confusion: whereas structural geologists commonly refer to brittle and plastic deformation (Rutter 1986), meaning failure by faulting (cracking plus sliding) and failure by intracrystalline plastic creep, modellers commonly refer to the same as ‘plastic’ and ‘viscous’, respectively, referring to the stress –strain relationship of faults (no deformation until the yield stress reached) and the ‘creeping’ deep viscous crust. The term ‘plastic’ is used in two contradictory ways by different groups. The use of the term ‘viscous’ for deep crustal deformation may also be a simplification as it implies that the exponent in the stress–strain rate equation is 1, which is not the case for most flow laws. The term ‘flow’ also carries too much baggage in terms of large-scale channel flow to be used to describe rheology, and the term ‘ductile’ also has its critics (e.g. Rutter 1986) who point out that ductile means distributed and includes brittle deformation by cataclastic flow. In this paper I will refer to failure along faults (cracking and sliding according to a Mohr–Coulomb criteria) as brittle, and to deformation by temperature- and lithologydependent intracrystalline plastic creep as ‘creeping’. Some authors use the term ‘stretching factor’ only to imply deformation by pure shear (e.g. Davis & Kusznir 2004). This I consider unhelpful as it may be difficult to determine accurately the exact mode of deformation, whereas it is generally simpler to measure crustal thinning and/or extension. Others describe whole crustal stretching factors, upper and lower crustal stretching factors, and mantle stretching factors. I will use the term ‘extension factor’ to describe the degree of thinning and extension of the whole lithosphere, and, where specified, of the various portions of that lithosphere, regardless of whether this occurs by ‘pure’ pure shear or by some more realistic mixed mode extension. Finally, as will become clear in the paper, the term continent–ocean transition (COT) both describes the process and the resulting structure, and should be preferred to other terms (ocean continent transition – OCT; continent–ocean boundary – COB; ZECM – zone of exhumed continental

79

mantle) which, respectively, have the wrong sense of transition, imply a sharp boundary or which may be too specific for general use.

West Iberia margin The West Iberian margin is the best-studied nonvolcanic rifted margin, having been the subject of part of a Deep Sea Drilling Project (DSDP) leg, three dedicated Ocean Drilling Program (ODP) legs, as well as numerous seismic experiments and submersible diving. Two main segments have been studied (Fig. 2); the Deep Galicia margin and the southern Iberia Abyssal Plain (IAP) margin (Fig. 1).

Deep Galicia margin The northernmost segment of the West Iberian rifted margin is the Deep Galicia margin located to the west of the Galicia Bank (Fig. 2) and conjugate to the Flemish Cap margin (Fig. 3). Comparison between the two conjugate margins shows that whereas crustal thickness reaches ‘normal values’ within approximately 100 km of the COT at the Flemish Cap margin, at the west Galicia margin, the Galicia Bank 100 km east of the COT is still characterized by relatively thin crust (15– 20 km) and is flanked by the Galicia Interior Basin to the east and the Deep Galicia margin to the west. The latter is characterized by well-defined tilted fault blocks of continental material that, west of approximately 128150 W (ODP Site 639), decrease abruptly in size and become underlain by the bright S reflection (Fig. 3). Above S, the amount of extension that can be deduced from the geometry of these fault blocks is considerably less that that implied by the degree of subsidence (Sibuet 1992) and of crustal thinning (Reston 2005). S is thought to be both a detachment fault and a tectonic CMB, forming the boundary between the tilted blocks and the underlying partially serpentinized mantle (Reston et al. 1996). Mantle rocks crop out at the sea floor 25 km to the west of the last sampled continental fault block (other probable continental blocks lie between), and can be traced along strike as a partially buried ‘peridotite ridge’. Synrift igneous rocks have not been sampled within the main drilling transect here, although possible synrift lower crustal gabbros have been sampled by submersible further to the north. Drilling during ODP Leg 103, as well as sampling by submersible and by dredging, has shown that the tilted fault blocks are capped by a deep-water clastic sequence, rather than true prerift sediments (e.g. shallow-water carbonates). The implications of this stratigraphy are discussed at the end of this paper.

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Fig. 2. Map of the West Iberian margin showing the location of the DGM and IAP margins, of ODP holes and of profiles discussed in the text.

join made at peridotite ridge Flemish Cap margin - crust thinning oceanward from ~30 km thick

10 km

12°15'W

serpentinized mantle (base gradational)

Galicia Bank margin - crust thinning west from ~15 km thick CMB

CMB

20 km

W

2400 shotpoint

2600

2800

ODP 639

E

Dive 11

6

6

8

8

S CMB

10

10

serpentinized mantle km 5 km base of postrift section

km early generations of faults

latest generation of faults, including S

base of late synrift (Aptian ?)

base of early synrift - Valanginian - Barremian?

Fig. 3. Deep Galicia margin as part of conjugate pair with Flemish Cap margin. Top: generalized crustal sections at 2:1 vertical exaggeration (after Funck et al. 2003 and Zelt et al. 2003) showing a contrasting crustal thickness of Flemish cap and Galicia Bank. The S reflector in bold may continue as CMB on the Flemish Cap side. Below: Prestack depth migration of profile GP101 shown at no vertical exaggeration (location approximately marked by box), showing latest fault blocks and tentative identification of earlier faults (short dashes). ODP Site 639 straddles 128150 W.

RIFTED MARGINS BY PROGRESSIVE EXTENSION

Iberia Abyssal Plain The southern Iberian Abyssal Plain (IAP) margin has been the subject of two ODP drilling legs (149 and 173) as well as numerous seismic experiments. The basement geology across the feather edge of continental crust of the IAP is well constrained by the presence of several ODP boreholes (Figs 1, 2 & 4). Moving from the continent towards the ocean, Sites 901 and 1065 sampled the top of two tilted fault blocks and recovered shallow-water clastics, believed to indicate the presence of continental basement underneath. Sites 1067, 900 and 1068 sampled the same block and recovered, respectively, anorthosites, amphibolites and tonalites (1067), sheared mafic granulites (900) and serpentinized peridotites (1068). These are interpreted to represent lower crustal rocks (1067 and 900) and the subcrustal mantle (1068). The H detachment (Krawczyk et al. 1996) appears to intersect top basement near Site 900, so that the CMB here might be a tectonic boundary similar to that inferred at the West Galicia margin further north in the region of S. Site 1069 sampled a fault slice that appears to be resting on the same peridotitic basement sampled at Site 1068. As continental sediments were recovered,

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it is thought that this fault slice is allochthonous (see below). Further to the west still, Site 899 recovered serpentinite breccias intercalated with sediments, perhaps indicating the presence of a nearby serpentinite high of some form (see below), and 897 recovered highly sheared partially serpentinized continental mantle peridotites. Finally, Site 1070, located just 15 km east of the J anomaly (unequivocal oceanic crust), recovered serpentinized mantle rocks and a lesser amount of gabbro. Although the main focus of the ODP drilling was the basement type, the south IAP drilling transect (legs 149 and 173) does constrain certain aspects of rift history. Tithonian clastics tilted within the fault blocks at sites 1065 and 901 were deposited in enclosed, relatively shallow and anoxic basins, close to vegetated land. These can be interpreted as the earliest synrift units, deposited before substantial extension had occurred (Whitmarsh & Wallace 2001). Later synrift units have, however, not been sampled, but the timing of crustal extension is constrained by Ar/Ar cooling data for samples from sites 1067 and 900, indicating that these deep crustal rocks were exhumed from more than 18 km (0.6 GPa) to approximately 4 km depth by 137 Ma (Wilson et al. 2001). If this

Fig. 4. Prestack depth migration of critical portion of profile LG12 (Krawczyk et al. 1996) and a structural interpretation. Note that early synrift sediments sampled at Sites 901 and 1065 are not obviously present in the intervening blocks, the presence of structures within the basement, interpreted here as segments of earlier faults, and the kink in the fault scarp west of Site 901, interpreted as two intersecting fault planes. H is a probable detachment fault at the top of serpentinized mantle. The location of the seismic on the transect is shown at the top: note that the top basement does not rise (and hence crust does not thicken) appreciably for over 50 km to the east.

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degree of exhumation applied to the whole crust, then in approximately 10 Ma between the onset of rifting in the Tithonian and the unroofing of these rocks, the crust would have been thinned by a factor of approximately 5. The final phase of crustal extension that produced the current geometries observed probably occurred shortly after this. Between the first oceanic crust, marked by oceanic magnetic anomalies, and the highly extended continental crust is a 120 km-wide zone of unroofed mantle, in which a series of basement highs were drilled (1068, 899, 897 and 1070). The sediment atop the highs represents a minimum age for the unroofing of that particular portion of peridotitic basement and is characterized by mixtures of serpentine breccias and mass flow deposits of Aptian–Barremian age. Intriguingly, the oldest of these deposits corresponds to the age of the first spreading anomaly (M3) observed further to the west, which can be taken as the onset of sea-floor spreading (base of the Barremian at 127 Ma according to Gradstein et al. 1994). The mass flow deposits may be related to the breakup process (Tucholke et al. 2007). To summarize, in the southern Iberia Abyssal Plain the evidence points towards major rifting and crustal extension occurring over perhaps 10 million years between the Tithonian and the Valanginian, followed by a phase of mantle unroofing over a further 10 million years before the onset of sea-floor spreading. As with the Galicia Bank margin, extension appears to have migrated oceanwards, but culminating in the unroofing of a far wider zone of mantle rocks. Although the overall time between the onset of rifting and of sea-floor spreading is similar at the two margin segments, approximately half of that time at the IAP margin only involved the unroofing and subsequent extension of mantle rocks. Hence, the rift duration leading to complete crustal separation may have only been about 10 million years. Both these margin segments thus exhibit tilted fault blocks, detachment faulting, serpentinization beneath the highly attenuated crust and a zone of mantle exposed at the sea floor during rifting. In this paper I will discuss how such margins might have formed, describing the structures that might be expected in each case, then comparing them with the observations from West Iberia.

Thermal and rheological evolution The thermal structure of the lithosphere is a major control on the rheology of the lithosphere, and hence an important consideration during progressive extension. A common misconception is that rifting leads to heating, and hence to thermal weakening of the rocks above a developing rift. Although rapid pure shear stretching (McKenzie 1978) leads to the steepening of the geotherm and increased

heat flow as hot rocks towards the base of the lithosphere are brought towards the surface, no actual rock gets hotter during uniform stretching at sediment-starved margins. Hot rocks are simply brought nearer the surface. Instead, adiabatic decompression means that even for instantaneous stretching the temperature of an individual rock actually drops marginally. For more realistic finite duration stretching (Jarvis & McKenzie 1980), some cooling occurs during rifting, leading to a cooling of those deep lithospheric rocks that have been transported towards the surface by the process of thinning, and consequently a less pronounced steepening of the geotherm.

Embrittlement Only where simple shear or strongly heterogeneous stretching occurs or where the margin is blanketed by thick sequences of thermally insulating sediment (e.g. Allen & Allen 1990) is some heating possible. For instance, some heating occurs in the hanging wall of a low-angle shear zone as heat from the exhumed footwall is transferred across the shear zone (e.g. Voorhoeve & Houseman 1988), but heterogeneous stretching is local in nature and is balanced by excessive cooling elsewhere (e.g. of the footwall in simple shear models). A rheological effect of extension more immediate than cooling is the pressure reduction at depth that results from lithospheric thinning. Decreasing the pressure leads to a reduction in the resistance to frictional sliding and fracturing, i.e. to brittle failure. Any simultaneous decrease in the temperature (due to cooling during slow rifting) serves to increase the resistance to creeping (e.g. dislocation creep). Consequently, rocks which once preferentially deformed by creep start to deform by brittle failure during the rifting process (Pe´rez-Gussinye´ & Reston 2001). For lithospheric parameters suitable for the West Iberia margin, Pe´rez-Gussinye´ & Reston (2001) showed that as extension factors increased, the upper creeping layer thinned, and the upper brittle –creeping transition shallowed until the entire upper and mid-crust became brittle at extension factors slightly below 3 (Fig. 5). If the spacing of faults is controlled by the thickness of the brittle crust (e.g. Scholz 1990), this implies that the spacing between new faults should decrease until extension factors approach 3. However, once the upper creeping layer has become entirely brittle, the spacing of new faults may increase again as they cut from the surface down to the deeper creeping layer. Note, however, that although the thickness of the crust prior to rifting may be reasonably well constrained, the boundary between the upper and lower crust is less well resolved.

RIFTED MARGINS BY PROGRESSIVE EXTENSION

0 10

0

yield stress (MPa) 200 400 600

(b) 0

middle and lower crust creep

WQ

Depth [km]

(a)

20 Depth [km]

AN 30 40 50

strong mantle lid

0

OL

120

10

m

110

r the geo

80

100

20

peridotites toward base lithosphere (here below ~ 80 km) are very weak and may be able to flow

200

400

OL

30

t = 6.5 m.y. β = 2.8

40

70

90

yield stress (MPa) 200 400 600

Middle crust brittle. Only lowermost crust creeps 10 strong mantle 20 lid

(c) 60

0

600

800

0

yield stress (MPa) 200 400 600 Whole crust now brittle OL t = 8.1 m.y. β = 3.6

30

T (°C)

original base of lithosphere

Fig. 5. The 1D rheological evolution of the lithosphere during pure shear, based on a model appropriate for west Iberia (see Pe´rez-Gussinye´ & Reston 2001 and references therein for the flow laws used: WQ, wet quartz; AN, anorthosite; OL, olivine) for a strain rate of 5  10215 s21. Base original lithosphere is a bold dashed line – lower lithosphere is very weak and may flow. (a) Starting model (geotherm shown by a dashed line), crust contains creeping zones within the mid crust and lower crust; (b) as extension proceeds, the mid-crust creeping zone becomes entirely brittle; (c) lower crust now also entirely brittle, faults cut down into the mantle. Note that lower crustal flow is not required for the entire crust to become brittle.

Furthermore, Pe´rez-Gussinye´ & Reston (2001) showed that for a West Iberia model, the entire crust should become brittle at extension factors of about 4 (Fig. 6), depending on rift duration and the rheology of the lower crust. Complete crustal embrittlement has two important consequences for lithospheric extension. The first is the disappearance of any weak decoupling zone within the middle or lower crust, allowing tighter coupling between the crust and the mantle. The second is the serpentinization of the uppermost mantle.

Coupling and lithospheric-scale simple shear The initial lithospheric strength profile in Figure 2 shows creeping zones both at the base of the upper crust and at the base of the lower crust. In these zones, strength decreases downwards so if the zones are thick enough, their lower portions may represent significant weak zones within the lithosphere.

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The pattern of decoupling predicted is somewhat more complex than has been modelled in most studies that include strain weakening (e.g. Huismans & Beaumont 2003). In particular, by having a strong lower crust bound above and below by weak zones, the upper crust, lower crust and uppermost mantle may to some extent deform independently (Coward 1986; Reston 1990). As the overall resistance to shearing in these weak zones increases with distance between the loci of upper crustal and mantle extension, large lateral offsets of these loci can be ruled out (Braun & Beaumont 1987), but local decoupling and differences in the loci and style of deformation is likely. In particular, the relatively strong top layer of the lower crust may deform initially by boudinage, leading to a degree of depth-dependent stretching within the crust, discussed more below. Low strain lozenges of lower crust have been observed within deep crustal high strain zones (e.g. see Reston 1990), and inferred from seismic reflection data (Reston 1988). However, during progressive extension, the mid-crustal weak zone gradually disappears and the top of the lower crust becomes ever more tightly coupled to the upper crust. Complete coupling between the upper and lower crust is likely before the entire mid-crust becomes brittle at about 150% extension, and the location of the lower crustal boudins and shear zones may well control the development of second-generation faults at this time (see below). This might limit further depth-dependent stretching within the crust. Further extension will also couple the crust to the mantle, as the base crustal creeping zone also becomes brittle. At this stage, the entire upper lithosphere should be strongly coupled and resemble the coupled models of Huismans & Beaumont (2003), which (depending on strain rate and strain weakening) can develop strong asymmetries that resemble lithospheric simple shear.

Serpentinization The second effect of complete crustal embrittlement is the formation of mantle serpentinites. Serpentinized mantle forming both top basement within the COT (e.g. Pickup et al. 1996) and underlying the feather edge of the crust (Boillot et al. 1989) are important characteristics of NVRMs. The presence of serpentinites beneath the thinned continental crust implies that water passes down through the crust to react with the underlying mantle, requiring (O’Reilly et al. 1996; Pe´rez-Gussinye´ & Reston 2001) that the entire crust is brittle: fluid flow in sufficient volumes only occurs along faults and fractures. Thus, only when the entire crust has become brittle (extension .300%) should serpentinization first occur by the passage of water down

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Rift duration [Ma]

7 βb

30

6

0

5

4

1

3 10

IAP

Alps

IAP

DGM

Labrador Sea

20

2

3

4

5

6

2 7 8

2 3 9 10 11

β Fig. 6. Plot in rift duration v. extension factor space of crustal embrittlement curve (bb – bold line), maximum thickness of the serpentinizing zone (thin solid contours – brittle mantle thickness between the CMB and the 500 8C isotherm assuming pure shear) and melt thickness (dashed contours after Bown & White 1995). See Pe´rez-Gussinye´ & Reston (2001) for details. DGM, Deep Galicia margin; IAP, southern Iberia Abyssal Plain margin; Lab, Labrador margins; Alps, rifted margins exposed in Alps. Note the maximum extension factor (b) plotted is 12. For melt thickness at b ¼ 50 see Figure 8.

through the brittle crust along active faults (Pe´rezGussinye´ & Reston 2001). This prediction agrees well with the thickness of the crust at the landward limit of the underlying partially serpentinized peridotites at the West Iberian margin (Pe´rez-Gussinye´ & Reston 2001). Serpentinization leads to a dramatic drop in the strength of the uppermost mantle, leading to the development of serpentinite detachment systems, a different style of decoupling, at the base of the crust. Unlike the decoupling by the creeping zones discussed above, the continuity of brittle faulting passing down from the crust into the mantle may also lead to continuity of structure even if parts of the brittle fault zone are weakened by serpentinization. Further extension is thus likely to still resemble asymmetric simple shear, leading to the complete separation of the crust along the brittle faults and the final unroofing of a wide region of mantle. The structural dip and sense of shear deduced for both the S and H detachments (Krawczyk et al. 1996; Reston et al. 1996; Manatschal et al. 2001) is consistent with these

structures contributing to the unroofing of mantle rocks observed oceanward of each. A similar detachment (P) has been observed beneath the Porcupine Basin (west of Ireland), where it cuts from the eastern flank of the basin, beneath the centre of the basin (where the crust beneath the post-rift appears to be only about 3 km thick) and cuts to depth under the western flank of the basin (Reston et al. 2004). This is of interest as it is an asymmetric structure, implying that equivalent mirror image structures of S and H need not lie under the conjugate Newfoundland margins. In summary (Fig. 7), thinning results in rheological changes within the crust, with the lower crust becoming progressively less prone to creep and more brittle. The entire crust becomes brittle at extension factors in excess of about 4, when water can reach the mantle, leading to serpentinization. The weak hydrated mantle rocks formed at the base of the fault blocks will allow the formation of detachments near the base of the crust, perhaps accompanied by a change from pure shear to simple shear and the unroofing of mantle rocks.

RIFTED MARGINS BY PROGRESSIVE EXTENSION

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Depth (km)

TIME / CONTINUING EXTENSION mid-crust brittle

0

entire crust brittle - seawater down to mantle

crustal separation

t

detachmen

10

Mantle exhumation

20 30 brittle crust

plastic crust

Mantle

Serpentinized peridotite

top lower crust

upper brittle-plastic transition

Detachment at top serpentinized mantle (e.g. H or S)

Fig. 7. Cartoon showing the effect of rheological evolution during progressive extension. Note the gradual embrittlement of the crust, the changing depth of detachment of the main faults (dashed line) from mid to deep crust, and the development of serpentinite detachments leading to final crustal separation once the entire crust has become brittle. Modified from Pe´rez-Gussinye´ & Reston (2001).

An important question to address when considering rifted margins is why some are characterized by voluminous magmatism and others by very little, if any. From studies of mid-ocean ridges it has been inferred (McKenzie & Bickle 1988) that the ‘normal’ temperature of the asthenosphere had it had been brought instantaneously (adiabatically) to the surface would have been approximately 1300 8C (the potential temperature, Tp). During rifting, the upwelling asthenosphere is likely to have undergone partial melting when its temperature exceeded the mantle solidus. When this occurred would depend on three main factors: the initial thickness of the lithosphere, the rate and amount of asthenospheric upwelling, and the initial temperature of the asthenosphere. Various authors (e.g. Bown & White 1995; Pe´rez-Gussinye´ et al. 2001) have calculated in 1D models the point at which the geotherm would intersect the solidus (the onset of melting) and also the melt thickness generated. Although using different parameterizations, these studies find that, depending on the rift duration, significant melting is not expected during the evolution of a margin similar to West Iberia until b . 5, where b is the extension factor (Fig. 6) and that total melt thickness is unlikely to exceed 4 km until after continental breakup (Fig. 8). This is, however, considerably more than is thought to have occurred at the West Iberia margin (e.g. Minshull et al. 2001; Reston & Phipps Morgan 2004) (Fig. 8). The same applies to the Alpine margins and other less wellconstrained examples (Reston & Phipps Morgan 2004). Similar results are generally obtained in

those 2D modelling studies where melt production has been explicitly included, e.g. Bowling & Harry (2001) showed that approximately 5 km of melt are expected before crustal separation can take place for a mantle potential temperature of 1300 8C. Thus, there is a general observation that at NVRMs, the amount of melt observed is less than that predicted by pure shear models for the

6

Melt thickness (km)

Melting and mantle unroofing

5

13

00

4 12

50

3

12

2

00

1

Alps

IAP 0 0

10

20

30

Rift duration (m. a.) Fig. 8. Melt thickness expected for various rift durations at an extension factor of 50 for various asthenospheric potential temperatures (after Minshull et al. 2001 using parameterization of Bown & White 1995). The thickness observed at the IAP (Minshull et al. 2001) and Alpine margins (Mu¨ntener & Herrmann 2001) is consistent with cool sublithospheric mantle.

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extension factor derived from crustal thinning. This, I term the melt discrepancy.

Temperature (°C) 400

600

800

1000

1200

1400

1600

Cool subcontinental geotherm Geotherms from seismic and heat flow Global: Archaean craton, stable platform, tectonic province

200

Eastern North America average western U.S. average

300

Kaapvaal kimberlite geotherms

unheated

Depth (km)

100

“plume-heated” at base lithosphere

mantle adiabats 400 1300

1100

Potential temperature Tp (°C)

1200

One solution is that the paradigm of a 1300 8C asthenosphere may be incorrect (Reston & Phipps Morgan 2004). The thickness of melt at volcanic margins is far greater, and is generally attributed to a Tp greater than 1300 8C in the upwelling mantle (e.g. a plume): could the lack of magmatism at nonvolcanic margins imply that elsewhere Tp beneath the lithosphere is less than 1300 8C? If the mantle beneath the continental lithosphere is even slightly cooler (Tp c. 1250 or 1200 8C, for instance), the volume of melt produced at effectively infinite extension factors by partial melting of dry peridotite is considerably reduced (Fig. 8), although small amounts of wet melting will take place. Some variation in sublithospheric temperature is, of course, to be expected if the asthenosphere is convecting (Lenardic & Moresi 2000) – convection is, after all, driven by temperature variations. Surprisingly little is known about the temperature structure beneath the base of the continental lithosphere, an issue complicated by the different definitions of where this base lies. In the following discussion, we concentrate on evidence of the temperature structure beneath the thermal lithosphere, i.e. in the presumably convecting region beneath the outer thermal boundary layer (Fig. 9). The first evidence comes from cratonic regions (Reston & Phipps Morgan 2004) via geothermobarometry on xenoliths brought up from the base of the lithosphere by kimberlites (e.g. Kohler & Brey 1990). The xenoliths occur in two families – sheared xenoliths heated at the base of the lithosphere by the mantle plume that initiated kimberlite magmatism and unsheared xenoliths from the entire lithospheric column. The former exhibit a Tp range between approximately 1200 and 1400 8C at a largely constant depth of about 200 km, indicating both the depth of the base of the lithosphere and the temperature range between unheated and plume-heated xenoliths there. In contrast, the unsheared xenoliths show that away from plume heating, the potential temperature at the base of the lithosphere was generally equal to or below a potential temperature of 1200 8C, and thus that the convecting subcratonic mantle has a Tp well below 1300 8C. Unless this differs from other subcontinental convecting mantle, the implication may be that this is also relatively cool. More generally, the temperature beneath the subcontinental lithosphere can be deduced from the depth to and thickness of the mantle transition zone (e.g. Gossler & Kind 1996), from seismic

Fig. 9. Subcontinental geotherms derived xenoliths in kimberlite pipes (cratonic) and from seismic þ heat flow (global and North American averages). Considerable variation in the seismic/heat flow geotherms above about 100– 200 km represents the varying temperature structure of the continental lithosphere, beneath which the geotherms are subparallel to mantle adiabats at potential temperatures of between 1000 and 1400 8C. Allowing for the error in such seismic geotherms, it is still clear that many sublithospheric geotherms from a variety of tectonic settings must lie well below the 1300 adiabat. Xenoliths within Kaapvaal kimberlites show two families: those following the steady-state geotherm down to a potential temperature of 1200 8C, and those which were heated to a higher potential temperature by the impact of the plume on the base of the lithosphere. These imply that the potential temperature at the top of the convecting interior was 1200 8C away from plumes. Other kimberlite provinces are similar or slightly cooler still. See Reston & Phipps Morgan (2004) for a complete discussion of these geotherms.

velocities (e.g. Ro¨hm et al. 2000; Goes & van der Lee 2002) and waveforms (Cammarano et al. 2005), and from studies of heat flow (e.g. Jaupart & Mareschal 1999). All of these suggest considerable variability in the potential temperature of the convecting mantle, but with the general indication that it may be considerably cooler beneath the continents than beneath the oceans (Fig. 9). The differences are sufficient to explain dramatic variations in melt volumes during rifting (Fig. 8). Differences in temperatures of the convecting interior immediately beneath the the continental and oceanic lithosphere

RIFTED MARGINS BY PROGRESSIVE EXTENSION

are an integral part of some models of mantle convection (e.g. Lenardic & Moresi 2000). Others have suggested that the oceanic ridge system may be supplied by a relatively thin layer of warm asthenosphere derived from mantle plumes and ponding beneath the thinner oceanic lithosphere (Phipps Morgan et al. 1995), a suggestion supported by the studies of transition zone thickness and seismic waveforms cited above.

Melt trapping within the lithosphere It is, of course, possible that the lack of observed magmatism at NVRMs is not owing to a lack of melting, but rather that considerable volumes of melt are trapped and solidify in the mantle lithosphere, a suggestion also made for those portions of slow-spreading mid-ocean ridges characterized by thin magmatic crust (Cannat 1996). A recent seismic experiment (Lizarralde et al. 2004) showed that beneath slow-spreading crust the velocity structure of the uppermost mantle was consistent with the presence of numerous deep gabbroic intrusions with an integrated thickness of about 1.5 km. A similar thickness trapped at depth at rifted margins could help explain the melt discrepancy. The recognition of intrusions in the mantle at rifted margins will be complicated by the difficulty not only in distinguishing between gabbros and serpentinites, but also by any intrinsic velocity contrast between exhumed subcontinental lithospheric mantle and its oceanic equivalent. Exhumed mantle rocks in the Alps do show evidence both of intrusion by gabbroic bodies and impregnation by basaltic melt (Mu¨ntener & Herrmann 2001); these are included in the estimates of melt thickness in the Alps (Fig. 8). However, even a small amount of melting may focus extension by weakening the thinning lithosphere through which they are percolating (Lavier & Manatschal 2003). The observation at the Galicia Bank margin of minor synrift gabbros (Scha¨rer et al. 1995) that were strongly sheared during further extension is compatible with deformation at depth being focused where partial melt is present. If melt lenses do act to focus deformation, this should increase the degree of exposure of igneous rocks (brought up in the footwall of the fault) and makes the scarcity of such rocks still more problematic.

Depth-dependent stretching (DDS) The volume of melt produced might also be reduced if the extension of the lithosphere is heterogeneous (Minshull et al. 2001), i.e. if depth-dependent

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stretching (DDS) takes place so that crustal thinning exceeds the total lithospheric thinning beneath the rift centre. More specifically, crustal separation may occur before complete lithospheric separation. The key in suppressing melt within the COT is thus to extend the crust in a narrower zone than the subcrustal lithosphere, perhaps particularly the deep lithosphere (see below). This is a very different sort of DDS to that proposed as an explanation for the extension discrepancy, where it has been argued that different crustal levels are extended by different amounts in different places. Apart from local heterogeneities associated with faults and localized shear zones, significant and systematic lithospheric-scale DDS may require the decoupling of the deeper lithosphere from the shallower structures. I suggest that lithospheric-scale DDS to suppress melting may occur through a combination of several processes (Fig. 10): by flow of the deepest lithosphere (e.g. see models of Huismans & Beaumont 2003), which is consistently weak during rifting (Fig. 5), by displacement along large structures cutting through the entire mantle lid (Fig. 5), leading to a lateral displacement of the locus of mantle and crustal extension (Reston 1993), by shear along the mid and lower crustal weak zones (e.g. Reston 1988, 1990, 1993), and by displacements along late-stage serpentinite detachments (Pe´rez-Gussinye´ & Reston 2001). The development of serpentine detachments (Fig. 7) occurs late in the rifting process and would appear to provide a mechanism for widening the region of mantle extension and focusing that of the overlying crust (leading to complete crustal separation as the blocks are pulled apart). However, serpentinization is not predicted to occur until extension factors exceed about 4, by which time the asthenosphere should have upwelled to within about 30 km of the surface and, for a Tp of c. 1300 8C, melting should have just started (Pe´rez-Gussinye´ et al. 2001). It is hard to see how the development of serpentine detachments can prevent further asthenospheric upwelling while still allowing the unroofing of a 120 km-wide expanse of mantle rocks, as observed west of Iberia (Fig. 1). If DDS is allowed by decoupling within the crust, it is likely to be most pronounced relatively early in the extension history, before the crustal weak zones disappear through crustal embrittlement and the crust becomes tightly coupled to the mantle (Figs 5 & 7). Such a decoupling mechanism has been inferred from seismic reflection data (Reston 1990, 1993) and is compatible with numerical modelling (e.g. Braun & Beaumont 1987), although doubts remain about the efficiency of the process over large length scales (Braun & Beaumont 1987). Reston (1990, 1993) proposed that this mechanism

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Fig. 10. Depth-dependent stretching controlled by shear in the lower crust and shear zones in the uppermost mantle from linedrawings of unmigrated seismic sections from offshore UK, and from numerical modelling. (a) DRUM profile showing mantle reflection dipping to a depth of approximately 75 km (after depth migration). This structure may be an extensional shear, detaching upwards in a creeping layer at the base of the crust (Reston 1993). (b) North Sea line NSDP-84-01, showing symmetric outward-dipping mantle reflections, detaching upwards at the base of the crust (Klemperer 1988; Reston 1993). In both cases, lower crustal reflectivity may be related to shearing as strain is transferred between the crust and the mantle (Reston 1990). (c) Numerical model of decoupled lithospheric extension, including viscous strain softening (Huismans & Beaumont 2003), showing highly sheared creeping layer at the base of the crust (only one crustal lithology modelled), and boudinage of strong mantle lid, together with shearing of deeper mantle. Dashed lines represent interpreted high-strain zones in the mantle for comparison with the seismic. (d) More extended coupled lithospheric model (also with viscous strain softening) showing how mantle shear zones both broaden the region of mantle extension over focused crustal thinning and allow the incipient unroofing of rocks from the lower lithosphere at the sea floor. Displacement of lower lithospheric mantle to the centre of the rift should suppress melt generation to some extent.

RIFTED MARGINS BY PROGRESSIVE EXTENSION

might be coupled with the localization of deformation within the uppermost mantle lithosphere: a dipping band of seismic reflections are observed on the DRUM profile from the base of the crust down to a depth of approximately 75 km (Reston 1993) (Fig. 6), roughly the thickness of the strong mantle lid (Fig. 5). Symmetric extensional mantle shear zones under the rift flanks and dipping away from the rift centre, such as those reported beneath the North Sea (Reston 1993) (Fig. 10), might then allow the region of mantle extension to be more widely distributed than that in the crust, and may further allow the subcrustal lithosphere to be displaced towards the rift centre. Such outward-dipping mantle structures are predicted to develop according to some numerical models (Harry & Sawyer 1992; Huismans & Beaumont 2003) and would help explain the onlap of the post-rift sequences onto the prerift of the basin margins (Reston 1993). However, numerical modelling is somewhat equivocal on whether the zone of mantle extension is likely to be more or less focused than that of the crust. The models of Huismans & Beaumont (2003) with an (unrealistically) weak decoupling lower crust show that crustal extension is more distributed and less focused than mantle extension; their coupled models where the lower crust is (unrealistically) strong do predict more localized crustal than mantle extension (Fig. 10). The behaviour of a model containing initially moderately decoupling zones in the mid and lower crust that gradually couple deformation more strongly with ongoing extension is unclear, although may show a gradual increase in the focusing of crustal deformation as discussed above. Perhaps the most promising form of heterogeneous extension throughout the rifting process is the idea that the lowermost lithosphere (initially beneath c. 80 km depth – see below) is stretched over a wider region than the stronger upper lithosphere (e.g. Minshull et al. 2001) (Fig. 10d). Unlike decoupling within the crust (early extension only) and movement along serpentinite detachments (late extension only), displacement of the lower lithosphere towards the rift zone may

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continue through most of the rifting process (Huismans & Beaumont 2003). The presence of subcontinental lithospheric mantle (e.g. Manatschal et al. 2001) directly beneath the post-rift succession implies that some form of depth-dependent stretching took place during rifting. Similar subcontinental mantle has been described in the Alps, where the rocks appear to come from a progressively deeper and hotter source moving oceanwards: although the mantle directly in contact with the crust is preserved (e.g. in the Malenco area where it has a T at the start of rifting of c. 600 8C: Mu¨ntener et al. 2000), Desmurs et al. (2001) report that the prerifting temperature of the peridotites sampled across the Err– Platta COT increase oceanward from 850 + 50 to 1050 8C. At the West Iberian margin, the trend is less obvious with temperatures generally in the range 1000– 1250 8C across the entire COT (Table 1, Fig. 1). Taking the thermal model of Pe´rez-Gussinye´ et al. (2001) this would place the exhumed mantle at original depths of 90– 120 km, indicating that DDS is probably important within the lower half of the lithosphere. Within the COT the lack of mantle samples that were at temperatures of between 600 (slightly above Moho temperature) and 800 8C before rifting implies that DDS does not systematically lead to the displacement of the upper two-thirds of the mantle lithosphere levels towards the COT. Mantle levels initially above about 80 km (Fig. 5) may be too strong to undergo systematic large lateral displacement towards the developing rift, and may instead fail by a boudinage process, bound by landward-dipping shear zones in the mantle and above by a decoupling lower crust (Fig. 10). The presence of inferred landwarddipping high-temperature normal shear fabrics in the peridotites sampled west of Iberia (e.g. Beslier et al. 1996) and of landward-dipping structures in the COT (Boillot & Froitzheim 2001; Reston et al. 2004) is entirely consistent with the development of such outward-dipping mantle shear zones. It seems likely that this process together with early and late contributions from crustal decoupling

Table 1. Summary of prerift/synrift temperatures of mantle peridotites sampled in the COT of the West Iberian margin. All results are from the IAP transect except for Site 637 from the DGM transect. Sites ordered from ocean (left) to continent (right). Note that there is no obvious landward progression to cooler temperatures (Opx, orthopyroxene; Cpx, clinopyroxene) Site

1070

897/899

T (8C) Opx: 800 –1250 880 –1000 shearing 1170 – 1230 Cpx: ,900 8C crystallization Ref. He´bert et al. (2001) Beslier et al. (1996), Seifert & Brunotte Cornen et al. (1996) (1996)

637 (DGM) 1000 – 1250 Beslier et al. (1996)

1068 Opx: 1200 –1280 Cpx: ,900 8C He´bert et al. (2001)

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and mantle serpentinization, respectively, may contribute to DDS of the lithosphere and hence to the reduction of melt generation. Although numerical models of lithospheric extension shows that the region of mantle extension may be wider than that of the crust (e.g. Braun & Beaumont 1987; Reston 1993; Huismans & Beaumont 2003), these studies have not yet been combined with melt generation. However, we can estimate the required heterogeneous stretching by assuming that melt generation is directly related to the rate and amount of asthenospheric upwelling. To generate only about 1–2 km melt at the IAP margin and less than 1 km at the Alpine margins, the upwelling should only be equivalent to that occurring during pure shear by a extension factor of approximately 9+2 (Fig. 6), implying that across the COT about 15 km thickness of (mantle) lithosphere should still be present. Minshull et al. (2001) came to a similar conclusion and pointed out that, although a degree of heterogeneous stretching could easily explain the lack of melt within a narrow region of exhumed continental mantle, it is hard to see how this could explain a paucity of melt in a zone of exhumed continental mantle with a width of up to 170 km as any layer of lithospheric mantle originally exhumed within the COT would be continually stretched and extended with ongoing crustal separation. Thus, the amount of synrift igneous rock in the COT may simply be underestimated as it is trapped at depth, or be reduced either by a cooler subcontinental geotherm or by depth-dependent stretching. All of these hypotheses have their merits and all are likely to contribute to the apparent melt discrepancy at rifted margins. There is ample evidence that the continental sublithospheric mantle may be cooler than the asthenosphere beneath the oceans; there is good evidence of depthdependent stretching of the lower lithosphere and for the development of serpentine detachments during late-stage rifting, and some melt is likely to be trapped at depth beneath rifted margins as it is beneath slow-spreading oceanic crust.

Faulting and the extension discrepancy Some (but not all) non-volcanic rifted margins are characterized by well-defined rotated fault blocks. (Similar blocks may be present at volcanic margins but are buried beneath thick sequences of lavas which obscure the underlying image.) Where fault blocks are well imaged (White 1990) the amount of extension can be determined from the geometry of the blocks, specifically from the amount that they have been rotated (e.g. Barr 1987). In effect the presence of the prerift sedimentary units can be used to determine the amount of

the extension, but accurate determination of what is prerift is essential. However, at some margins thinning appears to occur without significant faulting (e.g. the Exmouth Plateau margin: Driscoll & Karner 1998) and at others the amount of extension that can be inferred from the geometry of the faults at the feather edge of the continent is less than that indicated by the crustal thinning observed (e.g. Sibuet 1992). This is the so-called extension discrepancy. The rolling hinge model (Buck 1988) for the development of fault slices (which may well be applicable to late-stage faulting during the formation of the West Iberia margin: e.g. Reston et al. 1996) does not alone solve the extension discrepancy as it should still be possible to reconstruct the geometry of the hanging wall from the various fault slices if the prerift is present. Only if these slices were cutting through a previously thinned section should the reconstructed section fail to restore to full crustal thickness. This can be seen with restoration of depth sections from the West Iberia margin (see below). The West Iberia margin exhibits an apparent extension discrepancy. The structure of the south Iberia Abyssal Plain segment of this margin is well imaged by the prestack depth migration of a high-quality reflection profile LG12 (Krawczyk et al. 1996) (Fig. 4) and further constrained by seismic refraction data (Chian et al. 1999; Dean et al. 2000). The most obvious features of the seismic data are the L fault, a large partly listric fault that appears to flatten at a depth of about 14 km, and the H detachment, which can be traced westwards down from the 1065 high, beneath a couple of wedge-shaped fault blocks and back to the top of the site 900 high. At the DGM, the seismic signature (polarity, waveform and amplitude: Reston 1996) of the S detachment fault indicates that it forms a sharp boundary between the crustal rocks and underlying higher velocity partially serpentinized peridotites (reflection coefficient 0.2), a result confirmed by wide-angle data (Zelt et al. 2003). Similarly, the H reflector shows an almost identical waveform to the sea-floor reflection, itself shown by comparison with its multiple to be a reflection from a single interface; the reflection coeffcient of H is estimated to be 0.15 (Krawczyk 1995). Between L and H, another fault (here labelled K) can be traced to depth from near the west flank of the 1065 fault block. To the west of H, a reflection that marks the continuation of top basement underlies the 1069 high; it is thus thought that the 1069 fault slice is allochthonous and is separated from the underlying basement (serpentinized peridotite?) by a low-angle fault (F: Krawczyk et al. 1996; or HHD, Whitmarsh et al. 2001) that follows top basement for much of its length.

RIFTED MARGINS BY PROGRESSIVE EXTENSION

A restoration of the depth section of LG12 illustrates the extension discrepancy. This restoration differs from those of Manatschal et al. (2001) and of Whitmarsh et al. (2001) in several ways. First, I have restored movement along the fault K, ignored previously. Second, I have taken the CMB within the limits of Whitmarsh et al. (2000), constrained by wide-angle data. It is important to take the CMB and not the deeper Moho as the

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serpentinites between the two is petrologically part of the mantle not the crust. Third, I have made the restoration in two phases (Fig. 11), indicating the oceanward migration of the location of the extension. By initially removing the effect of fault F, which clearly cuts across and hence postdates the H detachment, I have also ‘undeformed’ the wedge-shaped fault blocks directly above the H detachment and in doing so restored H to a

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Final section. Movement along F caused flexural back rotation of H and of F due to footwall unloading. Possibly enhanced by serpentinization (volume increase) of mantle beneath H and F and by further faults to west. βf (stretching factor) during these last phases of faulting (H, K, L and then F): 37.4 km / 21.4 km = 1.75.

Fig. 11. Two-stage restoration of LG12 based on the structural interpretation shown in Figure 4. The restoration illustrates several key points: the fact that movement along the latest faults cannot explain the extreme thinning of the crust, the possible occurrence of polyphase faulting, the development of faults cutting across the entire crust, the development of serpentinite detachments, and the unroofing of mantle and deep crustal rocks. See text for discussion.

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consistent, if largely low-angle, westward dip. Despite these differences, the general result is similar to previous restorations, showing that the 1069 block was originally contiguous with the blocks overlying the H detachment and that the crust was already quite thin before the latest phase of faulting. Finally, I emphasize that this is only a partial restoration which aligns intrabasement reflections that may correspond to earlier faults (see below). The wedge-shaped fault blocks above H, and the resultant oceanward steepening of the bounding faults on the restored section, is consistent with a rolling hinge model, but is not evidence for the application of this model to more than the breakup of the crustal block above H. Although it is quite possible that the other faults marked in bold in Figure 11 were active sequentially (e.g. in a rolling hinge model with the locus of deformation migrating oceanward), the lack of cross-cutting relationships between the faults prevents confirmation of this hypothesis, so I have restored the movement along these faults in one go. The restored section has a crustal thickness of less than 6 km. Even using the deeper Moho of Manatschal et al. (2001), the crust was only about 8 km thick prior to the development of the latest faults. Clearly, the crust has not been solely thinned from a thickness of approximately 30 km by movement along these faults. The question is how the extra thinning has occurred. Proposed solutions to the extension discrepancy include lithospheric- (or

at least crustal-) scale simple shear, depth-dependent stretching and polyphase faulting.

Simple shear If the lithosphere was pulled apart by simple shear along a low-angle detachment (Wernicke 1985), the failure to recognize this detachment at either the upper or lower plate margins would result in a large underestimate of the amount of extension accommodated by faulting. At an upper plate margin, such a detachment may form a tectonic Moho and will not project up to a breakaway, making the recognition of the detachment particularly difficult: at a lower plate margin, the detachment fault should cut up through the crust towards a breakaway, and so should be easier to identify. As a result, attempts to explain the extension discrepancy by simple shear generally invoke such an invisible detachment to thin the crust beneath the upper plate. Apart from the rheological unlikelihood of cutting through the entire rheologically stratified crust with a through-going detachment at an early stage of rifting (e.g. Huismans & Beaumont 2003), another problem is that not all margins can be upper plate margins (the ‘upper plate paradox’: Driscoll & Karner 1998). The simple shear idea might, however, still apply if the detachment on the lower plate was simply not identified as such, so that the extension on this margin is also underestimated (Fig. 12). For instance, if it were to be cut by later faults

upper plate

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successor faults cut detachment + further thin lower plate. Lower crust exhumed and faulted

upper crust upper plate over-thinned by removal of lower crust

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Fig. 12. Cartoon illustrating how movement along a crustal-scale shear zone can explain an apparent extension discrepancy on the upper plate and on the lower plate if the nature of the detachment surface is not recognized, for instance if it is cut by later normal faults (note that owing to the embrittlement of the crust these can cut across lower crustal rocks). Only if the prerift is recognized and dated on both conjugate margins can such a possibility be excluded.

RIFTED MARGINS BY PROGRESSIVE EXTENSION

(e.g. successor faults, similar to those observed cutting detachment surfaces at mid-ocean ridges) the detachment might well not be recognized and the total amount of extension measured would correspond to that along the later, steep faults. This is, of course, an example of polyphase faulting (see below). The key here, as always, is to accurately identify the prerift sequences on both sides of the ocean and to distinguish these from those ‘prefaulting’ sequences deposited prior to the development to the visible faults.

Depth-dependent stretching in the crust Simple shear is just one form of depth-dependent stretching (DDS), which if occurring within the crust provides a possible explanation for the extension discrepancy. (It is not, however, the only solution, so DDS should not be used as synonymous with the extension discrepancy.) In DDS, the total amount of extension undergone by all the lithospheric levels must be the same at a regional scale to avoid space problems, but the lateral distribution of this extension may be different. Various forms of DDS has already been discussed above as a possible mechanism to explain the small amount of igneous activity associated with the development of NVRMs. It is also clear that some form of DDS is necessary to explain the unroofing of large expanses of subcontinental lithospheric mantle within the COT of some margins (West Iberia and the Alpine margins: Manatshal et al. 2001). However, as discussed above, the petrology of these mantle rocks indicates that they came from what was originally the lower part of the lithospheric mantle and as such do not provide evidence for DDS at higher lithospheric levels. This is important as only DDS within the crust can be invoked to explain the extension discrepancy, by thinning the crust more than observed from the faults imaged. In effect, explaining the extension discrepancy by DDS requires that the lower crust has somehow been displaced away from beneath the little faulted brittle upper crust, which may even be placed directly atop lithospheric mantle. This might in principal thin the crust by a factor of 3, assuming everything beneath the initial brittle upper crustal layer (c. 10 km thick, see Fig. 5) is removed without thinning the upper crust at all. Such a displacement might occur away from (Brun & Beslier 1996) or towards the margin (Driscoll & Karner 1998). The latter would seem more likely to lead to an inverse discrepancy as it would serve to maintain crustal thickness where the upper crust has been highly extended, and even completely separated, unless final breakup occurs where the lower crust has been removed, in which

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case the conjugate margin should show the inverse discrepancy as an apparent ‘lower plate margin’. In contrast, any postulated displacement of the lower crust toward the continent would have to be accompanied by opposite displacements of the mantle lithosphere toward the COT, where it is exhumed. Apart from crustal-scale simple shear, several mechanisms of crustal DDS are conceivable, including lower crustal flow and lower crustal boudinage (Fig. 13). Lower crustal flow is problematic as it has been shown that, at the start of rifting, the lower crust beneath a cool rift is too viscous to flow over large distances (Kusznir & Matthews 1988; Hopper & Buck 1998; McKenzie & Jackson 2002), although still capable of small-scale adjustment in response to the tilting of the fault blocks. Second, as discussed above, the rheology of the crust is not fixed, so that during progressive extension the crust becomes ever more tightly coupled to the mantle until, at large extension factors, even originally creeping rocks become brittle. The entire crust may be brittle by an extension factor of about 4– 5 (Pe´rez-Gussinye´ & Reston 2001), so that the lower crust should not be able to flow anywhere. Far from lower crustal flow being favoured at high extension factors as suggested by analogue models (e.g. Brun & Beslier 1996), a decoupled deformation within the lower crust must become progressively more difficult. The fundamental flaw of most analogue modelling is that the rheology of materials is fixed and cannot evolve in response to pressure, temperature, strain rate and fluid pressures. This hinders meaningful simulation of rift evolution at extension factors greater than about 1.5 when rheological changes start to become important. The implication is that at NVRMs, such as West Iberia, the extension discrepancy can neither be studied by such modelling nor be solved by lower crustal flow. More promising is the necking or boudinage of the strong core of the lower crust between weak decoupling zones above and below. A degree of such boudinage has been inferred from deep seismic reflection images (Reston 1988). Completely removing the lower crust beneath sections of the developing margin would require a far more extreme boudinage, and it is hard to see how the resultant differential subsidence and shear on the base of the brittle layer would not induce extensive faulting. Such boudinage is thus unlikely to explain those places where upper crustal extension appears negligible although the entire crust is strongly thinned (Driscoll & Karner 1998; Davis & Kusznir 2004), but may help to explain places where the amount of upper crustal thinning is substantial but nevertheless smaller than whole

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(a)

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IED ED COT UC

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lower crust displaced landward (c)

IED

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Fig. 13. Models for DDS within the crust that may give rise to an extension discrepancy. ED, extension discrepancy; IED, inverse extension discrepancy. (a) Lower crustal flow or displacement towards the COT (after Driscoll & Karner 1998). This should give an inverse discrepancy near the COT, unless the lower crust is itself subsequently faulted and misinterpreted as upper crust. (b) Lower crustal flow away from the COT (Brun & Beslier 1996). (c) Lower crustal boudinage along shear zones (Reston 1988). This should give local extension discrepancies between the boudins and inverse discrepancies above the boudins.

crustal thinning. If the locus of lower crustal boudinage influences the site of eventual breakup, a moderate extension discrepancy might thus result near the COT. Davis & Kusznir (2004) proposed that depthdependent stretching could develop in response to a divergent flow field associated with the onset of sea-floor spreading which would, in effect, shear the lithosphere and thus allow extension of the lower lithosphere over a wider zone than that of the upper crust (Fig. 14). This idea is attractive in that it provides a mechanism for extremely large-scale DDS linked to processes associated with the onset of sea-floor spreading, i.e. which may be restricted to margins. However, such a divergent flow field can only explain the extension

discrepancy if it leads to DDS of the different crustal levels. If, for instance, the lower crust is sheared away from the little extended upper crust, the displacement relative to the upper crust of the lower crust should be in the same direction as that of the underlying mantle, unless the kinematics of the deformation field are very complex. As the subcontinental mantle appears to be displaced towards the COT at NVRMs, the lower crust should also be (e.g. Driscoll & Karner 1998), which as discussed above may not provide an explanation for the extension discrepancy. Furthermore, it is clear that an active divergent upwelling can only affect the lower crust once it has affected the entire subcrustal lithosphere. The presence of mantle rocks within the COT dominantly from the lower half of the

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Fig. 14. Half of symmetric depth-dependent stretching models for the formation of rifted margins resulting from divergent upwelling, both passive and active (Kusznir et al. 2005). Although these models each predict some of the features discussed in the text, neither can duplicate all of the observations.

subcrustal lithosphere may mean that DDS has only affected the lower half of the mantle lithosphere and not the strongest mantle lid beneath the Moho. If this is the case, it cannot then have affected the lower crust and cannot help resolve the extension discrepancy. Despite these inherent problems, Davis & Kusznir (2004) and Kusznir et al. (2005) proposed that depth-dependent stretching explained the extension discrepancy at a variety of margins. Davis & Kusznir (2004) reproduced a small amount of depth-dependent stretching, but the amount was far less (maximum bwhole crust 2 bupper crust  0.2) than they had inferred at rifted margins (bwhole crust 2 bupper crust  1.8) to explain the extension discrepancy. Kusznir et al. (2005), explored the effect of active v. passive upwelling on the margin structure, including the unroofing of subcontinental mantle in the COT. The passive model (Fig. 14, top) predicts the unroofing of a zone of subcontinental mantle approximately 50 km wide (somewhat less than observed west of Iberia, but comparable to other NVRMs – Fig. 1). The model also predicts that a complete lithospheric section should be unroofed in the COT, with mantle lithosphere from all depth and temperature ranges. This is not observed: as mentioned above and discussed more below, the mantle rocks unroofed appear to come from the lower lithosphere. Also note that the thinning of the upper continental crust occurs rapidly at about 300 km and exceeds the total crustal thinning further oceanwards. Far from explaining the extension discrepancy, this model predicts an inverse discrepancy, which has not been noted. The active upwelling model in contrast (Fig. 14, bottom) does predict

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less upper crustal than whole crustal thinning between about 220 and 270 km, but not sufficient to explain anything more than a very slight extension discrepancy, and beyond 280 km the upper crust appears to be thinned more than the whole crust. Also note that this model can only explain the unroofing of a very narrow zone of subcontinental mantle (at depth of 305–320 km). There are other problems with crustal DDS as an explanation for the extension discrepancy. First, if the upper crust near the COT is extended far less than the whole crust because the LC has somehow been removed, where has the missing lower crust gone? Another way of looking at this is that if the upper crust near the COT is less extended than the deep crust, then somewhere else the upper crust must be extended considerably more than the lower crust. Such an inverse discrepancy is, however, never described, although required by models of DDS (e.g. Kusznir et al. 2005). I term this the ‘paradox of the missing lower crust’. As discussed above, if the lower crust has been excised and the upper crust placed directly against lithospheric mantle, it might in theory be possible to thin the crust considerably without significant faulting. The amount of crustal thinning that can occur without extending the brittle layer is simply dependent on the relative thickness of the whole crust and the brittle upper layer. In a warm lithospheric model with relative thick crust, appropriate for instance for the Woodlark Basin the upper brittle layer might be thinner and the whole crust thicker than shown in Figure 5, potentially allowing more crustal DDS, although whether this would actually occur is another matter altogether. However, for West Iberia, crustal DDS that removes everything beneath the approximately 10 km-thick brittle upper layer can only thin the crust by a factor of 3. Even after restoring the extension along the block-bounding faults West of Iberia (Fig. 11), the crustal section is considerably thinner than 10 km, requiring at least a supplementary explanation other than DDS beneath the brittle upper crust. A further problem of applying crustal DDS to the West Iberia margin is that, although some authors have suggested otherwise (e.g. Nagel & Buck 2004), there is no evidence that the lower crust is generally missing towards the COT of the West Iberia margin where the extension discrepancy is most pronounced. (The same holds for the Alps, where deep crustal prerift gabbros have been identified towards the COT, although their thickness apparently varies considerably through boudinage (Manatschal et al. 2001; Mu¨ntener & Herrmann 2001)). Although the entire crust is brittle, this does not exclude the presence of the lower crust as explained above. No unequivocal

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lower crust has been sampled along the drilling transect at the DGM, but basement rocks have only been sampled in one place there (granodiorites were recovered from the top of a fault block during Galinaute Dive 11), and the gabbros described by Scha¨rer et al. (1995) slightly further to the north have been interpreted as lower crustal. At the IAP margin, the anorthosites, mafic granulites and tonalities sampled at sites 1067 and 900 are typical lower crustal lithologies and record temperatures and pressures appropriate for a lower crustal origin (e.g. Manatschal et al. 2001), yet on the restored section are at a depth of only about 5 km – they were within 5 km of the surface prior to the latest phase(s) of faulting. Given that the brittle upper crust at the start of rifting had a thickness of about 10 km (Fig. 3), it is hard to see how such a fault slice could sample rocks from the deep crust without previously thinning the crust by brittle faulting, i.e. requiring polyphase faulting. As noted above, DDS can only realistically thin the crust without faulting beneath the brittle upper crustal layer, yet the restored thickness of the deep margins West of Iberia is less than the likely original thickness of this layer. I conclude that, although there is undoubtedly both an apparent extension discrepancy and lithospheric-scale DDS at many margins including West Iberia, it remains difficult to explain the extension discrepancy by DDS alone at the West Iberian margin. This does not imply, however, that it cannot be important at other margins, subject to the issues discussed above.

Polyphase faulting Even after restoring the extension along the blockbounding faults West of Iberia, the crustal section is considerably thinner than 10 km. This is not simply a function of pushing the toe of a large fault block underneath the restored section, as for instance at the IAP margin (Fig. 4), depth to top basement (and hence presumably crustal thickness) does not increase significantly for the next 80 km moving landward: there is no suitable large fault block. One possible solution is the occurrence of unrecognized polyphase faulting. Only if the prerift configuration can be recognized and distinguished from a prefaulting sequence (which may predate only the latest faults) and shown to be continuous at the top of the fault blocks on both sides of the rift –ocean can such a scheme be ruled out: without excellent stratigraphic control, an extension discrepancy cannot be proven to exist. A problem at most margins is that the sedimentary sequences are only sampled at the margin edge, i.e. in shallow water. Correlations with the units in the deep margin are partly carried out in terms of reflection

character, partly through jump correlations across large fault blocks/grabens and partly through the apparent relationship of units to faulting. This may well prove that certain units are prefaulting (or post-faulting) on a fault block scale, but does not imply that they are prerift or post-rift on a margin scale. Polyphase faulting is perhaps to be expected as the crust is progressively more and more extended. Simple mechanical considerations such as the Mohr–Coulomb criteria for shear failure (internal friction coefficient 0.85; cohesion, e.g. Scholz 1990) indicate that normal faults should form at angles of approximately 658 (Fig. 3). Once a fault has formed, the development of weak fault gouge (friction coefficient of perhaps 0.6) and the loss of cohesion along the fault may mean that the fault can stay active to angles of perhaps 358, allowing extension along faults to be accommodated by progressive back-rotation as in the domino model (Fig. 15). Beyond that in classical fault mechanics we would expect the fault to lock-up and a new fault to develop. The observation that most normal faulting earthquakes occur at dips between about 708 and 308 (e.g. Jackson 1987) supports this conceptual model. A rotation from about 658 to 358 for a simple domino model corresponds to an extension factor of about 1.6 (e.g. Barr 1987). Allowing the faults to initiate at a steeper angle of 708 and remain active to 308 increases the extension factor at which the faults should become inactive to 1.9. At higher values, we might expect faults to lock-up. These calculations are assuming that the fault block rotates rigidly, whereas it is quite probable that the block will internally deform, adding another component, the thinning of the block, that is not controlled by domino rotation. Thus, for the properties we might expect, faults should lock-up close to a extension factor of 2, and extension beyond such extension factors will require either the development of a new fault, or movement along an existing fault at low angle, or some other mechanism other than large-scale pure shear. Assuming that the former occurs, at extension factors beyond 2, movement will be largely accommodated by movement along a second generation of faults, which will rotate both the crustal blocks and the faults further (Proffett 1977). Furthermore, when these faults themselves become strongly rotated, they too should lock-up – at a extension factor of between 3 and 4. Beyond this a third generation of faults may be required, and so on. As discussed above, the spacing of a second generation of faults is likely to be less than that of the first phase, so that the second generation of faults is likely to cut the pre-existing blocks in a variety of places

RIFTED MARGINS BY PROGRESSIVE EXTENSION

97

Fig. 15. Basic domino model from brittle extension, showing fault-block rotation and synrift sedimentary geometries. Although such models are an oversimplification they do make some useful predictions and reveal some interesting characteristics.

(Figs 15 & 16), although perhaps the most likely is where the faulted crust is weakest, i.e. close to the base of the half-graben. In two dimensions, the second generation of faults may dip in the same way as the first direction, or in the opposite direction (Figs 16 & 17). In both cases, the predicted structure is very complicated, and may include: † old fault segments isolated within basement that may become increasingly hard to recognize; † local deep basins may develop, bounded by steep-sided highs. Such basins may be starved of external sediment input, but may collect the products of mass-wasting of highly fractured basement blocks and exposed early synrift sediments; † faulted exposures of synrift sediments; these are likely to be subject to reworking and masswasting, leading to the partial destruction of early synrift geometries.

Other predicted features when the second generation of faults dip in the same direction (Fig. 16) include: † fault surfaces intersecting at approximately 308; † old faults continue to dip in original direction (but at ever lower angles) until b3; a combination of the new and old fault surfaces may be interpreted as a single phase of faulting (partly concave up, partly convex-up), dipping at approximately 208; † old faults rotated to low angles and even to dip opposite to transport direction; † latest faults intersecting the original surface at approximately 908; † steep top basement surfaces and earliest synrift sediments–both of which are likely to be poorly imaged; † preferential reworking of the thicker portions of synrift sediment wedges, which are likely to occur in the footwall to the new faults.

98

T. J. RESTON synrift sediment

β ~ 2 - original faults lock-up, new faults develop

reworking of early synrift

mass-wasting of fractured basement

reworking of early synrift

β ~ 2.6 - early fault blocks dismembered. Rugged topography leads to mass-wasting / reworking of early synrift and isolation of small basins reworking of early synrift

mass-wasting of fractured basement reworking of early synrift

β ~ 4 - Early faults rotated to dip in opposite direction; original top basement and earliest synrift very steep (not imaged?); mass-wasting / reworking will have continued

Fig. 16. Domino model showing effects of further extension, with faults locking up and new faults cutting across. The resulting sedimentary and structural geometries are complicated even in such a simplistic model.

synrift sediment

β ~ 2 - original faults lock-up, new faults develop, dipping in opposite direction

mass-wasting of fractured basement

mass-wasting of fractured basement reworking of early synrift

reworking of early synrift

β ~ 2.6 - early fault blocks dismembered. Rugged topography leads to mass-wasting / reworking of early synrift and of fractured basement and isolation of small basins mass-wasting of fractured basement

reworking of early synrift

mass-wasting of fractured basement

reworking of early synrift

β ~ 4 - Original top basement ~ horizontal but widely-spaced; early synrift forms upward thickening fans; early faults rotated to steep dips (not imageable in basement). Mass-wasting / reworking will have continued

Fig. 17. Domino model in which second-generation faults dip in the opposite direction to initial faults. A whole new set of complications can result.

RIFTED MARGINS BY PROGRESSIVE EXTENSION

Other predicted features when the second generation of faults dip in the opposite direction (Fig. 17) include:

99

Extension discrepancy at the West Iberian margin As noted above, there is an extension discrepancy at the IAP margin (Fig. 4) as the restoration (Fig. 11) of the extension along the block-bounding faults gives a crust that was only about 6 km thick (or c. 8 km thick if Manatschal’s CMB is taken). Two other relevant observations can be made from the restored section. The first is that the restoration successfully places the lithologies drilled in an apparently logical structural sequence, with sites 1069, 1067 and 900 providing an almost vertical section through the thinned crust. However, as discussed above, the thinning of the crust from more than 30 to about 7 km cannot have been by ductile stretching, but must involve faulting not fully resolved. The mafic granulites drilled at site 900, however, occur perhaps 2 km above the Manatschal CMB but directly atop the Whitmarsh CMB, in which they must sample lowermost crust directly above a mantle transition zone. Furthermore, the presence of lower crustal rocks so close to the surface and to upper crustal sediments means that we cannot simply consider these fault blocks as slices through the upper crust transferred from the hanging wall to the footwall during extension along a rolling hinge model (Buck 1988). Similarly, the presence of deep crustal rocks in these small fault blocks is hard to explain using models in which the crust was simply thinned by the removal of the toe of one or more large fault blocks out to the east. Such an interpretation also conflicts with the inference that the crust to the east of Site 901 is not appreciably thicker for a distance of 50 km (Fig. 4). Second, after the restoration, weak basement reflections become aligned and appear to be associated with residual offset of the Whitmarsh CMB. I interpret this as an anastomosing network of earlier faults, evidence that more that one phase of faulting controlled the evolution of this margin. Unfortunately, uncertainities in the interpretation and in the direction of extension mean that it is

† different generations of faulting intersecting at approximately 908; † later faults intersecting the top basement at about 308, and rotating this to a lower angle; † apparently inverted wedges of early sediments, and the reversal of thickening directions; † preferential reworking of the thinner portions of synrift sediment wedges, which are likely to occur in the footwall to the new faults. The latter two effects may make identification of the synrift particularly difficult. These structures are, however, far simpler than those likely to occur in the real world, when polyphase faulting is likely to be 3D, is unlikely to be regularly spaced and subparallel, and may have variable polarity. Furthermore, the sketches shown in Figures 16 & 17 only show two stages of faulting, corresponding to a extension factor of, at most, 4. At the feather edge of the continents, crustal extension factors may in contrast exceed 10 as the crust is thinned to a few kilometres or less. In such cases, several other phases of faulting are to be expected, as observed in a simplified cross-section through part of the Betics (Booth-Rea et al. 2003) (Fig. 18) that has undergone post-orogenic extension. Here, a section of crust initially 20 km thick has been extended and thinned to 1 km by at least seven phases of normal faults. In the process, deep ductile crust has moved entirely into the brittle field and younger faults have cut and offset older faults. Although such structures may be identified after detailing mapping at outcrop scale, on seismic sections the recognition of polyphase faulting will be far more difficult on a seismic image, especially a time (i.e. distorted) section, and particularly when the margin is sediment starved, or when the early synrift is extensively reworked during later faulting, and the sedimentary geometries observed are likely to give a very incomplete picture of the rift evolution.

1 km

1 km

1

2

3

4

5

6

7

phase of faulting

Fig. 18. Simplified cross-section from the Betics (Spain), showing extensional fault geometries produced during thinning of this section from more than 20 down to 1 km (after Booth-Rea et al. 2003). Lithological information left off for clarity. How many of these faults would be recognized on a seismic section?

6.9

LC

5.2

6.1

3.7

8.0

6.7

4.2 4.9 6.1

6.9

5.2

7.8

5.8

5.2 5.2 5.7 6.7

100

110

Distance [km]

120

4 km streamer

120

1 km streamer

130

130

5.3 6.0 6.2 6.7

6.1

5.0

140

140

5.0

LC

Central Iberian terrane

7.8

4.0

8.0

150

multiples

6.9

MC

UC

East

20

15

10

5

0

0

12 14 16 18 20

14

16

18

20

Depth [km]

Fig. 19. Velocity section (numbers P-wave velocity in km s21) and seismic depth section of profile ISE17 across the Galicia Interior Basin (Pe´rez-Gussinye´ et al. 2003) at true scale. Note that this is a depth migration not a time migration so multiples do not come in at twice the travel time of the primary. Key here is the presence of late east-dipping faults (bold, solid) cutting and offsetting a gently east-dipping top basement (depth of 90– 110 km). Note also that beneath this east-dipping surface, high-velocity (lower crustal? – very light shading) rocks approach top basement. This can be explained if the gently east-dipping surface (dashed) is an earlier fault that has partially exhumed deep crustal rocks. Other gently east-dipping reflections are similarly interpreted (dashed lines).

22

10

12

8

6

10

8

6

4

90

100

5.0 5.2 5.6 6.0

4.9

7.0

2

80

70

90

6.9

6.4 6.7

4.7 5.6

Ossa Morena terrane

3.15 3.25

6.7

4

80

70

4.13.5 4.3 5.0

mantle

2.9 3.3

7.8

1.8

1.8 2.0 2.2 2.4 2.6 2.8 3.0 3.2 3.4 3.6 3.8 4.0 4.2 4.4 4.6 4.8 5.0 5.2 5.4 5.6 5.8 6.0 6.2 6.4

2-D VELOCITY STRUCTURE (OBS) + MCS INTERPRETATION

6.1

UC

5.2

km/s 1.5

2

0

25

20

15

10

5

0

West

Depth [km]

Depth [km]

100 T. J. RESTON

RIFTED MARGINS BY PROGRESSIVE EXTENSION

not possible to meaningfully restore the section further. That the CMB was already offset by faults when the crust had been thinned to approximately 6 km is completely consistent with the expected rheological evolution of the crust, which should become brittle at an extension factor of about 4. Crustal embrittlement may lead to mantle serpentinization, allowing faults such as H and F to remain active at low angles. Continued movement along these resulted in the dismemberment of the crust, final crustal separation and the exposure of the mantle forming the footwall to these faults as top basement. The restored section of LG12 thus requires some mechanism to have prethinned the crust, and provides evidence that this may have involved movement along an earlier phase of faults. There is independent evidence for at least two phases of faulting here. Formation microscanner data (Whitmarsh & Wallace 2001) show that the sediments drilled at Site 1065 were first tilted 158 to the SE during early rifting and then, after a pause of several million years, tilted to the east, consistent with movement along the latest westdipping faults. It is thus not surprising that a 2D restoration along the west-dipping faults does not explain all the observed thinning as these only controlled the latest phase of polyphase 3D extension.

The West Galicia margin Evidence for polyphase faulting is clearer at the West Galicia margin, where it is constrained by stratigraphic control, by the velocity structure and by the seismic image. This segment comprises the Galicia Interior Basin in the east, the Galicia Bank and to the west the Deep Galicia Margin (DGM). The structure of the GIB is constrained by a coincident reflection and wide-angle profile ISE17 (Pe´rez-Gussinye´ et al. 2003; Reston 2005) (Fig. 19). The wide-angle model shows a pronounced velocity layering beneath the Galicia Bank to the west; these layers can be traced to the east beneath the GIB, but appear to show that the locus of upper crustal and lower crustal extension are laterally offset, with lower crustal velocities at shallow depths (upper crust largely absent) at the west flank of the basin, and in contrast an absence of lower crustal material beneath the middle and west flank of the basement. The latest phase of faults defines the current fault blocks. However, although these faults dip to the east, top basement (marked by velocities well above 5 km s21) does not show consistent dips: in places it dips steeply to the west (intersecting the faults at c. 908), elsewhere it dips to the east. Both can be explained if the latest faults cut across an earlier set of east-dipping faults – where top basement dips to the east, it is the trace of an earlier

101

fault; where it dips steeply to the west, it is prerift top-basement that has been rotated to the west by two phases of east-dipping faults. This interpretation is illustrated by the restoration (Fig. 20), which clearly shows that two phases of faulting restores the crust that has acted in a brittle manner to approximately 24 km thickness and the entire crust to approximately 30 km. The restoration also illustrates how two-phases of faulting can explain the presence of high-velocity lower crust at top basement – this occurs where top basement is east-dipping and corresponds to the footwall of a major east-dipping normal fault. In this case, large east-dipping faults have lead to a lateral offset of the locus of upper crustal and lower crustal extension in a localized and specific form of depth-dependent stretching, which is only revealed by the presence of coincident reflection profiles and wide-angle data. To the west of the Galicia Bank is the DGM where Davis & Kusznir (2004) ruled out significant extension along unrecognized faults developed after the formation of the fault blocks imaged. However, they did not consider the possibility of earlier unrecognized faulting, for which available data provide ample evidence. The DGM has been the subject of one ODP leg, two submersible sampling campaigns and several high-quality seismic datasets. From these, a general stratigraphic framework has been constructed (Fig. 21) (Reston 2005). The salient points are that the seismostratigraphic unit 5a (terminology of Mauffret & Montadert 1988) tilted within fault blocks consists of deep-water deposits, and represents sediment deposited after the onset of rifting (as deep water) but before the latest phase of faulting (as tilted within the fault blocks). The synrift nature of 5a is supported by variations in its thickness. First, along an individual profile it appears to thicken across successive fault blocks to a maximum before being thinner in the next block. Second, its thickness also varies considerably between profiles. These characteristics imply that it was deposited in basins controlled by faults and that these basins did not simply run north –south in response to east –west extension but had components of extension oblique to the east –west profiles. The synrift nature of unit 5a implies that at least two phases of faulting occurred at this margin, one before or during deposition of unit 5a and one after. Furthermore, seismic unit 5a is diachronous, younging to the west across the drilling transect: sediment tilted within the blocks at site 640 are the same age as seismic unit 4a (post-faulting) further east at site 641 (Reston 2005). This diachroneity might be explained by an oceanward migration of extension during the second phase of faulting, or by an

102

T. J. RESTON "crestal collapse" graben accommodating flexure of the rollover margin

A: Present-day: simplified seismic interpretation (below postrift section) early synrift reworked

? ?

plastic deformation vertical grey bars - velocity boundary between upper and mid-crust from wide-angle data black-white dashed line - predicted boundary between upper and mid-crust from crustal restoration.

Moho

B: latest phase of faulting restored: end Valanginian ? mass-wasting

mass-wasting

"core complex". Mid-crustal rocks exhumed

"crestal collapse" graben accommodating flexure of the rollover margin ?

10 km

D: early faulting restored: pre-rift configuration (Tithonian) faults steeper away from suture 50

E: pre-rift restoration, including removal of 20% bulk stretching

initial low-angle due to reactivation of Variscan suture? 45

34

35 35¡

upper crust

0

58

2

55

45

45

35¡

4 6

mid crust

8 10 12 14 16 18 20 22 24 26

Depth [km]

Crustal blocks at depth shown rigid but probably deformed plastically during early phases of extension. Note however depth from which blocks ere exhumed.

Including a component of internal block stretching steepens initial faults considerably and increases depth of exhumation

original Moho ?

28

Fig. 20. Two-stage restoration of the section shown in Figure 19. Top: present situation. The bars show the location of the velocity boundary between the upper and lower crust, the through-going dashed line is an initially horizontal marker (see bottom). Note the close match between the two, illustrating that the restoration can explain the presence of high-velocity rocks near the top basement in the centre of the section (Fig. 6) by the exhumation of deep crustal rocks. Middle: restoring the first phase of extension brings gently east-dipping structures into alignment. Bottom; further restoration along these yields a tight reconstruction. Note only the portions of the crust that deformed by brittle faulting are shown; deeper levels deformed by plastic– ductile creep throughout evolution.

additional phase of faulting to the west of the 639– 641 high. If the former were the case, it should be possible to partially restore the section as one phase of migrating faulting (Fig. 22). Restoring the latest phase of movement along the blocks to the west of Site 639 produces a close fitting section less than 5 km thick: the quality of the fit suggests that faulting may have migrated oceanards for these faults, although it may equally well have occurred simultaneously (jump from Fig. 22a to 22d). However, the less than 5 km-thick section is considerably thinner than the Site 639 fault block – attempting to restore the section further by sliding the section west of 639 level with the top of block 639 produces a large misfit at the base of the blocks (Fig. 22e). Either the CMB is offset in a way suggesting that the restoration has gone too far, or the blocks must be very strongly distorted.

Removing further subseismic extension associated with the latest faults (Marrett & Allmendinger 1991) does not help unless it is both extreme and restricted to the small blocks west of Site 639. A solution to this misfit is to thicken the section west of 639 further along a set of intermediate age faults before attempting to restore movement along the older fault I at the west flank of the 639 block (Fig. 22f) (Reston 2005). Reston (2005) pointed out that this further restoration is only illustrative as the 3D geometry and sense of movement along the earlier faults is poorly constrained, but Figure 22 shows that it is feasible and apparently necessary. This interpretation is completely consistent with the diachroneity of seismostratigraphic unit 5a; west of 639 the latest block-bounding faults represent a later phase of faulting than that forming the west flank of the 639 block itself. Invoking an additional

RIFTED MARGINS BY PROGRESSIVE EXTENSION

112

latest faulting ?? Barremian- Aptian 5a sandstone sandstone ? 5a uncertain age

Aptian

rift durations

Sites 639 ,638 and 641

Site 640

Gal 11

breakup ?

WDGM

121 Barremian 127 Hauterivian 132 Valanginian 137 Berriasian 144 EDGM Tithonian + GIB Ma basement

deep-water limestone: Tithonian Valanginian?

5

granodiorite

B

103

latest fa

ulting d iachron ous synrift blocks rotated within is diach ronousfault

641: Aptian-Barremian 4a turbididtes, post local faulting 638: 250 m.sandstone Barremian , Hauterivian, Valanginian 5a 639: Valanginian marls, sandstone.

rift onset ? 5b 639: Dolomites - shallow water; Tithonian limestones -sandstones

pre-rift

B ? 639: rhyolites

seafloor 6 8

6

Depth

S

P

10

shot

2200

2300

2400

km

CMB? 2500

2600

2700

2800

Site 639

2900

Site Site 638 641

Dive Site 11

*

seafloor

Depth

km 10

serpentinized mantle

GP11

8

S

S

6

6

Depth 8

S

S

10

S5

S4

8

S6

km

10

GP101 6 8

seafloor 6

P

?

S

S

10

shot

Depth

6

km

8 10

1500

1600

8

km

serpentinized mantle

GP12

Depth

S

10

CMB 1700

Site 640

1900

2000

2100

2200

2300

seafloor 6

?

S

8

S

S

10

GP102 5 km

ODP Leg 103 sites

unit 4a (post- local faulting, pre-breakup)

post-breakup (unit 3) unit 4b (latest syn-faulting)

unit 5 (early synrift)

basement

Fig. 21. Summary of drilling (ODP Leg 103) and diving results across the Deep Galicia margin (DGM). Bottom: basic structure of the margin (in depth, modified after Reston et al. 1996) where the sampling took place. Note the division of the margin into seismostratigraphic units based on relationship to local faulting. Unit 5a: deep-water sediments rotated within the fault blocks and hence interpreted as early synrift – note that this sequence is not constant thickness, and thus not prerift; 4b: syntectonic wedges (nowhere sampled), deposited during motion on the adjacent fault; 4a: infill of half-graben topography after movement on a local fault but probably before breakup. Top: sampling results, defined in terms of the seismo-stratigraphic units: note how unit 5a is clearly diachronous – at Site 640 it is of the same age as the overlying unit 4a further east (Site 641). From this we can deduce at least three phases of faulting: one predating unit 5a across the entire section; one post-dating unit 5a in the east but predating it in the west; and one post-dating unit 5a in the west. Timescale (Gradstein et al. 1994; Pe´rez-Gussinye´ et al. 2003) shows duration of faulting in the Galicia Interior Basin (GIB), and the east and west DGM.

phase of faulting west of 639 also explains the dramatic drop in the size of the fault blocks west of 639. Finally, the latest fault blocks cut down to one of the earlier array of faults that formed a tectonic CMB and thus became a weak serpentine detachment (S): the last two phases of faulting at the DGM accompanied the serpentinization of the uppermost mantle (crust entirely brittle once thinned to 7 km) and the development of the S detachment.

The restoration with two phases of faulting is, however, still only a partial restoration. Movement occured along the faults that controlled the deposition of the deep-water units drilled at Site 639, and 639 has still yet to be restored; as noted above the geometry of these faults is not known and they may run oblique to the margin and the profiles. It is worth noting that the restored section has a similar thickness to the thinnest crust under

104

T. J. RESTON

Dive Gal-11

(a) 5 km

Proj. 640

638

IV

(b)

serpentinized mantle

base of postrift section

III base of late synrift (Aptian ?)

(c) base of early synrift - Valanginian - Barremian ? latest faults

(d)

II

S

earlier faults - Hauterivian?

I-IV

I

timing of deposition of latest synrift wedges

how much earliest synrift ?

F

(e)

serpentinizing mantle

130 % extension removed

space problem - light grey blocks need to be restored further before pushing back

Fig. 22. Partial restoration (modified from that of Reston 2005) of GP101 (Fig. 3), carried out sequentially (youngest faults to the west) Even restoring this section assuming an oceanward progression of faulting in a rolling hinge model it is clear that S was active at low angle. Misfit in E means that the section west of 639 needs to be further thickened before restoring movement along fault I. Reston (2005) suggested that movement along the interpreted earlier faults (black dashed) could solve this problem (F).

the GIB, where both the structural geometry and velocity structure require two phases of faulting. Thus, at the DGM, the stratigraphic and structural evidence indicates two phases of faulting (one for the entire DGM section, one just west of the 639 block) since the deposition of the early synrift section. The double restoration gives a crustal thickness of approximately 7 km, implying that all the faulting occurred after the onset of mantle serpentinization and was associated with

the development of the S reflector. Further landward (e.g. below the 639 high), the S reflector is not present. Within the GIB, the velocity structure and seismic image are consistent with two, but not with one, phase of faulting; the double restoration increases the thickness from about 8 to about 30 km. The rift history of the Galicia margin can be summarized. Rifting started in the Tithonian – Berriasian and two phases of faulting led to the formation of the Galicia Interior Basin and a similar

Fig. 23. Simplified 2D cartoon illustrating the key steps in the progressive extension of the lithosphere leading to continental break-up. (a) Early rift conditions up to the initiation of a second generation of faults (dashed). Initial decoupling zones at the base of the upper and midcrust (grey) thin, and the upper crust becomes increasingly coupled to the lower crust in the centre of the rift. Extension of the strong core of the lower crust and of the strong mantle lid (cross-hatching) is by necking/boudinage, the latter controlled by outward dipping shear zones that displace deep mantle rocks towards the rift axis. (b) An extension factor of 4 in the rift axis. The entire crust has now become brittle, mantle serpentinization (dark grey) has started and a weak detachment develops at the base of the crust. A third generation of faults cut down to this detachment (bold). (c) Crustal separation is imminent and will expose the underlying mainly deep lithospheric mantle. (d) After crustal separation, continued unroofing of the lithospheric mantle may require displacement along landward-dipping shear zones of such mantle from beneath the edges of the rift towards the COT. Complications owing to late igneous activity, 3D nature of extension and order of movement along faults (e.g. rolling hinge) are not shown for reasons of clarity.

RIFTED MARGINS BY PROGRESSIVE EXTENSION

Fig. 23.

105

106

T. J. RESTON

deep basin at the future DGM by the Hauterivian. Rifting now stopped within the GIB and on the west flank of the Galicia Bank, and focused instead within the deep basin to the west. Two more phases of faulting accompanied by the serpentinization of the uppermost mantle and the development of the S detachment here led to final crustal separation and the unroofing of mantle peridotites within the COT. Sea-floor spreading probably started in the Aptian.

Closing remarks As the preceding discussions have shown, I suggest that during progressive extension leading to continental breakup at a cool non-volcanic margin, such as west Iberia, lithospheric extension evolves through a number of main steps (Fig. 23). It is this evolution that distinguishes continental margins from continental rifts, which are subject to far smaller degrees of extension. Particularly important are polyphase faulting and the rheological evolution of the lithosphere (with implications for DDS, for crustal embrittlement, for mantle serpentinization and for final crustal separation along detachment faults rooting in serpentinized mantle). † Extension of rheologically layered lithosphere: faulting in the upper crust, and boudinage of both the lower crust (some component of crustal DDS) and of the uppermost mantle. Extension may focus where lower crustal and mantle boudinage are in phase. During this phase, DDS of the whole lithosphere may start through the displacement of the deepest lithospheric mantle towards the basin antiroot along outward-dipping shear zones. Symmetry at this stage is likely to be controlled by strain rate, and by heterogeneities in initial structure. † At an extension factor of 2, the original normal faults should lock-up and new faults develop (Fig. 23a). A failure to recognize such polyphase faulting may explain the extension discrepancy. At about the same time, or shortly after, the lower crust becomes increasingly coupled to the upper crust. When this coupling is complete, lower crustal boudinage is replaced by faulting, and DDS within the crust will be limited to local asymmetries associated with the largest offset faults. Coupling between the crust and mantle should increase, and displacement of lowermost lithosphere continues towards axis of rift. † Extension focuses strongly towards the site of eventual breakup as the crust thins to about one-quarter of its original thickness (Fig. 23b). At this stage, the second generation of faults lock-up and a third phase of faulting starts. This is accompanied by whole crustal

embrittlement and complete coupling of the crust to the mantle. The new faults may thus cut through entire crust into the mantle leading to serpentinization and the development of (asymmetric?) serpentine detachments (Fig. 23c). The serpentinization is effective in weakening the strong mantle lid. † Crustal separation and unroofing of deep subcontinental mantle oceanward of the last crustal blocks. Although melt generation might be expected during this stage, it is partially suppressed by a combination of mantle DDS and cool subcontinental mantle; continued unroofing of mantle may be controlled by successive landward-dipping shear zones (Boillot & Froitzheim 2001), which provide a mechanism for the displacement of deep sublithospheric mantel into the widening COT. † The onset of sea-floor spreading is probably controlled by either the final necking of the deep lithospheric mantle within the COT or by the influx of warm asthenosphere. In both cases sea-floor spreading is likely to substantially post-date crustal separation, as indicated by the presence of a broad zone of unroofed mantle within the continent–ocean transition (COT) at several margins (Fig. 1). At the West Iberia margin and at the preserved Alpine margins, we see evidence for all of the above processes. The apparent discrepancy between extension determined from crustal thinning or subsidence and that measurable from faults may be explained if extension is accommodated along multiple generations of faults that are incompletely resolved and which probably accommodated variable extension directions. As the amount and duration of extension increases towards the COT, so does the number of fault generations required: as not all of these can be fully imaged, let alone interpreted, an apparent extension discrepancy results. At the West Galicia margin, multiple phases of faulting are evidenced by the presence of deep-water sediments tilted within the fault blocks (and hence predating the faults that bound those blocks) and supported by structural reconstructions. At the IAP margin to the south, structural reconstructions and the tilting history of fault blocks also document the importance of polyphase faulting. This does not mean that at the West Iberian margin no crustal DDS occurred (e.g. through lower crustal necking and boudinage), just that it is not required to explain the extension discrepancy. Indeed, by analogy with the Alps and the basins offshore the UK, lower crustal boudinage probably did occur during the early evolution of the Iberian margin (Manatschal et al. 2001). Nor does it mean

RIFTED MARGINS BY PROGRESSIVE EXTENSION

that at other margins crustal DDS (including large-scale lower crustal flow) may not be required, but does suggest that polyphase faulting has to be rigorously ruled out before invoking these. At both the West Iberian and Alpine margins, the distribution and thickness of serpentinized mantle beneath the feather edge of the continental crust matches well the predictions of a simple numerical model for the embrittlement of the crust. The thickness of the partially serpentinized zone also matches the predictions of the model, particularly if a transition from pure shear to simple shear occurs when serpentinization starts (Pe´rez-Gussinye´ & Reston 2001). Detachment faults that appear to have been active at low angles, and which indicate a change from crustal-scale pure shear to crustal-scale simple shear, are observed and appear to follow the upper portion of the serpentinized mantle over much of their length. The presence of subcontinental lithospheric mantle unroofed within the COT implies heterogeneous stretching (DDS), which may be controlled by landward-dipping mantle shear zones similar to those imaged offshore the UK and beneath the West Iberian margin. Such DDS may have contributed to the reduction in melting at depth: again by analogy with the extended basins of NW Europe, this heterogeneous mantle extension probably started early during the rift history, and continued until final lithospheric separation and the onset of sea-floor spreading. Finally, the lack of significant igneous activity within the COT may be related to the presence of cool sublithospheric mantle in this region prior to the onset of seafloor spreading. Although there is undoubtedly both an apparent extension discrepancy and lithospheric-scale DDS at many margins, it remains difficult to explain such a discrepancy by DDS alone. At the West Iberian margin, polyphase faulting did occur and may suffice to explain the observed margin structure. However, other margins such as the Exmouth Plateau margin (Driscoll & Karner 1998) remain problematic. Here, the upper crust appears little extended, yet the margin has subsided several kilometres. Some other, as yet unknown, process may be important. This work has been supported by DFG grants Re 873-1, 873-3, 873-6, 873-7 and 873– 8, for which I am very grateful. Contributions by J. Hoffman, D. Klaeschen, C. Ranero, C. Krawczyk, J. Phipps Morgan and, in particular, M. Pe´rez-Gussinye´ made the work possible, although not all of these may agree with every conclusion reached. I would like to thank the organizers of the IMEDL workshop for organizing a superb meeting and especially G. Manatschal and O. Mu¨ntener for wonderful field excursions. Reviews by D. Sawyer, T. Minshull and N. Kusznir helped tighten the arguments.

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Roles of lithospheric strain softening and heterogeneity in determining the geometry of rifts and continental margins R. S. HUISMANS1,2 & C. BEAUMONT1 1

Geodynamics Group, Department of Oceanography, Dalhousie University, Halifax (Nova Scotia), Canada, B3H 4J1

2

Current address: Department of Earth Sciences, Bergen University, 5006 Bergen, Norway (e-mail: [email protected]) Abstract: Plane strain thermo-mechanical finite-element model experiments are used to investigate the effects of frictional–plastic strain softening and inherited weakness on the style of lithospheric extension. The model results are compared with the Newfoundland– Iberia conjugate rifted margins with the goal of understanding the lithospheric properties that controlled their evolution during rifting. Our proposition is that coupling between the plastic–viscous layering, acting together with frictional– plastic strain softening localized on inherited weak heterogeneities, can explain the initial wide rift and distributed rift basins that are later abandoned in favour of a narrow rift in which mantle lithosphere is exhumed to the surface. The models comprise uniform composition viscous and plastic layers in which focused deformation is nucleated on either a single weak ‘seed’ or a statistical white noise distribution of inherited strain. Strain softening of frictional–plastic layers acts as a positive feedback mechanism that creates localized shear zones from the inherited weak heterogeneities. The sensitivity of deformation to the choice of softening parameters and the type of inherited noise is examined in cases where the deeper part of the crust is either weak or strong. Lithosphere-scale models with a single weak seed exhibit a range of asymmetric and symmetric rifting modes that are mostly determined by the feedback between two primary controls, coupling between the plastic and viscous layers and strain softening. Decreasing and increasing the rifting velocity can change the mode, and asymmetry is strongest in models with low rifting velocities and a strong lower crust. Analysis of equivalent simple-bonded plastic– viscous two-layer models using the minimum rate of dissipation principle demonstrates that the mode selected depends on the division of the dissipation between the layers. Criteria developed on minimizing the total dissipation show how mode selection changes with increasing viscosity, or rifting velocity, from the: asymmetric plug or half-graben (AP) mode; through the symmetric plug or graben (PS) mode, to the distributed pure shear (PS) mode. Numerical models confirm these results. Models with statistical white-noise-inherited strain have similar modes to those with a single seed. In addition, modes with multiple sets of shear zones develop in the plastic layer for a range of intermediate parameter combinations. We believe that distributed noise in combination with a weak lower crust and slow extension can produce model results in accord with general features of the Newfoundland– Iberia conjugate margins; an initially distributed wide rift mode, followed by a late-stage narrow rift with a significant component of mantle exhumation.

Despite the large number of studies of continental rifts and rifted continental margins, many of which have produced high-quality data (Keen et al. 1987a, b, 1989; Mutter et al. 1989; Boillot et al. 1992; Sibuet 1992; Brun & Beslier 1996; Louden & Chian 1993; Dean et al. 2000; Funck et al. 2003, 2004; Hopper et al. 2004), we still lack a unified understanding of the mechanical and thermal processes that control their extensional geometry. This problem is compounded by the wide range of styles that include both non-volcanic and volcanic types, in which the latter exhibits direct evidence of voluminous mafic extrusives, and/or indirect evidence of intruded and underplated mafic magmas. In addition, the relative importance of active v. passive rifting, that is whether rifting is driven by active mantle upwelling

or whether mantle upwelling is a passive response to lithospheric extension (Salveson 1978; Sengor & Burke 1978; Turcotte & Emmerman 1983; Ruppel 1995; Huismans et al. 2001) remains to be determined. Important contributions to the problem of passive rifting leading to non-volcanic rifted margins defined possible end-member styles of lithospheric extension and their associated geometries, pure shear (McKenzie 1978), simple shear (Wernicke Burchfield 1982; Wernicke 1985) and combinations of these styles, for example Lister et al. (1986), referred to here as compound models (Fig. 1a–c). These kinematic models subsequently served as templates for the interpretation of observations. A primary focus in these interpretations has been whether rifting is symmetric, corresponding to the pure shear template, or

From: KARNER , G. D., MANATSCHAL , G. & PINHEIRO , L. M. (eds) Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 111– 138. DOI: 10.1144/SP282.6 0305-8719/07/$15.00 # The Geological Society of London 2007.

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Narrow Rift Mode (b) Asymmetric ‘Simple Shear’ Models

(a) Symmetric ‘Pure Shear’ Models Crust

Crust

Mantle

Mantle

Asthenosphere

Asthenosphere

(c) Compound Models Crust Mantle Asthenosphere

(d)

Wide Rift Mode

Crust Mantle

Asthenosphere

Fig. 1. End-member styles of rifting; symmetric, asymmetric and compound narrow models, and the wide rift mode. What controls these models of lithosphere extension; wide v. narrow, symmetric v. asymmetric? Can mode transitions be expected between asymmetric and symmetric modes, and between wide and narrow rifting?

asymmetric as predicted by the simple shear template (Fig. 1a, b). More generally, our lack of understanding of passive rifting at non-volcanic continental rifts and rifted margins can be expressed in terms of the following questions. † Can the geometrical styles of natural continental rifts and rifted margins be classified according to modes, such as the narrow, wide and core complex rift modes proposed by Buck (1991) (Fig. 1 a –d). † If so, how do the kinematic template styles of extension listed above, and their associated symmetries and asymmetries (Fig. 1), relate to the modes proposed by Buck (1991) (Fig. 1)? † If natural rifts can indeed be classified according to modes of extension, what controls the selection of these modes? † Conversely, if natural rifts exhibit various geometries that defy any simple mode classification, what controls this variability? Numerical modelling has the potential to provide at least partial answers to these questions. Model experiments can be designed to test the

role of a range of factors and processes on the styles of lithospheric extension. If the models are reasonable, albeit simplified, representations of their natural counterparts, the model results can indicate which processes are most likely to be important and how processes compete and tradeoff in different circumstances. We have chosen to focus on aspects of the rifting problem that we believe were important in the development of the Newfoundland–Iberia conjugate margin system and we address this specific application at the end of the paper. More generally, observations confirm that faulting and deformation on shear zones are important during rifting. At the small scale, deformation on faults and shear zones is inherently asymmetric. However, the question of the role of shear zones on the large-scale asymmetry of passive margins is still unanswered. This is partly a result of the lack of good constraints on the structure of the continental mantle lithosphere and its evolution during extension. Nevertheless, data do point to asymmetric geometries of conjugate margins (Boillot et al. 1992; Sibuet 1992; Louden & Chian 1999), and, for example, large-scale low-angle detachments and features

GEOMETRY OF RIFTS AND RIFTED MARGINS

interpreted to be high-angle shear zones that extend into the upper mantle (Brun & Gutscher 1992; Whitmarsh et al. 2001). There is some indication of patterns, or modes, in margin geometries. The significant variability in crustal-scale geometry that commonly occurs over quite short distances along strike on many margins poses an additional problem (see, for instance, the Iberia –Newfoundland margin system: Keen et al. 1987a, b; Torne et al. 1994; Funck et al. 2003; Dean et al. 2000; Pe´rez-Gussinye´ et al. 2003). Either this strike variability, or margin segmentation, is an inherent and perhaps predictable consequence of the mechanics of rifting in three dimensions, or it represents an effect of inherited inhomogeneous lithospheric properties, ‘noise’, in the system. For these reasons our general focus in the models described here is on processes that create shear zones and lead to mode selection, and on the effects of noise. We first assume that the lithosphere can be represented by a prototype uniform-layered model comprising frictional –plastic and viscous layers. We then consider the effects of different factors or processes on the style of extension of this model lithosphere. This approach is designed to answer the following particular questions. † What is the effect of strain-softening of the frictional –plastic parts of the system on the geometry during extension and does this feedback into mode selection? † Can the mode selection be explained by a general governing principle, specifically the minimum rate of dissipation of energy? † What is the effect of inherited heterogeneity, ‘noise’, in the model crustal properties on the style of extension? Does mode selection still occur in the presence of noise, or does noise lead to incoherent extension that is merely an expression of the noise? In section on ‘Thermo-mechanical models of lithosphere extension’ it is shown that strainsoftening leads to mode selection, and to asymmetric and symmetric styles of lithospheric extension. By analysing simple two-layer models we then demonstrate in the section on ‘Analysis of simple models using the principle of minimum energy dissipation’ that the minimum dissipation principle can be used as a guide to mode selection, at least in the numerical models. The section on ‘systems with statistical heterogeneity’ contains a preliminary treatment of the effects of distributed noise in the system on mode selection, and on how noise can lead to wide and narrow styles of lithospheric extension. Lastly, in ‘implications for the Iberia –Newfoundland conjugate margin system’ we ask whether what has been learned from the numerical models can be applied to the

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Newfoundland –Iberia conjugate margins to determine the controls on the tectonic style of this system. These sections are prefaced by brief reviews of processes that result in strain-dependent strength of rocks, the role and effect of inherited heterogeneity, and the two ways, deterministic and statistical, that this heterogeneity is introduced into the models. The section on ‘Thermo-mechanical numerical model’ describes the numerical models.

Strain-dependent strength of rocks and its representation in the numerical models The strength of faults and shear zones in the crust and mantle lithosphere on geological timescales is a controversial subject. They are interpreted to be either much weaker (Che´ry et al. 2001; Provost & Houston 2003) or of similar strength (Scholz 2000) to adjacent relatively non-deforming regions. Here, we accept that faults may be very weak and outline mechanisms that produce weak faults and shear zones. We then describe a simple parametric representation of these processes that we use in the numerical models. The frictional, or brittle, strength of dry rocks is represented to a first approximation by the Coulomb yield criterion for incompressible deformation in plane strain: ðJ20 Þ1=2 ¼ p sin f þ c cos f

ð1Þ

where (J20 )1/2 is the second invariant of the at yield deviatoric stress tensor, c is the cohesion, f is the internal angle of friction and p is the dynamic pressure (mean stress). Several mechanisms may contribute to strain softening, that is a reduction of the strength of the rock (deviatoric stress at yield, equation 1), with increasing strain but their relative and absolute importance is poorly constrained. At the scale of the crust, the brittle strength is dominated by the pressuredependent term (equation 1). Consequently, cohesion loss results only in a limited reduction of the strength of the brittle layer in the order of 2–10% depending on the ambient dynamical pressure and the depth of the brittle–ductile transition in the crust (Fig. 2a). Other mechanisms, such as formation of gouge and mineral transformations which reduce f with increasing deformation, can potentially result in much larger frictional strength reduction. For example, experimental work (Bos & Spiers 2002) on quartz–feldspar aggregates and their analogues has shown that mineral transformations in granitic rocks in the presence of water may produce mica-rich shear zones with very low effective internal angles of friction (f ¼ 38–108), resulting in a strength reduction of the order of 50–80% (Fig. 2b). Similar deformationassisted mineral transformation mechanisms have

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(a) Cohesion loss 0

C = 20 MPa

φ = 30o

Depth (km)

C=0 10

Brittle-ductile transition 20

30 0

100

200

300

Yield Strength (MPa)

(b) Mineral transformations 0

Depth (km)

φdry = 30

o

10

20

φ = 2–12

30 0

100

o

200

300

Yield Strength (MPa)

(c) Pore fluid pressure weakening 0

φeff(dry) = 30 Depth (km)

o

10

φ

eff

(hydro.) = 18

o

20

φeff(~lithostatic) = 1

o

30 0

100

200

300

Yield Strength (MPa)

Fig. 2. Summary of the causes of weakening in frictional-plastic rocks: (a) cohesion loss results in limited strength reduction; (b) mineral transitions can produce phyllosilicates that have very low internal angles of friction; and (c) high pore fluid pressures, pf, can reduce brittle strength very significantly. feff  308 (pf ¼ 0, dry strength), feff  188 (pf ¼ hydrostatic), feff  18 (pf ¼ lithostatic).

been documented in mafic lower crustal and upper mantle lithosphere rocks in shear zones in the Ivrea zone, which formed during extension (Handy & Stuenitz 2002). Transient and long-term fluid pressure variations may be significant factors in weakening brittle rocks (Sibson 1990; Rice 1992; Ridley 1993; Streit 1997; Ingebritsen & Manning 1999; Connolly & Podladchikov 2000). Pore fluid pressure, pf, acts to reduce the dynamic pressure and p in equation 1 is replaced by the effective pressure, p 2 pf. It can be seen that the rock strength can become very small when pf  p. An approximate way to consider the effect of pore fluid pressures is to introduce an effective internal angle of friction feff such that p . sinfeff ¼ (p – pf) sinf ¼ p(1 – l)sinf, where l ¼ pf/p is the pore pressure ratio. (Note that this definition differs from the Hubbert – Rubey pore pressure ratio for which l ¼ pf/poverburden, where poverburden is the weight of the overburden.) For example, using this definition of feff, rocks with f ¼ 308 have feff  188 in crust with hydrostatic pore fluid pressure. When the full range of pore fluid pressure variations that may occur during tectonic deformation is taken into account feff may vary from dry, feff  308, through the hydrostatic value to strongly overpressured, feff  08 (Fig. 2c). We use this effective internal angle of friction approach and adopt reference initial values of feff ¼ 308, 158 and 78 in order to investigate the sensitivity of the model results to a range of pore fluid pressure regimes. We also include a simplified parametric form of strain-dependent weakening in the numerical model experiments by linearly decreasing the internal angle of friction with the second invariant of the deviatoric strain, (I20 )1/2. In most of our models strain softening occurs in the range 0.5 , (I20 )1/2, 1.5. By using this approach we avoid the details of the physical mechanisms of strain softening, which are not well known in general but we can determine the sensitivity of the models to a range of possible strain-softening mechanisms. Strain-rate dependent weakening has also been used (Behn et al. 2002) to represent the rate dependence of frictional strength, which is observed in laboratory experiments. This approach characterizes the frictional behaviour of faults on seismic timescales (Scholz 1990). In contrast, our modelling addresses fault strength variations on geological timescales; therefore, we focus on the strain dependence of material properties.

Role of inheritance and heterogeneity We believe that the style of deformation of the lithosphere during rifting reflects the interplay

GEOMETRY OF RIFTS AND RIFTED MARGINS

of extension of pristine lithosphere with the superimposed effects of inherited structural heterogeneities, which can include both anomalously weak and strong regions. Models of extension of pristine lithosphere will display deformation modes that reflect the continuum nature of the system. Uniform pure shear is one such mode. However, the observed styles of rifts and rifted margins are highly variable, suggesting that inherited heterogeneities, ‘noise’ in the system, have a strong influence on its deformation. We cannot argue that the lithosphere is pristine. Therefore, it is critical to understand whether the noise overprints the inherent modes, or whether the underlying modes persist despite the noise. The noise is even more important if it contributes to the positive feedback effect of strain weakening, as discussed in the previous section. A small inherited weak heterogeneity, which is initially the focus of mildly enhanced strain, can become the seed for major localized deformation through this feedback mechanism. We consider two types of interaction between inherited ‘noise’ and strain softening. In sections on ‘Thermo-mechanical models of lithosphere extension’ and ‘Analysis of simple models using the principle of minimum energy dissipation’ we focus on the behaviour of deterministic model systems that include a single weak spot, or seed. This seed is taken to represent the weakest single zone of inherited damage that controls the strain localization process and, therefore, dominates over all other preexisting heterogeneities. A sensitivity analysis enabled us to conclude that weak seeds need to have a critical size and are most effective in localizing deformation when placed in the strongest part of the system, the upper mantle lithosphere in our models, thereby causing the greatest strength contrast with the surrounding material. From this analysis, we conclude that the system will preferentially select this type of seed over others, and we therefore focus on seeds placed in the upper mantle lithosphere. The approach that employs a single deterministic weak seed fails to consider that the weakest point may be ignored in favour of a distribution of inherited weak heterogeneities that link together to form even weaker shear zones. In the sections on ‘Systems with statistical heterogeneity’ and ‘Implication for the Iberia–Newfoundland conjugate margin system’ we consider the behaviour of systems characterized by a statistical distribution of inherited weaknesses that both compete and co-operate by linking to form shear zones. In this type of system, the continuum deformation activates the noise to create nascent-linked shear zones. Multiple shears may develop, evolve and compete. These two types of initial weakness, a single deterministic location or a statistical distribution, can be interpreted to be equivalent to natural lithospheric systems that are either dominated by one major zone of weakness or an

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evenly distributed background level of ubiquitous damage in the form of minor faults or shear zones that have not developed into penetrative shears at the crust or lithospheric scale.

Thermo-mechanical numerical model Model description We use an Arbitrary Lagrangian–Eulerian (ALE) finite-element method for the solution of thermomechanically coupled, plane-strain, incompressible viscous–plastic creeping flows (Fullsack 1995; Willett 1999; Huismans & Beaumont 2003). The model is used to investigate extension of a layered lithosphere with frictional–plastic and thermally activated power-law viscous rheologies (Fig. 3). For details on the numerical method see Fullsack (1995), Willett (1999) and Huismans & Beaumont (2003). When the state of stress is below the frictional – plastic yield the flow is viscous. We use both linear Newtonian viscous materials (see later sections in this paper) and temperature-dependent non-linear power-law rheologies (see later sections in this paper) based on laboratory measurements on ‘wet’ quartzite (Gleason & Tullis, 1995) and ‘dry’ olivine (Karato & Wu 1993), where the latter also includes the pressure dependence of viscosity. The effective viscosity, h, in the power-law model is of the general form:   Q þ Vp 0 h ¼ A1=n  ðI_ 2 Þð1nÞ=2n  exp ð2Þ nRT where (I˙20 )1/2 is the second invariant of the deviatoric strain rate tensor (121˙ij0 1˙ij0 )1/2, n is the power-law exponent, A is a scaling factor, Q is the activation energy, V is the activation volume (which makes the viscosity dependent on pressure, p), T is the absolute temperature and R is the universal gas constant. A, n, Q and V are derived from the laboratory experiments, and the parameter values are listed in Table 1. The reference parameter values for wet quartz listed in Table 1 lead to a weak viscous lower crust. Models of this type are described as weak crust models. In other models crustal viscosity, h(wet quartz), is everywhere increased by a factor of 100 to achieve a crust that is totally in the frictional –plastic regime. Models of this type are termed strong crust models. The viscosity scaling represents a simple technique that creates either strong frictional lower crust or moderately weak viscous lower crust without recourse to additional flow laws, each with its own uncertainties. The scaling can either be interpreted as a measure of the uncertainty in the flow properties of rocks

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Thermo-Mechanical Model Setup Strain softening Crust: Wet quartz Frictional plastic ρ0 = 2800 kg/m3 V

x,z = 0 km 35 km

Weak seed: 12 x 10 km von Mises plastic ρ0 = 3300 kg/m3 o σy = 107 Pa TMoho = 550 C

7

φ(ε)

1

0.5 ε 1.5

Strong lower crust

V 0

Weak lower crust

85 km 120 km

Moho

o

φ =1 φ =7

Vb

Lithosphere and sublithospheric mantle Dry olivine Frictional plastic (see right) 3 ρ0 = 3300 kg/m

Wet quartz

Depth

Vb

o

1330 C

o

Isothermal mantle, Temperature, Ta = 1330 C

o

1330 C x = 1200 km

Dry olivine

120 0

z = 600 km

o

250

0

250

Stress [MPa]

Fig. 3. Model geometry showing crust and mantle lithosphere layer thicknesses, the weak seed, their corresponding properties and the velocity boundary conditions, V, with Vb chosen to achieve a mass balance. Extension is driven by velocity boundary conditions and seeded by a small plastic weak region. Velocity boundary conditions are chosen to achieve a mass balance in the system. The model has a free top surface and the other boundaries have zero tangential stress (free slip). Whether materials deform plastically or viscously depends on the ambient conditions. At yield, flow is plastic. Below yield, deformation is viscous. Sedimentation and erosion are not accounted for in the model apart from surface smoothing resulting in small amounts of surface diffusion. Eulerian grid dimensions in horizontal and vertical dimension, nx, ny are, respectively, 401 and 151. The initial temperature field is laterally uniform, and increases with depth from the surface, T0 ¼ 0 8C, to base of crust, Tm ¼ 550 8C, following a geotherm for uniform crustal heat production, A ¼ 0.9 mW/m23 and a basal heat flux, qm ¼ 19.5 mW/m22. The temperature increases linearly with depth in the mantle lithosphere and the sublithospheric mantle is isothermal at Ta ¼ 1330 8C. The boundary conditions are specified basal temperature, 1330 8C, and insulated, q ¼ 0 mW/m22, lateral boundaries. Thermal conductivity k ¼ 2.25 W m21 8C and thermal diffusivity, k ¼ k/rcp ¼ 1  1026 m2 s21. Densities of crust and mantle, respectively, at 0 8C are r0 ¼ rc(T0) ¼ 2800 kg/m23 and r0 ¼ rm(T0) ¼ 3300 kg/m23 and depend on temperature with a thermal expansivity aT ¼ 3.1  10258C21, r ¼ r0 [1 2 aT(T –T0)]. On the right: initial (solid lines) and strain-softened (dashed lines) friction angle, f, 78 ! 18 representative strength envelopes of strong and weak crust models when V ¼ 0.3 cm year21. Frictional-plastic strain-softening behaviour is shown at the top.

where flow is dominated by quartz (e.g. ‘wet’ or ‘dry’) or to represent strong lower crust dominated by minerals that deform in the frictional regime for the conditions (temperature, strain rate) chosen for the model experiments. The plastic (frictional or brittle) deformation is modelled with a pressure-dependent DruckerPrager yield criterion which, with suitable adjustment of constants, is equivalent to the Coulomb yield criterion for incompressible deformation in plane strain. Yielding occurs when: ðJ20 Þ1=2 ¼ p sin feff þ c cos feff

ð3Þ

where (J02)1/2 ¼ (12 sij0 sij0 )1/2 is the second invariant of the deviatoric stress, c is the cohesion, feff is the effective internal angle of friction, defined earlier. With appropriate choices of c and feff this yield criterion can approximate frictional sliding in rocks.

Plastic flow is incompressible. Strain softening is introduced by a linear decrease of feff (1) over a range of strain where 1 represents the second strain invariant, (I20 )1/2 (Fig. 3, Tables 1 and 2). We have tested the dependence of the model results on mesh resolution and on the range of strain over which weakening occurs. Increasing the mesh resolution leads to narrower shear zones and to earlier localization because strain accumulates at a higher rate. Moderate changes in resolution do not, however, alter the main character of any particular model result. Also, the models do not show a strong sensitivity to the range of strain over which softening occur. The tests indicate that the major control on the model behaviour are the threshold at which strain softening starts and the amount of softening that occurs. The strong positive feedback between softening and strain accumulation implies that once initiated strain softening occurs rapidly.

GEOMETRY OF RIFTS AND RIFTED MARGINS

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Table 1. Parameters for the lithosphere-scale thermo-mechanical models. Power-law creep parameters wet quartz (Gleason & Tullis 1995) and dry olivine (Karato & Wu 1993) Parameter Rheological parameters Angle of internal friction Cohesion Power-law exponent wet quartz Activation energy wet quartz Initial constant wet quartz Activation volume wet quartz Power-law exponent dry olivine Activation energy dry olivine Initial constant dry olivine Activation volume dry olivine Universal gas constant Thermal parameters Diffusivity Heat production crust Moho temperature Asthenosphere temperature Thermal expansion coefficient Surface temperature Densities Crustal density at 273 ºC Mantle lithosphere density at 273 ºC Sublithospheric mantle density at 273 ºC Dimensions and boundary condition Base of crust Base of mantle lithosphere Base of upper mantle Extension velocity

Symbol

Value

f(e) amount of softening c nqt Qqt Aqt Vqt nol Qol Aol Vqt R

20  106 Pa 4.0 223  103 J/mol 1.10  10228 Pa2n s21 0 m3/mol21 3.5 540  103 J mol21 2.4168  10215 Pa2n s21 25  1026 m3 mol21 8.3144 J mol21 8C21

k A Tm Ta aT T0

1  1026 m2 s21 0.9  1026 W m23 550 8C 1330 8C 3.1  1025 8C21 0 8C

rc(T0) rm(T0) rm(T0)

2800 kg m23 3300 kg m23 3300 kg m23

V

35 km 125 km 600 km 0.06, 0.3, 0.5, 10 cm year21 full rate

In addition to solving the equilibrium equations for plastic–viscous flows, we also solve the thermal evolution in two dimensions. The mechanical and thermal systems are coupled though the temperature dependence of viscosity and density, and are solved successively during each time step. Initial conditions and other thermal properties are given in Figure 3 and Table 1.

18 –78 and 28 – 158

Thermo-mechanical models of lithosphere extension We start with an overview of a more comprehensive set of models from Huismans & Beaumont (2002, 2003). The models selected here have frictionalplastic strain softening, and deformation is nucleated by deterministic noise in the form of a

Table 2. Parameters for two-layer models Parameter Angle of internal friction Cohesion Viscosity lower layer Density Thickness of plastic upper layer Thickness of viscous lower layer Horizontal length scale Extension velocity

Symbol f(1) amount of softening c H r hp hv L V

Value 28 – 158 0 1  1021, 1022, 1023 Pa s 3000 kg m23 60 km 60 km 600 km 1 cm year21 (full rate)

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R. S. HUISMANS & C. BEAUMONT

single weak seed (Fig. 3). Most of the results can be classified as narrow rifts as defined by Buck (1991). We review the sensitivity of deformation style to the extension velocity and the strength of the middle and lower crust.

Sensitivity of model results to rifting velocity In order to demonstrate that rifting velocity is an important control on the extension mode we show results for a reference velocity of 0.3 cm year21 and for end-member variations of velocity of 0.06 and 10 cm year21 (Fig. 4). The models have strong middle and lower crusts, and include strain softening of the frictional –plastic rheology. At the reference velocity (Fig. 4b) strain softening in the frictional-plastic parts of the lithosphere results in strong localization in a single system of ‘faults’ (strictly shear zones), resulting in an asymmetric mode of extension. Middle and lower crust has been tectonically cut out and mantle lithosphere has been exhumed by large-scale frictional detachment. The proto-margins show distinct differences resembling an upper and lower plate conjugate margin pair. The main cause of asymmetry in the model is the frictional-plastic strain softening. After an initial phase of symmetric extension, feedback of strain softening results in preferential weakening of one of the two conjugate shears. The single weak shear zone remains the weakest part in the system and results in lithosphere-scale asymmetry. During later stages thermal advection produces a hotter, ductile lithosphere and a viscous necking style, and the final breakup phase is symmetric. However, asymmetry remains from the initial stages of the model evolution. At high rifting velocities (Fig. 4a) the style of extension is markedly different. The proto-margins are essentially symmetric. We have interpreted this model behaviour in terms of the increased role of viscous rheologies in the system and a stronger viscous coupling, where higher strain rates equate to higher viscous stresses at the base of the frictional layer (Huismans & Beaumont 2002, 2003). The viscous layer promotes a distributed, symmetric style of extension, and the tendency for localization and asymmetry given by the frictional–plastic strain softening is suppressed. The effect of a low end-member value for the extension velocity, v ¼ 0.06 cm year21, is illustrated in Figure 4c. The model is strongly asymmetric throughout its evolution. At this velocity thermal conduction is more efficient than thermal advection (Peclet number ,1). Consequently, the thermal evolution in the model lithosphere is close to the conductive limit, and the rift zone is relatively cool and deforms in the frictional regime. The

frictional weak shear zone therefore remains the single major weakness in the lithospheric system that allows for ongoing asymmetric extension.

Sensitivity of model results to strength of the lower crust We next examine the sensitivity of the model behaviour to the strength of the lower crust. The same frictional strain softening occurs as in the reference model, and in addition viscous flow following a wet quartz flow law now occurs in the weaker lower and middle crust (Fig. 5a). During the initial stages the crustal asymmetry is diminished because the conjugate frictional shears sole out in the weak ductile lower crust. Asymmetry in the lower lithosphere is still apparent in the shear zone cutting the mantle lithosphere. At later stages a set of synthetic frictional-plastic crustal shear zones form on either side of the central rift zone, thereby creating blocks bounded by frictional shear zones that sole out in the lower crust and that propagate beneath the rift flanks. During the late stage of deformation thermal advection results in nearly symmetric ductile necking of the lower lithosphere. The reference model result for a strong lower crust is given in Figure 5b.

Factors controlling mode selection and asymmetry during extension These results indicate three fundamental rift modes for the type of model described above: (1) lithosphericscale symmetric rifting; (2) lithospheric-scale asymmetric rifting; and (3) asymmetric upper lithosphere rifting concomitant with symmetric lower lithosphere rifting. Strain softening is the fundamental cause of the asymmetry and in a purely frictional-plastic model this would always lead to a phase of asymmetrical extension. In more complex models that include a viscous lower crustal layer the degree of asymmetry accompanying strain softening also depends on the rate of extension and on the viscous strength of the lower crust. The following factors need to be considered in regard to the efficiency of the strain-softening feedback in the frictional-plastic layer. † Geometrical hardening (Fig. 6d) may occur as the model evolves, for example by the rotation of shear zones, so that the active shears are no longer geometrically viable. This type of hardening opposes the continued increase in asymmetry and, depending on the amount and range over which strain softening occurs, even very weak shear zones will lock and be abandoned. † The extent to which the distribution of frictional-plasticity dominates the model

GEOMETRY OF RIFTS AND RIFTED MARGINS

119

Strong Crust, Sensitivity to Velocity High Velocity, V = 10 cm/a t = 12 Ma, Δx = 120 km

(a) 0 –20

o

550 C

z (km)

–40 –60 –80

o

1000 C –100 o

1200 C

–120 400

500

600

700

800

700

800

700

800

Moderate Velocity, V = 0.3 cm/a t = 40 Ma, Δx = 120 km

(b) 0 –20

z (km)

–40

o

550 C

–60 –80 o

–100

1000 C

–120

o

1200 C 400

500

600

Very Low Velocity, V = 0.06 cm/a t = 370 Ma, Δx = 222 km (c)

0 –20

z (km)

–40 –60

o

550 C

–80 –100 –120 400

500

600 x (km)

Fig. 4. Rift-mode sensitivity to extension velocity. Strong lower crust models with frictional-plastic strain softening, showing deformed Lagrangian mesh, velocity vectors and sample isotherms, for the dashed area in Figure 3. Model layers from top down denote upper and lower crust, strong frictional upper mantle lithosphere, ductile lower lithosphere and ductile sublithospheric mantle. Scaling of quartz viscosity makes all three upper layers frictional-plastic with the same f ¼ 78. (a) High rift velocity model, V ¼ 10 cm year21 at t ¼ 12 Ma and 120 km of extension. Note symmetric extension. (b) Reference rift velocity model, V ¼ 0.3 cm year21 at t ¼ 40 Ma and 120 km of extension. Note asymmetric extension. (c) Low rift velocity model, V ¼ 0.06 cm year21 at t ¼ 370 Ma and 222 km of extension. Note strong asymmetry. Animations of model evolutions can be found at: http://myweb.dal.ca/huismans/ jgr-animations.html.

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Effect Strength Lower Crust (a) Weak Lower Crust, V = 0.3 cm/a t = 40 Ma, Δx = 120 km 0 –20 o

550 C

z (km)

–40 –60 –80

o

1000 C

–100 –120

o

1200 C 400

500

600

700

800

600 x (km)

700

800

(b) Strong Lower Crust, V = 0.3 cm/a t = 40 Ma, Δx = 120 km 0 –20 o

z (km)

–40

550 C

–60 –80 o

–100 –120

1000 C

o

1200 C 400

500

Fig. 5. Rift-mode sensitivity to strength of the middle and lower crust. (a) Weak lower crust model, with h(wet quartz) at t ¼ 40 Ma and 120 km of extension. No scaling of quartz viscosity makes middle– lower crust ductile. Note the asymmetric crustal necking and symmetric lithospheric mantle necking. (b) Asymmetric strong middle– lower crust model at reference velocity at t ¼ 40 Ma and 120 km of extension is given for reference (same as in Fig. 4b). Animations of model evolutions can be found at http://myweb.dal.ca/huismans/ jgr-animations.html.

lithosphere. A range of behaviours exists between models where by frictional-plasticity is the dominant process in the extending lithosphere, which exhibit the most asymmetry, and those where geometrically larger ductile regions promote symmetry. † The relative integrated strengths, as opposed to geometries, of the plastic and ductile regions. For example, ductile strength increases with strain rate, producing higher stresses at higher rifting velocities, whereas the frictional

strength is independent of strain rate. This strain-rate dependence of strength explains the velocity sensitivity (Fig. 6d). The strong ductile region suppresses the asymmetric deformation preferred by the overlying frictionalplastic layer; ductile symmetry dominates over plastic asymmetry (Buck et al. 1999, 2003; Huismans & Beaumont 2002, 2003; Huismans et al. 2005). † The thermal state of the model, which also contributes to the asymmetry by determining the

GEOMETRY OF RIFTS AND RIFTED MARGINS

(c) Geometric Distribution Brittle & Viscous Materials

(b) Geometrical Hardening

A

Hp

B

isotherm marking the brittle-viscous transition

Hp

isotherm marking the brittle–viscous transition

Bending related hardening

z

Hb

brittle layer thickness

A

B

σ yield stress

σ yield stress

yield stress brittle lithosphere

121

z

Large proportion frictional z strength

Small proportion frictional strength

(a) Main Cause of Localization: Strain Weakening S brittle lithosphere

A

B

Hp A

isotherm marking the brittle–viscous transition

z

(d) Strain Rate Sensitivity Viscous Coupling . εA strain. . ε rate

ε

B

yield stress

brittle lithosphere

A

B

brittle yield stress

Hp

S

yield stress

yield stress

B

loss of cohesive strength

loss of frictional strength

z

(e) Thermal Advection/Conductive Cooling . total strain ε

εA

yield stress

A

B

brittle-viscous transition depth

A z

increase in yield strength at constant temperature

brittle yield stress

brittle lithosphere

Hb A

B

B

Brittle Ductile Transition

z

decrease in yield strength at constant strain rate

Fig. 6. Factors affecting localization of deformation (modified after Buck et al. 1999). (a) Strain softening is the main cause of localization of deformation. Factors that reinforce or counteract localization are: (b) geometrical hardening resulting from bending counteracts continued localized deformation; (c) the geometric distribution of the strainlocalizing plastic layer, and the non-localizing viscous materials determines the relative importance of plastic and viscous domains. A large viscous domain promotes symmetric or pure shear modes. (d) Brittle–ductile coupling. The viscous layer in the coupled system promotes distributed deformation through viscous shear stress along the interface of the coupled system, where at high strain rates associated high viscous stresses resist localized deformation in the plastic layer and suppress asymmetry. (e) The relative importance of thermal advection and conduction determines the relative distribution of plastic and viscous domains as the model evolves. For low extension velocities the plastic layer retains its thickness which promotes asymmetry.

relative distribution of frictional –plastic and ductile regions as the model evolves (Huismans & Beaumont 2002, 2003; Lavier & Buck 2002). When extension is slow, the thermal Peclet number is small and the lithosphere cools as it extends, thereby continuously renewing the frictional –plastic layers as the model evolves. This promotes asymmetric behaviour by comparison with models that have a large thermal Peclet number (Fig. 6a).

Analysis of simple models using the principle of minimum energy dissipation Approximate analytical theory The first series of models displayed a range of behaviours that we interpreted in terms of factors that reinforced or opposed asymmetric rifting. Although this approach does provide a certain level of insight, there is no governing principle that guides

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the overall interpretation. In this subsection we simplify the models to consider only the most basic behaviour of a two-layer system, and take a different approach to the interpretation by demonstrating the use of the minimum rate of energy dissipation as a guide to model behaviour (Bird & Yuen 1979; Sleep et al. 1979; Masek & Duncan 1998). Huismans et al. (2005) used this approach, subject to the restrictions noted by Bird & Yuen (1979), to understand the factors that control localized asymmetric modes of deformation v. distributed symmetric modes by comparing the minimum dissipation analysis with equivalent finite-element models, which are themselves based on minimum dissipation. We review the basic case of a two-dimensional, plane-strain system comprising a uniform, strainsoftening, frictional-plastic layer overlying and bonded to a uniform linear viscous layer. The analysis predicts the rate of internal energy dissipation (henceforth termed the dissipation for brevity) during the early stages of extension for three possible modes, the pure shear (PS), symmetric plug (SP) and asymmetric plug (AP) modes (Fig. 7). Details of the analysis are given by Huismans et al. (2005). For this simple system there are two main contributions to the rate of internal energy dissipation. The first is from extension of the plastic layer and is the same for each of the modes except when the dissipation is reduced by strain softening of the localized shears in the SP and AP modes. The second is from the viscous layer and its size ranges from the contribution owing to pure shear, in the PS mode, to a combination of pure shear and shear of a boundary layer at the top of the viscous layer when there is differential extension between the two layers in the SP and AP modes (Fig. 7). If we consider PS as a reference mode (Fig. 7), we can then ask under what circumstances a different, lower dissipation mode exists and will be selected. It can readily be understood that any overall reduction in dissipation depends on the trade-off between the ‘gain’ of reduced plastic dissipation owing to strain softening in the SP and AP modes v. the ‘penalty’ of increased dissipation owing to development of the viscous boundary layer in the SP and AP modes (Fig. 7). By considering these trade-offs, mode selection can be predicted as a function of the layer properties and the rate of extension. The latter is important because the rate of dissipation in the plastic layer depends on the strain rate, whereas that of the viscous layer depends on the square of the strain rate, so mode selection will depend on the extension velocity, V (Fig. 7). The total internal dissipation for each mode, WT, is calculated approximately for a given choice of the

(a) Pure Shear Mode (PS) hp

Plastic Layer: C(ε) φ(ε)

hv

Viscous Layer: η

V

V

(b) Symmetric Plug Mode (SP)

V

hb

Viscous Boundary Layer

V

Pure Shear

(c) Asymmetric Plug Mode (AP)

S

hb

2V

2L

Fig. 7. The total rate of dissipation is estimated for each of the following modes of deformation: (a) pure shear; (b) symmetric plug; and (c) asymmetric plug modes. The geometry and velocity field assumed are based on simple approximations to each of these modes. A plastic strain-softening layer is overlying and bonded to a linear viscous layer. Both layers with respective thicknesses hp and hv, and total length 2L, extend by boundary velocity V applied at each side. For localized symmetric and asymmetric plug modes, shear resisting localized deformation in the viscous layers is assumed to occur in a boundary layer of thickness hb. Deformation below the boundary layer follows pure shear. For more details on the analysis see Huismans et al. (2005).

model parameter values such as layer thicknesses, layer properties, the amount of strain softening and the extension velocity, and the results are plotted as a function of, for example, the viscosity, h, of the lower layer (Fig. 8). This figure shows the dissipation for the SP and AP modes. A mode transition occurs between the AP and SP modes where the curves cross at the transition viscosity, hT. In this case, the SP mode becomes the lower dissipation mode for h . hT. Mode transitions only occur when the dissipation curves cross; in other cases one mode systematically has a lower dissipation than another for all values of the viscosity. The total dissipation, WT, of each mode is equal at a transition, therefore W1Plastic þ W1Viscous ¼ W2Plastic þ W2Viscous

ð4Þ

where subscripts 1 and 2 successively represent the AP, SP and PS modes in our example. Solving

GEOMETRY OF RIFTS AND RIFTED MARGINS

(a) A

WT WTotal

S

WT

Mode Transition

ηt

η (b)

ηt = 3 hb ρ g hp2 (sinφpss – sinφss) Lv V Viscous penalty force

(c)

(d)

(e)

ηt Lv V hb Lv η t ε Α

=

=

ΔFv

Plastic gain force

2

3 ρ g hp (sinφpss – sinφss) 2

3 ρ g hp (sinφpss – sinφss)

=

ΔFp

Fig. 8. Mode-transition criteria based on the principle of minimization of dissipation. Main results and main equations are given here. Full derivation can be found in Huismans et al. (2005). (a) Typical plot of total ˙A dissipation for asymmetric plug mode (W I ) and ˙ SI ) of deformation. For viscous symmetric plug mode (W layer viscosities lower than hT the asymmetric plug modes provides the minimum dissipation mode and the converse for higher viscosities. (b) Transition viscosity, ˙A ˙S hT, which can be derived by solving W T ¼ WT, is a simple function of viscous and plastic layer length scales, strain-softening parameters and applied boundary velocity. Lv ¼ L 2 þ 4hb (3hv – hb))/L is a geometrical constant with viscous layer length scales. (c)–(e) The equation for transition viscosity can be rewritten in terms of two differential forces: (1) the viscous penalty force, D Fv, incurred by forcing localized shear in the viscous layer as a consequence of localization in the plastic layer; and (2) the differential plastic gain force, DFp, resulting from strain softening in the plastic layer. Transition between asymmetric plug and symmetric plug modes, and between symmetric and pure shear modes of deformation depends on the trade-off between these two differential forces. See Table 2 for values of parameters used.

the set of equations in equation 5 yields analytical expressions for the transition viscosities, hT1 and hT2, which then shows how the transitions depend on the model properties (e.g. Fig. 8, equation 1, for the AP to SP transition: see Huismans et al. 2005, for derivation of the equations). These expressions are relatively simple if the analysis is

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confined to the plastic and viscous contributions, as in equation 5, but become more complex when additional contributions, such as dissipation against gravity and layer bending, are included. It can also be seen from equation 5 that these additional terms can be ignored if their contribution to both modes at a transition is approximately equal. Under this circumstance they contribute equally to both sides of the equation and can be subtracted. The results from the approximate analysis (equation 5) demonstrate that modes are selected in the order AP, SP and PS as the viscosity of the lower layer increases and all other parameters are constant (Huismans et al. 2005). This order appears to be robust for the early phases of extension for parameter values appropriate for lithospheric-scale extension, and also applies to models in which the upper layer strain softens by loss of cohesion (Huismans et al. 2005). The effect of ignoring additional contributions to the dissipation means that the transition viscosity estimates, hT, are approximate, but the errors do not affect the selection order in this particular case (Huismans et al. 2005). The overall results of this analysis are summarized in Figure 9, which shows how the modes are distributed in relation to the plastic and viscous controls. NSS, PSS and SS are, respectively, the non-strain-softened, partially strain-softened and strain-softened strengths of the plastic layer (the corresponding values of f(1) appear in equation 1, Fig. 8), and hT1 and hT2 are transition viscosities between AP and SP, and SP and PS modes, respectively. Arrows show how models typically evolve during strain softening of the plastic layer, when the other properties of the system, such as extension velocity, remain constant. The mode selection and mode transitions can be understood from a physical perspective in terms of the competition between the ‘plastic gain’ and ‘viscous penalty’ contributions, defined above, to the dissipation. For example, for h . hT2 the mode selected is always PS (Fig. 9) because the viscous penalty incurred by the boundary layer dissipation in the SP or AP modes (Fig. 7) is always greater than the plastic gain achieved by strain softening of the localized shear zones. This result explains the perplexing observation that for this viscosity range the model will extend by pure shear even when the upper layer could develop weak localized shears. This result demonstrates that the properties of the bonded laminate are clearly different from those of its constituent layers. For hT1 , h , hT2 the strain-softened symmetric plug mode, SP2, is selected during plastic strain softening, and for h , hT1 the AP mode can be selected.

Plastic Strength, C(ε) or φ(ε)

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NSS

SP1

SP1

PS

Symmetric Plug

Symmetric Plug

Pure Shear

Plastic Strain Softening

PSS

SS

AP

SP2

Asymmetric Plug

Symmetric Plug

ηΤ1

PS

Pure Shear

ηΤ2

Viscosity Lower Layer Fig. 9. Rift-mode space. Evolutionary diagram of mode selection behaviour predicted by dissipation analysis and confirmed by numerical models in Figure 11. NSS, non-strain softened; PSS, partial strain softened; SS, strain-softened materials. At viscosities less then hT1 frictional-plastic strain weakening leads to the asymmetric plug mode. At viscosities larger then hT2 localization in the frictional-plastic layer is completely suppressed and distributed pure shear occurs. At intermediate viscosities, the symmetric plug mode is predicted.

However, as is explained below, the mode selection also depends on the way the modes evolve, that is, it is path dependent.

Simple two-layer numerical models The next step is to present finite-element model experiments designed to test the approximate analytical theory described in the previous section. The numerical model is the same as that used in the section on ‘Thermo-mechanical models of lithosphere extension’, but it now has the simple twolayer geometry (Fig. 10) and rheological structure (Table 2). A strain-softening frictional –plastic layer (f (1) ¼ 28–158) overlies and is bonded to the uniform constant linear viscous layer. The weak seed is placed at the base of the plastic layer. The results are shown after 40 km total extension at V ¼ 1 cm year21 for viscosities, h, ranging from 1021 to 1023 Pa s (Fig. 11). At low viscosity (Fig. 11a) the AP mode is selected. Strain softening has focused the plastic deformation onto one shear and the extension is highly asymmetric. When

h ¼ 1022 Pa s (Fig. 11b) asymmetry is suppressed and extension in the plastic layer occurs along two symmetric shears in the SP mode. At high viscosity (Fig. 11c) local deformation in the plastic layer is completely suppressed and extension of both layers occurs by the PS mode. Huismans et al. (2005) present a greater range of numerical experiments, but the three results (Fig. 11) serve to illustrate the basic behaviour.

Comparison of the analytical dissipation results with the numerical models The results of the numerical models are in qualitative agreement with the predictions of the dissipation analysis in that mode selection changes from the AP mode through SP to PS as the viscosity of the lower layer is increased (Fig. 11). The results also indicate that the transition viscosities, hT1 and hT2, fall between 1021 and 1022 Pa s, and between 1022 and 1023 Pa s, respectively, for the parameter values (Table 2) used in the models, which is in agreement with the equivalent analytical

GEOMETRY OF RIFTS AND RIFTED MARGINS

125

Numerical Model Setup 0

Plastic Layer

Weak Seed

V 60

V

Viscous Layer 120 km 0

600 km

Strain softening 15

o

φ(ε)

o

2

0.5

ε

1.5

Fig. 10. Geometry and boundary conditions for numerical experiments of extension in simple two-layer models. Upper strain-softening frictional-plastic layer and lower viscous layer are both 60 km thick. Side boundaries extend with velocity V ¼ 0.5 cm year21. Extension is seeded by a small plastic weak seed. The model has a free top surface, and the other boundaries have zero tangential stress (free slip). Sedimentation and erosion are not included in the model. There is, however, a small amount of numerical surface diffusion. The Eulerian grid has 401 and 121 nodes in the horizontal and vertical direction, respectively. Strain softening is achieved in a simple manner by a linear decrease of the angle of internal friction with deviatoric strain.

results. The finite-element model experiments not only confirm the approximate analysis for the initial mode selection, they also demonstrate that these modes persist as the model evolves. The results also confirm our assertion that the additional contributions to the dissipation, which were ignored in the analysis but are included in the full finite-element solution, do not change the order of the mode selection, thereby providing reassurance that the approximate treatment is acceptable. The results also indicate that other modes, not considered in the analysis, are not strongly preferred for this range of viscosities. We do, however, know that modes comprising a series of plugs, not just a single plug, have levels of dissipation intermediate between the SP and PS modes. In addition, it can also be shown that the mode in which the plastic layer totally detaches from the viscous layer is not selected for the range of parameters we consider because the plastic dissipation ‘penalty’ incurred by shear on the plastic detachment is much larger than any of the possible dissipation ‘gains’ in the viscous or plastic layers for this particular mode. The overall results can be interpreted in several ways, each of which offers a different insight into the mode selection. The first interpretation describes the mode selection in relation to increasing viscosity of the viscous layer. When the viscous dissipation is negligible, the AP mode is chosen because

this provides the fastest route to minimize the plastic dissipation by strain softening only one, not two, shear zones. When viscous dissipation is larger, the mode selected will change from AP to SP at the transition viscosity hT2, and to PS at hT2, as explained in the previous section. In the second interpretation the results are described in terms of the trade-off between two differential forces shown, for example, by the second equation in Figure 8. The left-hand side of the equation has dimensions of force per unit length of the model and is termed the differential ‘viscous penalty force’ incurred by deforming in the localized AP mode as opposed to the more distributed SP Mode. The right-hand side of the equation is termed the differential ‘plastic gain force’, and is equal to the reduction in the force in the plastic layer when deformed in the fully strainsoftened AP mode by comparison with the PS mode. Mode transitions occur when these two forces are equal. The physical insight provided by this requirement is that the system behaviour is determined not by the intrinsic properties of either the plastic or viscous layer alone, but instead it is determined by the trade-off between the differential penalty/gain forces. This interpretation can also expanded to consider the effect of other factors, for example, extension velocity as noted in the results of the section ‘Thermo-mechanical models of lithosphere

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Low Viscosity Lower Layer, V = 1 cm/a t = 4 Ma, Δx = 40 km

(a)

z (km)

0

φ(ε) = 15-2

o

60

η = 10

21

Pa s

120 0

600 km

Intermediate Viscosity Lower Layer, V = 1 cm/a t = 4 Ma, Δx = 40 km

(b)

0 z (km)

o

φ(ε) = 15-2 60

η = 10

120

22

Pa s

0

600 km

High Viscosity Lower Layer, V = 1 cm/a t = 4 Ma, Δx = 40 km

(c)

0 z (km)

o

φ(ε) = 15-2 60

η = 10

23

Pa s

120 0

x (km)

600 km

Fig. 11. Two-layer models of lithosphere extension. Upper layer strain-softening frictional-plastic, angle of internal friction f decreases from 158 to 28 as second invariant of strain ((I02)1/2 ¼ 1 ¼ 0.5–1.5) increases; cohesion c ¼ 0, lower layer uniform viscosity. Different panels are for different values of viscosity. Vertical side boundaries extend at 0.5 cm year21. Panels shown at 40 km of extension. Weak seed, f ¼ 28 in the centre at the base of frictional-plastic layer seed deformation. (a) Model with lower layer viscosity h ¼ 1021 Pa s. Note asymmetric mode is selected. (b) Model with lower layer viscosity h ¼ 1022 Pa s. Note symmetric mode is selected. (c) Model with lower layer viscosity h ¼ 1023 Pa s. Localization in the plastic layer is completely suppressed and the pure shear mode is selected.

extension’ (Fig. 4) (and also those of Huismans & Beaumont 2003) in which asymmetric, AP mode extension occurs when the rifting velocity is 0.3 cm year21, or less, but exactly the same system exhibits SP mode extension when the rifting velocity is 10 cm year21. As shown by equation 2 of Figure 8, the higher velocity modifies the viscous dissipation such that the differential viscous penalty force will be equal to the differential plastic gain force at a transition viscosity that is a factor of 33 (10/0.3 from above) smaller than that for the lower velocity, a result also implied by

equation 1 of Figure 8. This modified transition viscosity is much smaller than that for the lower extension velocity. Consequently, the rapid extension model is located in the SP part of mode space (Fig. 9), whereas the model with the lower extension velocity is located in the AP part. This result demonstrates that the transition viscosities change with model properties. At even higher extension velocities the mode would change to PS. It follows from Figure 8, equation 1 that variations in any of the parameters on the right-hand side of the equation, particularly the velocity, V, the

GEOMETRY OF RIFTS AND RIFTED MARGINS

thickness of the plastic layer, hb, and the initial and final strength of the plastic layer, rgh2b sin(fSP)/2 and rgh2b sin(fAP)/2 will modify the transition viscosity between the SP and AP modes. This result has the potential to provide insight into natural systems. The third interpretation uses the concept of ‘dominant’ and ‘subordinate’ rheologies introduced by Huismans & Beaumont (2003). This concept is qualitative but can be made more precise by considering the differential plastic gain and viscous penalty forces. Except at a mode transition (e.g. Fig. 8, equation 2) one differential force is larger and can be interpreted to imply that the corresponding rheology ‘dominates’, whereas the other is ‘subordinate’. For example, if the viscous differential force is larger, the viscous rheology dominates and forces the model to select the mode that minimizes the viscous dissipation, and the converse. The concept provides a useful physical interpretation of the model results. The preferred mode for the plastic layer acting alone would be AP, because this is the fastest way to minimize the dissipation of this layer. This mode is not selected when the viscous rheology is dominant because the ‘dominant’ rheology forces the mode selection. The same concept applies to the velocity sensitivity of the model results described in the section on ‘Thermo-mechanical models of lithosphere extension’ (Fig. 4). At V ¼ 0.3 cm year21, or less, the plastic rheology dominates and the mode is therefore AP, but at V ¼ 10 cm year21, the viscous rheology dominates and the mode is SP.

Systems with statistical heterogeneity Statistical model of inheritance from earlier tectonic deformation We now investigate systems that are characterized by the second type of heterogeneity, statistical noise in the initial strain field. The previous two sets of models used a single weak seed to localize the deformation. As noted in the subsection on ‘Role of inheritance and heterogeneity’ earlier, this approach fails to consider that the weakest point may be ignored in favour of a distribution of inherited weak heterogeneities that link together to form even weaker shear zones. Models with statistical heterogeneities in the initial strain field are designed to represent inheritance of deformation from previous tectonic phases. They then exploit the positive feedback link between the inherited strain distribution and strain softening. Specifically, the second invariant of the deviatoric strain, (I20 )1/2, is initialized with white noise that has a truncated Gaussian distribution with a

127

mean value ((I20 )1/2 mid) set to half the strain at which strain softening starts ((I02)1/2 min), and with a maximum value just below (I20 )1/2min. Therefore, there is no initial strain softening in the model, but the noise preconditions parts of the model to be on the threshold of strain softening. We have also explored other types of noise, in particular fractal noise with various fractal exponents, and have found that the first order of the model behaviour is similar for a range of noise types.

Simple two-layer models with statistical heterogeneity The results of models that are equivalent to those of the subsection on ‘Simple two-layer numerical models’, but instead have statistical noise in the initial strain field, are shown for viscosities of the lower layer ranging from 1021 to 1023 Pa s (Fig. 12). The models all extend at V ¼ 1 cm year21 and have exactly the same noise distribution. At low viscosity (Fig. 12a) the AP mode is selected. At intermediate viscosity (Fig. 12b) deformation is more distributed, strain softening having developed on several competing shear zones. At high viscosity (Fig. 12c) deformation in the plastic layer is distributed and characterized by a large number of shear zones. These behaviours can be interpreted using the minimum dissipation principle. The AP mode with highly asymmetric deformation can be adopted when the viscous penalty contribution to the dissipation is small (Fig. 12a), and in this case deformation is accommodated on only one plastic shear zone, thereby maximizing the plastic gain through rapid strain softening. At intermediate viscosity the mode is similar to SP (Fig. 11), but several shear zones have developed (Fig. 12b). Evidently, the total dissipation is best minimized by the activation of several plastic shears even though some will strain soften faster than others. In this case the major advantage is the ‘gain’ of reduced viscous dissipation by comparison with the AP and SP modes. Similar reasoning suggests the SP mode with a single plug should be selected for viscosities somewhere in the range 1021 Pa s , h , 1022 Pa s. At high viscosity, 1023 Pa s, the model (Fig. 12c) is dominated by the viscous rheology, at least in a relative sense. Some plastic gain has been achieved by strain softening on the plastic shear zones, but it is at the expense that many shears had to be created. The major gain is in the corresponding reduction in the viscous penalty because there are now only short, low strain rate, segments of the boundary layer flow between the multiple plastic shear zones, not long segments that accommodate major offsets as in the AP and SP modes.

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Low Viscosity Lower Layer, V = 1 cm/a t = 5 Ma, Δx = 50 km

(a)

z (km)

0

φ(ε) = 15-2

o

–60

21

η = 10 –120 0

Pa s 100

200

300

400

500

600

400

500

600

400

500

600

Intermediate Viscosity Lower Layer, V = 1 cm/a t = 8 Ma, Δx = 80 km

(b)

z (km)

0

–60

η = 10

22

–120 0

Pa s 100

200

300

High Viscosity Lower Layer, V = 1 cm/a t = 11 Ma, Δx = 110 km

(c)

z (km)

0

–60

–120 0

23

η = 10

Pa s 100

200

300 x (km)

Fig. 12. Two-layer models of lithosphere extension with distributed noisy heterogeneity. White noise with a Gaussian distribution is applied to the initial strain. Viscosity of the lower layer forms the only difference between models. (a) Asymmetric mode selected for low viscosity. (b) Multiple plug mode with several competing shear bands occurs for intermediate viscosities. (c) Pure shear mode occurs for high viscosity. Note that behaviour is similar to the single seed models in Figure 11.

The overall results indicate that this heterogeneous noise, which is below the strain-softening threshold, does not dictate the behaviour of the system and that deformation similar to the AP, SP and PS modes can occur depending on the viscosity of the lower layer. The difference is that modes with multiple plastic shear zones are now easily excited for the viscosity range hT1 , h , hT2 because the inherited strain heterogeneity places many regions of the model close to the threshold for strain softening. These new modes modify Figure 9 by occupying part of the mode space previously taken by the SP and PS modes, such that a new transition viscosity, to the multiple plastic shear modes, occurs between hT1 and hT2, and the value of hT2 is

increased. The multiple plastic shear modes will be characterized by an increasing number of shears with increasing viscosity of the underlying layer. This behaviour occurs because the viscous penalty is reduced by reducing the spacing between the plastic shears, as noted above, and in the limit of high viscosity, the new hT2, the overall deformation will approach the PS mode.

Lithosphere-scale extension with statistical crustal heterogeneity In this section we investigate the role of statistical heterogeneity in the initial strain field in

GEOMETRY OF RIFTS AND RIFTED MARGINS

lithosphere-scale models equivalent to those of the section on ‘Thermo-mechanical models of lithosphere extension’ to determine whether mode selection will be modified. We assume heterogeneity, due to earlier phases of deformation, is thermally annealed in the mantle where temperatures are high and, therefore, only apply statistical heterogeneity to the crust. Note that strain softening does not occur in the viscous (ductile) regime so that models with a strong (plastic) lower crust will experience strain softening in this region but those with a weak, ductile lower crust will not. For this series of models (I20 )1/2 is initialized with white noise that has a Gaussian distribution with a mean value ((I20 )1/2 mid) set to the strain at which strain softening starts ((I20 )1/2 min), and with the same Gaussian distribution. The other properties of these models are the same as the equivalent ones in the section on ‘Thermo-mechanical models of lithosphere extension’ except that the extension velocity is 0.5 cm year21. When the lower crust is strong (Fig. 13) extension is at first pure shear. When the incremental strain augments the initial statistical strain to the threshold for strain softening, strain softening occurs in the frictional-plastic crust and the SP mode is selected. The position of the plug depends on the particular realization of the white noise, a behaviour that is equivalent to randomly placing a weak seed. Localized deformation in the crust causes diffuse projections of the PS mode shears to develop in the upper mantle lithosphere and these forced shear zones eventually localize by strain softening in the upper frictional –plastic mantle. Feedback of strain softening results in the development of an asymmetric shear zone that penetrates most of the lithosphere (Fig. 13a) which facilitates exhumation of the upper mantle lithosphere (Fig. 13b) via the AP mode. During later stages, when thermal advection weakens the lithosphere, viscous necking becomes the dominant style of deformation (Fig. 13c). This model behaviour is very similar to its single weak seed equivalent (Figs 4b & 5b). The difference in the extension velocity, 0.5 v. 0.3 cm year21, has only a minor effect on results. A more important implication is that models with this particular set of properties behave in a similar manner and are not particularly sensitive to the difference between inherited statistical heterogeneity and a deterministic weak seed. The result is also consistent with the simple two-layer model with statistical heterogeneity and a low viscosity lower layer (Fig. 12a). Evidently, in all of these examples, which are essentially two-layer cases, the plastic gain achieved by selecting the AP mode outweighs the viscous penalty for low values of the effective mantle viscosity. For higher effective viscosities

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of the mantle lithosphere we may expect more distributed modes of deformation like those seen in Fig. 12b, c. The style of deformation for the equivalent model with the viscously weak lower crust (Fig. 14) is markedly different, both from the model with strong lower crust and, more significantly, from the equivalent model with a single weak seed (compare Figs 5a and 14). At first deformation occurs in an overall ‘PS-type’ mode (Fig. 14a). Sets of parallel and conjugate frictional shears form in the frictional upper crust, but strain localization does not occur in the upper mantle lithosphere which has no initial statistical heterogeneity. After approximately 160 km of extension, localization in the crustal layer feeds back into the frictional –plastic upper mantle lithosphere and seeds a localized shear zone (Fig. 14b). The remaining evolution is similar to the equivalent model with a strong lower crust except that the final necking is more symmetrical and there is somewhat less exhumation of the mantle (Fig. 14c). The significance of this model is that it displays a natural progression from an early distributed wide rift mode to a late localized narrow rift mode. Although the equivalent model with the single weak seed also exhibits decoupling and differential shear between the crust and uppermost mantle (Fig. 5a), the width of the zone of distributed crustal deformation is much narrower and tends to increase as rifting progresses.

Implications for the Iberia – Newfoundland conjugate margin system First-order characteristics of the Iberia– Newfoundland conjugate margin system In this subsection we explore whether what has been learned from the models above can be applied to extension of the Iberia –Newfoundland conjugate margin system. One of the main characteristics of the Iberia –Newfoundland conjugate margin is the high degree of variability in the geometry of the margins and the location of extension with position along the margins (Figs 15 & 16). This variability may result from the competition between local inherited structures and underlying physical processes that control the mode of extension. We first attempt to identify the general features of the system because these may lead to insight into the fundamental underlying processes responsible for its formation. Secondly, we address the potential causes of the variability. Cross-sections (Keen et al. 1987a, b; Torne et al. 1994; Dean et al. 2000; Funck et al. 2003; Pe´rez-Gussinye´ et al. 2003) shown restored to

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Fig. 13. Strong lower-crust thermo-mechanical model of lithosphere extension with white statistical noise applied only to the crust. The model selects the narrow rift mode with strong asymmetry. Note that model behaviour is similar to the equivalent single seed model (e.g. Fig. 4b).

approximate the configuration at the time of onset of sea-floor spreading (Fig. 16) illustrate the strike variability over a distance of only 100 km. We regard the following to be key features of the Iberia–Newfoundland conjugate margin system. (1) Extension and rifting occurred during a

prolonged period of approximately 80 Ma with at least three phases of rift activity (Mauffret & Montadert 1988; Tankard & Balkwill 1989; Tankard et al. 1989; Murillas et al. 1990; Driscoll et al. 1995; Stapel et al. 1996). (2) Initial rifting occurred in a wide rift mode over an extended

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Fig. 14. Weak lower-crust thermo-mechanical model of lithosphere extension white statistical noise applied only to the crust. Model evolution: (a) wide rift mode selection, only localization in crust, very widely distributed deformation in the mantle lithosphere; (b) and (c) narrow rift mode selection, feedback of localization in crust into mantle lithosphere. Note the natural progression from wide to narrow rift mode. Model behaviour is dissimilar with equivalent decoupled single seed model (Fig. 5a).

area that was at least 400 km wide (Keen et al. 1987a, b; Tankard & Welsink 1989; Manatschal 2004). (3) The system became a narrow rift late in its evolution and continental mantle lithosphere was exhumed during the terminal phase (Boillot

et al. 1988; Louden & Chian et al. 1999; Dean et al. 2000). (4) The system is a non-volcanic rift zone, the late stages of rifting and early sea-floor spreading being remarkably a-magmatic (Louden & Chian 1999). (5) Extension occurred at very

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Fig. 15. The Iberia–Newfoundland conjugate margin system. Map of the Cretaceous Central Atlantic restored to magnetic anomaly M0 time (c. 121 Ma) based on the reconstruction pole of Srivastava et al. (2000) with Newfoundland fixed to present geographic co-ordinates (modified after Preliminary Results Ocean Drilling Program, Tucholke et al. 2004). Northern, central and southern cross-sections, indicated with thick black lines, are shown in Figure 16.

low average velocities, an important point that we substantiate below. (6) Localized extension during the first two phases of rifting was limited to the crust with only minor thinning of the mantle lithosphere as evidenced by limited post-rift subsidence of the early rift basins (Stapel et al. 1996). (7) Asymmetries in crustal thickness, distribution of exhumed mantle lithosphere (Whitmarsh et al. 2001) and spatial variability of mineral chemistry of the exhumed mantle lithosphere (Muentener 2004) suggest that late-stage rifting was accommodated by an asymmetric shear that cut down into the mantle lithosphere (see also Whitmarsh et al. 2001; Manatschal 2004). Timing, duration and associated extension of the rift phases are not well constrained. We have constructed bounds on extension velocity and duration of each of the rift phases using various permissible rift-phase reconstructions and estimates of the total amount of crustal extension. The total crustal extension is estimated using the northern conjugate margin transect (Funck et al. 2003; Pe´rez-Gussinye´ et al. 2003) as a reference. Total extension of this part of the margin, prior to the onset of sea-floor spreading, ranges from 180 to 220 km, based on the cross-sectional area balance between an estimated original crustal thickness of 32 km and the present-day thickness distribution of the continental crust. This range overlaps with that based on a similar estimate of 100 –200 km derived from cross-sections in the central part of the Iberia– Newfoundland conjugate margin system (Minshull et al. 2001), although it is clear their estimate has a large uncertainty. This total extension must be apportioned among the different phases.

The timing of the rift phases can be determined approximately based on subsidence, from well data and seismic profiles (Tankard & Balkwill 1989), as a proxy for extension: Phase I:

218–205 Ma – minor relief formation in external basins; Phase II: 165–142 Ma – major relief formation in external basins (e.g. Jeanne d’Arc and associated basins, Galicia Interior Basin, Lusitanian Basin); Phase III: 142–122 Ma – Narrow rifting and breakup of the crust; Phase IV: 122–116 Ma – Asymmetric exhumation of the mantle lithosphere. Previous estimates of the extension velocity are of the order of 1.5 cm year21 full spreading velocity (Minshull et al. 2001). The approximate timing of rift phases and the value of 200 km total extension given above can be used to make simple estimates that give upper and lower bounds for the average extension velocity of the system. If the extension velocity was the same throughout all of the phases the rate was approximately 0.3 cm year21 full spreading rate. If all of the extension occurred during the latter two phases the estimated velocity would be approximately 0.75 cm year21. We regard this estimate to be an unreasonably large upper bound because phases I and II achieved significant normal faulting and basin subsidence. We therefore regard 0.5 cm year21 to be a representative upper-bound estimate of the average extension velocity during the phases when extension occurred.

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Fig. 16. Compilation of crustal conjugate cross-sections based on seismic reflection data, and gravity inversion for the Iberia Newfoundland margin system, and comparison to numerical model application shown in Figure 17. Section locations indicated in Figure 15. (a) Northern transect (Murillas et al. 1990; Funck et al. 2003; Pe´rez-Gussinye´ et al. 2003). (b) Central Transect (Keen et al. 1987a, b; Keen & de Voogd 1989; Dean et al. 2000). (c) Southern transect (Keen 1987a, b; Keen & de Voogd 1989; Torne et al. 1994). (d) Crustal section from numerical model application for comparison at the same horizontal and vertical scale as the crustal sections. (e) Numerical model after 230 km of extension at 46 Ma Note the formation of exterior basins, the thin crustal wedge on top of the mantle lithosphere and exhumed mantle lithosphere in the ocean–continent transition zone.

In summary, from the discussion above, a successful model needs to meet the following constraints: † an initially wide rifting mode with extension/subsidence focused in a number of distinct, widely distributed, fault-bounded basins, with structure that varies with position along the system;

† Initial extension in these basins was mostly limited to the crust; † Abandonment of the external basins in favour of late-stage narrow, possibly asymmetric, rifting; † exhumation of the mantle lithosphere during the final phases of breakup; † An overall slow extension rate probably close to 0.5cm year21.

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Towards a first-order model with the characteristics of the Iberia– Newfoundland conjugate margin system In the context of the models described in this paper, the progression from a wide to narrow rift can be met by models that inherit heterogeneous crustal noise and in which the crustal noise develops into a wide region of extension characterized by weak upper crustal shears and faults that bound distinct basins. Horizontal decoupling and shear in the lower or mid crust appear to be necessary, in addition to the inherited heterogeneity, (cf. Fig. 14) because the wide rifting mode does not develop when the lower crust is strong and precludes decoupling (cf. Fig. 13). The progression from a wide to a narrow rift occurs naturally in this type of model if the uppermost mantle lithosphere strainsoftens, leading to localization of the deformation (cf. Figs 13 & 14). It also appears that models with a strong, coupled, lower crust show the greatest potential to exhume mantle lithosphere to the surface during the narrow rifting mode (e.g. Figs 5b & 13c) and that this exhumation is commonly achieved by asymmetric extension (cf. Fig. 4b, c). These behaviours are also compatible with slow rifting velocities that favour AP and SP modes over distributed pure shear, as described in the section on ‘Analysis of simple models using the principle of minimum energy dissipation’. The model shown in Figure 14 has many of the required general characteristics displayed by the Newfoundland–Iberia conjugate margin system, as listed above. However, it is a model with a weak lower crust and, therefore, fails to exhume the mantle to the surface during the late narrow rifting phase. To date, we have been unable to create a selfconsistent model that produces both initial wide rifting and later stage cooling that strengthens the viscous lower crust sufficiently to create the necessary conditions for very narrow AP or SP rifting and mantle exhumation. We suspect this inability occurs because we consider models in which extension is uniform and continuous, and, therefore, there are no tectonically quiescent periods during which significant cooling can occur. We demonstrate the effect of cooling and strengthening of the extended crust using a composite model. The rift evolution starts in a model with inherited strain heterogeneity and a weak middle and lower crust. We then simulate the effects of cooling and late-stage embrittlement of the crust by suppressing the viscous flow in the crust during Phase 2. This composite behaviour is obtained by increasing crustal viscosity by a factor 100 during Phase 2 to achieve a crust that is totally in the frictional–plastic regime. The extension velocity is

0.5 cm year21. During the 24 Ma-long Phase 1, 120 km of extension is accommodated by localization on multiple sets of frictional shear zones in the upper crust (Fig. 17a). Deformation of the mantle lithosphere occurs in a distributed mode. During Phase 2, localization in the crustal layer feeds back into the upper mantle lithosphere where it causes localization in the frictional–plastic uppermost mantle. Deformation is asymmetric (Fig. 17b) and mantle lithosphere is exhumed during the final phases of breakup (Fig. 17c). To first order, this model result is compatible with the constraints compiled for the Newfoundland–Iberia natural system (Fig. 16), particularly when the significant variability among cross sections located close together is taken into account (Fig. 16a– c). It also provides a potential explanation for this variability in the geometry of the margins, in particular the initial irregular distribution and later abandonment of early rift basins that flank the margins if these are the consequence of deformation related to inherited crustal heterogeneity. We anticipate that an equivalent fully three-dimensional (3D) model, with inherited initial heterogeneous strain, would develop sets of anastamosing rift basins during the early phase of rifting and that the structure of these basins would vary significantly along strike of the system. These preliminary results lead us to the proposition that the early stages of rifting in margins of this type are disorganized, reflecting inherited lithospheric characteristics. Only later in the rifting does the system overcome this inheritance and become organized in the form of a narrow rift.

Summary and conclusions This article has focused on models of extensional processes with application to lithospheric extension during rifting and the formation of rifted continental margins. Even though rifts and rifted margins have been studied intensively for more than 30 years, a fundamental understanding of the factors that control the extensional process, the distribution of extension and the associated rift and rift margin geometry is still lacking. The purpose of the model studies reported here and in associated articles (Huismans & Beaumont 2002, 2003; Huismans et al. 2005) is to test a limited range of conceptual models that combine distributed viscous–plastic flows with feedback mechanisms, specifically plastic strain softening, that lead to localized deformation. This localization allows the models to represent faults and shear zones approximately. In our earlier research, the development of localized shears was in part deterministic and was triggered by a priori inclusion of small weak regions, ‘weak seeds’, in

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Fig. 17. Model application to the Iberia– Newfoundland conjugate margin system. (a) Initial weak lower-crust thermo-mechanical model of lithosphere extension with statistical noise applied only to the crust. Wide-rift mode selection, only localization in the crust, very widely distributed deformation in the mantle lithosphere. (b) and (c) Model restarted with a strong lower crust, simulating a phase of cooling and strengthening of the lower crust. Narrow rift mode selection, feedback of localization in the crust into the mantle lithosphere. Note the progression from wide to narrow rift mode and model selects a narrow rift mode with strong asymmetry and exhumation of mantle lithosphere during the final stages of rifting.

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the models. In this paper we have reviewed those results and expanded the research to include the effect of distributions of weak heterogeneities that take the form of white noise in the initial strain field and can be interpreted to represent strain, or in a wider sense damage, that is inherited. The effects of the viscous-plastic layered rheology combined with the inherited initial strain lead to more general representations of the type of deformation that can occur during lithospheric extension. We believe this approach, which combines a continuum model based on plastic and viscous rheologies with strain softening that results from a statistical distribution of inherited heterogeneity, provides a potential explanation for the general features of the Newfoundland–Iberia conjugate margins. These margins appear to have had an early phase of distributed disorganized crustal extension, and only later became organized into a narrow rift that exhumed continental mantle lithosphere during terminal extension prior to the onset of sea-floor spreading. We wish to thank the organizers of the IMEDL workshop for the invitation to participate. This research was funded through an ACOA–Atlantic Innovation Fund contract, and an IBM-Shared University Research grant. C. Beaumont was funded by the Canada Research Chair in Geodynamics. C. Beaumont also acknowledges support through the Inco Fellowship of the Canadian Institute for Advanced Research. Numerical calculations used software developed by P. Fullsack. We thank R. Buck, L. Lavier and an anonymous reviewer for the constructive and thorough reviews.

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The role of gravitational instabilities, density structure and extension rate in the evolution of continental margins E. BUROV Laboratory of Tectonics UMR 7072, University of Pierre and Marie Curie, Paris, France Abstract: Formation of rifted continental margins is associated with localized thinning and breakup of the continental lithosphere, driven or accompanied by the ascent of the lithosphere– asthenosphere boundary. Thinning creates sharp density and viscosity contrasts and steep boundaries between cold deformed lithosphere and hot upwelling asthenosphere, thus providing conditions for the development of positive (asthenosphere) and negative (mantle lithosphere) Rayleigh–Taylor (RT) instabilities. The evolution of many continental margins (e.g. Newfoundland margin and Iberian margin) is characterized by slow spreading rates. This allows the RT instabilities to grow at the timescale of rifting. The impact of positive RT instabilities (asthenospheric upwelling) is well studied. The negative RT instabilities, associated with mantle downwelling, remain an overlooked factor. However, these instabilities should also affect the rift evolution, in particular, they may cause mantle thinning or thickening below the rift flanks. Our numerical experiments suggest that the ratio of the RT-growth rate to the extension rate controls the overall rift geometry and evolution. Even if the effect of negative RT instabilities is more important for slow extension rates of 2  5 mm year21 (Deborah number, De , 1), it is still significant for 2– 3 times higher extension rates of 2  15 mm year21 (De , 10). The numerical experiments for extension rates of 2  15 mm year21 and mantle–asthenosphere density contrasts of 10–20 kg m23 demonstrate a number of structural similarities with continental margins characterized by low De (e.g. Flemish Cap and Galicia margin). In particular, rift asymmetry results from interplay between the RT instabilities and differential stretching at De , 1. Formation of interior basins occurs at De  1 –3. The best correspondence with the observed geometry of rifted margins is obtained for chemical density contrast of 20 kg m23 and extension rate of 2  15 mm year21, which is twice that of the averaged values inferred from the observations. This suggests that margins may initially (prebreakup stage) extend at higher rates than the average extension rates characterizing rift evolution. The influence of RT instabilities is strongly controlled by extension rate, density, rheology and thermal structure of the lithosphere; this implies that we need better constraints on these parameters from the observations.

Continental margins result from rifting characterized by large coefficients of extension 4–5 , b , 15 (e.g. Bott 1971; Salvenson 1978; Cloetingh et al. 1982; Buck 1991). Rifting processes involve laterally variable thinning of ductile layers of different densities. If a denser layer (mantle lithosphere) is located on top of a lighter layer (asthenosphere), then the system is gravitationally unstable and may develop negative Rayleigh –Taylor (RT) instabilities. This possibility specifically refers to rifting, passive or active, where hot light asthenosphere ascends to the surface and replaces colder denser lithosphere. In most situations, normal lithospheric mantle is approximately 10 –30 kg m23 denser than the underlying asthenospheric layer, basically due to its colder temperature but also due to compositional differences (e.g. Stacey 1992; O’Reilly et al. 2001; Poudjom Djomani et al. 2001; Turcotte & Schubert 2002). Some authors (e.g. Houseman et al. 1981) assume even much higher total density differences of up to 100 kg m23. The density differences of the order

of 20 kg m23 are commonly accepted for the Phanerozoic lithosphere, even though there is still a debate about whether it applies to the presumably Mg-rich and -depleted cratonic lithosphere. Irrespective, volumetric seismic velocities, which are generally considered as proxy for density, are systematically higher in the lithosphere mantle than in the asthenosphere. Depending on its viscosity the mantle lithosphere therefore has the potential to sink as the result of a RT instability (e.g. Houseman et al. 1981). The growth rate of the RT instabilities is directly proportional to the density contrast and inversely proportional to the thickness and the viscosity of the upper layer (if the viscosity of the lower layer is small). Depending on that, the instabilities may be rapid or slow compared to the tectonic deformation rates. In the first case, they will influence the rift evolution, but not in the second. The conditions when the RT instabilities are slow apply to the cases when the viscosity of the mantle lithosphere is higher than 1022 Pa s, when the density contrast between the lithosphere

From: KARNER , G. D., MANATSCHAL , G. & PINHEIRO , L. M. (eds) Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 139– 156. DOI: 10.1144/SP282.7 0305-8719/07/$15.00 # The Geological Society of London 2007.

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and the asthenosphere is small, or when the interface between the mantle and the asthenosphere remains flat (e.g. Houseman & Molnar 1997). This applies to oceanic and continental plates at

normal conditions that infer smooth boundaries and smooth vertical temperature profiles and practical absence of lateral heat transport (e.g. prerift geotherm in Fig. 1a).

Fig. 1. (a) Simplified cartoon showing the interaction between the asthenosphere and mantle lithosphere during rifting. Left: the mantle lithosphere below the rift flanks becomes gravitationally unstable owing to the negative density contrast with the hot asthenosphere and the loss of mechanical resistance resulting from lateral heat transfer from the asthenosphere below the rifted zone. The asthenosphere is positively unstable because of its positive density contrast with the embedding mantle lithosphere. As the viscosity of the mantle lithosphere is exponentially dependent on temperature, there is a relatively net ‘stability level’ that separates the regions of high viscosity from thermally weakened regions of low viscosity. These regions are subject to development of rapid RT instabilities. Right: simplified yield-stress prerift envelope of the continental lithosphere and typical prerift temperature profile. (b) Left: potential density contrast (reference density at 750 8C) and viscosity of olivine mantle, as function of temperature (left). Viscosity is shown for three representative values of the background strain rate (10214, 10215 and 10216 s21). Right: time until full detachment of the unstable mantle layer, as a function of initial layer thickness (olivine rheology), computed according to Conrad & Molnar (1999). The initial perturbation amplitude is w0. Note that for a normal depth– pressure–temperature profile of mantle lithosphere, the density contrast due to thermal expansion is nearly negated by the effect of compressibility (Turcotte & Schubert 2002). This is not the case of rifting when the mantle is heated without compensatory increase in pressure.

THE ROLE OF GRAVITATIONAL INSTABILITIES

Contrary to normal lithosphere, the rifted margins such as the Newfoundland and Iberian margin (Manatshal & Bernoulli 1999; Manatshal et al. 2001; Funck et al. 2003; Pe´rez-Gussinye´ et al. 2003; Hopper et al. 2004), as well as many ‘common’ continental rifts, constitute a perfect environment for development of gravitational instabilities owing to the presence of steep density boundaries and viscosity contrasts resulting from rifting (e.g. Huismans et al. 1998) (Fig. 1a). As mentioned above, most data suggest the existence of a compositional or chemical density contrast of 0–20 kg m23 between the mantle lithosphere and asthenosphere (e.g. Stacey 1992; O’Reilly et al. 2001; Poudjom Djomani et al. 2001; Turcotte & Schubert 2002), except probably the cratonic mantle (our settings consider, however, a noncratonic lithosphere). In addition to potential chemical density contrasts, strong thermal density contrasts of 30– 60 kg m23 arise when the mantle–asthenosphere boundary is locally uplifted owing to the rifting. In this case, horizontal heat conduction becomes at least as important as the vertical heat conduction. For the temperature contrasts of 500 –1000 8C, associated with moderate coefficients of extension b ¼ 2–5 (McKenzie 1978), horizontal thermal gradients may reach 10 –20 8C km21. The resulting lateral heat flow is of the order of 30–60 mW m22, which is higher than the normal mantle heat flow (30 mW m22: Jaupart & Marschal 1999; Turcotte & Schubert 2002). The consequences of the lateral heat transfer on the density structure during the rifting has been extensively studied in previous literature (e.g. Royden & Keen 1980; Stephenson et al. 1989). These studies have provided a number of important corrections to the most common one-dimensional (1D) thermal subsidence models (McKenzie 1978). The thermo-mechanical consequences of the lateral heat exchanges were also investigated (e.g. Keen & Boutilier 1995), yet mainly in terms of their impact on the mantle viscosity and on the development of positive asthenosphere instabilities under the rift axis (transition from passive to active rifting, post-rift doming). In particular, Huismans et al. (1998) conducted a detailed study of differential stretching of inelastic lithosphere in terms of its interaction with the unstable asthenosphere. Yet, the negative mantle instabilities were beyond the scope of this model because the experiments of Huismans et al. (1998) were limited to the upper 120 km of the lithosphere instead of at least 250 km needed to account for the mantle RT instabilities below the rift flanks. The asthenosphere has a low viscosity (5  1019 Pa s) and can flow at timescales of 1 ka. The associated heat advection is instantaneous compared to the rifting timescales that exceed 0.1 –1 Ma.

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For b ¼ 2, the temperature of the asthenospheric material advected to a depth of 70 –90 km below the stretched lithosphere is around 1330 8C. The temperature of the prerift mantle lithosphere at this depth is only 750–800 8C (e.g. Kirby & Kronenberg 1987), (Fig. 1a). Consequently, even for a moderate extension, the temperature contrast, DT, between the upraised asthenosphere and the surrounding mantle lithosphere is around 500– 600 8C. The associated thermal density contrast, Drt, is 50–60 kg m23 (Drt ¼ r0aDT, where r0 ¼ 3330 kg m23 and a ¼ 3  1025 K21). Drt sums up with the chemical density contrast Drc of 10 –20 kg m23 yielding important ‘available negative buoyancy’ (Conrad & Molnar 1999). The total density contrast may be 2–3 times higher than the contrast of 10 –20 kg m23 that is sufficient to drive mantle instabilities at time scales of 2 –10 Ma (Conrad & Molnar 1999). Of course, the initially higher density contrast vanishes as the heat propagates from the upwelling asthenosphere to the mantle lithosphere. This decrease in density contrast slows down the RT instability. Yet, as the temperature rises, the mantle viscosity decreases. The reduction of viscosity, in contrast to that of density, accelerates the RT instability. Hence, the impact of lateral heat propagation on the growth rate of instability depends on the ratio of the density change rate, @r(T)/@T, to the viscosity change rate, @m(T)/@T. The density change rate @r(T)/@T is linearly proportional to temperature: @r(T)/@T ¼ ar, where a is the coefficient of thermal expansion (a ¼ 3  1025 K21). The viscosity change rate is super-exponential: for dislocation creep, m(T) ¼ m0exp(Q/nRT), from where @m(T)/@T ¼ 2m0Q(nRT 2)21exp(Q/nRT), where Q is activation enthalpy (5  105 J mol21), n  3 and R is the universal gas constant (see Table 2). A 200 8C temperature rise from 750 to 800 8C (mechanical bottom of the lithosphere: Kirby et al. 1987) reduces the initial density contrast between the mantle and asthenosphere by 20 kg m23 (Turcotte & Schubert 2002). Yet, this temperature rise also reduces the effective viscosity by factor of 104 (!). Such reduction is sufficient to bring the system from a mechanically stable state (m ¼ 1023 – 1024 Pa s) to an extremely unstable state at m ¼ 1019 –1020 Pa s (e.g. Conrad & Molnar 1999). Consequently, temperature rise in the mantle near the lithosphere– asthenosphere boundary will accelerate the growth rate of RT instability rather then reduce it. The question on the significance of the negative RT instabilities for slow rifting is naturally related to the problem of possible variation of the extension rate during the rifting phase. The extension rates deduced for many rift systems are averaged over some important time spans. Yet, the mechanical

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response of the brittle –ductile lithosphere depends on the strain rate at each moment of time. If the extension rate changes during the active phase of rifting, the final rift structures may differ even for similar amounts of extension. No doubt, there is a large number of other factors that may crucially influence the evolution of rifted margins. In particular, inherited structures, heterogeneities, rheology or surface processes are certainly of great importance. Yet, their influence is well studied and discussed in the previous work (e.g. Kusznir & Karner 1985; Dunbar & Sawyer 1988; Bassi 1995; Burov & Cloetingh 1997; Huismans & Beaumont 2007). We thus focus our study on more overlooked factors: negative gravitational stabilities and initial density structure of the lithosphere.

Theoretical background A number of studies (e.g. Houseman et al. 1981; Fletcher & Hallet 1983; Bassi & Bonnin 1988; Houseman & Molnar 1997; Huismans et al. 1998) have investigated the development of RT instabilities in normal, extended or shortened continental lithosphere as a function of thermo-rheological parameters. Based on these studies, the first-order condition for instable behaviour is formulated as: Drgd=2tmax , d=dfold where Dr is the density contrast between the mantle and the asthenosphere, g is the acceleration due to gravity, d is the thickness of the lithosphere, tmax is its maximal mechanical strength and dfold is the characteristic decay length of ductile strength/viscosity. For mantle lithosphere (olivine) d/dfold is around 10 and Drgd/2tmax is negative or smaller than 3 (Fletcher & Hallet 1983; Bassi & Bonnin 1988). This indicates that the mantle lithosphere is instable and may need only a small perturbation to develop a gravitational instability. Rifting produces a very strong perturbation of the mantle– asthenosphere interface and increases density contrasts, thus almost inevitably leading to unstable behaviour of the mantle layer. The results of simple linear analysis demonstrate that the growth rate of instability is mainly controlled by the density contrast and the viscosity of the mantle lithosphere (Chandrasekhar 1961; Fletcher & Hallet 1983). Let us first consider a simplified case of a layer of Newtonian fluid with viscosity m, density rm and thickness d placed on top of a weak fluid layer (asthenosphere) of lower density ra. The growth rate 1/t of the developing instability w (vertical deflection of the layer) with initial amplitude w0 depends on m, on the wavelength of perturbation, l, on the thickness of

the layer, d, and on the density contrast, Dr (Chandrasekhar 1961): w ¼ w0 expðt=tÞ t ¼ 4mðDrgdÞ1 ðv þ ðshðv1 Þchðv1 ÞÞ1 Þ  ðv2 thðvÞ  shðv1 Þchðv1 ÞÞ1

ð1Þ

where v ¼ l/2pd ¼ (kd)21, k is a wave number. The most rapidly growing wavelength is l ¼ 2.568d, which corresponds to minimal growth time: tmin ¼ 13:04mðDrgdÞ1 : As can be seen, the growth time is directly proportional to the fluid viscosity; so that it is less than 1 Ma for viscosities of m ¼ 1020 Pa s (assuming d ¼ 100 km and Dr ¼ 20 kg m23, Fig. 1b). High (.1022 Pa s) viscosity is needed to keep mantle layer quasi-stable for 10–15 Ma. For viscosities of the order of 1020 –1021 Pa s the mantle may be destabilized at timescales of 1–10 Ma that are perfectly compatible with typical timescales of continental rifting. The viscosity of m ¼ 1020 Pa s is usually attributed to the transition zone between the mantle and asthenosphere, which suggests that this zone can be instable. A number of analytical studies have extended equation 1 to the cases of realistic non-linear rheologies (e.g. Conrad & Molnar 1999). These studies show that in the nonlinear case the growth rate is about n times higher than in the case of Newtonian rheology, where n is the power-law exponent n (n ¼ 3 for most olivine rheologies). In the case of rifted continental lithosphere, a multitude of factors such as the rheological stratification and rift geometry affect the viscosity, temperature and the growth rate of instability (Fig. 1a, which justifies the numerical approach developed in this study). For example, as the lithosphere thins, d decreases and as a result the growth rate and wavelength of new instabilities might decrease too. However, thinning results in heating of the mantle layer. Consequently, the effect of thinning on the growth rate is mitigated owing to temperatureinduced viscosity drop. The importance of RT instabilities during rift development can be characterized by the ratio of the characteristic growth time of the instability to the characteristic time of the rifting phase, tr, expressed by means of Deborah number, De: De ¼ tmin =tr ¼ 13:04m=ðDrgdtr Þ: If De  1 the instabilities do not effect rift evolution. If De  1 rift evolution is highly affected

THE ROLE OF GRAVITATIONAL INSTABILITIES

by RT instabilities. A more conventional way to define De would be to set it as a ratio of the Maxwell relaxation time tm to flow time. Yet, in our case, the development of gravitational instabilities appears to be a leading process, which becomes clear from an estimate of the ratio of tmin to tm: tmin =tm  13G=Drgd  5E=Drgd

tr ¼ L=ux where L is the width of the rift and ux is the rate of extension. Consequently, ð2Þ

We can further introduce the effective viscosity for ductile creep law: dð1nÞ=n

meff ¼ ek

ðA Þ1=n expðH=nRTÞ

where e dk ¼ Invk (eij))1/2 is the effective strain rate and A* ¼ 1/2A . 3(nþ1)/2 is the material constant, H is the activation enthalpy, R is the gas constant and n is the power-law exponent. So that: meff ¼ m0 expðH=nRTÞ

with dð1nÞ=n

m 0 ¼ ek

ðA Þ1=n :

Taking into account the thermal dependency of density and viscosity, De becomes: De ¼ 13:04 m0 expðH=nRTÞux =ððDrc

where G is the elastic shear modulus and E is the Young modulus. E  2.5G. For typical rocks E  1011 Pa. Thus, for all plausible values of Drgd (Drgd , 10 –100 MPa), tmin/tm varies from 1000 to 10 000. Hence, gravitational instabilities must be a leading long-term process and Maxwell relaxation is of secondary importance. Time of rifting, tr, can be expressed as:

De ¼ 13:04mux =ðDrgdLÞ:

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ð3Þ

þ ar0 DTÞgdLÞ

ð4Þ

where Drc þ ar0DT ¼ Dr, Drc is the compositional density contrast, a is the coefficient of thermal expansion, r0 is the density at reference temperature and DT is the temperature change in respect to the reference temperature. Expression 4 indicates that slow extension rates enhance the impact of the gravity-driven instabilities, whereas fast spreading rates reduce them. This clearly suggests that slow extension favours instabilities and, as a natural consequence, rift asymmetry.

Numerical model We conducted a series of numerical experiments to test the stability of the extending lithosphere as a function of the extension rate, density and thermorheological structure. In these experiments we introduced a lithospheric plate composed of crustal, mantle and asthenospheric rheological layers with brittle –ductile –elastic properties inferred from rock mechanics data (Tables 1 and 2). The model area (Fig. 2) is 500 km wide and 300 km deep (250  150 grid elements); the sub-Moho (35 km depth) lithospheric mantle has a thickness of 120 km, with thickness of the strong mechanical core of about 70 –80 km (Fig. 1a).

Table 1. Notations and physical values common for all experiments (Turcotte & Schubert 2002) Parameter s, t P u m 1˙ T hc d rm ra Drc g Cp DT a

Values and units

Definition

Pa, MPa Pa, MPa m s21, mm year21 1019 – 1025 Pa s s21 8C 7 km/35 km 120 km 3330 kg m23 3310 –3330 kg m23 0 – 20 kg m23 9.8 m s22 103 J kg21 8C21 250 8C 3  1025 8C21

Stress Pressure Velocity vector Effective viscosity Strain rate Temperature Moho depth Thickness of mantle lithosphere Reference mantle density Asthenosphere density Compositional density contrast Acceleration due to gravity Specific heat Temperature contrast Thermal expansion

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Table 2. Specific rheology and related thermal parameters. Compilation by Burov et al. (2001). r is density; Q, n, A are material-dependent parameters of ductile flow laws (Kirby & Kronenberg 1987; Kohlstedt et al. 1995). Other parameters from Turcotte & Schubert (2002). Parameter

Value

All rocks l, G Lame´ elastic constants (l ¼ G) f friction angle (Mohr –Coulomb rheology) C0 cohesion (Mohr–Coulomb rheology)

30 GPa 308 20 MPa

Specific upper and weak (quartz) lower crust properties r (upper crust) n (power-law exponent) A (power-law constant) Q (creep activation enthalpy)

2800 kg m23 2.4 6.7  1026 MPa2n s21 1.56  105 kJ mol21

Specific strong middle–lower crust properties (diabase or basalt) r n (power-law exponent) A (power-law constant) Q (creep activation enthalpy)

2980 kg m23 3.4 2  1024 MPa2n s21 2.6  105 kJ mol21

Specific mantle properties (olivine) r (lithosphere) n (power-law exponent) A (power-law constant) Q (creep activation enthalpy)

3330 kg m23 3 1  104 MPa2n s21 5.2  105 kJ mol21

Thermal model Surface temperature (0 km depth) Temperature at the bottom of thermal lithosphere Thermal conductivity of crust, k Thermal conductivity of mantle, k Thermal diffusivity of mantle, x Radiogenic heat production at surface, Hs Radiogenic heat production decay depth, hr Prerift thermotectonic age of the lithosphere, tq

0 8C 1330 8C 2.5 W m21 8C21 3.5 W m21 8C21 1026 m2 s21 9.5  10210 W kg21 10 km 150 Ma

Fig. 2. Model set-up. The model box has dimensions of 500  300 km (250  150 grid elements). The initial rheology (explicit brittle–elastic– ductile) and temperature profile is shown on the right-hand side of the model. The initial weak temperature anomaly (þ50 8C) is used to localize thinning in the middle of the model.

THE ROLE OF GRAVITATIONAL INSTABILITIES

The boundary conditions are: (1) lateral velocities applied at the box sides; (2) free surface at the top of the model; and (3) restoring hydrostatic forces proportional to the vertical displacement at the basement. As pointed out by Huismans et al. (1998), whatever is the type of rifting (active or passive), the application of kinematic boundary conditions may be justified by the restricted size of the model. These conditions represent either far-field tectonic forces or the integrated effect of large-scale basal drag owing to the regional asthenospheric upwelling. We use a ‘generalized’ initial geotherm that fits well the observed data for the Newfoundland– Galicia margin (Pe´rez-Gussinye´ et al. 2001; Manatschal 2004). This geotherm provides 550 8C at Moho depth and 1330 8C at the base of the lithosphere, and falls in the middle of the interval between 400 8C at Moho depth (cratons) and 700 8C at Moho depth (young lithospheres) (e.g. Burov & Diament 1995). The localization of rifting in the middle of the model is achieved through the introduction of a small thermal anomaly (þ50 8C) at the base of the mechanical lithosphere. This anomaly dissipates soon after the initiation of rifting and has no prolonged effect for the lateral stages of rift evolution. The FLAC-like code Par(a)ovoz v9 (derived from Poliakov et al. 1993) solves the mechanical and thermal equilibrium equations in a large strain mode with no practical limitations on the imposed rheology laws because of its explicit time-marching scheme. The implementation and application of this algorithm to basin modelling is explained in detail in the abundant FLAC literature (Cundall 1989; Poliakov et al. 1993; Buck & Poliakov 1998; Lavier et al. 2000; Burov & Poliakov 2001, 2003; LePourhiet et al. 2004). For this reason, we limit its description to basic principals. Par(a)ovoz is a large strain fully explicit timemarching Lagrangian algorithm that solves Newtonian equations of motion in the continuum mechanics approximation: ˙  divs  rg ¼ 0: krul

ð5Þ

These equations are coupled with the constitutive equations of practically any kind: Ds ¼ Fðs; u; ru; . . . T . . .Þ; Dt

ð6Þ

rCp @T=@t þ urT  k divðrTÞ  Hf ¼ 0

body forces: r ¼ r0 ð1  aDTÞ:

t ¼ S þ sn tgf where t is the shear stress and sn is the normal stress. In addition, linear cohesion softening is used for better localization of plastic deformation, 1p, (S(1p) ¼ S0 min (0, 1 2 1p/1p0), where 1p0 is 0.01). The ductile –viscous term is represented by nonlinear power law with three sets of material parameters (Table 2) that correspond to the properties of the respective lithological layers: upper crust (quartz), middle –lower crust (quartz –diorite), mantle (olivine): meff ¼ ek

ð7Þ

assuming adiabatic temperature dependency for density and a Boussinesq approximation for related

ð8Þ

Here u, s, g, k are the respective terms for velocity, stress, acceleration due to body forces and thermal conductivity. The over-dots refer to the time derivatives. The brackets in equation 5 specify conditional use of the related term: in quasistatic mode, the inertia is dumped using inertial mass scaling (Cundall 1989). The terms t, r, Cp, T, Hr, a designate, respectively, time, density, specific heat, temperature, internal heat production and thermal expansion coefficient. The terms @/@t, Ds/Dt, F are a time derivative, an objective (Jaumann) stress time derivative and a functional, respectively. In the Lagrangian framework, the incremental displacements are added to the grid co-ordinates allowing the mesh to move and deform with the material. This enables solution of large-strain problems locally using a small-strain formulation: on each time step the solution is obtained in local co-ordinates, which are then updated in the large strain mode. Solution of equation 5 provides velocities at mesh points used for computation of element strains and of heat advection urT. These strains are used in equation 6 to calculate element stresses, and the equivalent forces are used to compute velocities for the next time step. All rheological terms are implemented explicitly. The rheology model is serial viscous – elastic – plastic (Fig. 1a, right, Table 2). The plastic term is given by explicit Mohr–Coulomb plasticity (non-associative, zero dilatency angle) with linear Navier –Coulomb criterion that implies an internal friction angle f of 308 and maximal cohesion S of 20 MPa to fit the experimental Byerlee’s law of rock failure (Byerlee 1978):

dð1nÞ=n

and with the heat transport equations

145

ðA Þ1=n expðH=nRTÞ

where edk ¼ (Invk(eij))1/2 is the effective strain rate and A * ¼ 1/2A.3(nþ1)/2 is the material constant, H is the activation enthalpy, R is the gas constant and n is the power-law exponent (Table 2).

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The elastic parameters (Table 2) correspond to commonly inferred values from Turcotte & Schubert (2002). The reference density of the mantle lithosphere is 3330 kg/m23. The density of the asthenosphere has been varied from 3310 to 3330 kg m23.

Figures 3–6 summarize the results of numerical experiments. We have conducted several sets of experiments aimed at testing the influence of the extension rate and of the density structure on the interplay between the gravitational instabilities and tectonically induced deformation. The extension rate ux has been varied from 5 to 15 mm year21 at both sides of the system. The chemical density contrast between the mantle and the asthenosphere varied from 0 to 20 kg m23. For the experiments, we imposed a bilayer crustal structure that corresponds to a granite (quartz-dominated) upper crust with density of 2800 kg m23 and a diabase lower crust with density of 2980 kg m23 (Figs 1a & 2, Tables 1 and 2). The influence of the extension rate, the rheology and the density contrast is characterized by a single non-dimensional parameter, Deborah number (De, equation 4). Yet, owing to the depth and laterally variable viscosity, the estimated Deborah number De is highly approximate. It is computed from averaged viscosity for 300 km of extension.

the previous case, mantle instabilities develop at about 2 Ma after the onset of rifting. The mantle lithosphere is broken after 4–5 Ma of rifting. The main features of this case resume to: (1) more pronounced interior basins; (2) ridge-like geometry of the oceanized lithosphere with characteristic ridge-like increase of the depth to the isotherm of 550–600 8C with distance from the centre; and (3) important asymmetry of the margins. As in the previous case, there is a sharp transition from highly thinned (2–6 km thick) continental crust and mantle lithosphere across the margin to nearly an undeformed zone on the backside of the rift flanks. This sharp transition results from mechanical coupling of the mechanical layers in the rifted zone and from replacement of the weak continental crust by strong oceanized asthenosphere. There are two major factors defining formation of the interior basins: tensional and gravitational instabilities. For example, high tensile stresses due to faster extension rate and mechanical coupling of the crustal and mantle layers below the rift flanks locally increase the plate strength and thus favour extension in a new place. At slow extension rates (,2  5 mm year21) gravitational instabilities are more efficient. Tensional instabilities are efficient at high extension rate, which appears to be the case of this experiment, compared to the previous experiment. There are, of course, some intermediate situations when the role of both mechanisms is comparable.

De ¼ 0.3– 0.75, slow spreading rate (ux ¼ 2  5 mm year21)

De ¼ 1.5 – 10, high spreading rate (ux ¼ 2  15 mm year21)

The results of the experiments at ux ¼ 5 mm year21 and Dr ¼ 20 kg m23 are shown in Figure 3. This case is characterized by development of gravitational instabilities at 1–2 Ma and wide rifting, prior to pronounced narrow rifting, occurring after 6–7 Ma. After this time the mantle lithosphere continues to be thinned by gravitational de-blobbing concurrently with the formation of a rift basin. At b ¼ 5 (time 6– 7 Ma) the rift structure becomes strongly asymmetric and small interior basins form at the backside of one (or both) of the rift flanks. The rift asymmetry appears to be controlled by the gravitational instabilities that have different growth rates on opposite sides of the rift. As mantle viscosity is exponentially dependent on temperature, this asymmetry is reinforced due to viscosity variations caused by asymmetry in the thermal field that also results from the gravitational instabilities.

Figure 5 summarizes the results of the experiments at ux ¼ 15 mm year21 for three values of Dr (0, 10 and 20 kg m23). In all cases, the continental lithosphere is already broken up at 3 Ma after the onset of the extension (coefficient of extension, b  4–5). At 5.6 Ma mantle, upper and lower crust are widely exhumed at the surface (as it is actually observed, for example, beneath the Iberia Abyssal Plain: Whitmarsh et al. 2000). In case of zero and weak (10 kg m23) density contrast (Fig. 5a, b) the interior basins are less pronounced. In the case of higher density contrast (20 kg m23, Fig. 5c), secondary interior basins develop at rifted margins starting from 5 Ma and form a characteristic structure that resembles that of the rifted continental margins such as the conjugated Iberian and Newfoundland margins (Fig. 6). As discussed above, formation of the interior basins is obviously related to either tensional and/or gravitational instabilities (for extension rates of more than 2  15 mm year21 tensional instabilities dominate) and to the local mechanical coupling between the crustal and mantle layers. This coupling results from

Experiments and results

De ¼ 0.75– 1.5, intermediate spreading rate (ux ¼ 2  10 mm year21) The series of experiments at ux ¼ 10 mm year21 and Dr ¼ 20 kg m23 are shown in Figure 4. As in

Fig. 3. Results of the experiments for slow extension rate (ux ¼ 2  5 mm year21) and density contrast between the mantle and asthenosphere Dr ¼ 20 kg m23. Left: evolution of the material field (colour code: upper crust, salad-green; middle– lower crust, yellow; mantle lithosphere, blue; asthenosphere, marine-green; synrift sediment or otherwise reworked material, purple). Centre: temperature field. Left: logarithm of stress to strain ratio (Pa s), which is equivalent to the effective viscosity for the ductile domains. Note: (1) removal of a large part of the mantle lithosphere by RT instabilities leading to change of style of rifting; (2) strong asymmetric rifting at developed stages of extension; (3) breakup and oceanization of the lithosphere that commences at 9 Ma and is preceded by coupling of the crustal and mechanical layers within the lithosphere; (4) exhumation of small amounts of lower crust and mantle. Arrows show the position of the interior basins. Inserts correspond to 2 blow-ups of the framed areas. Purple (reworked) material most often directly overlies its source material, i.e. purple above the green layer means material is reworked from the upper crust. Purple above blue means either reworked mantle or sediment derived from the upper and lower crust.

THE ROLE OF GRAVITATIONAL INSTABILITIES 147

Fig. 4. Results of the experiments for ux ¼ 2  10 mm year21 and Dr ¼ 20 kg m23. See the caption to Figure 3 for notations. Note: (1) removal of a large part of the mantle lithosphere by RT instabilities leading to change of style of rifting; (2) strong asymmetric rifting at developed stages of extension (at the oceanization stage); and (3) exhumation of continental and oceanic mantle, and punctual exhumation of lower crustal material at ocean –continent limits. In this experiment, continental breakup and oceanization commence at 6 Ma. Arrows show the position of the interior basins. See the description of Figure 3 for other details.

148 E. BUROV

Fig. 5. (a) Results of the experiments for ux ¼ 2  15 mm year21 and Dr ¼ 0 kg m23. See the caption to Figure 3 for other notations. Note: (1) mitigated RT instabilities; (2) nearly symmetric rifting until developed stages of extension (oceanization stage); (3) less expressed interior basins; and (4) exhumation of continental and oceanic mantle, and exhumation of lower crustal material at ocean– continent limits. In this experiment, continental breakup and oceanization commence at 4 –5 Ma. See the description of Figure 3 for other details.

THE ROLE OF GRAVITATIONAL INSTABILITIES 149

Fig. 5. (b) Results of the experiments for ux ¼ 2  15 mm year21 and Dr ¼ 10 kg m23. Right: topography profiles. See the caption to Figure 3 for other notations. Note: (1) RT instabilities play a remarkable role, starting from 5 Ma; (2) nearly symmetric rifting; and (3) exhumation of continental and oceanic mantle, and exhumation of lower crustal material at ocean– continent limits. In this experiment, continental breakup and oceanization commence at 4 –5 Ma. See the description of Figure 3 for other details.

150 E. BUROV

Fig. 5. (c) Results of the experiments for ux ¼ 2  15 mm year21 and Dr ¼ 20 kg m23. See the caption to Figure 3 for other notations. Note: (1) RT instabilities below the rift flanks play an important role, starting from 2 Ma; (2) unstable extension leading to the formation of boudinage-like structures and interior basins; (3) nearly symmetric rifting; (4) formation of a strong segmented core in the middle of the rift separated by interior basins from the margin; and (5) exhumation of continental and oceanic mantle, and exhumation of lower crustal material at ocean– continent limits. At final tages remnants of the continental crust and mantle are spread in a thin layer (2 –6 km thick) over a wide area, some of them being replaced with oceanic material. In this experiment, continental breakup and oceanization commence at 2 –3 Ma. Arrows show the position of the interior basins. See the description of Figure 3 for other details.

THE ROLE OF GRAVITATIONAL INSTABILITIES 151

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THE ROLE OF GRAVITATIONAL INSTABILITIES

differential thinning accompanied by expulsion (squeezing) of the ductile crust from the thinned zones. As a result, the competent cores of mantle and crustal layers join together. This mechanical coupling locally increases the integrated strength of the lithosphere by a factor of 2 (e.g. Burov & Poliakov 2001), which favours thinning outside of the ‘hardened’ coupled zone. For higher extension rate, the rift structure is more symmetric both in the case of low- and highdensity contrast (Fig. 5a, b). This confirms the idea that high extension rate reduces the impact of gravity instabilities on the rift asymmetry. In the case of Dr ¼ 20 kg m23 (Fig. 5c) the vertical amplitude of gravity instability in the mantle part becomes important and reaches 20 –50 km after only 2 Ma of rifting. The lithosphere under both rift flanks thickens, leaving a narrowing stretched zone in between, which obviously helps to switch rifting from wide mode to narrow mode. The instability continues to develop at the exponential rate leading to de-blobbing of a part of the lithospheric mantle at the depth level corresponding to the effective viscosity of about 1021 Pa s (Fig. 5c). After this moment the mantle below the rift flanks is removed. The whole system remains roughly symmetric, with similar interior basins formed on both sides of the rift. Figure 6 compares the results of the experiment for Dr ¼ 20 kg m23 with the present-day structure of the conjugated rifted margins of Flemish Cap and Galicia (Funck et al. 2003). It reveals a number of important similarities: the geometry of the extended zones, the sharp transition from the extremely stretched crust and mantle to practically underformed lithosphere (e.g. the Flemish Cap), the geometry of the exhumed mantle and that of the exhumed lower crustal zones (e.g. Whitmarsh et al. 2000), the geometry and wavelength of the interior basins (Fig. 6b) (Galicia Interior Basin: Pe´rez-Gussinye´ et al. 2003). Even though the partial melting is

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not included in the present experiments, there is no difficulty in predicting, when examining Figures 5c, & 6a, b, the location of the potential zones of partial melting that are seemingly compatible with those interpreted from the seismic crosssections (Funck et al. 2003).

Discussion and conclusions † The experiments demonstrate the importance of the negative RT instabilities during the formation of ‘slow’ margins and continental rifts. These instabilities are highly effective in the range of 0,De , 10, or in the range of initial spreading velocities 0–2  15 mm year21. The RT instabilities are most efficient for De  1 and spreading rates 2  10 mm year21. The instabilities largely control rift symmetry and rifting mode. For example, the transfer from wide to narrow rift mode for De , 3 or ux  2  15 mm year21 may be associated with gravitational thinning and localized downwelling of the mantle lithosphere. In the absence of specific mechanisms of softening or inherited structures, rift asymmetry may result from gravitational instabilities and differential stretching for De , 1 or ux ,2  10 mm year21. As the temperature dependence of both the viscosity and the growth rate of instabilities is exponential, a minor variation in the thickness of lithospheric layers results in asymmetric rifting in case of small extension rates. Extension at rates of less than 2  5 mm year21 produces highly unstable asymmetric rifting. † Formation of the interior basins at rifted margins, such as the Galicia Interior Basin (e.g. Pe´rezGussinye´ et al. 2003), may be linked to the interplay between the tensional and gravitational instabilities and local mechanical coupling– uncoupling of the intracrustal and mantle layers for De,10 and b . 4– 5. At slow spreading

Fig. 6. (a) Comparison of the results of the experiment shown in Figure 5c (ux ¼ 15 mm year21, Dr ¼ 20 kg m23) with the present-day structure of the Flemish Cap margin (Funck et al. 2003). Note the similarities in the sharp transition from the deformed to undeformed continental lithosphere in the geometry of the extended crustal blocks including exhumed lower crustal and mantle material. Note that the actual margin has experienced major thermal subsidence and probably additional extension or other deformation after the rifting phase. Arrows show the position of the interior basins. The dashed line on the topography plot corresponds to the basement before water loading. See the description of Figure 3 for other details. (b) Comparison of the results of the experiment shown in Figure 5c (ux ¼ 15 mm year21, Dr ¼ 20 kg m23) with the actual structure of the Galicia margin that is conjugated with the Flemish Cap margin (Funck et al. 2003). Note the similarities in the geometry of the interior basins, in sharp transition from deformed thinned crust and mantle to undeformed continental lithosphere, in the geometry of the extended crustal blocks including exhumed lower crust and mantle material. Numerical resolution does not allow localized faulting to be reproduced, but comparison with Tirel et al. 2004 (based on the same numerical code) suggests that detachment faults appear in the zones of extra-thinned crust. Note that the actual margin has experienced major thermal subsidence and probably additional extension or other deformation after the rifting phase. Arrows show the position of the interior basins. The dashed line on the topography plot corresponds to the basement before water loading. See the description of Figure 3 for other details.

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rates (2  10 mm year21), gravity-driven instabilities dominate and generate interior basins for chemical density contrasts greater than 10 kg m23. At higher extension rates, interior basins are amplified by tensional instabilities, even for small compositional density contrasts. Prebreakup thinning of the crust and mantle first results in weakening of the continental lithosphere. Yet, in later stages (b . 4 –5) it leads to localized mechanical coupling of the competent crustal and mantle layers below the rift flanks, which locally increases the integrated strength of the lithosphere (e.g. Burov & Diament 1995). As a result, the interior zones (interior basins) become relatively weaker than the zones of mechanical coupling, and become a subject of more important stretching than the coupled zones. † The experiments for ‘high’ extension rate and Dr ¼ 20 kg m23 (215 mm year21) (Figs 5c & 6) demonstrate important similarities with the observations (Funck et al. 2003). In particular, the predicted geometry of sharp transition from zones of highly stretched crust (1–6 km thick) to practically undeformed crust and lithosphere below the rift flanks resembles that of the Flemish Cap and Galicia margin. The model predicts that the extension of the continental crust may occur in an extremely wide zone. In particular, after 9 Ma, b ¼ 5–15, the continental crust (Figs 5c & 6) thins from 30–40 km to only 1–2 km over a 80– 100 km-wide zone across the modelled margin. This crust is then punctually or totally replaced with the new oceanic material. The uplifted/ exhumed mantle underlies (depth of less than 2–5 km below the bottom) or directly replaces the ultra-thinned continental and oceanic material over a 70–100 km-wide continent– ocean transition zone (e.g. Fig. 6a). The wide zones of extended continental crust (stretched to a less than 5–6 km-thick layer) produced in the model are indeed observed across the Galicia margin. The model predicts important exhumation of the continental mantle and punctual exhumation of the lower crust, in between the exhumed upper crust and mantle. In the case of instable extension, this crust may periodically alternate with the mantle (e.g. serpentinized) and the upper crustal material. The predicted vertical crustal undulations resulting from gravitational and tensional instabilities may be associated with prominent detachment faulting, as was shown (using same numerical code) for domal uplift instabilities in Tirel et al. (2004). Yet, the numerical resolution of the present model is insufficient to reproduce localized faulting. However, the predicted patterns of accumulated plastic strain suggest that stretched and

undulated upper crust deforms along the detachment faults across the margin. In the model, interior basins are formed either on both sides of the rift (high extension rate of 2  15 mm year21) or only on one side (specifically for smaller extension rates ,2  10 mm year21). This process is governed by gravitational and tensional instabilities in multilayered lithosphere. † The experiments show that it is more difficult, although not impossible, to reproduce final structure of ‘slow’ rifted margins under assumption of low extension rates (,5 mm year21) during the entire duration of the phase. In particular, slow extension rates (,2  10 mm year21) are associated with unstable asymmetric rifting, whereas the conjugated slow margins such as the Flemish Cap–Galicia margin demonstrate less asymmetry than could be expected from the models, for reported extension rates of less than 5 mm year21. Consequently, it is not excluded that the extension rate could be higher at some prebreakup stages of rifting, and then slows down later to yield the reported average rate of 5 mm year21. This does not, of course, exclude the possibility that the extension rate could increase after the breakup. Indeed, most reliable (magnetic) data refer to sea-floor spreading stage after continental breakup (e.g. Srivastava et al. 2000). The extension rate at the prebreakup stage is much less constrained. One can admit that rift systems are subject to large prebreakup variations of the extension rate. This may explain the multitude of rifting styles even for comparable average spreading rates and coefficients of extension b. An alternative explanation may infer smaller density contrasts or stronger ductile rheology. Yet, in this case the formation of interior basins will be less probable. † It is necessary to stress the primary importance of obtaining better constraints on the major parameters controlling margin evolution: (1) the compositional density contrast between the mantle and asthenosphere; (2) the prerift thermo-rheological structure; and (3) the extension rate as a function of time. These parameters are difficult to derive from the available data, and their acquisition presents a new challenge for future geological and geophysical exploration of continental margins. † Of course, the results for narrow and wide rift might infer different reaction in terms of developing instability. We here have exploited only those initial geometrical and thermal configurations that represent in some way Flemish Cap–Galicia margins. The configurations, however, look to be quite generally applicable to slow extension zones.

THE ROLE OF GRAVITATIONAL INSTABILITIES The author thanks the reviewers – R. Govers, L. Moresi and M. Pe´rez-Gussinye´ – for highly constructive comments. I also thank G. Karner and the organizers of the IMEDL workshop for the invitation to participate and the following inspiring discussions. The numerical code Par(a)ovoz v.9 is a product of Paravoz v.3 (1996) by Poliakov and Podladchikov (Poliakov et al. 1993). This study has benefited from partial support by the MEBE program.

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M ANATSCHAL , G. & B ERNOULLI , D. 1999. Architecture and tectonic evolution of nonvolcanic margins: Present-day Galicia and ancient Adria. Tectonics, 6, 1099–1119. M ANATSCHAL , G., F ROITZHEIM , N., R UBENACH , M. & T URRIN , B. D. 2001. The role of detachment faulting in the formation of an ocean–continent transition: insights from the Iberia Abyssal Plain. In: W ILSON , R. C. L., W HITMARSH , R. B., T AYLOR , B. & F ROITZHEIM , N. (eds) Non-volcanic Rifting of Continental Margins: A Comparison of Evidence from Land and Sea. Geological Society, London, Special Publications, 187, 405– 428. M C K ENZIE , D. 1978. Some remarks on the development of sedimentary basins. Earth and Planetary Science Letters, 40, 25–32. P E´ REZ -G USSINYE´ , M., R ESTON , T. J. & M ORGAN , J. P. 2001. Serpentinisation and magmatism at nonvolcanic margins – The effect of the initial lithospheric structure. In: W ILSON , R. C. L., W HITMARSH , R. B., T AYLOR , B. & F ROITZHEIM , N. (eds) Non-volcanic Rifting of Continental Margins: A Comparison of Evidence From Land and Sea. Geological Society, London, Special Publications, 187, 551– 576. P E´ REZ -G USSINYE´ , M., R ANERO , C. R. & R ESTON , T. J. 2003. Mechanisms of extension at nonvolcanic margins: evidence from the Galicia interior basin, west of Iberia. Journal of Geophysical Research, 108(B5), 2245, doi:10.1029/2001JB000901. P OLIAKOV , A. N. B., P ODLADCHIKOV , YU . & T ALBOT , C. 1993. Initiation of salt diapirs with frictional overburden: numerical experiments. Tectonophysics, 228, 199– 210. R OYDEN , L. & K EEN , C. E. 1980. Rifting process and thermal evolution of the continental margin of Eastern Canada determined from subsidence curves. Earth and Planetary Science Letters, 51, 343 –361.

P OUDJOM D JOMANI , Y. H., O’R EILLY , S. Y., G RIFFIN , W. L. & M ORGAN , P. 2001. The density structure of subcontinental lithosphere: Constraints on delamination models. Earth and Planetary Science Letters, 184, 605– 621. O’R EILLY , S. Y., G RIFFIN , W. L., P OUDJOM D JOMANI , Y. & M ORGAN , P. 2001. Are lithospheres forever? Tracking changes in subcontinental lithospheric mantle through time. GSA Today, 11, 469. S ALVENSON , J. O. 1978. Variations in the geology of rift basins; a tectonic model. Conference Proceedings of Los Alamos Scientific Laboratory, 7487, 82– 86. S RIVASTAVA , S. P., S IBUET , J.-C., C ANDE , S., R OEST , W. R. & R EID , I. D. 2000. Magnetic evidence for slow seafloor spreading during the formation of the Newfoundland and Iberian margins. Earth and Planetary Science Letters, 182, 61– 76. S TACEY , F. D. 1992. Physics of the Earth. Brookfield Press, Brisbane. S TEPHENSON , R. A., N AKIBOGLU , S. M. & K ELLY , M. A. 1989. Effects of astenosphere melting, regional thermoisostasy, and sediment loading on the thermomechanical subsidence of extensional sedimentary basins. In: P RICE , R. A. (ed.) Origin and Evolution of Sedimentary Basins and their Energy and Mineral Resources. American Geophysical Union, Geophysical Monograph, 48, 17–27. T IREL , C., B RUN , J.-P. & B UROV , E. 2004. Thermomechanical modeling of extensional gneiss dome. In: Gneiss Domes and Orogeny. Geological Society of America, Special Paper 380, 67–78. T URCOTTE , D. L. & S CHUBERT , G. 2002. Geodynamics. Cambridge University Press, Cambridge. W HITMARSH , R. B., D EAN , S. M., M INSHULL , T. A. & T OMPKINS , M. 2000. Tectonic implications of exposure of lower continental crust beneath the Iberia Abyssal Plain, northeast Atlantic Ocean: Geophysical evidence. Tectonics, 19, 919–942.

A dynamic model of rifting between Galicia Bank and Flemish Cap during the opening of the North Atlantic Ocean D. L. HARRY1 & S. GRANDELL2 1

Department of Geosciences, Colorado State University, Fort Collins, CO 80523-1482, USA (e-mail: [email protected]) 2

Department of Geology, Adams State College, Alamosa, CO 81102, USA

Abstract: A finite-element model is used to simulate Late Jurassic– Early Cretaceous rifting between the Flemish Cap and Galicia Bank continental margins. The model results show that variations in the thickness of the continental crust on these margins at wavelengths greater than about 75 km can be explained as a consequence of the interaction of two pre-existing weaknesses in the lithosphere. A weakness in the crust, attributed to structural fabrics in the Variscan front, controls the location of crustal extension during the early stages of rifting in the model. This results in formation of a broad rift basin similar to the Galicia Interior Basin. A deep-seated weakness located 110 km further west, attributed to the thick crust beneath the central Variscan Orogen, controls the location of mantle necking. Extension in this region is initially diffuse, but accelerates and becomes more focused with time. Approximately 13 million years after rifting begins, the locus of crustal extension shifts from the region of pre-weakened crust into the region of pre-weakened mantle. This marks the end of subsidence in the Galicia Interior Basin and the onset of subsidence in the Flemish Cap and Galicia Bank marginal basins. Extension in these areas continues for another 12 million years before continental breakup. The asthenosphere does not ascend to depths shallow enough for decompression melting to begin until less than 5 million years before the onset of sea-floor spreading. The model predicts that all late-stage synrift magmatism during this period is limited to within 45 km of the rift axis, and production of melt thicknesses greater than 2 km is restricted to within 35 km of the rift axis. Mantle potential temperatures of 1250–1275 8C, 5 –30 8C cooler than normal, result in 3.1–4.5 km-thick oceanic crust at the time of breakup, in general agreement with the 2 –4 km thick crust observed adjacent to these margins.

Jurassic –Early Cenozoic rifting between the conjugate Iberia and Newfoundland continental margins progressed in a roughly south to north direction (Ziegler 1989; Srivastava et al. 1990). On the Iberian side, this resulted in three distinct rift margin segments. From south to north, these are the Tagus Abyssal Plain, Iberian Abyssal Plain and Galicia Bank segments. The Tagus and Iberia Abyssal Plain segments are both characterized by relatively wide regions of highly attenuated continental crust and what has been interpreted to be a broad area of exhumed subcontinental mantle (Pinheiro et al. 1992; Beslier et al. 1993; Whitmarsh & Sawyer 1993; Krawczyk et al. 1996; Minshull et al. 1998; Whitmarsh et al. 2000; Funck et al. 2003; Henning et al. 2004). The Galicia Bank segment exhibits a much more abrupt transition between relatively thick continental crust and oceanic crust, and a much narrower region of exhumed mantle (Boillot et al. 1980; Sibuet 1992; Whitmarsh et al. 1996). All segments of the Iberia margin, as well as the conjugate Newfoundland margin, are characterized by relatively little synrift magmatism prior to the onset of sea-floor spreading in comparison to most other rifted continental margins. Several numerical modelling studies have examined the

unusual structure of the rifted margin and the magmatic consequences of rifting on the Iberia Abyssal Plain and the conjugate portion of the Newfoundland margin (e.g. Bassi et al. 1993; Keen & Dehler 1993; Keen et al. 1994; Tett & Sawyer 1996; Harry & Bowling 1999; Bowling & Harry 2001; Minshull et al. 2001), but relatively little consideration has been given to the Galicia Bank –Flemish Cap segment and the differences between this segment and the more southerly parts of the margin. This paper describes the results of a finiteelement numerical simulation of Late Mesozoic and Early Cenozoic rifting on the Galicia Bank and Flemish Cap segments of the Iberia and Newfoundland conjugate continental margins. The modelling goal is to match the long-wavelength (.75 km) crustal thickness variations on the margins, the timing and duration of rifting, the magmatic history of the margins, the timing and location of regional shifts in extensional tectonism, and the regional subsidence history. This paper focuses on a conjugate transect between seismic profiles ISE-1 on the Iberian margin and SCREECH-1 on the Newfoundland margin (Fig. 1) (Funck et al. 2003; Henning et al. 2004; Hopper et al. 2004).

From: KARNER , G. D., MANATSCHAL , G. & PINHEIRO , L. M. (eds) Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 157– 172. DOI: 10.1144/SP282.8 0305-8719/07/$15.00 # The Geological Society of London 2007.

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The modelling technique used in this paper is based on the STRCH95 two-dimensional finite-element modelling program, which is a descendant of the STRCH program developed by Dunbar (1988). STRCH95 is used to model extensional processes on a lithosphere scale. The program allows for complex initial conditions to describe the (a)

16 W

14 W

00

12 W

10 W

8 W

-4000

0

-4

pre-extension structure of the crust and rheology and temperature of the lithosphere. The simulations discussed in this paper focus on how such preexisting features might have controlled the pattern of extension and shifts in the loci of extension with time on the Galicia Bank–Flemish Cap margins, and how such extension is accommodated

6 W

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42 N Fig. 1. Location map. (a) Iberia margin. (b) Newfoundland margin. Solid lines are locations of model transects. Dots are DSDP and ODP drill sites. Shaded regions are interior rift basins: GIB, Galicia Interior Basin; LB, Lusitania Basin; FCG, Flemish Cap Graben; FPB, Flemish Pass Basin; JDB, Jeanne d’Arc Basin; WB, Whale Basin; SWD, South Whale Basin; HB, Horse Shoe Basin; OB, Orphan Basin; CB, Carson Basin. Contour interval is 1000 m.

A FINITE-ELEMENT RIFTING MODEL

without producing significant amounts of synrift magmatism. In the following sections we describe the geological and geophysical observations that constrain the numerical models discussed in this paper and the modelling targets. We then describe the modelling algorithm. Finally, we describe the results and discuss their implications within the context of Galicia Bank– Flemish Cap rifting in particular and rifting on non-volcanic margins in general.

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(Minshull et al. 2001; Hopper et al. 2004). Absolute ages are based on the timescale of Gradstein et al. (2004). The first approximately 10– 12 million years of extension (Tithonian–Valanginian) involved the Galicia Interior Basin and Newfoundland Basin. The last approximately 13 –15 million years of extension was primarily focused seaward of the Galicia Bank and Flemish Cap.

Post-rift crustal structure

The history and structure of the Galicia Bank and Flemish Cap continental margins is described in detail elsewhere in this Special Publication. Here, we briefly summarize those elements of rift history and margin structure that contribute directly to constraining the numerical simulations presented in this paper.

Timing Extension on the Iberian Peninsula and Newfoundland occurred in several distinct phases. The first phase occurred in the Newfoundland and Galicia Interior basins during the Late Triassic –Early Jurassic periods (Fig. 1) (Tankard & Welsink 1987; Tucholke et al. 1989; Ziegler 1989; Murillas et al. 1990; Foster & Robinson 1993). This was followed by a period of tectonic quiescence lasting until Late Jurassic time. The second phase of extension began in the Late Jurassic epoch and lasted until late Valanginian time, when the locus of extension shifted seaward of the Flemish Cap and Galicia Bank (Tucholke et al. 1989; Murillas et al. 1990). This led to the third phase of extension, a period of lithospheric stretching and crustal thinning west of Galicia Bank and east of Flemish Cap in Hauterivian–Aptian time. This last phase of extension continued until continental breakup and formation of new oceanic crust between the time of sea-floor spreading anomalies M3 and M0 (Srivastava et al. 2000; Minshull et al. 2001; Funck et al. 2003; Hopper et al. 2004). In this paper we focus on the Late Jurassic – Early Cretaceous phases of extension. These phases involved an initial stage of extension in the Galicia Interior Basin, a later stage of extension seaward of the unextended continental crust on Flemish Cap and the slightly extended crust on Galicia Bank, and, finally, the onset of sea-floor spreading. The total duration of rifting (excluding the Triassic –Early Jurassic phase in the interior basins) lasted approximately 25 million years, beginning in Late Tithonian time (c. 146 Ma) and ending in the early Aptian age (c. 118 Ma)

A reconstruction of the Newfoundland and Galicia continental margins at the time of anomaly M0, near the onset of sea-floor spreading, is shown in Figure 2. Unextended portions of the continental crust on Iberia and Newfoundland are 30– 32 km thick (Mendes Victor et al. 1980; Banda 1988; Diaz et al. 1993). The prerift crystalline crust on the Iberia margin thins westward to less than 10 km-thick beneath the 100 km-wide Galicia Interior Basin (Pe´rez-Gussinye´ et al. 2003), increases to approximately 20 km beneath the Galicia Bank (Gonzalez et al. 1999; Pe´rez-Gussinye´ et al. 2003), and then thins progressively westward over a 100 km-wide region between the eastern edge of the Galicia Bank and the rift axis, where the oldest oceanic crust is emplaced. The prerift crust reaches a minimum thickness of 2– 5 km at its seaward limit (Chian et al. 1999; Gonzalez et al. 1999). On the Newfoundland margin, the crust is approximately 30 km thick beneath the relatively unextended Flemish Cap, and thins progressively eastward to approximately 1.5 km over an about 80 km-wide region between the eastern edge of the Flemish Cap and the rift axis (Funck et al. 2003). Mantle exposed in an approximately 10 km-wide region on the sea-floor west of Galicia Bank is thought to have been exhumed by latestage low-angle faulting during the last 1–2 million years of extension (Sibuet 1992; Pickup et al. 1996; Reston 1996; Fuegenschuh et al. 1998; Pe´rez-Gussinye´ et al. 2003). East

West Flemish Cap

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Depth (km)

Model constraints

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Distance from riftaxis (km)

Fig. 2. Cross-section of the Newfoundland–Iberia margins at the time of anomaly M0 along the trend of seismic transects SCREECH-1 on the Newfoundland margin (Funck et al. 2003; Hopper et al. 2004) and ISE-1 on the Galicia Bank margin (Henning et al. 2004). After Funck et al. (2003).

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Extension rate Extension rates on the Newfoundland and Iberia margins vary locally at the scale of a few tens of kilometres, probably reflecting different timing between movements on individual fault systems. In a regional sense, extension rates are best constrained on the Iberia Abyssal Plain. Here, regional extension rates are estimated to be approximately 10 mm year21 during rifting (Minshull et al. 2001) and 10 –14 mm year21 at the onset of seafloor spreading (Whitmarsh et al. 2001a; Russell & Whitmarsh 2003). Assuming symmetric spreading, this corresponds to a whole-extension rate of 20 mm year21 during rifting. This is consistent with area balancing of the cross-section in Figure 2. Restoration of the cross-section to a uniform crust thickness of 32 km requires an extension rate of 20 mm year21 to produce the observed amount of stretching in 25 million years.

Prerift lithospheric structure, rheology and thermal constraints Breakup between the Galicia Bank and Newfoundland occurred within a region roughly coincident with the trend of the Late Palaeozoic Variscan Orogen in this region (Capdevila & Mougenot 1988; Ziegler 1989). This is similar to the association of Mesozoic and Cenozoic rifting with the trends of other Palaeozoic orogens, including opening of the Gulf of Mexico along the Ouachita orogenic trend, opening of the North Atlantic along the Appalachian trend and opening of the Norwegian–Greenland Sea along the Caledonide trend (Fig. 3) (Ziegler 1989; Harry & Londono 2004). Focusing of extensional deformation within the Variscan Orogen was most probably a result of orogenic weakening of the lithosphere caused by the increased thickness of the crust beneath the orogen (Fig. 4). The crust in the central Variscan Orogen near the rift axis is estimated to have been approximately 34 –35 km thick prior to extension (Pe´rez-Gussinye´ et al. 2003), 2–5 km thicker than the crust in the unextended parts of the Iberia Peninsula and Newfoundland. Assuming the thermal parameters estimated by Tejero & Ruiz (2002) for the crust in the Duero Basin in the Iberian Peninsula and a simple two-layer rheological model following empirical ductile flow laws for quartz-diorite in the crust and wet dunite in the mantle (Table 1), this difference in crustal thickness would weaken the lithosphere in the central Variscan Orogen by up to 15% in comparison to the surrounding regions (Fig. 4). The strength of the nominal lithosphere predicted by this rheological model is 7.0  1012 N m21 (Fig. 4a), which is moderately weaker than the 8  1012 N m21 lithospheric strength

estimated by Tejero & Ruiz (2002) for the central Iberian Peninsula and is consistent with the modern (presumably steady-state) heat flow and thermal structure of the Iberian lithosphere estimated by Fernandez et al. (1998).

Magmatic history Both the Flemish Cap and Galicia Bank margins are flanked by unusually thin oceanic crust, ranging from 2.5 to 4 km in thickness (Whitmarsh et al. 1996; Hopper et al. 2004), suggesting a low volume of melt production immediately following continental breakup. On the Galicia margin the oceanic crust increases to normal thickness (approximately 7 km) within 15 km seaward of the oldest oceanic crust (Whitmarsh et al. 1996). At the estimated spreading half-rate of 10–14 mm year21 (Whitmarsh et al. 2001a), this suggests that the thermal regime of the mantle rapidly evolved to that of a typical mid-ocean ridge system within approximately 1–1.5 million years after the onset of sea-floor spreading. The melt production history after the onset of sea-floor spreading on the Newfoundland margin is more complicated. Here, the oldest oceanic crust is about 3–4 km thick and thins seaward to less than 1.3 km over an approximately 50 km-wide region (Hopper et al. 2004), indicating a magma-starved environment during the first approximately 3.5–5 million years of sea-floor spreading. Near-normal oceanic crust (.6 km thick) located 3 km seaward of the thinnest oceanic crust on the Flemish Cap margin (Funck et al. 2003; Hopper et al. 2004) indicates that the transition from magma-starved seafloor spreading to normal sea-floor spreading occurred in about 0.2–0.3 million years. Recovery of late synrift volcanic and plutonic rocks landward of the oldest oceanic crust on both the Newfoundland and Iberian margins suggests that the last stages of rifting were accompanied by minor amounts of magmatism, just prior to the onset of sea-floor spreading (Manatschal & Bernoulli 1999; Whitmarsh et al. 2001b; Tucholke et al. 2004). Regardless of the complexities of late rift-stage and early spreading-phase melt production, there is broad consensus that for most of the rift history the margins were amagmatic, or at least involved very little melt production. Volcanism that occurred prior to the onset of sea-floor spreading appears to have been limited in volume and space, being confined to within about 20 km of the location where oceanic crust is first produced.

Modelling targets On the basis of the above discussion, our numerical simulations of rifting between Galicia Bank and Flemish Cap are constrained by: (1) the prerift

A FINITE-ELEMENT RIFTING MODEL

161

Fig. 3. Late Permian reconstruction of the North Atlantic continents and trends of major Late Palaeozoic orogens (after Ziegler 1989).

structure of the crust, including a presumably weak and heterogeneous crust in the Variscan Orogen embedded between older strong Precambrian crust of Canada and the Iberian peninsula (Fig. 3); (2) the prerift crustal thickness, taken to be 32 km by comparison to the thickness of the modern unextended crust beneath Iberia and Canada; and (3) radiogenic heat production of the crust, taken to be similar to that of basement rocks exposed in Iberia (Table 1). The models seek to reproduce: (1) the post-rift crustal thickness variation across

the margins at length scales greater than approximately 75 km, particularly moderately extended crust beneath the Galicia Interior Basin, unextended crust beneath the Flemish Cap and slightly extended crust beneath Galicia Bank, and highly extended crust beneath the marginal rift basins; (2) the lack of magmatism during rifting; (3) the duration of rifting; (4) the amount of magmatism required to produce thin oceanic crust at the time of breakup; and (5) the subsidence history of the margin, including Late Jurassic–late Valanginian

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(b)

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35 km thick crust

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120 Fig. 4. Yield strength envelopes for lithosphere based on rheological and thermal properties in Table 1 calculated at a strain rate of 10215 s21. (a) Lithosphere with nominal 32 km thick crust. (b) Lithosphere with 35 km-thick crust. (c) Lithosphere with weakened upper crust. (d) Steady-state geotherm. See text for discussion.

subsidence in the Galicia Interior Basin and late Valanginian–early Aptian subsidence seaward of Flemish Cap and Galicia Bank. The finite-element model formulation used in these simulations does not explicitly account for the formation of detachment faults, and so does not reproduce exhumation of the peridotite ridge west of Galicia Bank.

Modelling method The finite-element modelling program used in this research, STRCH95, is the most recent version of the FORTRAN 90 computer program STRCH developed by Dunbar (1988). The STRCH95 algorithm successively solves the two-dimensional heat

Table 1. Fixed model parameters* Constraint Continental crust thickness Oceanic crust thickness Lithosphere thickness Surface heat production, a0 Heat decay exponent, D Thermal conductivity, K (crust) Thermal conductivity, K (mantle) Specific heat, Cp (crust) Specific heat, Cp (mantle) Thermal expansion coefficient, a (crust) Thermal expansion coefficient, a (mantle) Extension rate Brittle yield strength, S0 Slope of brittle failure strength, B Ductile creep coefficient, A (crust) Ductile activation energy, Qc (crust) Ductile creep exponent, n (crust) Ductile creep coefficient, A (mantle) Ductile activation energy, Qc (mantle) Ductile creep exponent, n (mantle) *Symbols are defined in the text.

Prerift model value 32–35 km 3–4 km 120 km 3.1 m Wm23 12 km 2.5 Wm21 K21 3.4 Wm21 K21 875 J kg21 K21 1250 J kg21 K21 3.1  1025 K21 3.1  1025 K21 20 mm year21 60 MPa 5 MPa km21 5  10218 Pa2n s21 219 kJ mol21 2.4 4  10225 Pa2n s21 498 kJ mol21 4.5

References Pe´rez-Gusenye´ et al. (2003) Tucholke et al. (2004) Fernandez et al. (1998) Fernandez et al. (1998) Fernandez et al. (1998) Tejero & Ruiz (2002) Tejero & Ruiz (2002) Tejero & Ruiz (2002) Tejero & Ruiz (2002) Tejero & Ruiz (2002) Tejero & Ruiz (2002) Russell & Whitmarsh (2003) Bowling & Harry (2001) Bowling & Harry (2001) Bowling & Harry (2001) Bowling & Harry (2001) Bowling & Harry (2001) Bowling & Harry (2001) Bowling & Harry (2001) Bowling & Harry (2001)

A FINITE-ELEMENT RIFTING MODEL

conduction–advection –generation equation: rc

@T ¼ r  KrT þ A @t

ð1Þ

and the Navier –Stokes equation for flow in a viscoplastic lithosphere r

@u ¼ rF  rP þ hr 2 u @t

ð2Þ

where u is the particle velocity, T is temperature, r is the density of the rock, c is the specific heat, K is thermal conductivity, A is the volumetric heat production rate, F is body force per unit mass, P is pressure and h is viscosity. The physical properties r, c, A and K for each element in the finite-element mesh are specified by the user. Viscosity is defined by the rheology of each element, and is pressure, temperature and strain rate dependent. Viscosity is determined by empirical relations derived from experimental rock deformation studies. Initially, a uniform strain rate, 1˙, is assumed throughout the model. An effective viscosity, heff, for each element is determined from the estimated stress s and strain rate: heff ¼

s : 1_

ð3Þ

The stress in the above equation is taken to be the weaker of either a ductile (power-law creep) or plastic (a linear pressure-dependent yield criterion) rheology:  1=n 1_ eQc =nRT A

ductile:

s¼

plastic:

s ¼ S0 þ Bz

ð4Þ ð5Þ

where z is depth and R is the universal gas constant (other parameters are defined in Table 1). Once the effective viscosity is determined, the Navier– Stokes equation is solved. This results in an updated estimate of the strain rates in the model, which are used to revise the estimated effective viscosity of each element. The process is repeated until the strain rate estimates on successive iterations agree to within a convergence criterion prescribed by the user. The strain rates are then used to determine velocities at each node in the mesh, and the geometry of the model mesh is updated by stepping forward the locations of each node in the mesh a finite amount of time (time steps of 10 000 years duration are typical). The heat equation is then solved to update the thermal structure, using the new mesh geometry as a boundary condition and the thermal state of the model at the previous time

163

step as an initial condition. The new model temperatures and the strain rates from the past time step are used to calculate a revised estimate of the element viscosities to begin the next time step. Accuracy of the model solution is assessed empirically by comparing model results obtained with coarse and fine meshes and coarse and fine time step sizes. For a given family of models, a coarse mesh and time step size is initially selected. The mesh size and time step size are then halved and the model re-run. Refinement of mesh size and time step size continues until successive models produce the same result. If the model becomes highly deformed (individual elements develop large aspect ratios) a singularity develops in the linear system of equations that define the finite-element problem being solved. When this occurs, the model is remeshed. During remeshing, domain boundaries that define different materials used in the model (e.g. the model edges and the contacts between different rock types) are preserved. Within each domain, new elements are defined that have an aspect ratio as close to unity as possible. Temperatures, pressures and strain rate information within the model are preserved during the remeshing process. An important limitation to the STRCH95 program is the way in which the asthenosphere is treated. STRCH95 is a lithosphere-scale modelling program. The program does not explicitly consider convective deformation and heat transport in the asthenosphere. Rather, the asthenosphere is treated as a boundary condition on the lithosphere. For the temperature condition, we consider constant heat flux, constant temperature or adiabatic temperatures at the base of the lithosphere. The mechanical boundary condition is prescribed by Airy’s theory of isostasy, essentially treating the asthenosphere as an inviscid fluid. In terms of deformation, this is unlikely to have a major impact on model behaviour at the scales considered here as the mantle at these depths is much weaker than most of the lithosphere. However, this simple treatment of the asthenosphere may have a significant impact on melt production predicted by the model. In particular, conductive cooling of the upwelling asthenosphere beneath the rift axis (Pedersen & Ro 1992; Bown & White 1995), which is not considered here, may diminish melt production during rifting in comparison to the model predictions. Consequently, the mantle potential temperatures required in these models to match the Iberia –Grand Banks magmatic history probably underestimate the steady-state potential temperature of the asthenosphere during rifting. We therefore consider the asthenosphere temperature in these models to be a lower bound on the true mantle potential temperature. We note, however, that this

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is a complex problem. Small-scale convection in the asthenosphere, also not considered in these models, may have the opposite effect, enhancing melt production during rifting and therefore leading to an overestimate of mantle potential temperatures in the models (e.g. Mutter et al. 1988; Boutillier & Keen 1999). The meshing algorithm in STRCH95 is not designed to handle highly deformed elements that result from mantle convection, so we are unable to fully address the feedback between asthenosphere thermal evolution and lithosphere deformation.

Model description A basic assumption behind the modelling strategy used here is that pre-existing strength heterogeneities in the lithosphere determine the location and distribution of loci of extension during rifting (e.g. Dunbar & Sawyer 1989). Weaknesses in the upper and middle crust (crustal weaknesses) tend to produce short-lived extensional provinces during the early stages of rifting, whereas deep-seated weaknesses (mantle weaknesses) control the location where lithospheric necking develops and, hence, the location of eventual continental breakup. If these weaknesses are offset from one another, interior rift basins (controlled by the crustal weakness) often form landward of the deep offshore marginal basin (controlled by the mantle weakness) (e.g. Harry & Sawyer 1992). The models presented here include both of these types of weakness in the lithosphere (Figs 4 & 5). Mantle weaknesses are simulated as regions of thickened crust, which is attributed to the root

beneath the Variscan Orogen prior to rifting. This is represented with a simple triangular-shaped region of increased crustal thickness. Key variables describing the mantle weakness are the width of the region of thickened crust and the amount that the crust is thickened. A crustal weakness is simulated by a region in which both the plastic and ductile yield strengths are decreased relative to the surrounding crust. In the plastic regime, this involves a reduction of S0 (equation 5). For ductile behaviour, both the temperature-dependent terms (Qc/n) and temperature-independent terms (A 21/n) are reduced by similar amounts (equation 4). The crustal weakness may be attributed to lithological variations in the upper and middle crust resulting from accretion of various exotic terranes during the Variscan Orogeny, structural fabrics associated with Variscan sutures between these terranes or to pre-existing faulting near the Variscan front (Martinez Catalan et al. 1997; von Raumer et al. 2003; Simancas et al. 2003). Key variables describing this weakness in the model are the amount by which the yield strength is decreased, the width and thickness of the weakened crust, and the position of the crustal weakness relative to the mantle weakness. All of these variables were iteratively adjusted to produce a model that best fits the geometry and rift history of the Flemish Cap and Galicia Bank margins. Other modelling parameters, including heat production, specific heat, thermal conductivity, extension rate and nominal thickness of the crust (outside the Variscan Orogen), were held fixed (Table 1). The criterion used to determine when rifting ends in the models is the requirement

T = 0°C

crust

Crustal weakness

Ux

Ux

Mantle weakness

mantle T = Tm

Isostatic Buoyancy Pressure Fig. 5. Finite-element model construction. Numbers at the top and sides indicate constant temperature and constant extension rate boundary conditions.

A FINITE-ELEMENT RIFTING MODEL

that the crust thin to approximately 2 km at the rift axis, which is similar to the thinnest continental crust on the Galicia Bank and Flemish Cap margins (Chian et al. 1999; Gonzalez et al. 1999; Funck et al. 2003).

Results Results of a typical model are shown in Figures 6 –8. The model simulates in general the present structure of the Galicia Bank–Flemish Cap conjugate margins. In particular, the model predicts extreme crustal thinning beneath conjugate offshore rift basins that form during the late stages of rifting, formation of an interior basin on the eastern shelf during the early stages of rifting, and a region of less extended crust between the offshore and interior rift basins (Fig. 7). Most of the models examined in this study produced these basic attributes. Exceptions were models in which either the crustal or mantle weakness was very small (typically involving a reduction of yield strength in the crustal weakness of less than 5% or a change in crustal thickness in the mantle weakness of less than 2 km). In those models (not shown here), the dominant weakness controls rifting from the outset, resulting in formation of a nearly symmetric rift basin centred on the dominant weakness. All other models that were examined, involving changes in yield strength in the crustal weakness ranging from 5 to 30% and changes in crustal thickness in the mantle weakness ranging from 2 to 5 km, produced results that were generally similar to that shown in Figures 6– 8. However, changes in the relative locations and magnitudes of the two weaknesses produces a wide variation in the widths and depths of the interior rift basin and deep offshore basins, the duration of extension in the interior basin, the time elapsed before breakup, the amount and pattern of crustal thinning, and the presence or absence of a region of relatively unextended crust between the deep offshore and interior rift basins. As discussed below, changes in these parameters also have a significant effect on the magmatic history of the model. The model shown in Figures 6–8 is based on the variables that provide the best fit to observations on the Galicia Bank–Flemish Cap conjugate margins (Table 2). During the first 10 million years, extension results in rapid thinning of the crust in the region encompassing the crustal weakness (Figs 6 & 7a). Crustal thinning also occurs above the mantle weakness, but at a much slower rate. Tectonic subsidence in the region of weakened crust begins almost immediately after the onset of extension, marking the first stages of formation of the interior rift basin (Fig. 7b). The region above the mantle weakness, where the deep offshore basin ultimately forms, does not subside below sea level until approximately

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5 million years later. The two basins are separated by a structural high that undergoes moderate crustal thinning (from 32 to c. 20 km) during the first 5 million years of extension, and only a minor amount of thinning (c. 3 km) afterwards. This pattern of crustal thinning and subsidence can be understood in terms of the distribution of extension, illustrated by the model strain rates (Fig. 8). Extension in the crust is broadly distributed across the region encompassing the crust and mantle weaknesses during the first 10 million years, but is concentrated most strongly in the area of the crustal weakness (Fig. 8). Strain in the mantle is more evenly distributed across the preweakened region, and is coupled to deformation in the middle and upper crust by narrow regions of relatively high strain. Offsets in the initially vertical element boundaries show that the high strain region in the lower crust behaves as a subhorizontal detachment during this time (Fig. 6). Because extension in the mantle is initially distributed over a wide region, asthenospheric upwelling is broad and the rate of lithospheric thinning is relatively slow. As extension progresses, strain in the mantle begins to focus strongly beneath the mantle weakness. After 10 million years this deformation begins to propagate upwards, creating two regions of focused extension in the upper crust (the interior rift basin and marginal basin) that are separated by a region of relatively slow strain forming the structural high between the two basins. At this time, extension and subsidence is waning in the interior rift basin and accelerating in the marginal basin. By 15 million years extension throughout the lithosphere begins to focus exclusively in the offshore region (Fig. 8). Crustal thinning and subsidence ceases in the interior rift basin and on the structural high (Fig. 7). Because extension is more focused, lithospheric necking accelerates (Figs 6 & 8). The lithosphere beneath the rift axis thins to approximately half its initial thickness 20 million years after the onset of extension, and proceeds rapidly to breakup in the next 5 million years.

Discussion The thickness of the crust in the model generally agrees with the long wavelength (.75 km) crustal thickness variations on the Flemish Cap and Galicia Bank margins along the SCREECH-1 and ISE-1 seismic profiles (Figs 2 & 7a). The duration of the main extensional episode in the interior rift basin is approximately 13 million years, in general agreement with the late Tithonian–Valanginian phase of rifting in the Galicia Interior Basin (Tucholke et al. 1989; Murillas et al. 1990). The model predicts that deep water depths (.c. 1 km) did not become established in the offshore rift basin seaward of Galicia Bank until approximately 12 million years after the onset

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Novia Scotia 0 Ma

Variscan orogen

Iberia

Weak upper crust Crust

0 32 km

Mantle

122 km 5 Ma

10 Ma

15 Ma

20 Ma

25 Ma

V.E.5:1

500 km

Fig. 6. Finite-element mesh for the model that best fits Galicia Bank– Flemish Cap rifting. Numbers on the left indicate time (million years) elapsed since onset of extension.

A FINITE-ELEMENT RIFTING MODEL

Crust Thickness (km)

(a)

40 35 30 25 20 15 10 5 0 –200 –100

0 Ma 15

10

5

20 25

0

100

200

300

400

5 –200 –100 0 100 200 300 Distance from rift axis (km)

400

Elevation (km)

(b) –1

0

5

0 10 15

1 2

20

3 25 Ma

4

Fig. 7. (a) Crustal thickness variation in the finite-element model. Bold line indicates observed thickness in the cross-section of Figure 2. (b) Elevation predicted by the finite-element model (does not include syn- and post-rift sediment loading).

of extension. Breakup is predicted in early Aptian time (c. 118 Ma), 25 million years after the onset of extension, in agreement with the estimated onset of sea-floor spreading at around anomaly M0 time (Srivastava et al. 2000; Minshull et al. 2001; Funck et al. 2003; Hopper et al. 2004). The magmatic consequences of rifting in the model are assessed in terms of the timing, volume and spatial distribution of melt produced by adiabatic decompression melting in the upwelling asthenosphere (McKenzie & Bickle 1988). As discussed previously, we neglect conductive heat loss from the rising asthenosphere, which may be significant in slow spreading systems such as the

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Newfoundland –Iberia rift (Pedersen & Ro 1992; Bown & White 1995; Whitmarsh et al. 2001a, b), so our estimate of melt production should be considered to be an upper bound. Key parameters in the melt calculation are the shape of the ascending asthenospheric diapir, which is dictated by the pattern of lithospheric thinning in the model, and the mantle potential temperature, which is a variable in the model. A series of models were developed using mantle potential temperatures ranging from 1250 to 1300 8C as a basal boundary condition (Fig. 5). Other parameters in these models were identical to those in Tables 1 and 2. Each of these models produced patterns of extension and subsidence very similar to that shown in Figures 6–8. The only significant differences were the predicted melt production history. Maximum melt thicknesses range from 3.1 to 6.5 km at the rift axis at the time of breakup (Fig. 9). The models using a 1250 and 1275 8C mantle potential temperature produced a melt thickness of 3.1 and 4.7 km at the time of breakup, respectively. This is in good agreement with the 3 –4 km-thick oceanic crust adjacent to the Newfoundland and Galicia Bank margins (Whitmarsh et al. 1996; Hopper et al. 2004), suggesting a mantle potential temperature during rifting somewhere between these two values. As synrift conductive cooling of the asthenosphere was neglected, this should be considered a minimum estimate of the asthenospheric potential temperature prior to and during rifting. It seems unlikely, then, that the mantle potential temperature was more than about 30 8C cooler than the 1280 8C global average calculated by McKenzie & Bickle (1988). In any case, the presence of normal 7 km-thick oceanic crust within about 15 km of the oldest oceanic crust on the Galicia Bank margin (Whitmarsh et al. 1996) requires that mantle potential temperatures reached approximately 1280 8C within about 1–1.5 million years after breakup, assuming a 10–14 mm year21 sea-floor spreading half-rate (Whitmarsh et al. 2001a; Russell & Whitmarsh 2003). In the 1250 and 1275 8C potential temperature models, melt production does not begin until 2.5

Table 2. Variable model parameters Parameters Mantle weakness width (km) Mantle weakness amplitude (km) Crustal weakness width (km) Crustal weakness thickness (edges) (km) Crustal weakness thickness (centre) (km) Strength reduction in crustal weakness (%) Distance between centres of weaknesses (km) Temperature at bottom of lithosphere (8C)

Best-fit model value 120 3 80 12 26 25 110 1250, 1275, 1300

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Fig. 8. Second invariant of the strain rate in the finite-element model. Strain rates range from 10215 s21 (dark blue) to 10210 s21 (red).

Melt Thickness (km)

A FINITE-ELEMENT RIFTING MODEL

8 1300 °C, 25 Ma

6

1275 °C, 25 Ma

4

1250 °C, 25 Ma 2

1300 °C, 25 Ma

0 –200 –100

0

100

200

300

400

Distance From Rift Axis (km) Fig. 9. Melt production predicted by models using mantle potential temperatures of 1300, 1275 and 1250 8C.

or 5 million years before the onset of sea-floor spreading, respectively. No melt is predicted further than 45 km landward of the rift axis in either model, and melt thicknesses in excess of 2 km is restricted to within 35 km of the rift axis. Minor late-stage synrift magmatic episodes have been documented further south on the Iberian margin (Manatschal & Bernoulli 1999; Whitmarsh et al. 2001b) and Newfoundland margin (Tucholke et al. 2004), and similar evidence in the form of small mid-ocean ridge basalt (MORB)-like volcanic edifices and gabbroic intrusions emplaced on and within exhumed subcontinental mantle has been reported on the Adrian fossil non-volcanic rifted margin exposed in the Swiss Alps (Manatschal & Bernoulli 1999; Muntener et al. 2000; Muntener & Piccardo 2003). We surmise that a minor amount of late-stage magmatism, within the range of volumes predicted by the models here, is likely to be typical of non-volcanic rifted margins such as Newfoundland–Iberia. If the mantle potential temperature at the onset of sea-floor spreading was about 1250 8C, as the model implies, what are the implications for the pre- and synrift thermal state of the asthenosphere? We consider the possibility that a layer of moderately cool mantle (c. 1250 8C potential temperature) separated the lithosphere from normal asthenosphere (1280 8C potential temperature) at the onset of sea-floor spreading. As mentioned above, this is at the low end of mantle potential temperatures required to account for the thickness of the oldest oceanic crust. How thick would this abnormally cool mantle be? The melt predictions in Figure 9 are obtained by integrating the melt fraction over the thickness of the melt column. A mantle with potential temperature of 1250 8C begins to melt when it reaches a depth of approximately 40 km according to the parameterization given by McKenzie & Bickle (1988). Hence, at the onset of sea-floor spreading, a 40 km-thick layer of relatively cool mantle would account for

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the 2–4 km thickness of the oldest oceanic crust. Extrapolating the amount of lithosphere thinning at the rift axis at this time (90% compared to the prerift thickness) to the sublithosphere mantle suggests an original thickness of approximately 400 km for this cool mantle layer. Although this simple analysis neglects the complexity of deep mantle flow during rifting, it does provide a firstorder test of the possibility that a moderately cool layer on the scale of the asthenosphere is consistent with the melt production and lithosphere thinning history of the margin. This is generally consistent with the hypothesis of Reston & Morgan (2004), who appeal to a cool prerift asthenosphere to explain the development of non-volcanic rift margins (although we note that our model invokes a much warmer lithosphere than they advocate). A problem with the scenario described above arises when we consider the time elapsed (1–1.5 million years) between the onset of sea-floor spreading and production of approximately 7 km-thick oceanic crust. The thinning rate of the model lithosphere beneath the rift axis after 25 million years of extension is 2.5 km Ma21. If the thinning rate of the cool layer in the asthenosphere is similar, then only 2.5– 3.75 km of cool sublithospheric mantle could be removed between the onset of sea-floor spreading and production of 7 km-thick oceanic crust. Reston & Morgan (2004) recognized this problem, and advocated along-strike movement of warm asthenosphere at the end of the rifting episode to account for the transition between amagmatic rifting and production of approximately 7 km-thick oceanic crust. An alternative is that the potential temperature throughout the asthenosphere in this area was close to the global average prior to rifting, and the moderate cooling needed to explain the low melt production at the onset of sea-floor spreading is a result of synrift conductive cooling within the upwelling asthenosphere diapir (e.g. Pedersen & Ro 1992; Bown & White 1995). In such a scenario, asthenosphere temperatures in the rift neck at the onset of sea-floor spreading would be quite low at the top and gradationally increase to about 1280 8C potential temperature at depth, in such a way that the total melt production would be similar to that of a moderately cool (1250 8C) asthenosphere. Because the cool mantle would be restricted to within the rift neck and concentrated primarily at shallow depths, this would allow for a rapid transition to normal mantle temperatures and mid-ocean ridge melt production volumes soon after sea-floor spreading began. This scenario is generally consistent with the spatial and temporal distribution of magma production observed on the margins, and does not require adhoc supposition of a pre-existing thin layer of anomalously low temperature in the asthenosphere.

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Summary Finite-element models simulating rifting between Galicia Bank and Flemish Cap invoke pre-existing weaknesses in the lithosphere to account for the shift in the locus of extension from the Galicia Interior Basin during the first half of the rifting episode to the deep offshore rift basin during the second half of the rifting episode. Two forms of weaknesses were examined: an upper mantle weakness, created by excess crustal thickening over a 120 km-wide region that is attributed to the central part of the prerift Variscan Orogen; and a crustal weakness created by reducing the yield strength of the upper and middle crust in an 80 km-wide region that is attributed to pre-existing structural fabrics in the eastern Variscan front. The model results were found to be robust over excess crust thicknesses ranging from 2–5 km and upper and middle crust weakening ranging from 5 to 30%. Under these circumstances, all models produced an early interior rift basin followed by a shift in extensional deformation to a deep offshore basin where continental breakup ultimately occurs. Details of the width and depth of the basins, the time at which deformation shifts from the interior rift basin to the offshore basin, and the pattern of crustal thickness at the time of breakup are determined by the relative positions, shape and magnitudes of the weaknesses imposed in the model. Model parameters in Tables 1 and 2 produced the general features of the last 25 million years of rifting on the Flemish Cap and Galicia Bank continental margins, including 13 million years of extension in the Galicia Interior Basin, 12 million years of extension in the Galicia Bank and Newfoundland Cap marginal basins, and formation of moderately extended crust on Galicia Bank. The model geometry at the time of breakup approximates structural features of greater than about 75 km wavelength on these margins. The non-volcanic nature of rifting, ultimate production of thin (2–4 km thick) oceanic crust and rapid transition to generation of normal thickness (7 km) oceanic crust within 1–1.5 million years after breakup requires a mantle potential temperature of 1250– 1275 8C, roughly 5–30 8C cooler than the global average at mid-ocean ridges. This modest amount of cooling is attributed to conductive cooling of the ascending asthenosphere during the late stages of rifting. At these mantle temperatures, the model predicts production of 3.1–4.5 km of new oceanic crust at the time of breakup. Synrift magmatism is limited to the last 5 million years or less of extension and to within 45 km of the locus of continental breakup. Melt thicknesses of greater than 2 km are restricted to within 35 km of the locus of continental breakup.

This research was supported by National Science Foundation grant OPP 0408475. We thank the organizers of the InterMargins Modelling of Extensional Deformation of the Lithosphere workshop held in Pontresina, Switzerland in July 2004 for providing a venue for discussions that led to this modelling study, and the many participants of the workshop who provided data used to constrain the models. S. Grandell received support from the Ronald E. McNair Postbaccalaurate Achievement Program for this research. The authors thank R. Buck, D. Mu¨ller and F. Tsikalas for their helpful reviews.

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B ALKWILL , H. R. (eds) Extensional Tectonics and Stratigraphy of the North Atlantic Margins. AAPG Memoirs, 46, 247–263. T UCHOLKE , B. E., S IBUET , J.-C. & K LAUS , A. & SHIPBOARD SCIENCE PARTY . 2004. Proceedings of the Ocean Drilling Program Initial Reports 210 (online). Available from World Wide Web: http://www-odp. tamu.edu/publications/210_IR/210ir.htm. VON R AUMER , J., S TAMPFLI , G. & B USSY , F. 2003. Gondwana-derived micro continents – the constituents of the Variscan and Apline collisional orogens. Tectonophysics, 365, 7– 22. W HITMARSH , R. B. & S AWYER , D. S. 1993. The ocean/ continent transition beneath the Iberia abyssal plain and continental-rifting to seafloor-spreading processes. In: WHITMARSH , R. B., SAWYER , D. S., KLAUS , A. & MASSON , D. G. (eds) Proceedings of the Ocean Drilling Program, Scientific Results, 149. Ocean Drilling Program, College Station, TX, 713–733. W HITMARSH , R. B., W HITE , R. S., H ORSEFIELD , S. J., S IBUET , J.-C., R ECQ , M. & L OUVEL , V. 1996. The ocean–continent boundary off the western continental margin of Iberia: Crustal structure west of Galicia Bank. Journal of Geophysical Research, 101, 28,291–28,314. W HITMARSH , R. B., D EAN , S. M. & M INSHULL , T. A. 2000. Tectonic implications of exposure of lower continental crust beneath the Iberia Abyssal Plain, Northeast Atlantic Ocean: Geophysical evidence. Tectonics, 19, 919–942. W HITMARSH , R. B., M ANATSCHAL , G. & M INSHULL , T. A. 2001a. Evolution of magma-poor continental margins from rifting to sea floor spreading. Nature, 413, 150– 154. W HITMARSH , R. B., M INSHULL , T. A., R USSELL , S. M., D EAN , S. M., L OUDEN , K. E. & C HIAN , D. 2001b. The role of syn-rift magmatism in the rift-to-drift evolution of the West Iberia continental margin: geophysical observations. In: W ILSON , R. C. L., W HITMARSH , R. B., T AYLOR , B. & F ROITZHEIM , N. (eds) Non-volcanic Rifting of Continental Margins: A Comparison of Evidence From Land and Sea. Geological Society, London, Special Publications, 187, 107–204. Z IEGLER , P. A. 1989. Evolution of the North Atlantic – An overview. In: T ANKARD , A. J. & B ALKWILL , H. R. (eds) Extensional Tectonics and Stratigraphy of the North Atlantic Margins, AAPG Memoirs, American Association of Petroleum Geologists, Tulsa, Oklahoma, 46, 111– 129.

A kinematic modelling approach to lithosphere deformation and basin formation: application to the Black Sea S. S. EGAN1 & D. J. MEREDITH2 1

Earth Sciences and Geography, School of Physical and Geographical Sciences, Keele University, Keele, Staffordshire ST5 5BG, UK (e-mail: [email protected]) 2

Grid Technology Group, Daresbury Laboratory, Warrington, Cheshire WA4 4AD, UK (e-mail: [email protected])

Abstract: A kinematic model of lithosphere deformation has been developed that integrates the following components: structural deformation of the crust and mantle lithosphere; thermal conditioning, perturbation and subsequent re-equilibration of the lithosphere temperature field; flexural isostatic adjustments; and surface processes, including both lateral and temporal variations in basin fill and bathymetry. This approach enables the forward modelling of extensional basin evolution in two and three dimensions followed by deformation due to subsequent extensional and compressional (i.e. inversion) events. The model has been applied to the Black Sea, which is one of the deepest basins in the world and yet it is poorly understood in terms of the mechanisms that have controlled its evolution. Although it is widely accepted that this basin was initiated by Mesozoic back-arc extension related to the subduction of the Tethys Plate to the south, most of the subsidence observed today occurred within the Palaeogene and Neogene (i.e. within the framework of the Alpine–Himalayan orogenic belt). The modelling approach described above has been used to test possible geological and geodynamic mechanisms that have controlled the subsidence history of the Black Sea. In particular, the investigation has focused on trying to explain the anomalously thick post-rift subsidence that occurred in the basin. Models assuming uniform lithosphere extension do not generate the observed thickness of sediment infill in the basin. Similarly, modelling of the compressional deformation around the edges of the basin structure does little to explain the large magnitude of subsidence within the centre of the basin. Model results show that the observed basin depths can be attained only when the total magnitude of deformation is constrained from crustal thickness changes rather than by fault displacement measurements.

Kinematic models define the dimensions of all components of a system as well as a set of relationships that describe how the parts in the system interact and evolve. This paper is concerned with the kinematic modelling of the geological and geodynamic processes that occur during the extension and compression of the continental lithosphere, but placing emphasis on how this deformation controls uplift and subsidence at the surface. The main advantage of the kinematic modelling of lithosphere-scale deformation is that it can be used to investigate the effects of processes that are poorly constrained by geophysical and field data, and to provide a better understanding of the structural, thermal and stratigraphic development of deep frontier basins where well control is poor. The first kinematic models of lithosphere extension and basin formation (e.g. McKenzie 1978; Royden & Keen 1980) combined the crustal thinning, thermal and local isostatic processes associated with the instantaneous stretching of the lithosphere to quantify surface subsidence or, sometimes, uplift. Although these models are now

considered relatively simplistic, they are based on concepts that underpin much of the lithosphere-scale basin modelling being carried out today. One of the main aims of this paper is to describe the individual processes that play a key role in deformation at lithosphere scale and how these processes have been integrated within a twodimensional (2D) and 3D kinematic modelling environment. This approach enables the forward modelling of extensional basin evolution followed by subsequent extensional or compressional deformation. The modelling approach has been applied to provide insights into the evolution of a specific case study in the geological record, namely the Black Sea. The evolution of the Black Sea region represents an interference of tectonic events over geological time in that most of the subsidence (up to 14 km) took place within the basin when the immediate surrounding regions were experiencing compressional deformation. Few modelling studies have focused on the evolution of the Black Sea such that the mechanisms that have controlled

From: KARNER , G. D., MANATSCHAL , G. & PINHEIRO , L. M. (eds) Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 173– 198. DOI: 10.1144/SP282.9 0305-8719/07/$15.00 # The Geological Society of London 2007.

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the subsidence history of the basin remains poorly constrained (Brunet & Cloetingh 2003). The main aim of this case study is to both test and apply 2D and 3D kinematic models of lithosphere extension to obtain a better understanding of the geological and geodynamic processes that have contributed to the subsidence history of the Black Sea basin. Models have been generated to investigate the evolution of the basin in the context of uniform lithosphere extension, compressional deformation and loading at the edges of the basin, and depthdependent lithosphere extension in the lower crust and mantle lithosphere. The modelling has also been applied to the Iberian margin as part of a benchmarking exercise for the comparison of different modelling approaches carried out as part of the 2004 InterMARGINS Extensional Deformation of the Lithosphere (IMEDL) workshop (see the Appendix).

Lithosphere extension and basin formation: historical background The deformational response of a rheologically layered lithosphere to applied tensional and compressional forces is highly variable. A wide range of basin-scale phenomena have been attributed to the varying mechanisms of lithosphere-scale deformation, such as the difference in width between extending regions (e.g. Buck 1991) and the generation of asymmetric rift basins (Coward 1986). Despite the complexity of lithosphere-scale deformation, crustal thickness changes are generally modelled by either simple shear (i.e. faulting and ‘brittle’ deformation: e.g. Wernicke 1985) or pure shear (i.e. stretching or squashing: e.g. McKenzie 1978) or by a combination of the two (e.g. Kusznir et al. 1987). The pure shear and simple shear mechanisms represent end-member lithosphere deformation processes and were used as a foundation for the early models of basin formation. One of the first numerical models to link deformation of the continental lithosphere to extensional basin formation, using a kinematic approach was the McKenzie or uniform lithosphere extension model (McKenzie 1978). This model, which is shown diagrammatically in Figure 1a, assumes that the whole of the lithosphere deforms by a pure shear or stretching mechanism quantified by the b or extension factor, equivalent to 1 plus the strain. The McKenzie model is essentially 1D in that it generates results in the form of subsidence curves (Fig. 1b), which show the rift and post-rift phases that characterize many extensional sedimentary basins. The rift phase is caused by the combined effects of crustal thinning, rift-induced heating owing to raising of the lithosphere–

asthenosphere boundary and isostatic compensation. It is followed by the post-rift or thermal subsidence phase primarily driven by the gradual cooling of the lithosphere back to an equilibrated state. The McKenzie model and various derivatives, such as the non-uniform lithosphere extension model (Royden & Keen 1980; Hellinger & Sclater 1983), have been applied to many extensional basins in the geological record (e.g. Barton & Wood 1984) and can explain successfully the subsidence history of those basins that have experienced a relatively uncomplicated evolution of rifting followed by gradual post-rift/thermal subsidence. There was another significant period of advancement in lithosphere-scale modelling during the mid –late 1980s, which was mainly driven by new insights into crustal and upper mantle structure being gained from the acquisition and interpretation of deep seismic data. For example, deep seismic reflection data acquisitioned by the BIRPS group from the United Kingdom (e.g. Brewer & Smythe 1984; BIRPS & ECORS 1986) reveals the important role played by crustal faults during basin formation. The importance of faulting in lithosphere-scale deformation was further emphasised in a qualitative model presented by Wernicke (1985). This model, which is mainly based on observations made from the Basin & Range province in the western United States, assumes lithospheric deformation to be controlled by brittle failure and faulting (i.e. simple shear) in the form of a large-scale shear zone that propagates through the entire lithosphere with a dip at a low angle. The model, therefore, effectively divides the lithosphere into an upper hanging-wall plate and a lower footwall plate. As a result of simple shear, upper crustal (thin-skinned) deformation is spatially offset from zones that have experienced significant lower crust and mantle thinning. Consequently, the model predicts that a basin should be characterized by absent or minimal thermal subsidence in the region that has experienced significant upper crustal deformation. The major weakness of this model is that it fails to account for basins that have post-rift thermal subsidence sequences overlying the regions of major fault controlled subsidence. Deep seismic reflection data (e.g. Snyder & Hobbs 1999), along with studies of lithosphere deformation (Wernicke 1985), lithosphere rheology (e.g. Kusznir & Park 1987) and seismological investigations (Jackson & McKenzie 1983), have been very important in revealing the large-scale layering and structure of the continental lithosphere within various tectonic regimes. In addition, these studies suggest that lithosphere extension occurs as a hybrid of fault-controlled deformation within the brittle layer of the crust and pure shear extension within the lower crust and mantle lithosphere that

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Fig. 1. The uniform lithosphere extension model (McKenzie 1978). (a) The model assumes that the lithosphere extends by pure shear and quantifies subsidence owing to crustal thinning and uplift caused by the raising of hotter material at depth nearer to the surface, along with the associated Airy isostatic compensation. The second part of the model simulates the thermal subsidence phase of basin evolution. (b) Subsidence curve generated by the McKenzie model showing basin depth following 100% lithosphere extension (i.e. b factor ¼ 2).

dominantly exhibit ductile rheological behaviour (Coward 1986; Kusznir et al. 1987). The next section describes how these developments have been extended in improved kinematic models that account more satisfactorily for the evolution of sedimentary basins.

Integrated kinematic modelling of lithosphere extension and compression This section describes the major processes that occur during lithosphere extension and compression along with some of the approaches that are used to numerically model them. These processes are then integrated into a kinematic model of continental lithosphere deformation.

Mechanical thinning or thickening of the crust owing to movement along crustal faults Extensional movement along crustal faults generates the foundations of basin structure. It also causes crustal thinning and, therefore, negative loading of the lithosphere, which responds by

flexural isostatic rebound. In contrast, reverse movement along crustal faults generates increased surface topography (i.e. crustal thickening) as well as constituting a downwards-acting load on the lithosphere, which responds by flexural isostatic subsidence. Fault configuration within the model is defined in terms of the number of faults required, and their relative spacing and orientation (i.e. synthetic or antithetic). All of the faults are assumed initially to terminate or detach at mid –lower crustal levels in order to maintain compatibility with evidence from deep seismic reflection data (e.g. Hall et al. 1984; BIRPS & ECORS 1986), seismological investigations (e.g. Jackson & McKenzie 1983) and rheological modelling of the lithosphere (e.g. Kusznir & Park 1987). There are several methods, mainly based on geometric constructions, that can be used to simulate deformation caused by fault movement. One the most popular of these constructions is the vertical shear or Chevron construction (Verrall 1982). It assumes that each vertical thickness of hanging wall is displaced laterally by the same amount of extension, usually defined in terms of a heave value. The next assumption is that any unsupported

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part of the hanging wall following extension deforms such that it collapses directly downwards onto the underlying fault and footwall, which are assumed to be rigid. Geometric methods such as the Chevron construction and derivatives, like the inclined shear method (White et al. 1986), are very suitable for numerical modelling applications because they give reasonably accurate results for lithosphere-scale deformation and are computationally efficient.

Mechanical thinning or thickening of both crust and mantle lithosphere by a regionally distributed pure shear mechanism Although the deformation caused by fault movement plays an important role during lithosphere deformation, seismic data and rheological studies indicate that the lithosphere is mostly too ductile below depths of about 15– 20 km to support major fault structures. Instead, the lower crust and parts of the mantle are thought to deform by movement along a network of shear zones that split and merge over relatively short distances (e.g. Reston 1988). One approach that can be used to model this type of regionally distributed deformation is pure shear, stretching under extension and squashing under compression (Kusznir et al. 1987). Both the lateral position and the width of the pure shear deformation can be defined in the model. The overall deformation by both faulting and pure shear is initially area balanced in the profile in order to avoid mass balance problems.

Disturbances caused to the lithosphere temperature field Lithosphere extension and compression causes an overall heating and cooling, respectively. For example, reverse movement along crustal faults causes the emplacement of relatively hot hangingwall material onto that of a cool footwall, creating a cooling of the geotherm at shallow depth. Compressional deformation by pure shear thickens the lithosphere and also reduces the geothermal gradient. Within the model the lithosphere temperature field is represented as a 2D or 3D grid. Each node on the grid is assigned a predeformational temperature. The temperature of each grid node is then modified according to the amount of deformation experienced.

Re-equilibration of the temperature field after deformation The lithosphere temperature field experiences gradual thermal recovery after a major tectonic

event. For example, in the context of extension the lithosphere cools back to an equilibrated state. The model calculates both lateral and vertical heat flux by the mechanism of conduction. The physical properties of the lithosphere are such that the thermal recovery process takes of the order of 100 million years but follows an exponential decay path.

Flexural isostatic compensation of tectonic loads Crustal thinning and thickening, and thermal perturbations caused by tectonic activity, impose loads on the lithosphere that have to pass through an isostatic filter to give a compensated amount of subsidence or uplift observed at the surface. The flexural isostatic response of the lithosphere to loading is calculated within the model in order to generate a realistic surface or basement profile and underlying crustal structure. The methodology used to model flexure assumes that the lithosphere behaves as a continuous elastic plate, which is in equilibrium under the action of all applied loads (Turcotte & Schubert 2002). The flexing properties of the lithosphere are defined by its flexural rigidity (Walcott 1970; Watts 2001), which is, in turn, set in the model by the parameter effective elastic thickness (Te).

Surface processes The model contains simple algorithms to simulate basin infill and the erosion of uplifts. Isostatic adjustments are calculated in response to these processes such that sedimentary infill enhances subsidence within a basin, whereas erosion has the combined effect of reducing the size of the uplifts and unloading the lithosphere, which responds by regional isostatic uplift or rebound (Egan & Urquhart 1993). A comprehensive description of all theoretical aspects of the above model are presented in Kusznir & Egan (1989) and Egan & Urquhart (1993), and will not be repeated here. Figure 2, however, illustrates some of the parameters included in the model calculations and examples of results. A typical starting condition for the modelling is shown in Figure 2a, which illustrates a regional crosssection of undeformed lithosphere. The crustal component of this lithosphere is assumed to be 35 km thick with a density of 2800 kg m23, while the mantle lithosphere is assumed to be 90 km thick with a density of 3300 kg m23. The modelled lithosphere is thermally conditioned with an equilibrated geotherm, which has a surface temperature of 0 8C and a temperature at the lithosphere–asthenosphere boundary of 1333 8C. All of these parameters can

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Fig. 2. An integrated kinematic model of lithosphere deformation. (a) Cross-section of undeformed lithosphere showing initial values of parameters used in the model calculations. This model can be applied to simulate; (b) lithosphere extension and associated basin formation; (c) lithosphere shortening and associated thrust-belt and foreland basin formation, or a combination of the two.

be varied according to the tectonic scenario to be modelled. The effects of extending this lithosphere are shown in Figure 2b. The model shows how the crust has accommodated extensional deformation by movement along a sequence of faults, while the mantle lithosphere has been deformed by regionally distributed pure shear. The boundary

between upper crustal deformation and that in the lower crust and mantle lithosphere is determined by a detachment depth, which is equivalent to the concept of necking depth presented by other authors (e.g. Kooi et al. 1992). The model shows a basement profile consisting of a sequence of closely spaced half-grabens with relative uplift of the footwall owing to flexural isostatic processes

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(Weissel & Karner 1989; Egan 1992). Extension has also caused heating of the lithosphere, which subsequently has cooled to generate gradual subsidence. The effects of this can be seen in the model by a post-rift stratigraphic sequence that blankets the underlying fault blocks and synrift sequence. Lithosphere shortening is represented by the model in Figure 2c, which shows a ‘piggy-back’ style of thrusting adjacent to a foreland basin that has been generated mainly by flexure in response to crustal thickening caused by compressional tectonics. In addition, the model is sufficiently versatile to include multiple extensional and compressional events.

Application of integrated kinematic modelling of lithosphere extension and compression to investigate possible subsidence mechanisms within the Black Sea basin The modelling approach described in the previous section has been applied to the Black Sea, which is a deep, semi-isolated marine basin located north of Turkey and south of Ukraine and Russia. The main aim of the work described in this section has been to use kinematic modelling to try and decipher the role of the main subsidence mechanisms within this poorly understood basin.

An overview of the regional structure and stratigraphy of the Black Sea basin The Black Sea basin covers an area of approximately 423 000 km2, has a present maximum bathymetry of 2200 m and comprises the western and eastern Black Sea sub-basins, which are separated by the Mid-Black Sea high (Fig. 3). The origins of the basin go back to a phase of back-arc extension during the mid-late Cretaceous when northwards subduction was occurring beneath modern-day Turkey representing the closure of Tethys Ocean (Hsu¨ 1977; Letouzey et al. 1977; Zonenshain & Le Pichon 1986; Go¨rur 1988; Okay et al. 1994). Extension continued throughout the upper Cretaceous, perhaps leading to continental separation and the formation of oceanic crust. However, most of the 12–14 km of subsidence experienced by the central parts of the basin took place when the surrounding regions were experiencing compressional deformation. This is demonstrated by Alpine– Himalayan thrust belts, including the Pontides, Balkanides, Crimean Mountains and Greater Caucasus (Fig. 3), that surround the Black Sea.

Regional cross-sections are presented (see Fig. 3 for location) to illustrate the structure and stratigraphy within the western and eastern Black Sea basins. The section presented in Figure 4a shows structure and stratigraphy from offshore Turkey into the centre of the western Black Sea. It is a section constructed from the depth-converted interpretation of a regional seismic line acquired by BP and TPAO. Interpreted horizons have been constrained by data from a well (Akc¸akoc¸a-1) located in the southern part of the section, which penetrated down into rocks of Cretaceous age at a depth of about 2000 m. The original seismic data are confidential, but a detailed description of its interpretation and analysis can be found in Robb (1998). The cross-section is dominated by major extensional faults, generated by an Albian– Aptian rift phase, which has produced half-graben structures separated by areas of relative footwall uplift. Go¨ru¨r (1988) relates this rift phase to the opening of the western Black Sea due to back-arc extension overlying a northerly dipping subduction zone. However, the magnitude of extension associated with this rift phase was not very great and amounts to about 3 km of heave on the faults observed in the section (approximately 2% overall extension). The deposition of middle–upper Cretaceous synrift sequences was followed by a broad, thin blanket of Paleocene deposits, which from its general distribution is interpreted as a post-rift thermal sag phase. The section becomes more complicated in the Eocene with the contractional reactivation of extensional faults and the development of new reverse faults at the basin margin. In contrast, the northern part of the cross-section shows that the centre of the Black Sea was subsiding very rapidly from the Palaeogene to the present, with the deposition of over 12 km of sediment. The Ukrainian Black Sea region is dominated by basin inversion as shown by the section in Figure 4b. The sequence of tectonic events that have generated the overall structure and stratigraphy exhibited in this section are, however, very similar to those already described in the context of the Turkish section. Again, there is evidence of rifting during the middle Cretaceous. The extensional evolution of this part of the basin was terminated during the Eocene by an inversion phase. Inversion continued into the Oligocene, generating uplift and erosion of the earlier syn- and post-rift sequences. A foreland basin-type sequence was generated in the Miocene as a response to the crustal thickening and flexure caused by the inversion tectonics. Finally, the Neogene sequence indicates a dramatic increase in subsidence towards the centre of the Black Sea, which has also caused the proximal parts of the shelf to subside flexurally. The deeper structure of the western Black Sea basin is a matter of controversy because of an

Fig. 3. Black Sea location map showing major tectonic elements (adapted from Robinson 1997) and location of cross-sections used for modelling study of the basin. The rectangle positioned within the eastern Black Sea shows the approximate extent of the 3D model (see Figs 10 & 11).

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Fig. 4. Sections showing structural and stratigraphic components across the western Black Sea (see Fig. 3 for location information). (a) Depth section across the Turkish part of the basin. (b) Two-way time (TWT) section across the Ukrainian part of the basin. (c) Composite TWT section across the central (western) Black Sea (adapted from Finetti et al. 1988) based on the interpretation of several regional seismic lines (vertical dashed lines indicate the boundaries between the individual seismic lines).

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absence of well data and inadequate seismic reflection data at depth. The section presented in Figure 4c illustrates the overall structural and stratigraphic style in the central part of the basin, and is based upon interpreted seismic data presented by Finetti et al. (1988). This section shows that the central western Black Sea basin is characterized by a flat-lying sedimentary sequence, which reaches depths of about 14 km. Although the seismic data do not show deep structure, Finetti et al. (1988) suggested that the basin-fill sequence has been deposited on flat-lying basement, probably basaltic in composition. Evidence from gravity and seismic refraction data (Neprochnov et al. 1970; Belousov 1988) also suggests that western Black Sea basement may consist of oceanic crust. There is, however, no direct evidence as to whether the density structure determined from these data is due to oceanic crust or to highly attenuated and/ or metamorphosed continental crust. Four regional-scale sections across the eastern Black Sea are presented in Figure 5 (see Fig. 3 for locations). These sections are derived from the interpretation of seismic reflection lines acquired by BP and TPAO that have been interpreted by BP personnel and subsequently modified (Meredith & Egan 2002). The original seismic data are confidential. Section D (Fig. 5a) shows structure and stratigraphy from onshore Turkey, into the centre of the eastern Black Sea and finally onto the Russian shelf in the NE of the region. As for the western part of the basin, the section is dominated by extensional faults. In particular, the Archangelsky and Shatsky ridges represent regionally uplifted footwall blocks to extensional faults that form the southern and northern continental slope regions, respectively. This interpretation is in agreement with Banks et al. (1997), who suggest that the Shatsky ridge defines the northern rift margin of the eastern Black Sea. Similar to the western Black Sea, fault displacement measurements indicate that the magnitude of extension was not very great and amounts to about 60 km of heave on the faults observed in section D. This equates to an overall magnitude of extension of approximately 13% (i.e. b ¼ 1.13). Thickness variations across the oldest horizons in the section suggest that rifting occurred later than in the western sub-basin, possibly some time in the Paleocene or Eocene. This timing correlates with a late Paleocene –early Eocene rift event suggested by Banks & Robinson (1997) and Spadini et al. (1997), or early –middle Eocene rifting inferred by Kazmin et al. (2000). The onshore and shelf regions of the eastern Black Sea become more complicated at the end of the Eocene with evidence of compressional deformation. Compressional tectonic features are best developed in the NE part of section D (Fig. 5a)

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where the Tuapse fold belt, which forms the southern limit of the Great Caucasus mountain belt, shows reverse movement along basement faults. Immediately adjacent to this thrust belt is the Tuapse Trough, which is superimposed onto the footwall block forming the Shatsky ridge. The deeper structure of the central part of the eastern Black Sea basin is a matter of controversy because of the general poor quality of seismic reflection data at depth. The sections presented in Figure 5 show a similar structural and stratigraphic style to the western Black Sea in that the central part of the basin is characterized by a ‘layer-cake’ stratigraphy, which is present to depths of over 12 km.

The western Black Sea basin – application of a 2D kinematic approach The three cross-sections described in Figure 4 have been combined to produce a regional cross-section that extends from the Ukrainian to Turkish margins of the western Black Sea (Fig. 6). Possible processes that have contributed to the subsidence history exhibited by this section are explored below using the kinematic modelling approach described in the previous section on ‘Integrated kinematic modelling of lithosphere extension and compression.’ Has the western Black Sea evolved due to uniform lithosphere extension combined with loading from the surrounding thrust belts?. Figure 7 shows a model of the western Black Sea using the methodology described in the previous section. The overall dimensions of the model, along with major fault structures, are based upon the regional north –south section across the western Black Sea presented in Figure 6. In addition, deformation by faulting has been balanced to regionally distributed stretching at depth to represent uniform lithosphere extension. The model results show clearly that it is not possible to reproduce Black Sea subsidence with extensional deformation constrained by the magnitude of faulting. The magnitude of observed deformation as determined from extensional movement along fault structures amounts to 6.75 km total heave (equivalent to b of approximately 1.01) across the section. In addition, this deformation is mainly confined to the basin margins and can only account for overall subsidence in these parts of the basin. However, the model shows less than 1 km of subsidence in the central part of the basin, which is mainly produced by thinning of the lower crust, thermal subsidence (modelled for a 100 million years time period) and sediment loading in response to pure shear deformation. This amount falls far short of the 14 km of total subsidence indicated by seismic data.

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Fig. 5. North–south sections across the eastern Black sea (see Fig. 3 for locations) showing a dominance of compressional tectonics within the Turkish and Russian onshore and near-shelf regions. Extensional tectonics dominates both northern and southern continental slope regions. The deep basin, however, shows a thick, flat-lying sedimentary sequence that is mostly devoid of structure.

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Fig. 6. Composite section (adapted from Robb 1998) across the western Black Sea based on the combination of the individual sections presented in Figure 4.

It is possible that a significant component of post-rift subsidence was caused by compressional tectonics. More precisely, the loading generated by the thrust belts surrounding the Black Sea may have caused the development of foreland basins, which constructively interfered to produce one large basin. Cobbold (1993) refers to this type of structure as a ‘push-down basin,’ while Graham (1995) calls them ‘Central Asian type basins,’ which have the characteristics of being ‘relatively symmetric, relatively fixed and commonly contain enormous thicknesses of sediment.’ The effects of loading from surrounding thrust belts have been modelled in the context of the western Black Sea. The model profile presented in Figure 7c represents the extensional evolution of the basin followed by compression at the margins occurring at a modelled time of 50 Ma to represent the onset of deformation due to the Alpine–Himalayan orogeny. The magnitude of both extensional and compressional deformation is constrained by the fault movement measured from the cross-sections in Figures 4 & 6. A lack of data prevents widening of the model to reproduce compressional deformation observed in the Pontides and Ukraine. However, it is very evident that the flexural subsidence generated in response to the formation of the thrust belts deepens the original extensional basin at its edges, rather than in the centre. The model also shows how crustal thickening has caused uplift of pre-, syn- and postrift sequences at positions 10 and 500 km along the section. These uplift structures can also be observed in the data, but they are not subaerial. In addition, the model does not generate sufficient flexural subsidence in the centre of the basin. Crustal thinning beneath the western Black Sea. Although there appears to be a lack of extensional faulting within the western Black Sea, it cannot be

discounted that there are more extensional faults than resolved by the seismic data and that the magnitude of deformation assumed within the models presented so far is an underestimation. The model results presented in Figure 8 are also based on uniform lithosphere extension. However, the magnitude of extension has been increased significantly compared to the models presented in Figure 7. This is because the magnitude of deformation has been calculated using crustal thickness rather than basement faulting. The cross-section presented in Figure 8a has been adapted from that in Figure 6 to show the relative thickness of syn- and post-rift stratigraphy. The boundary between these sequences is placed at the division between the Cretaceous and Paleocene, as shown in Figures 4 & 6. The section also shows an estimation of the depth to basement (i.e. base synrift). Although this boundary cannot be accurately determined from the data available, it is assumed to occur at the base of the Upper Cretaceous (Go¨ru¨r 1988; Spadini et al. 1996). An estimation of Moho topography is also displayed and is based on commercial gravity data and published material (e.g. Neprochnov et al. 1970, 1974; Belousov 1988). The prerift thickness of the crust in the Black Sea region is open to debate. However, Spadini et al. (1996) propose that the western Black Sea formed by the rifting of the Moesian-European platform, which has a characteristic crustal thickness of 35 km. From the section presented in Figure 8a it is possible to determine the magnitude of crustal thinning or thickening relative to a regional average and, in turn, to calculate a sequence of b values (i.e. 1 þ strain: McKenzie 1978) across the basin. This b distribution has been used to constrain the magnitude of extension and compression evident in the model presented in Figure 8b, which assumes uniform deformation of the lithosphere based entirely on a pure shear mechanism. The

Fig. 7. Model profile based on the cross-sections presented in Figures 5 & 6. (b) is a vertical exaggeration of (a) to show the details of basin structure and its infill. The model represents uniform lithosphere extension, whereby deformation by faulting is balanced to regionally distributed stretching at depth. Although it is possible to explain the overall magnitude of subsidence at the basin margins by this mechanism, model results show that it is not possible to reproduce the large amount of subsidence observed in the central part of the basin. (c) Model profile investigating the possibility that Black Sea subsidence can be explained by loading caused by adjacent thrust belts.

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Fig. 8. (a) Cross-section adapted from Figure 6 to show syn- and post-rift stratigraphy, top basement and Moho depth. Thinning or thickening of the crust (Co0 ), relative to a regional average of 35 km (Co), has been used to quantify the magnitude of deformation across the profile in terms of a b distribution. (b) Model based on the magnitude of deformation calculated in (a) assuming uniform lithosphere deformation. The model shows overall subsidence that is comparable with that observed in the western Black Sea. However, the relative thicknesses of syn- and post-rift sequences do not match observations.

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model shows an overall basin depth that is almost comparable with that across the western Black Sea. Maximum subsidence predicted by the model is 13.5 km. The value of elastic thickness for the Black Sea region at the time of rifting and during post-rift evolution is open to speculation. Model calculations have assumed an effective elastic thickness (Te) of 5 km at rifting and 10 km for the post-rift phase based on a sensitivity test of the parameter in terms of it generating the most realistic combination of basin geometry and subsidence. These values are also compatible with a number of rift settings (e.g. Barton & Wood 1984; Kusznir et al. 1991; Van Wees & Cloetingh 1994). Although the model includes the effects of infilling the basin with sediment (density ¼ 2500 kg m23), a bathymetry, rising to a maximum depth of 3 km in the centre of the basin, has been maintained so as not to overestimate the loading and subsidence caused by sediment infill. The main difference between the model profile and observed data concerns the proportions of syn- and post-rift stratigraphy. The model shows a synrift sequence that reaches a thickness of 3.5 km. This value is not compatible with the data interpretation in Figure 8a, which suggests that the basin fill is mainly post-rift and that the synrift sequence is of the order of 1–2 km in thickness. Possible effects of subsurface loading. In this subsection we investigate the possibility as to whether the subsidence history of the western Black Sea basin can be explained by depthdependent extension, such that the deformation of the mantle lithosphere by stretching exceeds that in the crust due to faulting. The importance of depth-dependent stretching was identified by the numerical modelling of two-layer lithosphere stretching carried out by Royden & Keen (1980) and Hellinger & Sclater (1983). More recently, depth-dependent stretching has been used to explain the discrepancies between upper and lower lithosphere extension within continental margin settings (e.g. Roberts et al. 1997; Kusznir et al. 2004). The model profiles presented in Figure 9 have again used a faulting configuration based on the regional north– south section across the western Black Sea, but the extension has been increased in the lower crust and mantle lithosphere (i.e. below the assumed detachment depth at 20 km). The model presented in Figure 9a has assumed extension in the lower crust and mantle lithosphere to be given by a maximum b value of 8.5 (at x ¼ 300 km), which is compatible with the overall magnitude of crustal thinning. b is decreased following a sinusoidal distribution to no deformation over a lateral distance of 210 km either side. The pattern of subsidence exhibited in this model is

similar to that exhibited in the regional section. More specifically, the model shows a very thin synrift phase, analogous to the small thickness of middle –upper Cretaceous synrift deposits observed in the basin, and a large thickness of post-rift deposition, analogous to the thick Palaeogene and Neogene sequences observed in the basin. However, overall basin depth is too shallow compared to the observed data for the basin. The modelling approach used to generate Figure 9a assumes a fixed detachment depth of 20 km, whereby the upper crust above this depth is thinned by faulting as observed from the data, while the lower crust is thinned by pure shear according to the estimated thinning of the crust (see Fig. 8a). The assumption of a fixed detachment depth limits the magnitude by which the crust can be thinned overall and explains the relatively low basin subsidence generated in the model profile. The model presented in Figure 9b attempts to reconcile the observed magnitude of fault-controlled extension with the observed attenuation of the whole crust. The scenario modelled is one where the detachment or necking depth is allowed to progress towards the surface rather than being fixed at mid-crustal levels. The detachment zone may migrate throughout the evolution of a rift or during multiple rifts because the thickness of the brittle and ductile zones would constantly change during deformation in response both to the changing thickness of the crust and to temperature perturbations (Kusznir & Park 1987). In addition, Driscoll & Karner (1998) show an example of how an eastdipping detachment model of the northern Carnarvon Basin, Australia has been used to account for the observed discrepancy between the amount of upper crustal and lower lithosphere deformation. The model presented in Figure 9b shows how thin early fault-related synrift sediments are generated during an initial stage of rifting. Depth-dependent stretching then thins the crust (from its base) during a second rift stage. Maximum basin subsidence is about 13.5 km, which is comparable to that observed in the western Black Sea. Similarly, the model exhibits a very thin synrift stratigraphic thickness, large post-rift thickness and a realistic bathymetry. These results suggest that enhanced mantle extension may have played a role in the evolution of the western Black Sea.

The eastern Black Sea basin – application of a 3D kinematic modelling approach A weakness with 2D basin models is that the Cartesian co-ordinate that extends perpendicular to the plane of the section is assumed to be both constant and infinite in extent in terms of its physical

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Fig. 9. (a) Model based on depth-dependent stretching, whereby the lower crust and mantle lithosphere have been extended more than the upper crust. This model is comparable with the data in that it shows a relatively thin synrift sequence overlain by a thick post-rift sequence. (b) Model that reconciles the observed magnitude of fault-controlled extension with the observed attenuation of the whole crust. The scenario modelled is one where the detachment or necking depth is allowed to progress towards the surface rather than being fixed at mid-crustal levels.

properties. As a result, lateral discontinuity in structure and thermal perturbation outside the plane of section is ignored. This has significant implications on the flexural compensation of loads and on the true geometry of the plane of the modelled section. In three dimensions, flexural interactions modify the spatial pattern and extent of basement subsidence and/or uplift. A 3D lithosphere-scale basin model has been developed in order to investigate how regional interactions between a range of geological and geodynamic processes control basin subsidence and stratigraphy through time. This model has been applied to provide insights into the regional subsidence history of the eastern Black Sea. The methodology developed is similar to that in the 2D modelling whereby a range of parameters can be independently controlled and defined in order to simulate basin development through time whilst accounting for multiple rift phases, progressive sediment infill and reproduction of bathymetry. The following subsection describes this modelling

approach and the additional considerations required when modelling geological and geodynamic processes in three dimensions. † A set of crustal thickness profiles (e.g. Fig. 8a) are constructed initially. † For each model time step, the tectonically induced subsidence (thermally and mechanically derived) is calculated and flexurally compensated for by assuming an initial water infill. Loading owing to changes in crustal thickness, thermal effects, sediment infill and bathymetry are compensated for by using a 3D flexural algorithm modified after Hodgetts et al. (1998). † Perturbation and/or re-equilibration of the lithosphere temperature field is calculated for each model time step. † In order to simulate the observed proportion of sediment infill, the basin is then filled with sediment to a potential maximum bathymetry. The extra flexural subsidence generated from the initial sediment infill is then optionally filled

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with sediment or water ensuring that variations in the infill density are compensated for at all times. † If necessary, this procedure is repeated using a different potential maximum bathymetry until an approximate representation of the proportion of sediment infill is obtained. Owing to the lack of data generally available on the regional palaeobathymetry, this method provides an alternative approach to filling the basin to a ‘known’ palaeobathymetry and also provides an estimate of how the basin’s sediment –water interface has evolved through time. A detailed description of the theoretical aspects of this 3D modelling approach can be found in Meredith (2003). The cross-sections presented in Figure 5 have been used to provide ‘seed-lines’ for a 3D model that covers an area of 52 000 km2 of the Turkish and central regions of the eastern Black Sea (see Fig. 3 for location). A b profile has been calculated for each of the cross-sections based upon variation in crustal thickness. b values have been calculated at 4 km intervals along each profile and, in turn, placed into a 3D spatial model as to coincide with the relative positions of each cross-section. Figure 10 shows a 3D model representation of the eastern Black Sea assuming evolution by uniform lithosphere extension of a 35 km-thick crust. The model has been simulated over a period of 65 Ma. Effective elastic thickness is assumed to be 5 km during instantaneous rifting, evolving to 10 km during the post-rift subsidence phase. The model also reproduces and regionally compensates for temporal changes in both stratigraphic thickness and bathymetry. Post-rift sediments have been subdivided into early- and late-stage sequences in order to replicate the Paleocene –Middle Miocene sequence and the anomalously thick Upper Miocene–Quaternary sequence. Total subsidence is shallower in the 3D model than is observed. Maximum basin depth is 12 km, whilst maximum depth in the model is 8.5 km. A lack of subsidence is further depicted in the model by the elevation of the Archangelsky and Andrusov Ridges, which, according to the data, have accumulated significant thicknesses of Quaternary sediment. This reduction in basin depth in the 3D model partly results from the increase in out-of-plane flexural support provided by the MidBlack Sea High. This is illustrated in Figure 10 by the increase in basement elevation when moving westerly toward the Mid-Black Sea high. Variations in prerift lithospheric configuration, including crustal thickness, can strongly control the total magnitude of subsidence produced during

and after extension. Estimation of the prerift crustal thickness remains widely open to conjecture owing to the long succession of rifting and collisional events that have preceded eastern Black Sea basin development. If the prerift crust had been thicker than the commonly assumed thickness of 35 km, the total subsidence derived from the apparent magnitude of extension would be enhanced. According to Sengo¨r et al. (1985), the prerift crust of western Turkey was thickened by early Palaeogene compression to an estimated thickness of 50 –55 km. Kazmin et al. (2000) also propose that the prerift basement of the eastern Black Sea was analogous to that in the eastern Pontides, which is known to have undergone strong Paleocene compression and thickening (Okay & Sahintu¨rk 1997). This evidence suggests that the thickness of the prerift crust may have been in the order of 40 –45 km. Consequently, models have been produced in order to investigate rifting of a thickened prerift crust. Figure 11 shows a 3D model representation of the eastern Black Sea assuming evolution by uniform lithosphere extension of a 45 km-thick prerift crust. Raising the initial crustal thickness value to 45 km produces extension across the whole model and a maximum b value of 4.2 in the central basin. Model results indicate that rifting of a thickened crust generates sufficient subsidence to house the great thicknesses of post-rift sediments including the present-day deep bathymetry. Three-dimensional reproduction of the basin depth, and syn- and post-rift stratigraphic architectures are directly comparable with that observed. In addition, model results suggest that the anomalously thin synrift sediments and thick post-rift sequences could have resulted from underfilling of the basin due to prolonged sediment starvation followed by late stage post-rift sedimentation.

Discussion There has been much speculation within the literature as to the cause of the large magnitude of subsidence within the Black Sea. In this study we have used 2D numerical modelling of lithosphere extension and compression to test some of the subsidence mechanisms proposed for the western part of the basin. Readers should refer to Meredith & Egan (2002) for the application of a similar approach to the eastern Black Sea. Model results show clearly that it is not possible to explain subsidence in the Black Sea by uniform lithosphere extension coupled with sedimentary infill, when the magnitude of extension is constrained by the observed faulting of basement. Furthermore, it is very unlikely that the lack of extensional structures evident within the central part of the basin can be entirely

Fig. 10. Application of the 3D kinematical modelling approach to the eastern Black Sea. Extension of a 35 km-thick crust simulates the main structural elements and a synand post-rift stratigraphy that has a near-comparable ratio to those observed. The basin model, however, is shallower than observed partly due to the increased flexural support produced by the Archangelsky Ridge, which is orientated obliquely to the direction of the data sections.

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Fig. 11. Three-dimensional model of the eastern Black Sea. Extension of a thicker (45 km) prerift crust maximizes synrift subsidence, and produces a syn- and post-rift stratigraphy that is directly comparable in terms of thickness and architecture to that observed in the eastern Black Sea.

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attributed to a lack of resolution within the currently available data. Walsh et al. (1991) have shown that typically 10–40% of total fault-related extension may not be observed on regional-scale seismic profiles. However, the difference between magnitude of observed fault controlled extension and overall deformation of the crust is over 200% in the context of the Black Sea. It is now widely accepted that the relative thickness of syn- and post-rift sequences within a basin can be influenced by the nature of the extensional deformation experienced by the lithosphere. More precisely, the relatively thick post-rift sequence within the Black Sea could be explained by depthdependent extension whereby the magnitude of deformation within the mid–lower crust and mantle lithosphere is much greater than that experienced by the upper crust. Modelling of this deformational regime (Fig. 9) is compatible with the subsidence history of the Black Sea in terms of the overall magnitude of subsidence as well as generating a relatively small amount of synrift subsidence followed by a much thicker post-rift sequence. It is also compatible with the suggestion that the Black Sea is floored by oceanic crust because enhanced mantle extension raises the lithosphere asthenosphere boundary to a shallow depth resulting in adiabatic melting within the mantle lithosphere and the potential to generate several kilometres of melt thickness (McKenzie & Bickle 1988). It is beyond the capabilities of the modelling used in this study simulate the formation of oceanic crust. In addition, there is little constraint from data or the published literature on how oceanic crust may have formed in the Black Sea region or, indeed, what its relationship is with the surrounding (continental) margins. However, the possible occurrence and mechanisms of generation of oceanic crust within the Black Sea requires further investigation. The observed discrepancy between the amount of upper crustal extension by fault displacements and overall crustal thinning in a variety of sedimentary basins has been identified by a number of authors (e.g. Moretti & Pinet 1987; Ziegler 1992; Driscoll & Karner 1998; Kusznir et al. 2004). Evidence of crustal stretching is often absent or indicates only less than 10–15% stretching (Artyushkov 1992). A number of problems exist with regards to the validity of the enhanced extension mechanism. In particular, space problems arise if there is exaggerated stretching of the lower crust and mantle lithosphere (e.g. Rowley & Sahagian 1986). A number of possible mechanisms may provide a solution to this apparent complication. Royden & Keen (1980) have argued that magmatic intrusion into the crust may account for some of the inherent space problems. According to Brun & Beslier (1996), mantle exhumation at passive margins and in the core complex of

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intracontinental basins may also accommodate the space problem associated with enhanced extension of the mantle lithosphere. Watcharanantakul & Morley (2000) have also recently argued that enhanced stretching could be accomplished by subduction roll-back and slab steepening of a subducting plate and have applied this mechanism to the syn- and post-rift modelling of the Pattani Basin, Thailand. In addition, lower crustal flow has also been advocated as a mechanism for explaining the observed discrepancy between the amount of upper crustal extension and extent of crustal thinning (e.g. Bertotti et al. 2000). At geological timescales, weak ductile lower crust may behave as a ductile fluid channel able to flow from areas of high lithostatic stress to areas of lower lithostatic stress (Block & Royden 1990; Westaway 1998). However, predicting the flow behaviour of the mid and lower crust remains uncertain, and in terms of numerical modelling requires further investigation. A 3D kinematic modelling approach illustrates the importance of simulating regional-scale basin subsidence and flexure in three dimensions in order to have an appreciation of out-of-plane flexural support of loads. Application of the 3D kinematic modelling approach to the eastern Black Sea suggest that extension of a thickened (40–45 km) prerift crust is necessary in order to generate the necessary amount of subsidence. Indeed, a number of studies support the occurrence of compressional tectonics related to Tethys subduction occurring in the region before the onset of (back-arc) extension. Consequently, it is argued that extension of a continental crust 5–10 km thicker than average is a potential mechanism that may certainly account for the overall subsidence and basin depth.

Summary A kinematic model of lithosphere deformation has been developed that includes the complex interaction, in two and three dimensions, of structural, thermal, isostatic, rheological and surface processes to simulate the evolution of extensional basins and thrust belt– foreland basin couplets. Application of this kinematic modelling approach to the western Black Sea indicates that it is not possible to reproduce subsidence within the basin when the magnitude of lithosphere extension is based on the amount of fault-controlled deformation. However, modelling of the Black Sea in which the magnitude of deformation is constrained using crustal thinning/thickening generates amounts of total subsidence that are comparable with that observed. These models rely on a depth-dependent extension mechanism to reconcile the observed (small) magnitude of faulting with overall attenuation of the

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crust. In addition, a 3D modelling approach has been applied to the eastern Black Sea and shows that the magnitude of total subsidence is significantly reduced when accounting for a realistic bathymetry, infill and regional flexure, even when the overall magnitude of deformation is constrained by the observed crustal thinning. The observed subsidence has to be accounted for by additional subsidence mechanisms, such as the extension of thickened crust. The main conclusions from the application of the modelling is that for highly extended margins and basin structures the observed basin depth can only be attained when the total magnitude of crustal thinning is constrained from crustal thickness changes rather than by fault-displacement measurements. Nevertheless, reconstruction of the seismically visible faults simulates a comparable basement topography to that observed. Consequently, while it may be regarded that upper crustal faults account for the apparent magnitude of extension at the level of the basement, processes occurring deeper in the crust and mantle may thin the crust without appreciably deforming the faulted basement topography. Processes such as lower crustal flow, penetrative pure shear, lateral flow of weak ductile crustal layers between stronger layers, and a component of subseismic scale deformation may combine to produce whole crustal thinning without significant mismatches in upper and lower crustal stretching factors. Clearly, these processes require further investigation both through modelling and observation.

Appendix: InterMARGINS Extensional Deformation of the Lithosphere (IMEDL) benchmarking exercise: application of a kinematic modelling approach to the Iberian margin Introduction In this section the 2D kinematic modelling approach described in the section on ‘integrated kinematic modeling of lithosphere extension and compression’ in the main part of the paper has been applied to the Iberian margin as part of a benchmarking exercise carried out for the 2004 InterMARGINS Extensional Deformation of the Lithosphere (IMEDL) workshop. The aim of this benchmarking exercise was to establish the strengths and weaknesses of different modelling approaches (i.e. dynamic, kinematic, inverse, etc.) when applied to the same case study consisting of a combined section across the Iberia and Newfoundland margins, including their respective ocean–continent transitions. The database made available for the exercise consisted of multichannel seismic reflection data,

refraction data, published material and Ocean Drilling Program (ODP) drilling data sets. Our main aim, as part of this benchmarking exercise, was to show how a kinematic modelling approach can be applied to provide insights into the geological and geodynamic processes that control overall subsidence history and basin architecture of a passive margin setting.

Data constraints The modelling carried out has focused on the Iberia margin and is mainly based on a regional seismic reflection line, Iberia Seismic Experiment – 1 (ISE-1), which was made available for the IMEDL benchmarking exercise. The geological evolution of the Iberian margin is described in a number of publications (e.g. Pinheiro et al. 1996; Whitmarsh et al. 1996) and will not be repeated here. Seismic line ISE-1, which was acquired in 1997, is shown in Figure A1a. It starts about 11 km to the west of the Portuguese– Spanish coastline, crossing the Galicia Interior Basin, the Galicia Bank, the deep Galicia Basin and the Peridotite Ridge. The section has been interpreted to show the sea bed, top basement and crustal faults (Fig. A1b). In order to start building the first model, the magnitude of heave has been quantified for each of the faults that have been identified (Fig. A1c). The heave along each fault has been measured as the distance from the footwall cut-off to the point of contact between basement and the fault. No attempt has been made to reconstruct the footwall tops to take into account material lost through erosion so it is acknowledged that the heave values measured are likely to be an underestimate. Total extension, as given by the sum of all heave values, is 104.5 km, which approximates to about 50% extension (b ¼ 1.5).

Application of kinematic modelling One of the main observations to arise from the modelling of the Black Sea basin in the main part of the paper is that it is not possible to use the observed magnitude of faulting to constrain deformation throughout the lithosphere. This is partly because the data acquisition and interpretation process does not reveal all of the faults that exist at a variety of scales within the crust. However, there is a dilemma in that many basins within the geological record show a mismatch between the deformation due to faulting and overall thinning of the crust that is too great to explain by a lack of identification of faults within the available data. This mismatch is apparent across the Iberian margin where the b value determined from the fault heave values is about 1.5, which leads to postextensional crustal thickness of about 23 km. In contrast, seismic refraction data (Whitmarsh et al. 1996) indicate that the crust is about 2– 4 km thick at the western edge of the section, which is indicative of b values of 8 or 9. One possible scenario that can reconcile the observed low magnitude of fault deformation with high attenuation

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Fig. A1. (a) Iberia Seismic Experiment – 1 (ISE-1) regional seismic reflection line that was made available for the 2004 IMEDL benchmarking exercise. The x-scale is defined in CDP locations, which are 12.5 m apart, giving the section a total length of about 325 km. (b) The section has been interpreted to show the sea bed, top basement and crustal faults. (c) The magnitude of heave has been measured for each of the faults that have been identified.

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Fig. A2. Model based on reflection line ISE-1 assuming a migrating detachment depth (4 –15 km). It shows a realistic basement geometry and overall pattern of relative uplift and subsidence across the margin.

of the crust determined from geophysical data is a migrating detachment or necking depth (Braun & Beaumont 1989). This depth determines the relative importance of upper crustal thinning due to faulting relative to lower crustal thinning due to a more regionally distributed process. It is possible therefore that during the evolution of the Iberian margin there has been a change from a deep detachment depth and therefore dominance of crustal thinning due to faulting to a shallow crustal necking depth where there is a dominance of lower crustal thinning. The driving mechanism for this migration of the detachment depth can be explained by an increase in the geothermal gradient as the rifted margin evolves (e.g. Spadini et al. 1996). Figure A2 shows a model profile based on a migrating detachment depth. The deformation due to faulting is based on that interpreted from the ISE-1 section (Fig. A1b, c) and assuming a fault detachment depth of 15 km, which is compatible with evidence from deep seismic reflection data (e.g. Hall et al. 1984), seismological investigations (e.g. Jackson & McKenzie 1983) and rheological modelling of the lithosphere (e.g. Kusznir &

Park 1987). The crust is assumed to have a predeformational thickness of 35 km (Pinheiro et al. 1996; Manatschal 2004). A second phase of extension is then modelled with the necking depth positioned at a depth of 4 km, and the lower crust and mantle lithosphere is gradually thinned from no deformation in the east to a maximum B value of 8 at the western edge of the section in order to reproduce the observed attenuation of the crust. The model has been allowed to run for a simulated time of 200 million years to represent the first phase of rifting in the lower Jurassic. For simplicity, the basin has been assumed to fill to sea level with sediment during rifting, whereas the post-rift phase has been water filled to avoid overloading and overdeepening the basement. The elastic thickness of the lithosphere is taken to be 5 km at rifting, which is compatible with a number of active rift settings (e.g. Barton & Wood 1984; Kusznir et al. 1991; Van Wees & Cloetingh 1996) gradually rising to 20 km at the end of the model evolution in response to cooling of the lithosphere. The model shows distinct uplift of the footwall blocks covered by a broad blanket of thermal subsidence.

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Fig. A3. (a) Composite section across the Iberian margin showing structure and stratigraphy with respect to depth (adapted from Pinheiro et al. 1996). The location map shows the near proximity of this section to seismic line ISE-1. (b) Modified model (cf. Fig. A2) whereby any subsided part of the basin following rifting has been half-filled with sediment and allowed to subside flexurally. Any remaining subsidence is filled with water. This model shows a close comparison to the structural and stratigraphic cross-section presented in Figure A3a.

The resultant model profile exhibits both a realistic basement geometry and overall pattern of subsidence across the margin. A published section across the Iberian margin by Pinheiro et al. (1996) has been used to make a detailed comparison with the model results (Fig. A3a). This section shows the variation of structural and stratigraphic components with respect to depth and has some

overlap with seismic line ISE-1. A comparison between the western parts of this section and the model profile in Figure A2 shows that the modelled basement depth is reasonably accurate with the tops of the fault blocks at depths of 4– 5 km and basement within the graben structures at depths of 6– 8 km. However, there are mismatches between the model results and interpretation. In particular,

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the modelled water depth is too low and the sediment thickness is too great. This incompatibility can be reduced by decreasing the amount of infill of the basin. A new model has been generated in Figure A3b using an identical set of parameters to that shown in Figure A2, except that any basin subsidence generated immediately following rifting has been half-filled with sediment, while any remaining component flexural subsidence has been filled with water. Although the modified model shows a close match to the western part of the section, the Galicia Bank region is too deep in the model. This mismatch can be explained by an overestimation of the magnitude of fault-controlled extension in the region. Alternatively, a closer match could also be obtained by modifying the flexural strength of the lithosphere. The eastern part of the model shows a reasonably accurate geometry and depth profile across the fault blocks, apart from the exaggerated footwall uplift structure at the edge of the profile. This feature is not present in the geological section possibly owing to either erosion or by additional extension off-section to the east.

Summary A kinematic model of lithosphere extension has been included within a benchmarking exercise to compare the strengths and weakness of different modelling methods in the context of the Iberia margin. Twenty four crustal faults have been interpreted in seismic section ISE-1 across the Iberia margin, exhibiting a total horizontal extension of about 105 km (b ¼ 1.5). This magnitude of fault-controlled deformation has to be reconciled with the overall attenuation of the crust to generate a realistic magnitude of subsidence, which can be explained by a gradually shallowing necking depth. We would like to thank BP and TPAO for providing data for this study. S. Egan is also grateful for the opportunity to have taken part in the 2004 IMEDL workshop. The authors would also like to thank N. Christie-Blick, A. Goodliffe and an anonymous reviewer for their constructive comments during the review process of this manuscript.

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Early kinematic history of the Goban Spur rifted margin derived from a new model of continental breakup and sea-floor spreading initiation D. HEALY1 & N. J. KUSZNIR2 1

Department of Earth Sciences, University of Durham, Durham DH1 3LE, UK (e-mail: [email protected])

2

Department of Earth & Ocean Sciences, University of Liverpool, Liverpool L69 3GP, UK Abstract: Deformation of continental lithosphere leading to breakup and the initiation of seafloor spreading can be described in terms of an upwelling divergent flow field within the continental lithosphere and asthenosphere. A new model (SfMargin) of rifted continental margin formation using this upwelling divergent flow field has been developed that successfully predicts lithosphere depth-dependent stretching and mantle exhumation at rifted margins. The kinematics of this flow field has important consequences for the bathymetric (subsidence) and thermal history of continental margin lithosphere. We apply the new model to a profile through the Goban Spur on the eastern Atlantic continental margin. Forward modelling yields parameter values describing the kinematics of the flow field leading to continental breakup and sea-floor spreading initiation that are consistent with the known amagmatic history of this margin. We employ a grid search method to systematically explore model parameter space and provide an assessment of the sensitivity of the model to the kinematic flow field parameter values. The preferred forward model parameters coincide with low values of misfit with respect to observed present-day bathymetry and free air gravity anomalies. Our results reveal the early kinematic history of the Goban Spur margin as characterized by a ratio of upwelling rate to half-spreading rate of unity, consistent with that expected for the initiation of sea-floor spreading at a non-volcanic rifted margin and only a modest amount of prebreakup lithospheric stretching (b ¼ 1.5). These results conform with recent observations of depth-dependent stretching and serpentinized mantle exhumation at this amagmatic rifted continental margin.

The early kinematic evolution of rifted continental margins plays a major role in the burial and thermal histories of the sediments deposited thereon. Recent mantle convection modelling by Nielsen & Hopper (2002) has shown the importance of the mantle upwelling velocity in determining the volume of magmatic material produced during continental breakup and sea-floor spreading initiation. Existing models of rifted continental margin formation (e.g. Le Pichon & Sibuet 1981) have been based on an extrapolation of the McKenzie (1978) intracontinental rift-basin model, with stretching factors (b) tending to infinity as the continental lithosphere necks and finally ruptures. However, more recent observations of depth-dependent stretching at rifted continental margins (Driscoll & Karner 1998; Davis & Kusznir 2004; Kusznir et al. 2004) and serpentinized mantle exhumation (Pickup et al. 1996; Manatschal et al. 2001; Whitmarsh et al. 2001) are not explained by existing models of rifted continental margin formation. A new model of rifted continental margin formation (Kusznir & Karner 2007) has been developed based on the assertion that thinning of continental lithosphere leading to breakup and sea-floor intiation

occurs in response to an upwelling divergent flow field within continental lithosphere and asthenosphere. An upwelling divergent flow field provides the kinematic basis for ocean ridge models (e.g. Spiegelman & McKenzie 1987; Su & Buck 1993), but can also be applied to the thinning of continental lithosphere leading to breakup and the initiation of sea-floor spreading. In the new model of rifted continental margin formation proposed by Kusznir & Karner (2007) the upwelling divergent flow field within continental lithosphere and asthenosphere, which leads to thinning, breakup and sea-floor spreading initiation, is defined kinematically (Fig. 1a). The model also includes a contribution to lithosphere thinning from prebreakup depth uniform (pure shear) lithosphere stretching. The modelling described in this paper seeks to recover values of two key kinematic parameters that exert a major influence on the early history of rifted continental margins: the ratio of the upwelling velocity (Vz) to the divergence velocity, or half-spreading rate (Vx), and the prebreakup pure shear lithosphere stretching factor (b). Using these optimum parameter values the new rifted margin model can then be used to predict crustal

From: KARNER , G. D., MANATSCHAL , G. & PINHEIRO , L. M. (eds) 2007. Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 199– 215. DOI: 10.1144/SP282.10 0305-8719/07/$15.00 # The Geological Society of London 2007.

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(a)

x

vx

vx

z=0

vz z

x=0 (b)

(c)

Steady-state breakup flow Vx

Vx Vz

Upwardly propagating breakup flow Vx = 0

Vx = 0

Vx = 0

Continental crust Continental lithosphere

Vx > 0 Before breakup Vx

Sea Floor Spreading

Vz

Before breakup Vx

Vz

Vx

Asthenosphere

Sea Floor Spreading

Vz

Vx Crust

Lithosphere

After breakup

After breakup

Asthenosphere

Fig. 1. Parameters used to describe different effects of model flow fields in the SfMargin model on predicted lithosphere structure. (a) The 2D model, half-space symmetrical about x ¼ 0, showing stream lines for upwelling divergent flow. Vx is the horizontal velocity (or half-spreading rate) and Vz is the vertical velocity in this corner flow solution (Batchelor 1967). (b) For steady-state flow during breakup, the upwelling divergent flow field is immediately active over the whole thickness of the lithosphere from t ¼ 0. (c) For upwardly propagating flow during breakup, the divergent flow field progressively moves up through the lithosphere section. Before the zone of divergent flow propagates to the surface Vx, above the zone of divergent flow (dotted line), is 0 and greater than 0 within the divergent zone streamlines for a pure shear flow field typical of the McKenzie rifting model. (b) Lithosphere cross-section at rupture produced by pure shear flow. Note the lack of mantle exhumation or depth-dependent stretching predicted by this model. (c) Streamlines for the upwelling divergent flow field implemented in SfMargin. The prescribed velocities, Vx and Vz, are constant along the top and left sides, respectively, and control the corner flow solution. (d) Lithosphere structure at breakup predicted by upwelling divergent flow where Vx ¼ Vz. Note the exhumation of lithospheric mantle at the ocean– continent transition. (e) Streamlines for an upwelling divergent flow field where Vz  Vx. (f) Predicted lithosphere structure when Vz  Vx. Note the depth-dependent stretching at the continental margin.

and lithospheric thinning, subsidence and heat flow histories for the margin transect, and is of direct importance to hydrocarbon exploration in these frontier areas. The method employs observations derived from publicly available bathymetry and gravity anomaly datasets (Sandwell & Smith 1997; IOC, IHO & BODC 2003). The technique does not depend on detailed geological data such as palaeobathymetry indicators derived from hydrocarbon wells (cf. White 1993; White & Bellingham 2002), and can therefore be applied to

any continental margin in advance of a potentially expensive drilling program.

A new model for the formation of rifted continental margins The formulation of the new model of rifted margin formation (SfMargin) is described in greater detail in the companion paper by Kusznir & Karner (2007). The upwelling divergent flow field is described using analytical stream functions defined by

KINEMATIC HISTORY OF THE GOBAN SPUR

an isoviscous corner flow model (Batchelor 1967). SfMargin uses a coupled fluid mechanics and thermal advection –diffusion solution to describe the upwelling divergent flow field of continental breakup and sea-floor spreading. This flow field is applied to a two-dimensional (2D) template of continental crust, continental lithosphere and asthenosphere, with initially uniform thickness and temperature distribution, and this flow field is symmetrical about the spreading (z-) axis (Fig. 1a). Fluid flow boundary conditions are imposed as kinematic constraints along the surface (top) and spreading axis (left) edges of the model space. The fluid flow field is used to advect lithosphere material and lithosphere and asthenosphere temperature. Note that the upwelling divergent flow field is applied after an optional amount of instantaneous prebreakup pure shear stretching. Thermal solutions are computed using a finite-difference method. Local (Airy) isostasy is assumed in order to calculate bathymetry from crustal thickness and lithosphere temperature. Crustal and lithosphere parameter values used in this study are listed in Table 1. Gravity anomalies are computed from the bathymetry, crustal structure and lithosphere temperature. The upwelling divergent flow field can be imposed on the modelled section in two distinct ways. Steady-state breakup flow imposes the divergent flow field over the full vertical thickness of the model from the beginning (Fig. 1b). As a consequence, all parts of the model are subjected to non-zero values of Vx from t ¼ 0. In contrast, upwardly propagating breakup flow starts with the divergent flow field in the asthenosphere only and then progressively migrates up through the lithosphere section (Fig. 1c). At any instant above the zone of divergent flow, the half-spreading rate Vx is 0. When the divergent flow field reaches the surface, the flow field resembles the steady-state breakup flow field. Initiating the divergent flow field in these different ways has important

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consequences for the evolution of the rifted margin lithosphere. The key kinematic parameters controlling the fluid flow field are the vertical upwelling velocity, Vz, the horizontal half-spreading rate, Vx, and the prebreakup lithospheric stretching factor, b. Modelling by Nielsen & Hopper (2002) shows how early flow velocities during continental breakup and seafloor spreading initiation can differ from the later steady-state kinematics of sea-floor spreading. The SfMargin model allows for a short transient period (typically 2 –20 Ma) of different velocities at the early stage of sea-floor spreading. This paper focuses on recovering optimum values for the ratio of these early velocities and the prebreakup stretching factor, using observations of present-day bathymetry and gravity.

Application of SfMargin to the Goban Spur profile The Goban Spur region is considered to be a typical non-volcanic (amagmatic) and relatively sedimentstarved rifted continental margin (Horsefield et al. 1993; Watts & Fairhead 1997; Bullock 2004; Bullock & Minshull 2005). Figure 2 shows the location of the Goban Spur margin and the bathymetry and gravity observations along the profile modelled in this study. Depth-dependent lithosphere stretching has been observed for the Goban Spur rifted margin (Davis & Kusznir 2004) and mantle exhumation has been inferred (Bullock 2004; Bullock & Minshull 2005). Goban Spur shows (Fig. 3) a region of the order of 75 – 100 km width adjacent to the ocean –continent transition in which stretching of the upper continental crustal stretching is substantially less than that of the whole crust and lithosphere (Davis & Kusznir 2004). Extension and thinning of the upper crust has been measured using fault heaves from seismic reflection data; of the whole crust using

Table 1. Model parameters used in this study Symbol

b Vx Vz

Description Initial crustal thickness Initial lithosphere thickness Initial asthenosphere temperature Density of crust Density of mantle Density of sea water Goban Spur rifted margin age Prebreakup stretching factor Half-spreading rate Upwelling rate

Value and units 30 km 125 km 1300 8C 2800 kg m23 3300 kg m23 1000 kg m23 120 Ma cm year21 cm year21

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Fig. 2. Location maps for the modelled Goban Spur profile. (a) Bathymetry data from GEBCO (IOC, IHO & BODC 2003) and (b) gravity data from Sandwell & Smith (1997) are shown from the UK continental shelf out to the mid-Atlantic ridge. Our modelled profile runs parallel to the fracture zones, i.e. parallel to the inferred direction of ocean opening.

crustal basement thinning derived from crustal structure utilizing wide-angle seismology and gravity studies; and of the whole lithosphere from post-breakup subsidence using flexural backstripping of postbreakup stratigraphic data. A detailed description of stretching and thinning measurements and their analysis to determine depth-dependent lithosphere stretching is given in Davies & Kusznir (2004). We apply the new model of continental breakup and sea-floor spreading initiation to recover the kinematic flow field parameter values for this margin using the SfMargin to model observed bathymetry and gravity anomaly profiles. The bathymetry values are taken from GEBCO (IOC, IHO & BODC 2003) and the free air gravity data is from Sandwell & Smith (1997). Our selected profile across this margin is parallel to the fracture zones in the oceanic crust (Fig. 2) and we infer that this line is parallel to the oceanic opening vector. Observations are extracted from the published datasets along the profile line at a constant, arbitrary

horizontal spacing (c. 20 km). Both the bathymetry and gravity observations are affected by noise, i.e. features not modelled by SfMargin, particularly over the oceanic part of the profile. Figure 3 shows the bathymetry and gravity observations along the profile modelled in this study. Note the spikes in bathymetry and gravity anomaly between 250 and 1000 km horizontal offset from the ridge, caused by unmodelled roughness on the ocean floor. We discard observations close to (,250 km) the mid-Atlantic ridge as too noisy. The key features of a rifted continental margin are clearly present, including a distinct continental slope and sharp shelf break in the bathymetry, and a coincident dipole in the free air gravity anomaly.

Model sensitivities to kinematic parameter values The results of any forward modelling are inherently non-unique, so we have explored the SfMargin

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Fig. 3. Depth-dependent lithosphere stretching has been observed on the Goban Spur rifted continental margin. Stretching and thinning factors for the upper crust are significantly less than those of the whole crust and lithosphere within 75– 150 km of the ocean– continent transition (adapted from Davis & Kusznir 2005).

model sensitivities to the kinematic parameter values through a suite of constrained models.

much sharper OCT. These geometric differences in the subsurface architecture are reflected in the distinct patterns of thinning factors shown in Figure 4g, h.

Initiation of upwelling divergent flow In the new model, the upwelling divergent flow field is applied after an optional amount of instantaneous pure shear lithosphere stretching. The flow field can be applied instantaneously to the whole thickness of the lithosphere (steady-state breakup flow) or allowed to migrate from the base of the lithosphere to the surface (upwardly propagating breakup flow). Figure 4 shows the effects of these two distinct flow regimes on the model predictions compared to the observations from Goban Spur. In each of these models, b ¼ 1.5 and Vz/Vx ¼ 1. The predicted bathymetry and gravity profiles (Fig. 4a–d) are similar for both breakup flow regimes, and match the observations fairly closely. The most important difference lies in the predicted lithosphere crosssections (Fig. 4e–f). For steady-state breakup flow, continental lithospheric mantle is exhumed at the ocean–continent transition (OCT), whereas upwardly propagating breakup flow generates a

Varying the prebreakup stretching factor, b Holding the value of the velocity ratio Vz/Vx constant at 1.0, we varied the model value of b (prebreakup stretching factor) to explore its effect on the model bathymetry and gravity predictions. For steady-state breakup flow, the effect of varying b can be seen in Figure 5. Increasing the value of b from a physical minimum of 1.0 for no stretching to 5.0, when the lithosphere has been thinned to one-fifth of its original thickness, leads to a marked steepening of the continental slope (Fig. 5i) and a broadening and deepening of the gravity dipole (Fig. 5j). Overall, a similar progression is observed in the case of upwardly propagating breakup flow (Fig. 6). In comparison to the equivalent models using steady-state breakup flow, the net effect of upwards propagation of the flow field, for a given b, tends to flatten the continental slope (compare Fig. 6g to 5g) and reduce

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Fig. 4. Model results comparing a steady-state breakup flow field with an upwardly propagating flow field. In all cases the margin age is 120 Ma, prebreakup stretching b ¼ 1.5 and the Vz/Vx ratio for the early sea-floor spreading phase was set at 1.0. (a) Model bathymetry (red) compared to observations (blue) for the steady-state flow field. (b) Bathymetry for the upwardly propagating flow field. (c) Gravity anomaly (free air over sea, Bouguer over land) for the steady-state case. (d) Gravity anomaly for the upwardly propagating case. (e) Predicted lithosphere cross-section for the steady-state case. Base lithosphere (red), Moho (blue) and base upper crust (green) are shown for the continent. Note the exhumed lithospheric mantle at the ocean– continent transition (OCT). (f) Cross-section for the upwardly propagating flow field case. Note the lack of mantle exhumation. (g) Predicted thinning factors (1 2 1/b) for the upper crust (green), whole crust (blue) and whole lithosphere (red) for the steady-state breakup flow case. (h) Thinning factors for the upwardly propagating flow field case.

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Fig. 5. Model sensitivity to changing values in prebreakup pure shear b on the predicted bathymetric profiles and gravity anomalies, for the case of steady-state breakup flow. Model results shown in red, observations in blue. Note that b ¼ 1 (a, b) implies zero prebreakup pure shear and b ¼ 1 (i, j) means that pure shear stretching ruptured the lithosphere before the upwelling divergent flow field started. With the Vz/Vx ratio held constant at 1.0, higher values of b lead to a steeper, more angular continental slope in the bathymetry signal and a broader, higher amplitude dipole in the gravity anomaly at the OCT.

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Fig. 6. Model sensitivity to changing values in b for the case of upwardly propagating breakup flow. Compared to the corresponding results for steady-state breakup flow (Fig. 5), the predicted profiles have a greater wavelength and a lower amplitude at the OCT.

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the amplitude of the gravity dipole at the OCT (compare Fig. 6h to 5h).

Varying the velocity ratio, Vz/Vx We now consider the effect of varying the velocity ratio of the 10 Ma early sea-floor spreading phase, with a constant prebreakup stretching factor of 1.5.

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For steady-state breakup flow, the model sensitivity to variations in Vz/Vx from 0.5 to 5.0 is shown in Figure 7. The lower limit of 0.5 is close to the analytical lower limit of the stream-function cornerflow solution of 2/p for passive sea-floor spreading (Phipps Morgan 1987). Progressively higher values of Vz/Vx generate shallowing and eventual uplift (Fig. 7g) of the continental margin, and a

Fig. 7. Model sensitivity to changing values in the ratio Vz/Vx for the early sea-floor spreading phase (duration: 10 Ma) on the predicted bathymetric profiles and gravity anomalies, for the case of steady-state breakup flow. Model results shown in red, observations in blue. Bathymetric profiles in the left column, gravity anomalies in the right column. With a constant prebreakup b ¼ 1.5, higher values of Vz/Vx predict shallowing (compare 7a with e) and eventual emergence (see g) of the continental crust at the OCT, and a degradation of the gravity dipole.

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corresponding flattening of the gravity dipole at the OCT. Model sensitivity to changing the velocity ratio in the case of upwardly propagating breakup flow is shown in Figure 8. In comparison to steady-state breakup flow, higher values of the

velocity ratio (Fig. 8g, h) predict significant uplift of the margin and a very distinct and complex gravity anomaly at the OCT, reflecting a continental margin with a laterally heterogeneous density distribution.

Fig. 8. Model sensitivity to changing values in the ratio Vz/Vx for the case of upwardly propagating breakup flow. At low values of the ratio (a –d) there is little effect on the profiles compared to the corresponding steady-state breakup flow models. However, at higher values (e– f), the predicted uplift of the margin increases and the gravity anomalies become very distinct.

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Exploring parameter trade-offs with a multidimensional grid search Least-squares method in a grid search algorithm The SfMargin model encapsulates a complex geodynamical process and displays a non-linear relationship between the various parameters and the predicted data. To further assess the validity and robustness of our forward model results, and to explore potential trade-offs among the model parameter values, we ran a grid search over discrete ranges of parameter values. We vary two-parameter values in the grid search, the ratio of Vz to Vx and b, with all other parameters held at constant values (see Table 1). The velocity ratio varied between 0.5 and 5.0, covering a range of viable upwelling and half-spreading rates. The prebreakup stretching factor b parameter is constrained to the range 1.0 – 5.0, with 1.0 as a physical minimum (no stretching) and 5.0 exceeding the likely maximum for the lithosphere beneath Goban Spur. The SfMargin model is executed for each combination of parameter values in this 2-space, and the results are stored as a suite of trial models. To compare the results of a trial model with synthetic observations we compute the mean square misfit, m: m¼

N 1X ðobsi  triali Þ2 N i¼1

ð1Þ

where N is the number of data points, obsi is the i-th observation and triali is the i-th prediction. For a trial model represented by a single combination of parameter values, we calculate the misfit between model predictions and a set of observations. Note that a grid search algorithm makes no a priori assumptions about the distribution of misfit minima and we apply no smoothing or filtering to the data or the misfit function.

Investigating kinematic parameter trade-offs with a synthetic model To assess the interplay of the kinematic flow field parameters we used the least-squares grid search method with synthetic observations generated by a previous run of the forward model. Perfect, zero noise data should result in a zero misfit for the trial model at the appropriate combination of parameters. Using a previous forward model with parameter values Vz/Vx ¼ 2 and b ¼ 3, we generated synthetic bathymetry and free air gravity observations. We calculate separate misfits for both bathymetry and gravity datasets for each trial

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model. Figure 9 shows contour plots of the mean square misfit calculated for bathymetry (Fig. 9a) and gravity (Fig. 9b) datasets using a steady-state breakup flow field. For these synthetic observations, the grid search method successfully located a single misfit minimum of zero within the parameter constraints at precisely the parameter combination used by the original forward model. Furthermore, the location of this minimum was the same for both bathymetry and gravity datasets. The same process was repeated using an upwardly propagating breakup flow field, and the contoured misfit surfaces for bathymetry and gravity data are shown in Figure 9c, d, respectively. A single misfit minimum of zero was located at the correct forward parameter combination for each dataset. These misfit surfaces generated using synthetic observations allow us to explore the trade-off between the kinematic parameter values of the Vz/Vx ratio and b. For both breakup flow fields, the low misfit region (pale colours in Fig. 9a–d) is an elongate narrow valley with its long axis lying along b ¼ 3 (the forward model parameter value used to generate the synthetic observations). The misfit is therefore more sensitive to variations in b compared to Vz/Vx.

Grid search method applied to observations from Goban Spur The least-squares grid search method has also been applied to observed bathymetry and gravity data from the Goban Spur profile. For these uncorrelated and noisy data, we use a weighted mean-square misfit defined as: m¼

N 1X ðobsi  triali Þ2 s2 N i¼1

ð2Þ

where s is the standard deviation of the measurements. For the bathymetry data we took s as +250 m (IOC, IHO & BODC 2003), and for the gravity data we took s as +5 mGal (Sandwell & Smith 1997). Weighting the least-squares misfits for bathymetry and gravity observations allows us to form the sum and define a (dimensionless) composite misfit. Figure 10 shows the contoured misfit surface computed from the bathymetry (Fig. 10a) and gravity (Fig. 10b) observations measured along the Goban Spur profile. The composite misfit (weighted bathymetry misfit þ weighted gravity misfit) is also shown (Fig. 10c). The misfit surface topography is broader and flatter in all of these plots when compared to their synthetic equivalents (compare Fig. 9a, b with Fig. 10a, b). The noise on the bathymetry and gravity observations (see

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Fig. 9. Plots of contoured mean square misfit from the grid search. The misfit was calculated with respect to synthetic observations generated by a previous run of the forward model with kinematic parameter values of Vz/Vx ¼ 2 and b ¼ 3. Within the constrained parameter ranges a single misfit minimum is clearly present for both bathymetry and gravity datasets. A grid search using a steady-state breakup flow field generated the misfits shown in (a) and (b), with the corresponding plots for the upwardly propagation flow field shown in (c) and (d). The misfit minimum occurs at the same parameter combination and is precisely 0 in all cases. Trade-offs between the parameters are discussed in the text.

profiles in Fig. 3) degrades the misfit throughout the model parameter space, although the misfit minima are slightly sharper for the case of upwardly propagating breakup flow (compare Fig. 10a–c with Fig. 10d–f). In some plots there is no single misfit minimum, although there are distinct regions of low and high misfit. Using the measured data, the preferred parameter combination for Goban Spur closely coincides with the regions of lowest misfit (paler grey shades in Fig. 10). The plots shown in Figures 9 & 10 show that it may be possible to invert measured bathymetry and gravity observations to obtain the initial kinematic flow field parameters for a given rifted continental margin. The presence of misfit minima within the model parameter space, even for noisy observations, should enable us to apply formal minimization methods (e.g. Powell’s method, downhill simplex,

etc.) to locate the optimum model parameter combination. Work is under way to develop and apply this technique to other datasets, for example measured upper crustal stretching values, and to other rifted continental margins, for example Iberia, Grand Banks and Flemish Cap.

Reference frame issues In calculating the misfit for a given trial model, we transform the predicted data values into the reference frame of the observations by applying a horizontal (x-) shift and a vertical (z-) shift. The x-shift is calculated from correlation (e.g. Press et al. 1992), by selecting the lag value (x-offset) with the highest correlation between the trial model predictions and the observations.

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Fig. 10. Plots of contoured mean square misfit from a grid searches using observed bathymetry and gravity data from Goban Spur. For the case of steady-state breakup flow shown in (a) – (c) the misfit surface has relatively low relief, although lower values of misfit are concentrated towards the lower left corner of each plot, corresponding to low values of Vz/Vx and b. For upwardly propagating breakup flow shown in (d) – (f) pronouned relief exists on the misfit surface, implying higher sensitivity to the parameter values of this flow field. High values of Vz/Vx are ruled out by the high misfits to the top of each plot, and for realistic values of b for the Goban Spur (,2) a low value of Vz/Vx is more likely.

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In general, we can never know the exact horizontal velocity (half-spreading rate) history of a margin and this correlation technique allows us to bring the model predictions and observations together in a systematic fashion.

Discussion and summary A newly developed forward model, SfMargin, explains recent observations at rifted continental margins and serves as a predictive geodynamic tool for the hydrocarbon industry operating in frontier areas. Using publicly available bathymetry and gravity data from the Goban Spur margin, forward modelling yields parameter values for the early kinematic history of this margin consistent with observations of depth-dependent lithosphere stretching. A sensitivity analysis among the kinematic parameters supports the results of the forward modelling. A least-squares grid search method employing synthetic observations allows us to assess the trade-offs between the parameter values. The same grid search method applied to observations of bathymetry and gravity measured along the Goban Spur profile

yields a misfit minimum almost coincident with that derived from forward modelling. The observation of depth-dependent stretching at Goban Spur and other non-volcanic and volcanic margins (Davies & Kusznir 2004; Kusznir & Karner 2007) rules out a dominantly depth-uniform (pure shear) mechanism for thinning continental lithosphere leading to continental breakup and sea-floor spreading initiation. Instead, the dominant process reponsible for continental lithosphere thinning leading to seafloor spreading initiation is an upwelling divergent flow field within continental lithosphere and athenosphere. Models using this mode of continental lithosphere thinning are able to satisfactorily predict observed bathymetry and gravity anomalies. While a contribution to continental lithosphere thinning from a depth-uniform (pure shear) mechanism may exist, it is not believed to be the dominant lithosphere deformation process leading to breakup. A region of exhumed serpentinized mantle of the order of 70 km wide, separating thinned continental crust and oceanic crust, has been inferred for at Goban Spur margin from seismic refraction and reflection studies by Bullock & Minshull (2005), and is summarized in Figure 11. The exhumation

Fig. 11. Seismic refraction and reflection studies on Goban Spur suggest the existance of a region of exhumed serpentinized mantle of the order of 70 km wide separating thinned continental crust and oceanic crust (from Bullock & Minshull 2005).

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of continental lithospheric mantle is predicted by the steady-state pattern of breakup flow shown in Figure 4e; in contrast, the upwardly propagating flow field pattern model shown in Figure 4f predicts a much sharper ocean –continent transition with no exhumation of continental lithosphere mantle. Intuitively, the upwardly propagating flow field pattern seems more physical and likely to occur; however,

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Fig. 12. Development in time of continental lithosphere thinning and bathymetry predicted by the upwelling divergent flow model of continental breakup and sea-floor spreading initiation compared for steady-state and upwardly propagating upwelling divergent flow patterns. The model with steady-state flow field (a)– (c) ruptures continental crust before continental lithospheric mantle, and leads to the exhumation of a broad region of continental lithosphere mantle. The model with an upwardly propagating flow field (d)– (f) ruptures continental lithospheric mantle before the continent crust, and results in no exhumation of continental lithospheric mantle.

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mantle exhumation during ultra-slow sea-floor spreading at the Gakkel and South West Indian ridges (Dick et al. 2003; Michael et al. 2003). Another important difference between the steady-state and upwardly propagating flow field pattern models (Fig. 4) of continental lithosphere thinning is in the time sequence of the rupture of continental crust and lithospheric mantle. The steady-state flow field predicts that the continental crust ruptures before the continental lithosphere mantle, while the upwardly propagating flow field predicts that the continental lithospheric mantle ruptures before the continental crust. These are illustrated in Figure 12. At present it is not known which model best describes the natural process: new observations are required to answer this question. Further work will employ measured upper crustal stretching factors as observations, and model the inferred conjugate margin in the western Atlantic around Flemish Cap. SfMargin has been developed as part of the NERC Ocean Margins iSIMM (integrated Seismic Imaging and Modelling of Margins) project. The iSIMM team comprises researchers from the Universities of Liverpool and Cambridge, Badley Geoscience and Schlumberger, and is funded by a consortium of NERC, DTI, Agip, BP, Amerada Hess, Anadarko, ConocoPhillips, Shell, Statoil, and WesternGeco. iSIMM investigators are R.S. White, N.J. Kusznir, P.A.F. Christie, A.M. Roberts, A. Chappell, J. Eccles, R. Fletcher, D. Healy, N. Hurst, Z. Lunnon, C.J. Parkin, A.W. Roberts, L.K. Smith, R. Spitzer and V. Tymms. The GMT software of Wessel & Smith (1998) was used to produce some of the figures. The authors gratefully acknowledge discussions with iSIMM team colleagues, constructive reviews by G. Houseman, J. Nunn and an anonymous reviewer and the editorial patience of G. Karner, all of which significantly improved this manuscript.

References B ATCHELOR , G. K. 1967. An Introduction to Fluid Dynamics. Cambridge University Press, Cambridge. B ULLOCK , A. D. 2004. From continental thinning to seafloor spreading: a geophysical study of rifted margins southwest of the UK. PhD thesis, University of Southampton. B ULLOCK , A. D. & M INSHULL , T. A. 2005. From continental extension to seafloor spreading: crustal structure of the Goban Spur rifted margin, southwest of the UK. Geophysical Journal International, 163, 527– 546, doi:10.1111/ j.1365-246X.2005.02726.x. D AVIS , M. & K USZNIR , N. J. 2004. Depth-dependent lithospheric stretching at rifted continental margins. In: K ARNER , G. D., T AYLOR , B., D RISCOLL , N. W. & K OHLSTEDT , D. L. (eds) Rheology and Deformation of the Lithosphere at Continental Margins. Proceedings of NSF Rifted Margins Theoretical Institute, Volume 1. Columbia University Press, New York, 92–136.

D ICK , H. J., L IN , J. & S CHOUTEN , H. 2003. An ultraslow-spreading class of ocean ridge. Nature, 426, 405– 412. D RISCOLL , N. W. & K ARNER , G. D. 1998. Lower crustal extension across the Northern Carnarvon basin, Australia: Evidence for an eastward dipping detachment. Journal of Geophysical Research, 103, 4975– 4991. H ORSEFIELD , S. J., W HITMARSH , R. B., W HITE , R. S. & S IBUET , J.-C. 1993. Crustal structure of the Goban Spur rifted continental margin, NE Atlantic. Geophysical Journal International, 119, 1– 19. IOC, IHO & BODC. 2003. Centenary Edition of the GEBCO Digital Atlas. Published on CD-ROM on behalf of the Intergovernmental Oceanographic Commission and the International Hydrographic Organization as part of the General Bathymetric Chart of the Oceans, British Oceanographic Data Centre, Liverpool. K USZNIR , N. J. & K ARNER , G. D. 2006. Continental lithospheric thinning and breakup in response to upwelling divergent mantle flow: application to the Woodlark, Newfoundland and Iberia margins. In: K ARNER , G. D., M ANATSCHAL , G. & P INHEIRO , L. M. (eds) Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 389– 419. K USZNIR , N. J., H UNSDALE , R. & R OBERTS , A. M. 2004. Timing of depth-dependent lithosphere stretching on the S. Lofoten rifted margin offshore mid-Norway: pre-breakup or post-breakup? Basin Research, 16, 279–296. L E P ICHON , X. & S IBUET , J. C. 1981. Passive margins: a model of formation. Journal of Geophysical Research, 86, 3708– 3720. M ANATSCHAL , G., F ROITZHEIM , N., R UBENACH , M. J. & T URRIN , B. 2001. The role of detachment faulting in the formation of an ocean–continent transition: insights from the Iberia Abyssal Plain. In: W ILSON , R. C. L., W HITMARSH , R. B., T AYLOR , B. & F ROITZHEIM , N. (eds) Non-volcanic Rifting of Continental Margins: Evidence From Land and Sea. Geological Society, London, Special Publications, 187, 405– 428. M C K ENZIE , D. P. 1978. Some remarks on the development of sedimentary basins. Earth and Planetary Science Letters, 40, 25–32. M ICHAEL , P. J. & L ANGMUIR , C. H. ET AL . 2003. Magmatic and amagmatic seafloor generation at the ultraslow-spreading Gakkel ridge, Arctic Ocean. Nature, 423, 956– 962. M U¨ LLER , R. D., R OEST , W. R., R OYER , J.-Y., G AHAGAN , L. M. & S CLATER , J. G. 1997. Digital isochrons of the world’s ocean floor. Journal of Geophysical Research, 102(B2), 3211–3214. N IELSEN , T. K. & H OPPER , J. R. 2002. Formation of volcanic rifted margins: are temperature anomalies required? Geophysical Research Letters, 29, 2022– 2025. P HIPPS M ORGAN , J. 1987. Melt migration beneath midocean spreading centers. Geophysical Research Letters, 14, 1238– 1241. P ICKUP , S. L. B., W HITMARSH , R. B., F OWLER , C. M. R. & R ESTON , T. J. 1996. Insight into the nature of the

KINEMATIC HISTORY OF THE GOBAN SPUR ocean-continent transition off West Iberia from a deep multi-channel seismic reflection profile. Geology, 24, 1079–1082. P RESS , W. H., T EUKOLSKY , S. A., V ETTERLING , W. T. & F LANNERY , B. P. 1992. Numerical Recipes in FORTRAN: The Art of Scientific Computing. Cambridge University Press, Cambridge. S ANDWELL , D. T. & S MITH , W. H. F. 1997. Marine gravity anomaly from Geosat and ERS 1 satellite altimetry. Journal of Geophysical Research, 102(B5), 10039– 10054. S PIEGELMAN , M. & M C K ENZIE , D. 1987. Simple 2-D models for melt extraction at mid-ocean ridges and island arcs. Earth and Planetary Science Letters, 83, 137–152. S U , W. S. & B UCK , W. R. 1993. Buoyancy effects on mantle flow under midocean ridges. Journal of Geophysical Research, 98(B7), 12,191–12,205.

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The boundary between continental rifting and sea-floor spreading in the Woodlark Basin, Papua New Guinea A. M. GOODLIFFE1 & B. TAYLOR2 1

Department of Geological Sciences, University of Alabama, Tuscaloosa, Alabama, USA (e-mail: [email protected])

2

School of Ocean and Earth Science and Technology, University of Hawaii, Honolulu, Hawaii, USA

Abstract: Seismic reflection, bathymetry, acoustic imagery and magnetic data are presented that encompass the boundary between rifting of the Papuan continent and westward propagating seafloor spreading in the Woodlark Basin. West of the spreading tip, the southern margin is characterized by large fault blocks, which were tilted to the south on north-dipping normal faults during the current rifting phase, and graben further south where previous rifting failed. The northern margin is devoid of large offset normal faults and its subsidence from near sea level requires synrift flow of the lower crust. The margin asymmetry primarily reflects across-strike differences in the prerift geology, morphology and rheology. Horst and graben with 3 km relief below sea level occur adjacent to the spreading tip. Although heat flow data imply greater crustal thinning to the south of Moresby continental seamount, seismicity shows that deformation is currently focused on the Moresby normal fault that bounds its northern side. Based on logging results from ODP Leg 180 and semblance analysis, depth conversion of seismic reflection data show that Moresby Fault has a dip 298, which is compatible with an earthquake slip plane of 308–338. Directly east of Moresby Seamount, the first spreading segment is identified on the basis of magnetic, bathymetry and sidescan data to comprise two ring dykes in the west and Cheshire Seamount in the east. A seismic reflection line crossing the western ring dyke shows an intrusive body that cuts Moresby Fault at depth. Currently, intrusive rocks have only reached the surface near the ring dyke, leaving flat-lying synrift sediments, sills (?) and surface lava flows overlying the main intrusive body. The boundary between continental rifting and sea-floor spreading in the Woodlark Basin is spatially abrupt, although the processes overlap in time. The progression from rifting to spreading is characterized by a decrease in sedimentation as the margins are progressively thinned, subside below sea level, are eroded less and trap sediments in proximal basins. The post-rift sedimentation near the continent– ocean boundary is hemipelagic and drapes breakup topography without a distinct breakup unconformity. Synrift sediments, deposited above the 8.4 Ma rift onset unconformity and prior to breakup, are characterized by rotated sections (with parallel or diverging reflectors) as well as by ponded sections that onlap bounding fault blocks where local extension has ceased. Splayed reflectors representing sediment deposition on rotating fault blocks do not characterize the full synrift section, but only those times and places where extension is localized.

The transition between continental rifting and seafloor spreading marks the boundary between two fundamentally different modes of strain accommodation. How this transition takes place is poorly understood at all scales. Much of our present knowledge is based on observations from old and cold passive margins where the location of the conjugate margin is ambiguous and the basement is buried by thick sediments or voluminous magmatism (Holbrook & Kelemen 1993; Driscoll & Karner 1998). Active rifting–spreading boundaries are found only in a small number of locations. The Red Sea and the Gulf of Aden, both in cratonic continental settings, are currently forming as a result of rifting of the Arabian Peninsula from Africa (Cochran 1981; Martinez & Cochran 1988). The

Woodlark Basin and the Gulf of California (Moore 1973; Taylor et al. 1995, 1999; Fletcher & Munguia 2000) are also the result of continental breakups, but the tectonic history is made more complex by the interaction with recently active subduction systems. Several kinematic models have been proposed to explain basin-scale variations in features such as continental margin width, breakup style and spreading centre segmentation by varying the amount of extension prior to the initiation of sea-floor spreading (Courtillot 1982; Vink 1982; Martin 1984). Finite-element models have been used to match the observed data by varying the thickness and composition of the crust, the geotherm, the extension rate, and the strength and distribution of pre-existing

From: KARNER , G. D., MANATSCHAL , G. & PINHEIRO , L. M. (eds) Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 217– 238. DOI: 10.1144/SP282.11 0305-8719/07/$15.00 # The Geological Society of London 2007.

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weak zones (Bassi et al. 1993; Dunbar & Sawyer 1996; Harry & Bowling 1999; Lavier & Buck 2002; Corti et al. 2003; Van Wijk & Blackman 2005). Modelling studies have been hampered by a lack of accurate kinematic constraints on parameters such as strain rate, duration of rifting preceding rupture, and prerifting continental width and crustal thickness. The evolution of rift architecture through time is typically described as a transition from broadly distributed extension, where large crustal blocks rotate along faults that sole out at depth or transition to ductile deformation, to a narrow zone of focused extension that allows rapid lithospheric thinning and melt generation (Cochran 2005). In contrast, South Atlantic passive continental margins (Karner et al. 2003) are characterized by the stacking of regional synrift sag basins. Large, late-stage rift faults in the South Atlantic margins do not control the regional development of synrift accommodation and are thus a small part of the extension process. Late-stage subsidence during the rifting– spreading transition is usually explained by interpreting the margin being studied as the upper plate in a simple shear model. This has resulted in the ‘upper plate paradox’ (Driscoll & Karner 1998), where both conjugate margins have been interpreted as the upper plate. This paradox can be avoided by using lower crustal and upper mantle deformation to create subsidence during late-stage continental extension (Karner et al. 2003). Subsidence may also originate from cooling and contraction of anomalously hot asthenosphere in the region of deformation (Buck 2004) or be the result of an upwelling divergent flow field within continental lithosphere and asthenosphere (Kusznir et al. 2004). The processes involved in the initiation and growth of a spreading centre in a continental rift have been examined on a large scale at several locations. In the Red Sea, continuous sea-floor spreading in a welldefined axis in the south becomes discontinuous to the north, forming discrete deeps interpreted as cells of sea-floor spreading (Bonatti 1985). The regular spacing of the deeps and the similar spacing of discontinuities in the southern Red Sea suggest that both may be caused by asthenospheric upwelling, the wavelength of which is controlled by the ratio of the viscosities and the layer thicknesses. Once a diapir reaches the surface at the centre of a segment it continues to grow along strike, eventually merging with its neighbours to create a continuous spreading centre. During rifting, adjacent segments may be offset by accommodation zones that eventually form the site of oceanic transform faults (Cochran & Martinez 1988). In the Woodlark Basin, two distinct modes of sea-floor spreading initiation have been identified. One involves propagation of sea-floor spreading into the rifting continent; the other nearly simultaneous nucleation of

sea-floor spreading along more than 100 km-long rift segments (Taylor et al. 1999). The former results in a continent–ocean boundary (COB) that is highly oblique to early margin structures, abyssal hill fabric and magnetic isochrons. In the later case the COB, abyssal hill fabric and magnetic isochrons are all parallel. Other studies of the geological and dynamic processes by which sea-floor spreading initiates in a continent (Rosendahl 1987; Hayward & Ebinger 1996) demonstrate that many of the details are highly debated. One shortcoming of all these studies is that they have not determined how the spreading centre first ruptures the surface and how this location is related to the late-stage continental faulting. The Woodlark Basin is perhaps the only place where the COB has been located regionally at a kilometre scale. Sediment cover is generally thin, and the opening kinematics, strain history and margin subsidence have been intensely investigated (Taylor et al. 1995, 1999; Goodliffe 1998; Martinez et al. 1999; Taylor & Huchon 2002). This affords us an unprecedented view of the detailed processes involved in: (1) continental rifting directly preceding the initiation of a spreading centre; (2) the mechanics of spreading centre emplacement; and (3) the early evolution of a spreading centre.

Spreading history and Euler poles Rifting of the Papuan Peninsula started at approximately 8.4 Ma (Taylor & Huchon 2002), and led to separation of the once contiguous Woodlark and Pocklington rises that bound the Woodlark Basin (Fig. 1). The oldest oceanic lithosphere in the Woodlark Basin, formed during magnetic chron 3A.1 (Taylor et al. 1999) (c. 6 Ma: Cande & Kent 1995), is found at the eastern end of the basin adjacent to the southern margin; any older oceanic lithosphere has been subducted under the Solomon Islands. Reconstructions based largely on fracture zone trends (Goodliffe 1998) show that since magnetic chron 3A.1 the majority of basin opening can be described as a rotation about a single Euler pole (1478 þ 18–28E, 9.38S + 0.28S, 4.2348 Ma21 Fig. 2), implying that rifting initiated simultaneously throughout the basin. Abyssal hill fabric, magnetic isochrons, a major low-angle fault bordering Misima Island (Hill 1995), faults bounding Milne Bay graben (Jongsma 1972) and Goodenough Fault, a structure bounding the southern side of Goodenough Basin (Fang 2000) (Figs 1 & 2), are all co-polar with this Euler Pole. Sea-floor spreading propagated in a stair-step fashion from the east into rifting crust at an average rate of approximately 140 km per Ma21 (Taylor & Exon 1987), reaching Moresby Seamount, where the rifting to spreading transition is

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Fig. 1. Regional setting of the boundary between continental rifting and sea-floor spreading in the Woodlark Basin. The conjugate Woodlark and Pocklington Rises have separated over the last 8.4 Ma as a result of rifting the Papuan Peninsula and the westward propagation of sea-floor spreading (accreted oceanic crust is outlined by a solid black line; the spreading centre is depicted by a black line with white centre). Dashed black lines indicate active strike-slip motion. Teleseismicity is shown by black circles.

currently found. The 500 km-long active spreading centre, located on the basis of bathymetry and acoustic imagery, comprised five main segments at the start of the Brunhes chron (0.78 Ma). An up to 228 counter-clockwise reorientation of the spreading centre at 80 ka (Goodliffe et al. 1997) in response to a shift in the Euler pole at 0.52 Ma (Goodliffe 1998) (Fig. 2) resulted in the subdivision of many of the spreading segments into shorter segments that are now oblique to the Brunhes isochron. The modern Euler pole (128S, 1448E, 2.4378 Ma21: Fig. 2) is defined largely by the offset of pre-existing fracture zones and is close to the Australia– Woodlark Euler pole (11.288S, 147.6.28E, 2.82 + 0.188 Ma21) (Wallace et al. 2004) determined by simultaneously inverting GPS data and earthquake slip vectors with transform orientations. Spreading segment 1 extends from the current spreading tip at 1518400 E–152851.60 E and comprises three subsegments, 1a, b and c (Fig. 2). Seafloor spreading on this segment nucleated between 1518450 E and 152851.60 E at approximately 0.8 Ma (Taylor et al. 1995; Goodliffe 1998). At the resolution of the available data and analysis techniques the COB is coincident with the Brunhes isochron

and parallel to great circles about the pre-0.52 Ma Euler pole (Fig. 2). The change in Euler pole resulted in the subdivision of segment 1 into two shorter segments and an approximately 88 rotation. Unlike segments 2– 5, segments 1b and 1c are closer to co-polar with the pre-0.52 Ma Euler pole and the Brunhes isochron than to the current pole. To the north of segment 1, the seismically active Egum Graben (Figs 1 & 2) cross-cuts normal faults in the northern margin of segment 1 and has a trend that is co-polar with the present Euler pole. Although this may be coincidental, it may mark the reactivation of a favourably trending structure by the change in Euler pole. This might also explain the concurrent sea-floor spreading and continental rifting at this longitude (Taylor et al. 1995). The fault bounding Egum Graben has been interpreted as a continuation of the strike-slip fault separating the eastern Woodlark Basin from the Solomon Sea (Taylor et al. 1999). This transtensional fault system may bound a third (Trobriand) plate, including the margin to the north of Moresby and Egum Graben, Egum Atoll and Woodlark Island (Davies et al. 1984), and form a triple junction with

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Fig. 2. Map of major faults seen at the surface in the vicinity of the rifting–spreading transition (a). The axes of the active neovolcanic zones are shown by thick solid lines and the limit of oceanic crust is dashed. Sea-floor magnetization illuminated from the north by bathymetry (b). Oceanic crust formed since the start of the Brunhes chron (0.78–0 Ma) on segment 1 has a positive magnetization. At the resolution of the data the continent– ocean boundary and the 0.78 Ma isochron are coincident. Small and great circles about the present Euler pole (128S, 1448E, 2.4378 Ma21, solid grey lines) and the pre-0.52 Ma Euler pole (1478 þ 18–28E, 9.38S + 0.28S, 4.2348 Ma21,

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spreading segment 1a and the Papuan rifts. Given the complex tectonic setting of segment 1a it is not clear that it is opening around the same Euler pole as the spreading segments further east.

The continental margins proximal to the initiation of sea-floor spreading Using swath bathymetry data and three representative six-channel seismic lines (see Appendix 1 for details of the data used and processing steps), we first describe the margin structures that surround sea-floor spreading segment 1. The margin of the Woodlark Basin south of segment 1 and Moresby Seamount (Fig. 2) is characterized by large fault blocks, trending on average east–west but locally NE–SW, that are typically tilted to the south and overlain by isolated sedimentary basins that record varying degrees of rotations (Fang 2000) (Fig. 3). Poor basin interconnectivity and an absence of sampling make attempts at dating fault motion purely speculative. A scarcity of seismicity in this region (Abers et al. 1997; Ferris et al. 2002) (Fig. 1) indicates that there is currently little fault motion. An analogy has been made to oceanic rift propagation, with many of the fault blocks rotated to the NE about vertical axes by shear between laterally offset centres of extension (Mutter et al. 1996). Six-channel seismic (SCS) line MW9304-30 (Figs 2 & 3) crosses the current focus of deformation to the west of Moresby Seamount where sea-floor spreading has not yet initiated. At the southern end of the line there are isolated sediment ponds that record tilting to the south of at least two large crustal blocks. The northward tilting of an adjacent progradational sequence illustrates that deformation is locally more complex. A graben centred on SP900 that marks the deepest point at this longitude widens to the east to become the South Moresby Graben (SMG). Further north a series of three fault blocks tilted to the south on north-dipping normal faults sit in the westward extension of Moresby Seamount and North Moresby Graben (NMG). Some complex basement geometries (for example, at SP1560) suggest locally more complex deformation, possible involving unresolved south-dipping faults. A fourth ridge (SP1900) is an uplifted horst block with a thick synrift sedimentary sequence. To the north, the largely undeformed northern margin with its thick synrift sedimentary sequence dips to the south.

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The stratigraphy here is contiguous with that imaged to the east and sampled during ODP Leg 180 (Taylor & Huchon 2002). North of SP2500, prominent intrabasement reflectors below the synrift sediments dip to the north. These are believed to originate from sills within the approximately 66 Ma diabase and metadiabase basement (Brooks & Tegner 2001; Monteleone et al. 2001). SCS line MW9304-50, 18.5 km to the east, crosses the SMG, Moresby Seamount, and NMG. Owing to the near coincidence with this line of multichannel seismic (MCS) line EW9510-1366 and the ODP Leg 180 drilling transect, the stratigraphy and basin history of Moresby Seamount and the northern margin is very well known at this longitude. In summary, this involved gradual subsidence from sea level and tilting to the south of the northern margin, with early focusing of extension to the north and south of Moresby Seamount and its subsequent isostatic rise between late bounding faults (Taylor & Huchon 2002) (Fig. 4). It is important to note that, although the northern margin is bounded in the south by the major north-dipping Moresby Fault, the wavelength of subsidence is too great to be due to that fault alone. The absence of large-offset normal faults on the northern margin means that other mechanisms such as flow of the lower crust must be called on to explain its subsidence. The ODP drilling transect did not extend to the southern section of SCS line MW9304-50, leaving the details of the evolution of this part of the basin unexplored. SMG (Fig. 2) is deepest directly to the south of the peak of Moresby Seamount. Here it is subdivided into two parts separated by an north– south trending ridge that plunges to the north. To the east of the ridge the graben floor reaches 3150 m. Low-amplitude positive magnetization implies that rifting may have been accompanied by magmatism extending west from spreading segment 1b. The graben floor is subhorizontal and sediment covered, with the exception of localized regions of high acoustic backscatter in the north indicative of fault scarps (Fig. 2c). A 4  0.1 km linear feature on the sea floor at 1518420 E, 98570 S is interpreted as a lava flow. To the west of the northplunging ridge the graben floor is characterized by generally smooth sea floor reaching depths of 3000 m. A somewhat rougher sea floor close to Moresby Seamount is likely to be the result of debris flows. Consistently high values of heat

Fig. 2. (Continued) dashed grey) are shown. The 0.78 Ma boundary is co-polar with great circles about the pre-0.52 Ma Euler pole, whereas Egum graben is co-polar with great circles about the current Euler pole. Bathymetry illuminated from the north and overlain by acoustic imagery (c) shows the location of the neovolcanic zone (black ¼ high acoustic backscatter) relative to the main fault trends in the vicinity of the rifting– spreading transition. The location of the six-channel seismic lines described in this paper are shown along with ODP sites 1108, 1117 and 1118.

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Fig. 3. Six-channel seismic (SCS) lines MW9304-30 (top), MW9304-50 (middle) and MW9304-70 (bottom) across the continental margins in the vicinity of the rifting– spreading transition. The rotation of back-tilted fault blocks is recorded locally by the fanning of sediment packages in half graben that form small isolated basins. The southern margin is dominated by these tilted fault blocks, whereas the northern margin is largely unfaulted, but subsided and south tilted. From SCS line MW9304-30 in the west to MW9304-70 in the east there is a focusing of extension first at North and South Moresby Graben and then on spreading segment 1a. The blue line marks the base of the synrift sediments. ODP Site 1118 is projected from its true location 1.5 km to the east. ODP Site 1117 is projected from its true location 0.7 km to the west. The location of the SCS lines is shown in Figure 2.

flow, reaching a peak of 254 mW m22, were recorded in SMG (Goodliffe et al. 2000). This is in contrast to a maximum of 100 mW m22 recorded in NMG. SCS and MCS data in SMG (Fig. 3) do not satisfactorily image basement. Nevertheless, gravity data indicate a similar amount of crustal thinning as NMG (Goodliffe et al. 1999b). To the south, between SP500 and SP1200 on SCS line MW9304-50, north-dipping normal faults bound crustal blocks that have been tilted to the south. Further to the south at SP390 a southdipping normal fault forms the northern margin of a shallow rift basin, the eastward extension of which is seen in the form of a series of southdipping faults on SCS line MW9304-70. The southern side of this rift basin was not imaged

owing to its proximity to the reefs that top Pocklington Rise. Further east, between 1518500 E and 1528300 E, a scarp dipping approximately 308S is seen in the bathymetry (Figs 1 & 2). This fault may form the northern boundary of a basin seen to the west of Misima Island. Still further to the south a large (.160 km long and c. 40 km wide) basin is bounded by reefs (Fig. 1). The absence of seismicity and the location of this basin away from the focus of current activity suggest that it is an early (failed) rift that predates the rifting that led to breakup and the initiation of sea-floor spreading on segment 1. To the north of SP250 on SCS line MW9304-70 there are at least three large fault blocks that have been tilted to the south by motion on north-dipping

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normal faults. North of SP800 fault geometries become unclear, possibly because of extreme fault block rotation and subsequent cross-cutting by the late-stage north-dipping faults indicated by the small sea-floor offsets. Locally, the sea floor is underlain by horizontally stratified sediments. This close to the western end of segment 1b, the presence of lava flows, dykes and sills cannot be ruled out. Basement and fault plane reflectors on the small ridge, continuous with Moresby Seamount to the west and centred on SP1600, suggest that it is underlain by south-tilted fault blocks. Directly to the north the sea floor has high acoustic reflectivity and is interpreted to be floored by recent lavas (Fig. 2c). Small ridges on the sea floor at SP1830 and SP1920 are part of an approximately 4 km-diameter circular feature on the sea floor. Below the region of high sea-floor reflectivity there are no coherent reflectors. At SP2195 and SP2350 major south-dipping normal faults downdrop the thick synrift sedimentary sequence of the northern margin to the south. As in SCS lines MW9304-30 and MW9304-50, the margin further north is tilted to the south and shows no evidence of major fault offsets. The SCS lines, especially MW9304-30 and MW9304-50, show marked margin asymmetry (Fig. 3). The northern margin slopes gradually to the south with little evidence of large-offset normal faults until the NMG is reached. In the south large fault blocks are rotated. At face value this could imply a simple shear origin for the margin (Wernicke 1985; Lister et al. 1991), but a detachment fault of sufficient size to explain the wavelength of the subsidence has not been identified. More probably, the structural asymmetry is a function of the preexisting continental morphology–geology–thickness (forearc basin v. mountainous continental arc and back-arc) combined with lower crustal flow (Goodliffe et al. 1999a). Between 1518550 E and 1528500 E the fabric on the southern and northern rifted margins close to the continent–ocean boundary (COB) is broadly parallel to the N848W-trending COB (Fig. 2). This is consistent with the initiation of sea-floor spreading almost simultaneously along segment 1 (Taylor et al. 1999). The normal faults bounding the margin north of segment 1 are cut by Egum Graben. This more than 100 km-long structure trends N738E from Moresby Seamount and is the focus of ongoing seismic activity with many .Mw6 earthquakes (Taylor et al. 1995, 1996; Abers et al. 1997). To the east the rift narrows and shallows towards the platform surrounding Woodlark Island. At 1528130 E Egum Graben is 17.5 km wide and bounded by scarps with approximately 408 dips. The main rift floor, at a depth of 2500 m, is 600 m below the rift shoulders. A

nested 4.1 km-wide and 50 km-long graben on the southern side of Egum Graben locally reaches a depth of 2600 m and terminates in the west against Cheshire Seamount. There are small edifices to the south of Egum Graben at 1528100 E, 98500 S that may be volcanic in origin, but are no longer active (Fig. 2c). At 1528400 E, 108180 E the southern margin is cross-cut by a WNW-trending COB (Fig. 2). Seafloor spreading on segments 1 and 2 propagated past one another, forming an overlapper that surrounded and rotated counter-clockwise a sliver of continental crust (centred at 1528400 E, 108100 E) (Goodliffe 1998). Since 80 ka, sea-floor spreading on segment 2 has rotated and jumped to the south. Propagation of the new spreading into the margin has created a small re-entrant at 1528450 E, 108250 S (Fig. 2c). Kinematic reconstructions at the longitude of Moresby Seamount (Fang 2000; Taylor & Huchon 2002) of faulting and subsidence north of 108S show that many of the faults (with dips ranging from 308 to 608) that are active today were active from at least 5.5 Ma (Fig. 4). Between 5.5 and 1.6 Ma, extension on the major faults only added up to 7 km, compared to more than 12 km between 1.6 Ma and the present day. However, similarities in sedimentation rate between ODP sites 1108, 1109, 1118 and 1114 (Takahashi et al. 2002) show that these sites were in a common tectonic setting with no large depositional barriers separating them until at least the late Pliocene (Robertson & Sharp 2002). Although the rift basins in this region were growing, fault growth was being outpaced by sedimentation. The post-1.6 Ma increase in extension rate at the longitude of Moresby Seamount (Taylor & Huchon 2002) and a dramatic decrease in sedimentation on the northern margin at 1.2 Ma (Takahashi et al. 2002), implying that the rapid subsidence of NMG and uplift of Moresby Seamount had cut off sedimentation to the north (Sharp & Robertson 2002), both indicate strain localization on the rifts surrounding what would become sea-floor spreading segment 1. When sea-floor spreading initiated on segment 1 at about 0.8 Ma, the neovolcanic zone extended only from 1518450 E to 1528400 E (Goodliffe 1998). Beyond these limits, strain accommodation and crustal thinning was still not sufficiently focused for spreading to initiate. To the east strain localization was achieved by propagation of the neovolcanic zone, forming a spreading centre that overlapped with segment 2 (Goodliffe 1998). The neovolcanic zone did not propagate to the west, remaining stalled at 1518480 E for approximately 0.7 million years. Strain accommodation on the faults directly to the west of the spreading tip must therefore have still remained distributed, at

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least within the confines of the segment 1 rift. This produced the horst and graben terrain that is centred on Moresby Seamount and continues westward to the north of Normanby Island (Figs 1 –3). Within this faulted terrain, earthquakes indicate that current deformation is localized on the NMG (Fig. 1), probably at least since the nucleation of spreading on segment 1a about 0.1 Ma.

Faulting directly ahead of the westernmost spreading segment NMG has been a focus of MCS surveys and ODP Leg 180 drilling aimed at understanding continental breakup and the evolution of a shallow-angle normal fault. We reproduce here just four of the tens of MCS lines across the structure (Figs 5–7). In addition, a line drawing interpretation of MCS line EW9510-1366, which passes through the primary transect of drill sites, is shown in Figure 4 and all of the EW9510 MCS lines are published (Goodliffe et al. 1999b). The MCS data show that NMG is an asymmetric graben bounded to the south by Moresby Fault and to the north by an antithetic normal fault system. The width of the graben between the sea-floor traces of the bounding faults varies from approximately 12 km NNE of Moresby Seamount to more than 20 km on either side, varying as a function of the curvilinear to zig-zag fault patterns (Figs 2 & 5). MCS line EW9510-1369 crosses the NMG near its narrowest (Fig. 5) and images Moresby Fault continuing at depth beneath the antithetic fault and a downdropped keystone basement block (Fig. 6). Basement drilled on both sides of NMG beneath the rift-onset unconformity (Fig. 4) comprises diabase and, in the footwall of Moresby Fault at ODP Site 1117, gabbro with a crystallization age of 66 + 1.5 Ma (Brooks & Tegner 2001; Monteleone et al. 2001). Heavily faulted Pliocene–Pleistocene synrift sediments cap Moresby Seamount, which rises to almost 150 m bsl (below sea level) (Taylor & Huchon 2002) (Figs 3 & 4). The northern border fault of the NMG is a growth fault, dipping approximately 548S, with basement offsets of about 1.5 km but surface offsets of only approximately 150 m (Figs 4 & 6). Whereas its growth matched that of Moresby Fault early in the development of the rift, it has been far exceeded in the Quaternary by the faults that bound Moresby Seamount (Fig. 4). The NMG is contiguous with the trough containing spreading segment 1a. It shares a common northern border fault with Egum Graben, but has a deeper axis. From the western end of the neovolcanic zone the sea floor climbs 150 m over a distance of approximately 5 km before descending to the deepest part of the NMG at 3250 m (Fig. 5) where

Fig. 5. Bathymetry (a) in the vicinity of the rifting– spreading transition. Key features, including two ring dykes on the seafloor at the western end of segment 1a, are labelled. Cheshire Seamount is a volcanic construct at the eastern end of segment 1a. Sea-floor magnetization (b) for the same area (see Fig. 2 for the colour scale). Profiles along which the horizontal gradient was calculated are shown in grey. Black dots mark the location of gradient maxima. Insets show the magnetization (blue) and horizontal gradient (green) along the profiles A1– A2 and B1 –B2. The separation of the maxima is a measure of the width of the causative body assuming a tabular geometry.

the hanging-wall sediments are tilted south and have collapsed the most (Fig. 4). ODP Site 1108 (Fig. 2) was drilled in the bathymetric saddle on MCS line EW9510-1374 (Fig. 6). It and other drill sites at the foot of the Moresby Fault could not reach the fault at depth owing to the presence of coarse, unconsolidated talus shed from the slope. The synrift sediments north of the NMG are approximately 1 km thick and thin northwards, but those within the graben are more than 2 km thick and subparallel to basement (Figs 4 & 6). The graben basement is almost horizontal on MCS lines EW9510-1369 and EW9510-1374 but dips gently south on MCS lines EW9510-1366 and EW9910-11. Basal onlap is rarely observed. An exception occurs south of SP7200 on MCS

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Fig. 6. MCS lines EW9510-1369 (top), EW9510-1374 (middle) and EW9910-11 (bottom) across the North Moresby Graben to the west of the rifting– spreading boundary (see Fig. 5 for locations). The MCS lines image the shallow-angle Moresby Fault and the high-angle, graben-bounding, antithetic fault system (on which the profiles are aligned). The blue line marks the base of the synrift sediments and the location of ODP Site 1108 is shown. For clarity, only some of the larger-offset normal faults are drawn. The NMG widens eastward; only the western line, EW9510-1369, images the Moresby Fault passing under the antithetic fault.

Fig. 7. MCS line EW9510-1369 (see Fig. 6a for time section) converted to depth using a velocity model for the synrift sediments derived from ODP results and semblance velocities. A velocity of 4.5 km s21 is assigned to basement at the base of the synrift sediments. Below this a gradient is applied such that the velocity is 6 km s21 at 8 s TWT. In this depth section the apparent dip of Moresby Fault is approximately 298 and of the antithetic fault is 548.

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line EW9910-11 (Fig. 6) where sediment onlap to the north is likely to have been the result of basement tilting to the south associated with early motion on the Moresby Fault. The thick NMG sedimentary fill is cut by myriads of small-offset, north- and south-dipping, high-angle normal faults. For clarity, most of these faults are omitted on the interpretations in Figure 6, but the surface traces of the larger ones are shown in Figure 2. Sediment reflectors lose coherency as Moresby Fault is approached, usually above the point at which the basement reflector intersects the fault. This is probably the result of fault drag, extensional collapse, fluid flow and deposition of slump material. Moresby Fault was first proposed to be active on the basis of teleseismic waveform inversions of normal faulting earthquakes in the vicinity of Moresby Seamount that had shallow-dipping nodal planes (Abers 1991). At least two of the focal mechanisms demonstrate normal faulting at dips of 258– 358, consistent with EW9203 MCS data (Abers et al. 1997). Evidence of water multiples in a water depth of 3 km indicated that the location of one of the largest earthquakes, on a plane dipping 31.68 + 1.38N, is beneath the deepest part of the NMG and coincident with the Moresby Fault. An initial estimate of the dip of Moresby Fault made using MCS lines EW9510-1369 and 1374 was 278 + 38 towards 0158 (Taylor et al. 1999). Velocity data collected by logging and laboratory measurements during ODP Leg 180 allow a more accurate determination of the dip of Moresby Fault. Here we present a depth-converted version of MCS line EW9510-1369 using realistic velocities for the synrift sediments in NMG. This line has a number of desirable characteristics: (1) Moresby Fault and the south-dipping antithetic fault intersect at depth; (2) a distinct and continuous reflector is imaged along the length of the Moresby Fault; and (3) in two-way travel-time both faults have a distinct kink that is due to velocity pull-up where they descend below the keystone basement block that should be removed by depth conversion. Using the velocity model derived in Appendix 2, MCS line EW9510-1369 was converted to depth (Fig. 7). Moresby Fault has an average apparent dip of 298N and that of the antithetic fault is 548S. The average strike of the antithetic fault system is east– west (Figs 2 & 5) and hence its apparent dip on this line is true. In contrast, Moresby Fault has a sinuous fault trace (Figs 2 & 5), depth contours indicate a curviplanar form and varying strike down-dip (Taylor et al. 1999), and the direction of slip may have changed during its long history (Fig. 4). If an upper fault strike of 121.58, determined from 11 200 m-spaced EW9910 MCS lines across the NE face (Fig. 5), is

used then the apparent dip equates to a true dip of 388. However, if the long-term slip paralleled small circles about the main Euler pole of opening (i.e. N78E, Fig. 2) then the Moresby Fault dip in this direction is 29.28, and in the 0158 direction used by Taylor et al. (1999) the dip is 29.98. Fault dips of 308 –338 are compatible with the slip plane of the large earthquake located beneath the NMG (Abers et al. 1997). That the approximately 308 – 338 Moresby Fault is active has been the subject of much controversy. Theory states that if the coefficient of friction is that of typical rock (0.6–0.8) and s1 is near vertical then a normal fault will dip near 608 (Anderson 1951) or 458 for the case of finite slip (Forsyth 1992; Buck 1993). Lower-angle faults either may be explained by rotation of the stress field (Yin 1989; Parson & Thompson 1993), weakening through near lithostatic pore-fluid pressures within the fault zone (Byerlee 1990) or fault gouge materials with a lower coefficient of friction (Hill & Thatcher 1992). A primary goal of ODP Leg 180 was to penetrate Moresby Fault and investigate the mineralogy and stress state to determine why this fault is active. Although the active segment of Moresby Fault was not penetrated, 12 m of talc –chlorite–serpentinite clayey fault gouge with coefficients of friction of 0.21 –0.3 (Kopf et al. 2003) was recovered at Site 1117 along a segment of the fault exposed at the sea floor. Underneath, a poorly recovered approximately 50 m-thick section of mylonites (recording plastic deformation) and ultracataclastites (recording intermittent plastic and brittle deformation) was recovered from the quartz-gabbro basement. Analysis of samples from Site 1117 and other proximal sites has demonstrated the importance of frictional heating and hydrothermal fluid flow in the evolution of the Moresby Fault (Kopf et al. 2003). Channelling of fluids along this fault is evinced by the greater permeability of the fault gouge, especially in the direction parallel to the tectonic fabric, than the overlying sediments (Kemerer & Screaton 2001; Kopf 2001; Stover et al. 2001). The continuation of the fault gouge along the Moresby Fault beneath the NMG is indicated by inverse models of the MCS data that include a 4.3 km s21, approximately 33 m-thick layer along the fault with isolated pockets of a 1.7 km s21 lower velocity material (Floyd et al. 2001). Given the friction coefficients of the fault gouge and the presence of fluids in the fault zone, either of which alone would make motion on the Moresby Fault feasible (Abers 2001), strain accumulation on this fault is no longer an issue. The initiation of motion along an approximately 308 fault remains problematic, for although full flexural modelling of the evolution of rift systems shows the importance of rotation from initially high to low dips (Buck

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1988), rotating the Moresby Fault to higher dips simultaneously rotates the large, contemporary, southdipping normal fault within Moresby Seamount to low dips (Fig. 4) (Taylor & Huchon 2002). The abundant pre-existing structures in this once convergent regime (Davies & Jaques 1984) may play a role.

Morphology and structure at the point of continental breakup Sea-floor spreading segment 1a occupies the southern half of the eastern NMG and obliquely truncates the SW end of Egum Graben. It is identified by a narrow band of high-amplitude positive magnetization and a coincident region of sea floor with high acoustic backscatter extending from 1518400 E to 151852.250 E (Figs 2 & 5). At the eastern end of the segment, the 8 km-diameter Cheshire Seamount, bisected by an east–west cleft, rises more than 850 m above the surrounding sea floor to a depth of 1750 m. Recent lava flows extend down-slope to the east but are unlikely to be underlain by extensive intrusives given the reduced amplitude of the magnetization. Directly to the west of Cheshire Seamount, and at depths of 3050 and 3250 m, are two ringshaped structures about 4 km in diameter that are in the centre of the Brunhes magnetization high (Figs 2 & 5). Around both constructs, high acoustic backscatter reveals lava flows that have pooled against the base and extended a short distance up the slopes of the surrounding fault-bounded ridges. Hyaloclastites were found in the sediments 5 km to the NW at ODP Site 1108 (Shipboard Scientific Party 1999b). They are high Na2O (4.3%), high TiO2 (2.8%) basaltic andesites (SiO2 ¼ 53%) that are light rare earth element (LREE)-enriched (chondrite-normalized La/Yb ¼ 3.5) (Lackschewitz et al. 2001). Similar rocks have been dredged from Cheshire and Franklin seamounts; the former also includes dacites (Binns & Whitford 1987; Dril et al. 1997). Other rocks from the west end of segment 1a are normal mid-ocean ridge basalts (N-MORBs) with indications of low degrees of melting (at MgO ¼ 8%, Na2O ¼ 3.1%, FeO ¼ 8.1% and TiO2 ¼ 1.14%) (Binns & Whitford 1987). Thus, the mantle sources of segment 1 lavas include N-MORB, arc-like and enriched components (Dril et al. 1997). The extent of the intrusive body that fed the segment 1a lavas can be derived from the magnetization anomaly. Assuming that the intrusion is broadly tabular, the edges of the body will be coincident with peaks in the horizontal derivative of the anomaly (Goodliffe 1993). A series of north–south profiles across the magnetization anomaly were extracted and the horizontal derivative calculated and maxima plotted. Five profiles centred on two locations (A1–A2, B1 –B2: Fig. 5b) were stacked to remove

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noise before the maximum horizontal derivative was calculated. The horizontal derivative for the western part of segment 1a displays distinct peaks that lie on either side of a body that is between 4.1 and 5.7 km wide. Closer to the Cheshire Seamount the calculated width of the body increases to 7 km, which suggests that segment 1a propagated to the west. Using the current Euler pole and an intrusive body width of 5 km we calculate that sea-floor spreading began on the western part of spreading segment 1a by 0.11 Ma. This is considered an upper limit as some of that width may originate from overprinting rather than strain accommodation. For example, if spreading initiation coincided with the 80 ka reorientation of the entire Woodlark Basin spreading centre (Goodliffe et al. 1997) then only 3.6 km of the intrusion would represent strain accommodation. MCS line EW9910-13 crosses the western end of segment 1a, providing a snapshot of the development of this young accretionary plate boundary (Fig. 8). On this MCS line, the Moresby Fault and the antithetic fault that together bound the NMG are readily identified, as are the thick synrift sediments above the keystone basement block. Between SP8300 and SP8500, however, there is a region of incoherent reflectivity below 5 s two-way time (TWT) against which the Moresby Fault, the synrift sediments and the keystone basement reflectors terminate (Fig. 8). We interpret this incoherent region to be the expression of the Brunhes axial intrusion. It correlates closely with, but is about 800 m north of, the limits of the intrusive body determined from the magnetization anomaly (Figs 5 & 8). Those limits exactly match the position of three sea-floor ridges that, with the bathymetry data, we interpret as volcanic piles above a ring dyke (northern two ridges) plus another dyke to the south, fed from the intrusion below (Fig. 8). Above Moresby Fault at the southern end of MCS line EW9910-13 there are few coherent reflectors, perhaps owing to deposition of sediments in the form of slumps originating on Moresby Seamount (probably similar to talus found at ODP Leg 180 Site 1108) (Shipboard Scientific Party 1999a) (Fig. 8). To the north of the intrusive body the synrift sedimentary sequence is offset at SP7925 by a south-dipping normal fault. This fault is imaged on MCS line EW9910-11 to the west and continues to the east as the major northern bounding fault of Egum Graben. The synrift sequence is interpreted to be continuous with and have a depositional history similar to that sampled to the west during ODP Leg 180 (Shipboard Scientific Party 1999a). The evidence presented above indicates that a major intrusive body underlies the centre of

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Fig. 8. Uninterpreted (top) and interpreted version (bottom) of MCS line EW9910-13 (see Fig. 5 for the location). The currently forming continent –ocean boundary involves igneous intrusions penetrating the shallow-angle normal fault of an asymmetric rift graben. Moresby Fault is cut by an approximately 5 km-wide intrusive body that is overlain by subhorizontal sediments that have been cross-cut by a ring dyke.

segment 1a. With a width of 5 km it is assumed that it is today accommodating extension in a similar fashion to a mature spreading centre, although the intrusion has only broken the surface in the form of localized dykes. If the intrusive body is similar to a mature spreading centre it may take the form of a sheeted dyke sequence that is analogous to those observed at ophiolites and inferred to exist below modern spreading centres. Although there is no evidence of an active magma chamber, the sheeted dyke sequence may be underlain by massive

gabbros. Although there are extensive lava flows at the surface, the MCS data suggest that they overlie an approximately 500 m-thick (assuming a velocity of 1800 m s21) extension of the synrift sedimentary sequence seen to the north. As the intrusive body does not reach the surface, extension above it must be transferred into the overlying synrift sediments. The well-stratified reflectors show little evidence of brittle faulting, implying either a dominance of soft-sediment deformation or the offset of extension to either side of the intrusion. Much of the sediment

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in this bathymetric low may have been deposited rapidly and recently.

Discussion Breakup model We present a simplified model for the formation of the current rift-spreading boundary in the Woodlark Basin (Fig. 9). The initial condition, prior to continental extension, would resemble the central

(a)

50 km

(b)

Trobriand Forearc VE = 4.5

2 0 -2 -4

Woodlark Rise

MS

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Failed Rift

VE = 4.5

2 0 -2 -4

kilometres

Pocklington Rise

Papuan Peninsula today, with an orogenic mountain chain of 1–3 km elevation, crust about 40 km thick, arc volcanoes spaced along the northern coastal plain and a broad shallow-marine forearc basin south of the Trobriand Trench (Taylor et al. 1999). At 8.4 Ma (Taylor & Huchon 2002) this inherently asymmetric region started to undergo north–south extension, lowering the topographic profile (Fig. 9a) and exhuming the metamorphic core complexes of the D’Entrecasteaux Islands (Fig. 1) to the west. Successive generations of

kilometres

Papuan Peninsula

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(c)

25 km

VE = 1.5

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kilometres

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2 0 -2 -4 -6 -8 -10

North Moresby Graben Moresby Fault Dyke

Antithetic fault

(d)

VE = 1

(e)

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Ring dyke

(f )

VE = 1

Fig. 9. The evolution of continental rifting through sea-floor spreading in the western Woodlark Basin. The Papuan Peninsula (topography based on a north– south profile across the Papuan Peninsula at 148.88E: (a) has undergone broadly north–south extension, subsiding to current depths at the longitude of Moresby Seamount (b). Southern graben represent former loci of extension that are now failed rifts. The southern margin is characterized by large fault blocks rotated by high-angle faults. The northern margin has subsided with little evidence of large-offset faults (c). At about 1.2 Ma extension focused on South and North Moresby Graben. More recently, extension has focused at the NMG and, ultimately, to Moresby Fault. A single dyke was intruded into the fault and the overlying sediments causing motion on the up-dip segment of Moresby Fault to cease (d). Progressively growing in width (e), the intrusion is today about 5 km wide but has only reached the surface in the form of localized dykes (f) and flows.

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normal faults thinned the palaeo-Papuan Peninsula, resulting in subsidence below sea level to form the Pocklington Rise. To the north, the palaeo-Trobriand forearc (now Woodlark Rise) has tilted to the south and subsided by up to 3 km, reversing the northward-directed Miocene drainage pattern (Fig. 9b). One model for the growth of the metamorphic core complexes in the D’Entrecasteaux Islands (Fig. 1) suggests that, with the aid of fluids liberated from the subducted Solomon Sea Plate, the lower crust of the northern margin is more capable of ductile flow than the margin to the south (Martinez et al. 2001). The paucity of upper crust normal faults in the northern margin implies that crustal thinning must be primarily achieved through means other than brittle faulting. The bulk of the thinning may take place through the removal of the lower crust (Driscoll & Karner 1998; Buck 2004; Kusznir et al. 2004). In contrast, brittle failure of the upper crust is more prevalent to the south. The margin asymmetry need not reflect lithospheric-scale simple shear, but instead is a function of the prerift geology, morphology and rheology. The broadly distributed extension of the palaeoPapuan Peninsula included locally focused extension that produced large graben in the southern margin that are now failed rifts (Figs 1 & 9b). Tilted fault blocks overlain by rotated sediment packages in broadly equal-sized graben (SCS line MW9304-30, Fig. 3) show that extension only focused on the axial graben very recently, even though results from ODP Leg 180 (Taylor & Huchon 2002) show that the latter formed early in the rift history (Fig. 4). A broad heat flow anomaly, that is centred in the northern part of the SMG but also covers the NMG, reflects maximum crustal thinning across this axial system (Goodliffe et al. 2000). Today, as evinced by seismicity (Fig. 1), extension has largely ceased on the southern faults, and is concentrated on the shallow-angle north-dipping Moresby Fault (Taylor & Huchon 2002). MCS lines EW95101366 (Fig. 4), EW9510-1369, EW9510-1374 and EW9910-11 (Figs 6 & 7) reveal the final phase of current rifting prior to breakup. MCS lines EW9510-1366 and EW9910-11 show synrift sediments in the NMG dipping to the south owing to slip on the Moresby Fault, whereas the keystone basement block and overlying sediments in the NMG are less rotated on the intervening lines. The stage where extension is primarily accommodated by accretion is represented by MCS line EW9910-13 (Fig. 8). To reach this stage, we envision that Moresby Fault was cut by a dyke swarm that gradually grew in width (Fig. 9d, e). When the first dykes cut through the fault and were emplaced into the overlying synrift sediments, motion on the up-dip portion of Moresby Fault ceased and/or shifted upwards into the sediments. A transfer fault/zone may have formed to connect accretion on spreading segment 1a to normal

faulting on Moresby Fault in the NMG. The NE flank of Moresby Seamount may be such a transfer fault; its NW–SE trend is consistent with small circles about the current Euler Pole (128S, 1448E, 2.4378 Ma21, Fig. 2b). If the width of the intrusive body represented only strain accommodation through accretion, the Moresby Fault would exit the intrusive body at the same height on the northern side as it entered on the southern side. Although there is some ambiguity in the location of the continuation of Moresby Fault to the north of the intrusion, this is clearly not the case (Fig. 8). Although we maintain that the majority of the intrusive body represents strain accommodation through accretion, which formed the associated magnetic anomaly (Fig. 5), a portion of the width may represent overprinting of the original country rock, perhaps through a process similar to stoping. Alternatively, the intrusion may not be continuous, but instead have intervening unresolved slices of country rock between dykes. For the purposes of our schematic model, we have arbitrarily assigned approximately 25% of the intrusion width to these processes (Fig. 9e, f). As the intrusion grew, sediments and possibly sills were added to the centre of the graben. A ring dyke was intruded and reached the sea floor (Fig. 9e) leading to eruptions that produced the extensive lava flows observed (Fig. 2c). Continued extension, vertical propagation of the intrusive body and horizontal growth of the ring dykes will soon construct an edifice at the sea floor on the western (as already exists on the eastern) part of spreading segment 1a. As shown here, the currently forming continent– ocean boundary involves igneous intrusions penetrating the principal (and shallow-angle) normal fault of an asymmetric rift graben. Further east in the Woodlark Basin, breakup along spreading segment 2 occurred within symmetric graben bounded by high-angle faults (Taylor et al. 1999). But in each case the magnetic, bathymetric and seismic data reveal that the boundary between continental rifting and sea-floor spreading in the Woodlark Basin is locally very sharp—notwithstanding that regionally the processes overlap in time. This is in stark contrast to breakup of the Newfoundland and Iberia conjugate margins associated with ultra-slow spreading in the early north Atlantic where transition zones up to 150 km wide separate clearly defined oceanic and continental crust (Tucholke et al. 2004).

Margin stratigraphy and unconformities MCS imaging calibrated by ODP drilling results allow us to observe the stratigraphic details of continental breakup in the western Woodlark Basin. How do these images and results compare to stratigraphic models of continental margin development?

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Falvey (1974) proposed a model to explain typical patterns of margin uplift and subsidence as a function of the thermal, erosional and metamorphic processes attendant on splitting a continent by emplacing and then cooling asthenosphere. He observed that two angular unconformities commonly separate pre-, syn- and post-rift sediments. He proposed that ‘the breakup unconformity is caused by erosion during the final uplift pulse associated with prebreakup upwelling in the mantle’. He recognized that ‘this unconformity is more localized than the rift-onset unconformity, being difficult to define in troughs’ and that ‘continental breakup and the onset of sea-floor spreading is probably not instantaneous’. The existence of two unconformities bounding synrift strata has found widespread application in rifted margin studies (Driscoll et al. 1995) even though, in the particular example Falvey used to illustrate the general model (the Otway Basin of South Australia), the proposed rift-onset unconformity actually correlates with the time of breakup and initiation of marine sedimentation, and the real rift-onset unconformity occurred much earlier (Hegarty et al. 1988). In the case of the western Woodlark Basin, there is a regional rift-onset unconformity. It truncated mafic igneous basement at ODP sites 1109 and 1118, produced a 4 Ma hiatus in the shallow-marine forearc basin sediments at ODP Site 1115 (Figs 3 & 4), and generated the late Miocene ‘unconformity A’ in the Nubiam-1 well and region north of the D’Entrecasteaux Islands (Francis et al. 1987; Taylor & Huchon 2002). The drilling and MCS studies show the reversal of the northward Miocene forearc drainage as the Papuan continent was stretched and thinned, and the northern margin (Woodlark Rise) progressively subsided southward toward the active rifts, from paralic to shelf to the present bathyal water depths (Fig. 4) (Taylor & Huchon 2002). Concurrently, the former landmass to the south was faulted and thinned to produce the graben and reefs of the Pocklington Rise (Fig. 1). Early synrift sediments were subsequently tilted on rotating fault blocks, but later sediments of the southern margin were ponded in half-graben as the early rifts in the south were abandoned, even as deformation continued further north focused on the axial rifts (Figs 3 & 9). Therefore ‘synrift’ sediments (those deposited between rift onset and breakup) can be characterized by rotated sections (with parallel or diverging reflectors) as well as by ponded sections that onlap bounding fault blocks where local extension has ceased. Splayed reflectors representing sediment deposition on rotating fault blocks do not characterize the full synrift section, but only those times and places where extension is localized. Also, in the common case where breakup propagates, the time represented by the synrift sediments and the timing of any breakup unconformity will vary

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along strike: by 6 Ma in the Woodlark Basin. Furthermore, as Falvey (1974) noted, sedimentation in the deep water parts of margins is likely to be conformable across the time of breakup. In the Woodlark Basin, the progression from rifting to breakup and then initial spreading is characterized by a decrease in sedimentation as the margins are progressively thinned, subside below sea level, are eroded less and trap sediments in proximal basins (Figs 3 & 4). The post-rift sedimentation near the continent– ocean boundary is hemipelagic and drapes breakup topography. A distinct breakup unconformity is not observed.

Conclusions † The continental margins asymmetry at the point of current breakup – high-angle normal faults bounding tilted crustal blocks in the south in contrast to a wide tilted margin in the north with only small-offset normal faults – cannot be explained by the observed shallow-angle northdipping detachment between them. Rather, the margins asymmetry is a function of prerift geology, morphology and rheology. Removal of the lower crust through ductile flow is the preferred mode of crustal thinning of the Woodlark Rise, the palaeo-Trobriand forearc. In contrast, on the Pocklington Rise to the south, where the high mountains of the Papuan Peninsula existed 8.4Ma ago, brittle faulting of the upper crust is more prevalent. † Initially, major upper crust normal faulting was distributed from the southern margin to what is now South and North Moresby Graben. Sequentially, the locus of faulting narrowed northwards, leaving large graben (foci of former extension that are now failed rifts) and tilted fault blocks in the southern margin. † At about 1.2 Ma, extension focused at NMG and SMG where rapid subsidence was accompanied by rapid uplift of the intervening Moresby Seamount. This was a precursor to the initiation of sea-floor spreading at approximately 0.8 Ma on segment 1. Extension ceased on the large faults bounding tilted blocks directly to the south. † Sea-floor spreading on segment 1a initiated at approximately 0.08 Ma in the along-strike continuation of the NMG and more specifically through the north-dipping Moresby Fault. Ahead of the spreading tip, extension has focused on the active Moresby Fault (Abers et al. 1997; Taylor & Huchon 2002). Although motion on this normal fault was once controversial owing to its low dip, fault gouge with a low coefficient of friction and evidence of high pore fluid pressures make motion on this fault unremarkable (Taylor & Huchon 2002). Depth conversions of MCS line EW9510-1369 based

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on ODP results and semblance velocities (Goodliffe et al. 1999b) indicate that the true fault dip is 298 and compatible with an earthquake slip plane of 308 –338 (Abers et al. 1997). † The normal faults bounding the northern side of spreading segment 1, co-polar with great circles about the pre-0.52 Ma Euler pole (Fig. 2), are crossed by the Egum Graben. The faults bounding the northern side of Egum Graben form the antithetic faults in the NMG and may be continuous with a major transform fault system that separates the Solomon Sea and Woodlark lithosphere to the east. If so, this would make the current rifting–spreading transition a triple junction between the Trobriand, Woodlark and Australian plates. Alternatively, the fact that the bounding faults of Egum Graben are co-polar with great circles about the present-day Euler pole suggests the possibility that this graben represents a new rifting trend, perhaps reactivating pre-existing weaknesses. † MCS line EW9910-13 records an early evolutionary step in the initiation of a spreading centre. An intrusion that cuts an along-strike continuation of Moresby Fault is connected to the surface via a ring dyke through which lavas have erupted onto the sea floor. Thus, the currently forming continent–ocean boundary involves igneous intrusions penetrating the principal (shallowangle) normal fault of an asymmetric rift graben. The magnetic, bathymetric and seismic data reveal that the boundary between continental rifting and sea-floor spreading in the Woodlark Basin is spatially very sharp – notwithstanding that temporally the processes overlap. † Synrift sediments (those deposited between rift onset and breakup) can be characterized by rotated and/or diverging sections, as well as by ponded sections that onlap bounding fault blocks where locally extension has ceased. A breakup unconformity is not generally observed at the deep-water boundary between rifted continental and accreted oceanic crust, but rather there is a decrease in sedimentation as the margins are progressively thinned, subside below sea level, are eroded less and trap sediments in proximal basins.

Appendix 1: Data and processing The data presented herein come primarily from three marine geophysical surveys conducted between 1993 and 1999 (Goodliffe et al. 1999b) and include singlechannel seismic (SCS) data, multichannel seismic (MCS) data, swath bathymetry, acoustic imagery and magnetic data. Generic mapping tools (GMT) (Wessel & Smith 1995) was used comprehensively throughout this study both for display and analysis.

Bathymetry and acoustic imagery data HAWAII-MR1 swath bathymetry and sidescan data from a 1993 R/V Moana Wave survey cover the western Woodlark Basin from 1508 500 E to 1548 450 E. Hydrosweep multibeam bathymetry data from R/V Maurice Ewing cruises in 1995 and 1999 provided additional detail in the region of Moresby Seamount. Beyond the coverage of the swath datasets, wide-beam profiler data from the National Geophysical Data Center (NGDC) database and digitized soundings from Australian Navy charts were incorporated. Land topographic data are from the Shuttle Radar Topography Mission (SRTM) 90 m grid cell dataset. Final bathymetry grids were produced with a 0.0018 grid size.

Magnetic data Magnetic data were collected along north–south 5 nm-spaced tracks during the regional 1993 R/V Moana Wave survey and during R/V Maurice Ewing surveys in 1995, 1999 and 2000 in the vicinity of Moresby Seamount. As the magnetic anomalies formed along an east–west spreading centre are highly skewed, the shipboard magnetic and bathymetry data were gridded at 0.018 and used to derive a solution for the intensity of magnetization of the sea floor following the technique of Macdonald et al. (1980). The magnetization inversion assumes a 1 km-thick source layer conforming to the sea floor. As this solution produced a reversal pattern in qualitative agreement with that from the total field anomalies, no annihilator (Parker & Huestis 1974) was added to the solution. Areas shallower than 1.2 km were masked to satisfy the requirement that the magnetic source does not come within one grid cell of the observation surface (Parker & Huestis 1974). Although the resulting magnetization map can be interpreted directly using the geomagnetic polarity timescale (Cande & Kent 1995), the oceanic crust in the region of interest is mostly formed during the Brunhes chron.

Marine seismic data The marine seismic data used in this study comes from three surveys. In 1993 six-channel seismic (SCS) data were collected during an R/V Moana Wave survey in the Western Woodlark Basin using a 150 m-streamer and a 120 cubic inch (inch3) airgun source. Shots spaced at 37.5 m yielded twofold 12.5 m-spaced CMPs. A 1995 R/V Maurice Ewing survey included a site survey for ODP Leg 180 (Taylor et al. 1996; Goodliffe et al. 1999b). To better image the Moresby Fault and document the subsidence of the northern margin of the Woodlark Basin, a 196-channel 4.9 km-streamer was deployed. The first line, designed for deepest penetration, was shot using a tuned 8460 inch3 20 airgun array. Shots every 20 s yielded a shot spacing of approximately 50 m and 48-fold 12.5 m-spaced CMPs, a record length of

BOUNDARY BETWEEN RIFTING AND SPREADING 16.384 s and a sample interval of 4 ms. Additional lines were shot using 10 airguns alternating in the same 20 airgun array. Shots every 12 s yielded a shot spacing of approximately 25 m and 60–75-fold 12.5 m-spaced CMPs with a record length of 9.116 s and a sample interval of 4 ms. In 1999 the R/V Maurice Ewing shot 13 dip lines oriented NNE– SSW, perpendicular to the strike of the fault (Taylor et al. 1999) and one strike line. Eleven of the dip lines, spaced 200 m apart, crossed Moresby Graben in the vicinity of ODP Site 1108. The two remaining dip lines crossed the western tip of the spreading centre to the east. Using a 48-channel 1.2 km-streamer, a 1395 inch3 tuned six airgun array was fired every 25 m, giving a 24-fold CMP interval of 12.5 m. The record length was 8192 ms and sample interval was 2 ms. The marine seismic data were processed using ProMAX. Data processing included trace editing, filtering, prestack and post-stack deconvolution, velocity analysis, inside and outside mutes, normal moveout corrections, stack and migration. The multichannel seismic (MCS) data presented herein have also undergone partial prestack migration.

Appendix 2: Velocity model and depth conversion Four sites were drilled and logged in the synrift sedimentary section. At Site 1108, sampled to a depth of 480 m, density and porosity measurements showed that 385 m of sediments have been removed from the top of the section through faulting and/or erosion (Shipboard Scientific Party 1999b). The termination of reflectors at the sea floor south of SP420 on MCS line EW9510-1374 and south of SP400 on MCS line EW9510-1369 (Fig. 6) confirms that this is a reasonable interpretation and any synrift sediment velocity model must take this into account. To the north of the NMG logging and laboratory measurements of velocity were obtained at sites 1109, 1115 and 1118. Depth v. velocity curves were derived for each of the three sites, refined using vertical seismic profile checkshots, and verified through the generation of synthetic seismograms (Goodliffe et al. 2001). Site 1118, on the margin just north of the NMG (Figs 2 & 4), sampled 860 m of synrift sediments. Numerous MCS lines (Goodliffe et al. 1999b) show that the sediments sampled on the northern margin can be correlated with those in NMG close to MCS lines EW9510-1369 and EW9510-1374. Here, the offset across the northern bounding fault is small, enabling the velocity– depth curve for Site 1118 to be used in the graben. As we only have synrift sediment velocity information to a depth of 801 m, the best estimate of interval velocity for deeper synrift sediments in the graben comes from semblance analysis on line EW9510-1369 where the strata are continuous and subhorizontal. The composite interval velocity curve predicts a velocity of approximately 3.5 km s21 at about 2.5 km below the sea floor (the approximate sediment thickness in the graben). Shipboard laboratory measurements of the basement

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diabase consistently gave velocities of approximately 6 km s21 (Shipboard Scientific Party 1999a). As fracture porosity of this lithology is often high, we assigned a velocity of 4.5 km s21 to the basement at the base of the synrift sediments and applied a gradient such that the velocity would be 6 km s21 at 8 s TWT directly below the centre of the graben. Extrapolating our one-dimensional velocity model to two dimensions in most places entailed hanging the sediment and basement velocity functions from the sea floor and basement, respectively. However, south of SP400, where reflectors terminate at the sea floor, the composite velocity curve was suspended at 3.54 s TWT, effectively increasing the sediment velocity at the water–sediment interface and closely matching the velocities of the sediments recorded at Site 1108. The water column was assigned an average velocity of 1494 m s21.

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Nature of the continent – ocean transition zone along the southern Australian continental margin: a comparison of the Naturaliste Plateau, SW Australia, and the central Great Australian Bight sectors N. G. DIREEN1, I. BORISSOVA2, H. M. J. STAGG2, J. B. COLWELL2 & P. A. SYMONDS2 1

Continental Evolution Research Group, University of Adelaide, Adelaide, SA 5005, Australia (e-mail: [email protected]) 2

Geoscience Australia, Canberra, ACT 2601, Australia

Abstract: We document the interpretation of three crustal sections from coincident deep seismic reflection, gravity and magnetic data acquired on Australia’s southern margin: one section from the Naturaliste Plateau and the Diamantina Zone; and two in the Great Australian Bight (GAB). Interpretations are based on an integrated study of deep multichannel seismic, gravity and magnetic data, together with sparse sonobuoy and dredging information. All interpreted sections of the margin show a transition from thinned continental crust, through a wide continent ocean transition zone (COTZ). In the GAB the transition is to slow sea-floor spreading oceanic crust that dates from breakup in the Campanian (c. 83 Ma); in the Naturaliste–Diamantina margin the earliest oceanic crust is undated. The COTZ on these margins is geologically and geophysically complex, but interpretation of all data, including dredge hauls, is consistent with the presence of a mixture of modified continental lower crust, breakup related volcanics and exhumed continental mantle. Serpentinized detachment faults are not well imaged, but have been inferred from high-amplitude magnetic signatures interpreted to arise from magnetite associated with the hydration of peridotites. Alternative models for the structure of the COTZ, involving either mafic underplating or aborted sea-floor spreading, have been explored, but are considered unlikely on this margin. Similarity in the final architecture of these margins has major implications for the nature of rifting in the Southern Rift System, and may point to the entire 4000 km-long system being non-volcanic in character. Second-order differences in geometry and morphology of the two areas studied are unlikely to be a function of strain rate. Instead, they probably reflect complexities owing to the multiple tectonic events that occurred during final Gondwanide fragmentation. The most dramatic of these is the impact of hotspot activity in the Kerguelen Plateau, which commenced some 50 Ma prior to final breakup in that sector.

Non-volcanic, or magma-poor, continental margins are defined by features such as a wide, largely amagmatic, continent–ocean transition zone (COTZ – defined below); unroofed mantle peridotite rocks exposed at the palaeo sea floor; relatively flat-lying, serpentinized detachment faults; typically slow (55– 20 mm year21) to ultra-slow (20– 1 mm year21) full spreading rates upon initiation of sea-floor spreading; and low volumes of magmatic rocks (Boillot & Froitzheim 2001; Whitmarsh et al. 2001). These type of margins have been documented in relatively few places around the globe. The principal example is off Iberia (Pickup et al. 1996; Whitmarsh et al. 2001) and its Newfoundland conjugate (Hopper et al. 2004). Other examples have been described in the Labrador Sea (Chian &

Louden 1994; Chalmers & Pulvertaft 2001), Woodlark Basin (Robertson et al. 2001), Great Australian Bight (Sayers et al. 2001) and East Antarctica (Colwell et al. 2006). Of these examples, the few studied with integrated geophysical data – usually deep multichannel seismic reflection (MCS), seismic refraction, gravity and magnetics – exhibit significant geometric and geophysical complexity both across and along the margin, and especially in the COTZ. We follow Colwell et al. (2006) in defining two terms that are often used confusingly in studies of rifted continental margins. † Continent –ocean transition zone (COTZ): a region on the continental margin that lies between the outboard edge of highly attenuated,

From: KARNER , G. D., MANATSCHAL , G. & PINHEIRO , L. M. (eds) Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 239– 263. DOI: 10.1144/SP282.12 0305-8719/07/$15.00 # The Geological Society of London 2007.

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unequivocal continental crust and the inboard edge of unequivocal oceanic crust. The COT includes both sedimentary and magmatic components in proportions that vary both along and across the margin, and may include areas of failed sea-floor spreading. † Continent–ocean boundary (COB): this marks the inboard edge of unequivocal oceanic crust. The complexity of non-volcanic margins includes exposure of magnetized (Zhao 2001) exhumed continental mantle peridotite–serpentinite complexes, and volcanic rocks at or near the sea floor (Abe 2001; Desmurs et al. 2001; Hebert et al. 2001) within the COTZ. Exhumation of serpentinized peridotite along extensional fault footwalls may lead to the creation of linear magnetic anomalies that may be misinterpreted as sea-floor spreading anomalies, as shown by Sayers et al. (2001). Another documented complexity of ‘amagmatic’ margins is the capping of attenuated continental crust in the COTZ by thick, magnetized basalt piles sourced from the locus of break-up (Colwell et al. 2006). These can, again, lead to the creation of apparently linear magnetic anomalies that may be mistaken for those associated with seafloor spreading. These kinds of complexities mean that there is potential for the outer parts of some non-volcanic margins to have remained unrecognized in the geological record – especially those where the distribution of continental and oceanic crust has been interpreted only on the basis of potential field data. This possible bias in interpretations may mean that the percentage of continental margins classified as non-volcanic margins may have been underestimated. This, in turn, could lead to a diminution of the importance of the mechanical processes that dominate non-volcanic margins in many recent generalized models for rifting and breakup. In this paper we present a new integrated interpretation of deep MCS, gravity and magnetic data across the southern margin of the Naturaliste Plateau and adjacent Diamantina Zone (the Naturaliste –Diamantina margin), and a new interpretation of data from the central Great Australian Bight margin. Both of these margins form part of the Mesozoic Southern Rift System (SRS: Stagg et al. 1999) and exhibit the characteristics of nonvolcanic rifted margins, including a wide COTZ (Sayers et al. 2001). The purpose of this study is to evaluate the degree of along-strike complexity within a single non-volcanic rifted margin system, at scales commensurate with the evolving palaeoplate boundary. This complexity, and its distribution, form important constraints on evaluating ranges of parameters

that are incorporated into thermo-mechanical models for continental rifting and breakup.

Study area The western and southern margins of Australia (Fig. 1) are two arms of a triple junction that formed during the final stages of fragmentation of Gondwana. Following a series of Palaeozoic –Mesozoic terrane-calving events (Metcalfe 1996), Greater India started to break away from Australia– Antarctica in what is now the Argo Abyssal Plain (AAP), commencing in the latest Middle Jurassic (Callovian) (Markl 1978a; Mihut & Mueller 1998). The rupture then propagated southwards through the Gascoyne, Cuvier and Perth margins during the Early Cretaceous (Valanginian– Hauterivian) (Markl 1978a, b; Veevers & Li 1991; Symonds et al. 1998). This process left an evolving free plate boundary, with Australia – Antarctica fringed by a new margin. The northern sectors (AAP to the Wallaby–Zenith Fracture Zone) of this system are largely volcanic in character (Symonds et al. 1998). The character of the Perth AP margin, south of the Wallaby–Zenith Fracture Zone, is not well constrained (Bradshaw et al. 2003), but may be volcanic on the basis of distribution of onshore volcanic rocks such as the Bunbury Basalts (Coffin et al. 2002). From the Middle Jurassic onwards, during rifting and breakup of Greater India from Australia– Antarctica, the proto-Southern Rift System was also undergoing mechanical extension and synrift sequences were deposited (Norvick & Smith 2001). At the eastern end of the Southern Rift System, meridional rifting between Lord Howe Rise and eastern Australia was underway with final breakup of this margin taking place at approximately 85 Ma (Santonian) (Gaina et al. 1998). Continued rifting of the Australian –Antarctic landmass led to eventual breakup between them at approximately 83– 80 Ma (Campanian) (Tikku & Cande 1999; Sayers et al. 2001) in the central Great Australian Bight, approximately contemporaneous with breakup on the eastern Australian margin. Whatever the ultimate cause, breakup along Australia’s southern margin appears to have occurred owing to ridge propagation from west to east along the Southern Rift System (Mutter et al. 1985), but with very slow initial oceanic spreading (Cande & Mutter 1982; Sayers et al. 2001). In the central Great Australian Bight (GAB) and conjugate Wilkes Land sector, breakup produced a complex non-volcanic margin system, marked by a wide COTZ with distinctive conjugate basement ridges (Sayers et al. 2001; Colwell et al. 2006).

COTZ ALONG AUSTRALIA’S SOUTHERN MARGIN

241

120°E 100° E

CA WZ

P

140°

FZ

E

150

E 80° S 20°

°E

PAP BRd

Figur e 2.

7

99/0 GA1

GAB

99/0

DZ

OB

GA -1

GA-187/01

S

40°

5

NP

SB

S I R

KP S

60°

BR WL

AL S

70°

0

1000 km

05-2

93-1

Fig. 1. Regional setting for this paper with the study area shown by the box and seismic lines annotated. Abbreviations: AL, Ade´lie Land; BR, Bruce Rise; BRd, Broken Ridge; CAP, Cuvier Abyssal Plain; DZ, Diamantina Zone; GAB, Great Australian Bight; KP, Kerguelen Plateau; NP, Naturaliste Plateau; OB, Otway Basin; PAP, Perth Abyssal Plain; SB, Sorell Basin; SIR, Southeast Indian Ridge; WL, Wilkes Land; WZFZ, Wallaby– Zenith Fracture Zone. Fine contour, 3 km-isobath; dashed lines are oceanic fracture zones.

Breakup in this sector also appears to have resulted in two broadly symmetrical margins (Colwell et al. 2006). These features are interpreted to comprise exhumed serpentinized mantle and magmatic products from decompression melting of the upper continental mantle (Sayers et al. 2001). Farther east, in the Duntroon–Otway and Sorell basins, rifting kinematics became very oblique to the east– west axis of spreading (Willcox 1990; Willcox & Stagg 1990). Sea-floor spreading magnetic anomalies are thus very difficult to recognize in this sector (Roots et al. 1985; Finlayson et al. 1998). The Naturaliste Plateau (NP) (Figs 1 & 2) is a large, rectilinear, continental margin plateau off the SW coast of Western Australia. It sits at the locus of intersection of the mainly Early Cretaceous volcanic margin of Western Australia, and the Late Cretaceous, largely non-volcanic margin of southern Australia. Its nature and possible origins have been investigated by (Petkovic 1975a, b; Jongsma & Petkovic 1977; Nicholls et al. 1981; Munschy 1998; Beslier et al. 2001), and its possible provenance connections to onshore geological features explored in Collins (2003). The NP lies in moderately deep (.2 km) water, and incorporates sedimentary sequences that probably range in age from at least Late Jurassic to Recent; however only Late Cretaceous and younger successions have been recovered so far by

very limited Deep Sea Drilling Project (DSDP) drilling (Hayes et al. 1975). Since the 1970s, there has been debate about whether the basement to these sedimentary sequences is composed of continental (Petkovic 1975a); or oceanic crust (Coleman et al. 1982). Recent dredging on the SW margin of the plateau has recovered Cambrian-age continental rocks (Beslier et al. 2004). In addition to the ambiguous nature of the NP itself, seismic reflection transects and dredging south of the plateau (Chatin et al. 1998; Munschy 1998; Beslier et al. 2001, 2004) have identified a zone of anomalous ocean floor – the Diamantina Zone (DZ) (Figs 1 & 2). The DZ is a wide (c. 200 km), anomalous bathymetric and gravity feature, which trends generally eastwards to the GAB in the east (Fig. 1), where it becomes increasingly buried by sediment. The DZ has highly variable topography, with bathyal depressions and ridges with relief of up to 1000 m over distances of less than 10 km. Dredging within the DZ (Chatin et al. 1998; Beslier et al. 2004) recovered samples of peridotite that they interpreted to have distinctive continental lithospheric affinities, indicating that at least part of this anomalous zone is unroofed continental mantle, exposed at the sea floor. Although these recent discoveries have added to our knowledge of this region, there is to date no reliable, interpreted crustal section through the

242

N. G. DIREEN ET AL. 120°E

130°E

110°E

PB

106°E 30°S

So uth

We st e r n A u s t r a l ia

PAP

A u s tr

257

136°E

al ia

GC ELT55-DR12

BB LC

258 NF

NP

Z

4

264

DIAM

ANT

MD80-DR3

INA

MD110-DR06

MD110-DR08

ZON

MD110-DR01

MD110-DR05

MD110-DR02

MD110-DR04

MD110-DR11

E

MD110-DR03

1 GA-187/0

40°S

0

4

PoB

GAB

MD110-DR07

MD110-DR10

MD110-DR09

0.2 1

GA-1 99/0 5 GA-1 99/0 7

MD110-DR11

2 3

Bathymetry (km)

05-293

500 km

Surveys 187 & 199

-2

DSDP site

Dredge station

Fig. 2. Study area for this paper; location shown in Figure 1. Bathymetry in km, with 5 km isobath deleted because of extreme complexity of the Diamantina Zone. Also shown are the locations of seismic profiles and potential field models discussed in this paper, dredge stations (listed in Table 1) and DSDP sites. Abbreviations: BB, Bunbury Basalts; GAB, Great Australian Bight; GC, Gawler Craton; LC, Leeuwin Complex; NFZ, Naturaliste Fracture Zone; NP, Naturaliste Plateau; PAP, Perth Abyssal Plain; PB, (onshore) Perth Basin; PoB, Polda Basin. The complexity of the bathymetry in the Diamantina Zone is reflected in the seismic profile illustrated in Figure 3.

Naturaliste Plateau and its southern flanking abyssal domain(s). Because the nature of the deep crust and mantle near the NP remains enigmatic, there is not yet a self-consistent, well-tested model for how this feature relates to Gondwana breakup in this region, nor a clear genetic link to either the older Western Australian margin to the north or to the southern Australian margin along strike to the east. A further complication in the study of the fragments related to Cretaceous Gondwana breakup is the presence of overprinting volcanics related to the Kerguelen plume. Volcanics from the Kerguelen Plateau have been dated as Barremian–Albian (c. 118 –100 Ma) (Coffin et al. 2002). Geochemical arguments have been used to correlate these volcanics with the Bunbury Basalts on the SW Australian mainland (Coleman et al. 1982; Ingle et al. 2002) (Fig. 2), just inshore of the Naturaliste Plateau. The latter are dated between approximately 132 and 123 Ma (Valanginian –Hauterivian) (Coffin et al. 2002). Basalts of the Rajmahal Traps (c. 118 Ma) are also correlated with the Kerguelen plume activity on geochemical grounds (Coffin et al. 2002; Ingle et al. 2002). The timing of breakup to the west and north of the Naturaliste Plateau is uncertain. It has tentatively been interpreted as Valanginian (c. 140 Ma) on the basis of a pronounced regional unconformity in regional seismic data and drilled within the southern Perth Basin immediately onshore to the east of the NP (Bradshaw et al. 2003). The oldest tentative sea-floor spreading magnetic anomalies identified

in the Perth AP to the north of the NP are M10 N (Hauterivian; c. 134 Ma: Cande et al. 1989) or M11 (Markl 1978a), which tends to support a Valanginian breakup age. The age of breakup to the south of the Naturaliste Plateau is difficult to constrain. Linear magnetic anomalies recognized directly south of the Naturaliste Plateau and Diamantina Zone begin at anomaly 20, in oceanic crust of Middle Eocene age (Tikku & Cande 1999). This corresponds to a period of fast spreading in the Southeast Indian Ridge System that commenced elsewhere at anomaly 22 time (Tikku & Cande 1999). Older magnetic anomalies have tentatively been identified several hundred kilometres SE of the Naturaliste Plateau, south of the Leeuwin Complex (Fig. 2). These anomalies are inferred to be responses of ultra-slow spreading oceanic crust of chron 34 age and younger (Mutter et al. 1985; Tikku & Cande 1999).

Previous studies of the Naturaliste Plateau Lamont Doherty Earth Observatory conducted early seismic reflection surveys of the NP using the Robert D Conrad and Eltanin (Burkle et al. 1967; Eltanin 1978). These results were augmented by Geoscience Australia (then BMR) Continental Margins Survey 18, which acquired a regular grid of fair-quality, six-channel seismic lines, and six

COTZ ALONG AUSTRALIA’S SOUTHERN MARGIN

moderate-quality, 24-channel seismic lines recorded by the Shell Development vessel Petrel (Boeuf & Doust 1975). Most of these early data were acquired over the Naturaliste Plateau proper, with only limited coverage of the plateau flanks and adjacent deep-water basins. In 1997 Geoscience Australia survey 187 recorded a single, good-quality seismic line from the fast-spreading, Eocene-age oceanic crust, northwards across the Diamantina Zone (DZ) to the crest of the Naturaliste Plateau (Borissova, 2002) (Figs 1 & 2). Limited single-channel seismic reflection data over the DZ was also acquired by the French vessel Marion Dufresne in 1994 (Munschy 1998). There is very little direct geological information about the nature of the basement rocks of the Naturaliste –Diamantina margin. Two DSDP wells were drilled into the Naturaliste Plateau in 1972: DSDP sites 258 and 264 (Hayes et al. 1975), but neither hole reached basement, bottoming in pre-Cenomanian volcaniclastic conglomerate and middle –late Albian terrigenous detrital clay and glauconitic sandstone, respectively. Two basement dredges, one on the NW margin of the NP and one on the southern margin, were recovered by the Eltanin and by Marion Dufresne cruise 80, respectively (Fig. 2). The Eltanin dredge recovered manganiferous-coated sediments and conglomerate. The Marion Dufresne dredge from the southern scarp of the plateau recovered 77 kg of mixed high

243

metamorphic grade gneisses, granite, diorite and prerift (480 Ma) gabbros (Beslier et al. 2001, 2004). These were interpreted as being evidence of continental basement beneath the NP, in accord with the previous interpretation of (Petkovic 1975b). A number of dredge hauls within the DZ were also recovered during Marion Dufresne cruise 110, and by the Diamantina. These sites are shown in Figure 2, and the dredge hauls are summarized in Table 1. These dredging results have been variously interpreted by Nicholls et al. (1981), Coleman et al. (1982), Chatin et al. (1998), Munschy (1998) and Beslier et al. (2001). Nicholls et al. (1981) interpreted the single 25 kg dredge haul by the Diamantina in 1977, from a seamount in the DZ. The rocks consisted of a mixture of lherzolites, clinopyroxenites, subsidiary serpentinites and minor breccia comprising fragments of the above, plus a ‘granitic’ fragmental assemblage (qz –ksp– plag–bt –hbl). These rocks were interpreted by Nicholls et al. (1981) as similar to rock suites dredged from the Iberian nonvolcanic margin (Boillot et al. 1980) – now interpreted as serpentinized peridotite allochthons and detachment faults (Pe´rez-Gussinye´ et al. 2001; Whitmarsh et al. 2001). Coleman et al. (1982) interpreted dredged material from the Eltanin, which was taken from the northern flank of the NP. They correlated these highly altered conglomeratic cobbles as

Table 1. Summary of dredge hauls from the Naturaliste Plateau (after Borissova 2002) Survey

Year

Dredge

Latitude

Longitude

Description of Haul

EL55

1972

DR12

232.9100

110.9700

DIM

1977

DR11

236.6333

112.2666

MD110

1998

DR01

237.0600

119.1500

MD110

1998

DR02

237.6160

119.9000

MD110 MD110

1998 1998

DR03 DR04

237.3390 237.3483

118.2288 117.8166

MD110 MD110 MD110 MD110 MD110 MD110 MD110

1998 1998 1998 1998 1998 1998 1998

DR05 DR06 DR07 DR08 DR09 DR10 DR11

237.0465 236.3000 235.6666 236.3938 236.6167 236.0116 235.0933

117.0788 115.9667 115.0500 112.2338 112.2417 111.4132 110.3500

MD80

1994

DR3

236.5000

106.9000

MD80 MD80

1994 1994

DR5 DR7

234.3000 234.3000

103.1333 103.1333

Slabs of manganese mat with conglomerate Blocks of ultra-mafic rocks covered by manganese crust Dolerite, basalts (fresh), gabbro, sediments, cherts and nodules Peridotite, lherzolite, gabbro, pyroxenite, basalt Basalt, dolerite, gabbro Olivine-phyric basalt and plagioclase-phyric basalt (Vesicular) basalt, manganese crusts Lherzolite, harzburgite, basalt Gneiss and granite (Vesicular) basalt Peridotite, gabbro, harzburgite Basalt Metamorphic rocks, granite, diorite, gabbro, basalt Highly altered peridotites in a manganiferous matrix Slightly porphyritic alkaline basalts Slightly porphyritic alkaline basalts

244

N. G. DIREEN ET AL.

redeposited material from a protolith similar to the Bunbury Basalt. Coleman et al. (1982) raised the possibility that, contrary to previous interpretations of the NP as a continental plateau (Heezen & Tharp 1973; Petkovic 1975b), the origin of these rocks did not preclude a plateau of oceanic magmatic affinity. Chatin et al. (1998) interpreted alkaline basalts dredged from the inner part of the DZ and highly altered peridotites from its outer part. The alkaline basalts were recovered from the top of a large tilted block, and are similar in composition to some basalts drilled on the Kerguelen Plateau and, particularly, to the basalts from the Kerguelen Islands themselves. Because alkaline basalts never occur at mid-ocean ridges, Chatin et al. (1998) concluded that the DZ is not underlain by typical oceanic crust. Peridotites were also recovered at four Marion Dufresne dredge sites in the DZ on igneous basement ridges (Chatin et al. 1998; Royer & Beslier 1998). To the west of the Naturaliste Fracture Zone (1108E) the geochemical signature of the ridge rocks reflects very small degrees of partial melting, indicating rapid mantle exhumation (Chatin et al. 1998). Farther to the east, dredged peridotites exhibit even more similarity with peridotites sampled from the Iberian margin, with intrusive textures overprinted by high-temperature mylonitic and porphyroclastic fabrics (Beslier et al. 2004). These new sampling results indicate extreme crustal thinning in the Diamantina Zone. Beslier et al. (2001) estimated that emplacement of the peridotites in the Diamantina Zone took place between 90 and 84 Ma, on the basis of 40 Ar/39Ar ages on magmatic amphiboles and biotites from these rocks. In this paper we investigate the two alternatives (continental versus oceanic) for the affinity of the NP basement, and examine the structure of the DZ immediately to the south, by integrating information from dredging results with deep seismic reflection and potential field data. We refer to this broad margin as the Naturaliste – Diamantina margin.

Data Geophysical data from the Naturaliste –Diamantina margin used in this study come from Geoscience Australia survey 187. This consists of coincident shipboard free air gravity and magnetics, together with deep (16 s TWT record length) multichannel seismic reflection data, recorded along a single, 438 km-long, north–south line (line GA-187/01; Fig. 3). The seismic data are 40-fold, and were acquired with a 4500 cubic inch airgun array and

a 320-channel, 4000 m-streamer. Navigation for the survey was by differential GPS positioning. Owing to the absence of sonobuoy or OBS data from this area, velocity models for migration and depth conversion were developed from smoothed stacking velocities. Stacking velocities from the acquired data were picked at 4 km intervals, and an accurate water bottom (WB) was picked from TWT profiles, based on the stacked time migrated data. A time-variant smoothing function was then applied to the velocity picks from the WB down. The smoothing function was designed by determining the number of velocity functions to include in the smoothed velocity function (centre point of the weights) and the weight applied to each velocity. The larger the weight, the more bias applied by that velocity to the final result. The final smoothing function applied is described in the Appendix. Because of the generally uniform high velocities in the deeper parts of the section, the degree of vertical and horizontal resolution decreases with depth (Fig. 3), and the errors in depth conversion and migration increase. This means that the deeper parts of the seismic section are only a guide to the depths, which have been further assessed using potential field data and modelling. The gravity data modelled are 1967 Geodetic Reference System Free Air Anomalies calculated at mean sea level, derived from shipboard gravity meter measurements. These have been drift and noise filtered, and Eo¨tvos corrected. The final RMS noise envelope of the data is estimated to be approximately 1 mGal. The magnetic data are residual anomalies (i.e. with the International Geomagnetic Reference Field removed), recorded from a surface-towed magnetometer. Magnetic anomaly data have been corrected for the magnetometer-ship offset. No diurnal corrections have been applied, but the data have been filtered using a three-point de-spiking filter.

Results In order to examine the crustal architecture of these problematic but crucial areas of the Southern Rift System, we have created a 2.5D integrated forward model of the top 30 km of crust along the deep-seismic reflection transect GA-187/01 from the central Naturaliste Plateau, across the Diamantina Zone, to the Australian–Antarctic Basin (Fig. 4a).

Seismic interpretation Line GA-187/01 has been divided into a series of morphotectonic domains (Fig. 3). These comprise

Fig. 3. Uninterpreted and interpreted seismic profile GA-187/01 from the Naturaliste Plateau and Diamantina Zone, showing the major tectonic elements on this part of the margin. Also shown are the projected locations of dredge stations MD110-DR10 and MD110-DR11 (details in Table 1). Boxes show the locations of detailed seismic sections in Figure 5.

Pressure (MPa)

Magnetic anomaly (nT)

FAA (mg)al

(a)

50 25 0 –25 –50 –75

Observed

Model

750 500 250 0 –250 –500

Model

Observed

750 740 730 720 710

SP 2000

3000

4000

5000

6000

7000

8000

0

G2, G3

A

5

R2 Layer 2A

E2

B

B

Depth (km)

F2

E1

Layer 2B M

10

F1

A

D

G4

R1

G1

G5

SM

15

Moho

20

Mantle

0

50 km

Line GA-187/01

(b)

0 –20

FAA (mgal)

25

A

C

B

R3

G6 Layer 3

9000 C

Observed

–40

Model

–60

Magnetic anomaly (nT)

–80 Model

100 0

Observed

–100 SP 2000

–200 0

4000

5000

6000 C

A

Layer 2A

5 Depth (km)

3000

E1

U2

I2

B

A

T U1 H T

Layer 2B I1

10

I3

Layer 3

R2

M

U3

BI.W R3

R1

SM1

15

Br.W

GC SM5

S

SM4

SM2 SM3

20

Magnetic anomaly (nT)

(c)

FAA (mgal)

25

Mantle

0

o Moh

50 km

Line GA-199/05

0 –20 Observed

–40 –60

Model

–100 0

Model

–100

Observed

SP 2000

3000

4000

5000

0

A B

5

Layer 2A

C

A

E1

E2

Depth (km)

H

10

I2

I1

Layer 2B

T

U

Layer 3

M

BI.W

R

Br.W

S

15

SM1

M o

oh

Mantle

20 25

GC

SM2

0

Line GA-199/07

SM3

50 km

05-293-7

Fig. 4. Reverse of A3 foldout (a) Potential field model for line GA 187/01. Top panel observed and calculated Free Air gravity anomaly (mGal); middle panel observed and calculated residual IGRF corrected magnetic anomaly (nT); bottom panel, lithostatic loading curve (MPa). (b) Potential field model for line GA 199/05, panels as above, without lithostatic loading curve. (c) Potential field model for line GA 199/07, panels as above without lithostatic loading curve. The properties of all model bodies are listed in Table 2.z

2.10 2.10 2.10 2.10 2.10 2.20

2.52

2.92 2.78 2.60

2.60 2.65

2.65 2.65 2.60 2.70

D

E1 E2 F1

F2 G1

G2 G3 G4 G5

2.10

1.03 2.00 2.10

r (103 kg m23)

0.006/1.3/N 0.08/1.5/R 0 0.045/1.2/R

0.07/1.8/N 0

0.08/1.8/R 0.01 0.04/1.5/N

0

0.03/3/R 0.09/3.6/R 0.08/2.7/R 0.01/1/R 0.07/3/R 0.06 /3/R

0.06/2/R

0 0 0

k (SI)/ Q/polarity

Naturaliste–Diamantina (GA-187/01)

(from N– S)

C

Sea water A B

Body

0.06/2.5/R 0.03/1.6/R 0.07/2/R 0.04/2/R

Br W 2.45 U1 2.67 U2 2.71 U3 2.75 2.60

M 2.85

0.035/2/R

–

0.01/1.8/R 0.02/2/R

T: 2.35 Bl.W 2.46

GC 2.80

0

0

0 0

H: 2.12

2.00

1.03 1.60 1.84

k (SI)/ Q/polarity

Central GAB (GA-199/05) r (103 kg m23)

Table 2. Properties of modelled bodies in Figure 4

M 2.88

GC 2.80

2.60 2.50

U 2.71

Br. W 2.48

T: 2.42 Bl.W 2.50

H: 2.12

2.12

1.03 1.60 1.85

r (103 kg m23)

0.03/1.5/R

0.03/1/N

0.04/2/R 0.03/3/R

0.02/2/R

0.08/4/R

0.03/2/R 0.02/2/R

0

0

0 0 0

k (SI)/ Q/polarity

Central GAB (GA-199/07)

(Continued)

Proterozoic basement of variable age and evolution. M are continental rafts modified by later intrusion

Altered stacked volcanics and volcaniclastics

Hammerhead Supersequence in GAB Tiger Supersequence in GAB Blue Whale Supersequence in GAB Bronze Whaler Supersequence in GAB Tilt blocks of upper crustal sequences intruded by mafic sills (Altered) mafic volcanics

Dugong Supersequence in GAB Upper Hammerhead Supersequence in GAB Wobbegong Supersequence in GAB

Comments

COTZ ALONG AUSTRALIA’S SOUTHERN MARGIN 245

2.72 2.80

2.90 2.90

2.70

2.90

2.55

2.60

2.85 3.30

R2 R3

M

SM

Layer 2A

Layer 2B

Layer 3 Mantle

3

r (10 kg m23)

0 0

0.04–0.08/ 2 – 3/R 0

0.02

–

0.07 0.07/1.2/R

0.02/2/R 0.045/1.3/R

k (SI)/ Q/polarity

Naturaliste– Diamantina (GA-187/01)

G6 R1

Body

Table 2. Continued

0.01/1.5/N 0.01/1/N 0.04/1.2/R 0.04/1.6/R 0.004– 0.05/ 1 – 2.6/N&R 0 – 0.03 N&R

SM2 3.00 SM3 3.05 SM4 3.25 SM5 3.25 2.47 – 2.55, average 2.50 2.75

0 0

0.04/1.5/R 0.04/1.3/N 0.02/2/R

I2 2.71 I3 2.67 SM1 2.90

2.85 3.30

0.06/2/R 0.03/1.3/N 0.03/1/R 0.01/1/R

0.008/1/N

k (SI)/ Q/polarity

R2 2.73 R3 2.83 S 2.75 I1 2.70

R1 2.85

r (103 kg m23)

Central GAB (GA-199/05)

2.90 3.30

2.72 – 2.75

2.47 – 2.50

SM2 3.25 SM3 3.25

SM1 3.05

I2 2.70

S 3.00 I1 2.67

R 3.05

r (103 kg m23)

0 0

0.001– 0.06/ 1 – 4/N&R 0.008– 0.04/ N&R

0.04/2/N 0.027/2/N

0.03/1.7/N

0.022/1/N

0.01/1/N 0.08/1/N

0.02/1/N

k (SI)/ Q/polarity

Central GAB (GA-199/07)

Variably faulted and altered basaltic layer Sheeted dyke layer. Variable densities owing to alteration on extensional faults (e.g. Sayers et al. 2001)

Altered and/or partially melted upper mantle

?Serpentinite Migmatized continental crust or crust heavily intruded by dykes

Variably serpentinized peridotite, gabbro and basalt basement ridges

Comments

246 N. G. DIREEN ET AL.

COTZ ALONG AUSTRALIA’S SOUTHERN MARGIN

the Naturaliste Plateau (sensu stricto) at depths of 2200–3000 m, its southern slope and the COTZ as defined above. The total width of the COTZ is about 250 km. The COTZ is further subdivided into an inner COTZ, and an outer COTZ that incorporates the Diamantina Zone. The southern end of the line terminates on the rugged, fast-spreading oceanic crust of the Australian–Antarctic Basin. Seismic details of the principal morphotectonic domains are shown in Figure 5.

Naturaliste Plateau (sensu stricto) (SP 8958 – SP 7700) The Naturaliste Plateau (NP) is characterized by a planated acoustic basement in which a number of small, fault-controlled basins have developed (e.g. Fig. 5e). The faults bounding these basins are generally planar, ENE oriented and preferentially dip SSE. The sedimentary fill in the basins is up to 1.2 s TWT (c. 2 km) thick and their upper surface is marked by an erosional unconformity at a similar depth to the surrounding basement. The basins are 10–30 km wide and up to 120 km long (Borissova 2002). The deepest imaged sediments in these rift infra-basins are at approximately 5 s TWT. These sediments have never been sampled by drilling or dredging. The deeper parts of these basins display thick (0.5 s TWT) high-amplitude, continuous-dipping laminar reflectors. These are interpreted as stacked synrift lava flows, and intercalated volcaniclastics, as they show growth against the faults. These may be equivalents of the late Jurassic basaltic rocks of the Casterton Beds in the deep Otway Basin (Finlayson et al. 1998) in the eastern part of the Southern Rift System. Basement and the rift section are unconformably overlain by post-rift sedimentary section up to 0.5 s (c. 0.5 km) thick that is characterized by highcontinuity, laminar reflections, typical of an open marine environment. On the basis of ties to DSDP wells 258 and 264, and onshore petroleum wells into the southern Perth Basin (Bradshaw et al. 2003), these upper sediments are interpreted as post-Valanginian in age.

Southern slope (SP 7700 –SP 7550) The southern slope of the NP is a narrow, steep (sea-bed gradients up to 158) series of scarps that are underlain by major, south-dipping basement faults. The slope is draped by a thin post-rift cover contemporaneous with the sediments overlying the crest of the plateau. The fault scarp was dredged by the French Marion Dufresne cruise 110, described above. On the basis of this dredge haul, the basement to the NP rift basins is

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interpreted as high-grade metamorphic rocks, possibly equivalent to the onshore Leeuwin Complex (Collins 2003).

Inner continent – ocean transition zone (SP 7750 – SP 6000) Morphologically, the inner COT forms a 40 – 50 km-wide terrace at depths of 4800–5000 m adjacent to the southern slope of the NP. This terrace is bounded by the steep planar fault scarp of the NP slope to the north and to the south by a large basement high that stands approximately 1000 m above the adjacent terrace (Fig. 5d). Basement beneath the terrace is down-faulted northwards and is covered by a post-rift sedimentary fill about 0.5 s (c. 0.5 km) thick. The basement high outboard of the terrace contains many high-amplitude reflections, some with steep dips. Because of its seismic signature, this high may be a densely faulted, continental basement horst intruded by mafic magmatic rocks; however, basement in this zone has not been sampled.

Outer continent– ocean transition zone (SP 6000 – SP 2700) The basement in the outer COTZ is highly faulted, and is considerably deeper than that in the inner COT, generally at depths of 7– 8 s (Fig. 5c). The sediment cover in this zone is only a few hundred metres thick. The basement has been sampled in two locations (MD110 –DR08 and MD110-DR10, Figs 2 & 3) with both dredges recovering vesicular alkaline basalt of continental affinities (Chatin et al. 1998). The DZ comprises ridges separated by a very deep basin at a depth of about 5600 m and underlain by up to 2.5 s (c. 3– 4 km) of sediment. Basement ridges in this province were sampled in several locations (including sites outside the study area) yielding basalts and upper mantle rocks including peridotite, lherzolite, harzburgite and pyroxenite (Nicholls et al. 1981; Chatin et al. 1998; Royer & Beslier 1998; Beslier et al. 2004). Details of the internal structure of these ridges is shown in Figure 5b, c. In Figure 5c, beneath the thickest sedimentary pile from SP 3400 to SP 3600, lie a series of tilted and rotated fault blocks separated by planar, south-dipping normal faults. The upper parts of these blocks are characterized by regular, high-amplitude reflectors. The large planar fault that breaches the sea floor near SP 3720 separates these blocks from a large ridge with a chaotic internal structure and many highamplitude events. Figure 5c shows details of the internal structure of this ridge. Of note are several

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Fig. 5. Detailed seismic sections from line GA-187/01; locations of these details are shown in Figure 3. (a) Fast-spreading oceanic crust of the Australian– Antarctic Basin. (b) Anomalous oceanic crust of the Australian– Antarctic Basin. (c) Diamantina Zone in the outer COTZ. (d) Diamantina Zone in the inner COTZ. (e) Cretaceous half-graben on the Naturaliste Plateau; dashed line shows erosional unconformity of mid-Cretaceous age separating synrift and post-rift sediments.

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Fig. 6. (a) Uninterpreted and (b) Interpreted TWT migrated seismic profiles for line GA-199/05. Interpretation follows those of Totterdell et al. (2000) and Sayers et al. (2001). Morphotectonic elements are indicated. The TWT section for seismic line GA-199/07 is very similar in structural style to GA-199/05, and has been omitted for reasons of space.

wedges of ‘chattery’ irregular high-amplitude reflectors, often with a semi-continuous upper reflector, overlain by draping sediments. These wedges are interpreted as piles of lavas and volcaniclastic materials. The wedges are underlain by discontinuous, short reflectors with occasional high-

amplitude and steeply dipping reflectors. These are interpreted as possible faulted and intruded mafic igneous basement. Overall, this seismic character is similar to that observed by Sayers et al. (2001) on the southern side of a basement ridge in the Great Australian Bight (Fig. 6). Sediments in the

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deep basin south of the ridge have never been sampled.

Oceanic crust (SP ?2700 – SP960) The southern end of line GA-187/01 is characterized by rugged, southwards-shallowing basement that is almost devoid of sediment cover (Figs 3 & 5a). This basement shows no internal reflectivity, and from SP 960 to around SP 1700 is interpreted to be normal oceanic crust, generated during the fast seafloor spreading between Australia and Antarctica that commenced in the Eocene. The oceanic crust from SP 1700 to the outer edge of the COTZ (around SP 2700) is anomalously deep, and more highly faulted than expected for normal spreadingrate Eocene crust. It also has significant volcanic build-ups (e.g. around SP 2600). This crust is associated with difficult to interpret magnetic anomalies (cf. Tikku & Cande 1999), and is presently undated. We tentatively interpret this crust as a zone of slow-spreading crust that possibly formed in the interval C34 or C33 to C17 in the Late Cretaceous. Another alternative is that it is a narrow zone of crust associated with the earlier (north–south) breakup in the Australia–Antarctica–India separation; that is, it may be Early Cretaceous in age.

Potential field interpretation Gravity field The regional gravity field (Fig. 4a) is an order 2 surface ranging from approximately 0 mGal over oceanic crust at the southern end of the line to a minimum 250 mGal in the DZ, and rising to þ20 mGal over the NP. The gravity data divide broadly into three domains. Domain 1, over oceanic crust, contains entirely positive Free Air Anomalies (FAA) of up to þ20 mGal, relative to the local field. Domain 2, broadly coincident with both the inner and outer COTZ, is entirely negative, ranging from a high of 220 mGal to a low of 280 mGal. Wavelengths of the anomalies within this domain average approximately 80 km. Domain 3 is coincident with the Naturaliste Plateau and its southern slope, and is characterized by entirely positive FAA, from þ10 to þ 30 mGal. These have wavelengths of 20– 40 km over the sediment-filled half-graben incised into the NP basement.

Magnetic field The regional magnetic field (Fig. 4a) departs from the IGRF at both the northern and southern ends of the transect by up to 2200 nT.

The magnetic field broadly divides into four domains. Domain 1, from SP 962 to SP 2100, comprises a broadly negative domain with sinusoidal anomalies (c. 2200 to 2100 nT), over clearly identifiable oceanic crust. Domain 2, from SP 2100 to SP 4600, is a zone of low frequency (.60 km wavelength) anomalies varying between 0 and – 300 nT, around the morphological and basement depression of the Diamantina Zone. Domain 3, from SP 4600 to SP 7200, is an entirely positive series of magnetic anomalies, up to þ700 nT, with two superimposed wavelengths, 100 and 20 km, overlying elevated basement in both the outer and inner COTZ. Domain 4, from SP 7200 to SP 9000, coincides with the southern slope and plateau region of the NP. Highs of up to þ300 nT overlie exposed basement, and lows of up to 2400 nT overlie the interpreted half-graben filled with ?Early Cretaceous sedimentary sequences.

Modelling methods and assumptions The interpretation process described here was designed to quantitatively validate the time-based, depth-converted interpretation of the seismic data against coincident potential field data. Wherever possible, features interpreted on the seismic data have been used to constrain the geometries of the various bodies used in the potential field models. This means that modelled bodies in the top 2 km of section below the water-bottom (or 5 s TWT beneath NP to 10 s TWT beneath the DZ) are comparatively well-constrained by the seismic interpretation and changes in reflection character. However, below this, the seismic data quality is relatively poor owing to high noise and multiples, and no seismic interpretation was attempted. In these regions, which include the middle crust to the Moho, bodies have only been modelled in order to fit the overall gross shapes of the potential fields and to be consistent with the limited velocity information. Consequently, in the mid to deep crust, the potential field models should be considered a guide to the gross geology but not to finer scale structural detail. Information on the velocities and densities for the upper part of the section, which correlate with equivalent sequences in the southern Perth Basin, have been derived from (Iasky 1990). The DZ and NP deep section have been given density attributes using some sonobuoy P-wave velocity estimates from Australian and international surveys (Talwani et al. 1979), and the stacking velocities derived during the processing of the seismic data. Stacking velocities were smoothed using a five-point filter, with the degree of smoothing being increased with the depth below sea bed in order to minimize the

COTZ ALONG AUSTRALIA’S SOUTHERN MARGIN

instability of the stacking velocities with increasing depth. All velocity estimates were converted to densities using the equation of Ludwig et al. (1970). Limited information on magnetic physical properties has been derived from analysis of dredge samples and DSDP sites described above. Information about the magnetic properties of the continental basement was derived from sampling traverses in the Leeuwin Complex of SW Western Australia undertaken during this study (Fig. 2). The detailed geology of the Leeuwin Complex is described in Collins (2003). Investigations were made of fresh outcrops in wave-cut platforms of all the major lithotypes in the complex. These include several types of amphibolites, granites and granitic gneisses of differing ages. Magnetic susceptibilities for these rocks were acquired using a Fugro GMS-2 hand-held magnetic susceptibility meter. At least 10–12 readings were acquired at each site to build up a statistical estimate of surface susceptibilities. Susceptibilities of these rocks range from 5  1025 SI (granite) to 4500  1025 SI (amphibolites), with metre-scale variability between these limits owing due to the complex metamorphic layering. The median value over five separate sites in the Leeuwin Complex was 525  1025 SI and the average value was 1100  1025 SI. These data take no account of possible remanent effects or anisotropy of magnetic susceptibility within these rocks. The magnetic properties remain the least-constrained part of the validation process. This method was designed only to produce a set of starting parameters for the initial forward model, as the properties are perturbed during modelling. As the geometry is generally well constrained by the reflection seismic data down to at least basement depths, the main unknowns in modelling are the densities and magnetic parameters of the bodies, and the geometries of bodies in the deep crust. The gravity and magnetic fields of the petrophysically-attributed geometric models were then concurrently forward modelled. This was achieved by giving the sections limited (75 km) strike extents, and then calculating and integrating the 3D gravity and magnetic anomalies sourced by bodies in the section. This approach was considered reasonable, given the generally 2D strike of the geological features imaged at the scale of the seismic section. However truly 3D features, such as some of the volcanic build-ups in the DZ, may not be accurately modelled using this method. Other assumptions in modelling include the following. † The background density contrast of the model is that of standard continental crust: 2.67 103 kg m23 (Hinze 2003).

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† The inducing magnetic field, based on the IGRF, has a strength of 60386 nT, an inclination of 2718 and declination of 278. † The body responses are modelled in their correct x, y and z locations, not projected into a plane. The solutions are 3D analytic for polyhedra, based on the algorithms of (Coggon 1976) and (Lee 1980). † The calculations do not take account of sphericity of the Earth over the distance of the baseline. Induced errors due to this assumption, based on comparison of similar 2.5D and spherical 3D models by Takin and Talwani (1966) are of the order of 3 mGal and are not expected to exceed 5 mGal. Given the noise envelope of approximately 1 mGal, and the dynamic range of the data (c. 70 mGal), this is not considered significant. † All densities (r) are quoted in 103 kg m23; all magnetic susceptibilities (k) are quoted in SI units (Table 2). Magnetic induction is assumed to be the principal method of generating magnetic anomalies in continental crust; in these bodies the dimensionless apparent Ko¨nigsberger ratio (Qa), the ratio of the remanent field to the inducing field, is set to zero. In oceanic crust and many of the inferred igneous bodies within the COTZ, Qa is non-zero, implying a contribution from magnetic remanence. In the absence of detailed magnetic measurements on drill cores from the NP and DZ, this contribution is assumed (for simplicity) to be aligned (sub-) parallel to the inducing field. Thus, the true values of Q might be significantly different to those modelled if there is a high angle between real remanent vectors and the current inducing field vector.

Modelled solution The final modelled solution for the geometry of the Naturaliste –Diamantina margin (Fig. 4a) has a unitless RMS error (see Direen et al. 2001) between the calculated and observed fields of 3.7 for the gravity, and an RMS error for the magnetics of 4.9. Model bodies A –C beneath the NP correspond to the post-Albian sedimentary section. The densities of these bodies range from 2.00  103 to 2.30  103 kg m23. A and B are non-magnetized, but the bodies labelled C have strong negative polarity remanence. The magnetic susceptibilities of these bodies vary from 0.03 to 0.09 SI, with Qa from 1.5 to 3.6. These properties are consistent with the seismic interpretation of stacked mafic flows and volcaniclastics within these sequences. A major density change (from 2.10  103 to 2.52  103 kg m23) takes place at the transition

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from postrift to prerift (?Mesozoic) sediments beneath the NP interpreted at the boundary between bodies C and D. This corresponds to a change in seismic character from high-amplitude, high-continuity, laminar reflections in the shallow section, to more discontinuous reflectors at depth. The base of body D is relatively unconstrained, with no corresponding seismic reflection boundary with the underlying bodies, owing to the strong seabottom multiple. The upper part of the NP basement may consist of deformed Jurassic –Permian sediments, based on regional correlation with the onshore southern Perth Basin (Iasky 1990). The modelled thickness of Mesozoic –Palaeozoic section on the NP (c. 5200 m) is consistent with reported thicknesses of these sediments onshore (4000–6000 m: Iasky 1990). Bodies E and F, within the COTZ (Fig. 4A, SP 2560–SP 7500; Fig. 7), are coincident with two sets of irregular, high-amplitude reflectors. E, which overlies F, contains ‘chattery’, short, discontinuous reflectors, and is interpreted as post-breakup (?Albian) mafic volcanic flows. These bodies have variable properties, with densities ranging from

2.78  103 to 2.92  103 kg m23, and susceptibilities of 0.01–0.07 SI, with a low intensity (i.e. Qa is  1) reverse polarity magnetization. The bodies labelled F are coincident with more continuous to mounded irregular, high-amplitude reflections. These are interpreted as weathered mafic volcanic build-ups capping basement highs, and may date from the breakup. They are less dense (2.60  103 kg m23) than body E, with moderate –high magnetic susceptibilities (0.04 –0.07 SI) and have normal magnetic polarity. These interpreted mafic flows and intrusives modify both interpreted crystalline continental basement (body G) and exhumed upper mantle (bodies R1-R3). The crystalline basement, where it is imaged above the strong water-bottom multiple, is seismically chaotic, with many diffractions, probably indicating internal faulting and folding. This is consistent with its correlation with the polydeformed Leeuwin Complex rocks exposed onshore (Collins 2003). Beneath the NP, basement has been modelled as a weakly magnetized, predominantly felsic body (2.65  103 kg m23 0.005 SI). According to the indicative measured densities

Fig. 7. Potential field model for line GA-187/01 with seismic depth section. Bodies annotated as in Figure 4a. Note strong water-bottom multiple beneath the Naturaliste Plateau at c. 8 km, SP 7500– SP 9600.

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from samples in the Leeuwin Complex, a felsic crust composed of 55% granite, 35% felsic gneiss and 10% amphibolite is consistent with this density and magnetization. A body with increased magnetic susceptibility (0.08 SI) is also modelled on the escarpment fault to reproduce a prominent magnetic anomaly at that location. Beneath body F, in the outer COTZ, the crystalline basement is less dense (2.60  103 kg m23) and non-magnetic. A basement body is also interpreted to lie in deep water (c. 5500 m) beneath the sedimentary section in the outermost DZ (Fig. 4A, SP 2600– SP 3500; Fig. 7), coincident with interpreted tilted and rotated fault blocks separated by planar, seaward-dipping normal faults (outermost body G). This modelled block is denser (2.70  103 – 2.72  103 kg m23) than the continental crust modelled inboard, and also significantly magnetized (0.02–0.045 SI). It is therefore likely that this crust, which contains planar high-amplitude reflections, may be intruded by mafic sills that increase both the density and magnetization. We speculate that this crust is a combination of foundered, modified Mesozoic crust (equivalent to body D) overlying highly-thinned continental basement. This crust (G) is separated from the outer edge of continental basement by a normally-faulted basement high (bodies R2 and R3). This feature contains internal continuous irregular reflectors, and some steeply dipping strong, high-amplitude reflections. The basement high is modelled with four dense (2.80  103 – 2.90  103 kg m23) and highly magnetized (0.02, 0.05, 0.07 and 0.07 SI) bodies with a reversed component of magnetic remanence. We interpret the basement high to be comprised of a complex mixture of magnetic serpentinized peridotite, lherzolite and gabbroic intrusions, capped by basalt or ophicalcite breccia. This is consistent with dredge hauls by the Diamantina and Marion Dufresne cruise 110 made on this high, as well as its internal seismic reflection character. Outboard of the basement ridge and deep basin complex, oceanic crust (Layer 2A) is interpreted from about SP 2715 southwards. The oceanic crystalline crust interpreted from the reflection seismic data has been modelled using a variable three-layer structure. As no internal crustal layering is discernible in the reflection seismic data, the layers that are used as the basis for the potential field model are derived from seismic stacking velocities. The upper layer is bounded at its top by the rugose basement surface. All blocks in this layer are modelled with the same density (2.55  103 kg m23), but the magnetic susceptibilities vary from 0.04 to 0.07 SI, and the apparent Ko¨nigsberger ratios vary from 2 to 3. No linear magnetic anomalies have been picked in this area. Because a large part of this zone of oceanic crust is likely to be slow

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spreading crust, any magnetic anomalies in the plane of the section are likely to be the result of interference effects between a number of normal and reversely magnetized bodies. We have modelled the crust with individual bodies to fit the observed magnetic data, as at present we have no way of identifying actual blocks of crust that may interfere to produce the apparent magnetic pattern. Underlying this is a single higher density layer (Layer 2B) (2.60  103 kg m23) with negligible magnetization, interpreted as relatively altered dolerite dyke swarms and/or layered gabbro (oceanic layer 2B–3). This is, in turn, underlain by a uniformly denser layer (2.85  103 kg m23) of non-magnetized gabbro and peridotite (Layer 3). A lithostatic loading curve has been generated for the base of the model (30 km), to test whether the model exceeds the reasonable yield strength of the upper mantle. This curve was generated by dividing the model into approximately 430 columns of 1 km width in the x and y planes, and calculating the integrated weight at the base of each column (Funck et al. 2004). The average pressure at the base of the model is 733 MPa, with deviations no larger than + 26 MPa. This figure lies well within the compressional failure envelope of about 2 GPa (Watts & Burov 2003), indicating that this upper lithospheric configuration is at least isostatically stable. In summary, this model is consistent with the interpretation that the entire continent–ocean transition zone of the Naturaliste– Diamantina margin is comprised of thinned, extended continental crust, based on integrating the seismic reflection interpretation with potential field modelling and dredge constraints.

Alternative models Two alternative models for the structure of the Naturaliste –Diamantina margin have been generated (Figs 8 & 9). The first of these (Fig. 8) tests whether or not it is possible to fit the observed potential field data using a mafic magmatic source for the basement to the NP, as suggested by Coleman et al. (1982). The model shows that it is possible to fit the data with mafic underplating beneath the NP (densities: A, 2.70  103 kg m23, 0 SI; B, 2.90  103 kg m23, 0.02 SI); load calculations at the base of the model (30 km) indicate an average lithospheric load of 769 MPa, with deviations of up to + 26 MPa. These figures are higher than the loads for the preferred model, but are not in excess of the likely yield strengths at this depth (e.g. Watts & Burov 2003). However, the required thicknesses of mafic material are far in excess of what is reasonable. For example, the total thickness

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50 FAA (mgal)

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Fig. 8 Alternative potential field model – GA 187/01. Test of underplating beneath the Naturaliste Plateau. Panels as for Figure 4. Bodies: A 2.70  103 kg m23, 0 SI; B: 2.90  103 kg m23; 0.02 SI.

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Fig. 9 Alternative potential field model – GA 187/01. Test of aborted sea-floor spreading in the inner COTZ/DZ. Panels as for Figure 4; no loading curve was calculated for this model, as the load is between that predicted for the two extreme models above. A: (Layer 2A) 2.55  103 kg m23, 0.03 and 0.05 SI, and Qa of 1.5 and 3. B (sheeted dykes), 2.70  103 kg m23, 0 SI, C (gabbro): 2.90  103 kg m23, 0.02 SI.

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of approximately 20 km of mafic material is far larger than those described from the Ontong Java or Kerguelen plateaus, which are the largest known igneous provinces on the globe (Coffin & Eldholm 1992, 1993a, b, 1994, 2001). In addition, such a massive pile of mafic volcanics would have undoubtedly produced strong metamorphic overprinting of the thin sedimentary pile deposited above it, given the short interval between cessation of Albian hotspot volcanism in the Kerguelen system (Frey et al. 2000) and the oldest sediments intersected on the NP (Hayes et al. 1975). Such metamorphic effects are highly unlikely, given the observed velocities within the NP sequences; therefore, this model was discounted as being geologically unlikely. A second alternative, presented in Figure 9, is that the inner COTZ is comprised of slow spreading oceanic crust, separating a continental fragment in the DZ from the NP. In this model, the additional panel of oceanic crust has Layer 2A densities of 2.55  103 kg m23, magnetic susceptibilities of 0.03 and 0.05 SI, and Qa of 1.5 and 3. Again, while this alternative model can produce a credible fit to the observed data, with ‘reasonable’ property assignments, the extreme thickness of lower oceanic crust required in the COTZ (up to 13 km) is unlikely to be geologically correct. In an extensional non-volcanic margin such as the Naturaliste – Diamantina margin, where oceanic crust and continental mantle are likely to have been structurally thinned, rather than thickened volcanically, oceanic crust is more likely to be thinner than the ‘normal’ approximately 8 km. Thus, we believe it unlikely that normal oceanic crust produced by seafloor spreading occurs between the NP and the DZ.

Great Australian Bight margin The integrated modelling described above shows that it is likely that the area between the southern flank of the Naturaliste Plateau, through the Diamantina Zone, to the fast-spreading oceanic crust of the Australian–Antarctic Basin is a wide (c. 200 km) COTZ with a high degree of structural complexity. In this respect, it has similarities to the Great Australian Bight margin, located approximately 2000 km along strike to the east (Figs 1 & 2). Early investigations of this margin were reported by Bouef & Doust (1975), Ko¨nig & Talwani (1977) and Talwani et al. (1978, 1979). However, those studies lacked high resolution, deep MCS data, and were thus heavily reliant on unconstrained inferences from potential field datasets and refraction sonobuoys. More recent studies (Totterdell et al. 2000; Sayers et al. 2001) have included deep, high-resolution MCS data, combined

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with sonobuoys and potential field data, leading to more robust interpretations. The modelled structure of the Great Australian Bight margin is shown in Figure 4b, c. These models are based on deep seismic lines GA-199/ 05 and GA-199/07 from the Geoscience Australia Survey 199, acquired in 1997. Acquisition parameters of these data, and interpretations of the seismic reflection characteristics and velocity structure, are described in Sayers et al. (2001). These lines have been processed and modelled in similar fashion to that described above for line GA-187/ 01. Note that line GA-199/05, which is also presented in Sayers et al. (2001), has been partially reinterpreted, and the age of the sequences is now consistent with the interpretation of Totterdell et al. (2000). Details of this reinterpretation of the sequences stratigraphy are briefly described below, as appropriate. However, in broad terms the tectonomorphological divisions of the GAB margin remain as described by Sayers et al. (2001), and are: † a zone of highly attenuated continental crust; † a basement ridge beneath the inner edge of the deep ocean basin; † a continent–ocean transition (COT) that extends from the basement ridge to the continent–ocean boundary; and † faulted, slow-spreading oceanic crust. These structural subdivisions are very consistent along strike and are recognized on a number of profiles (e.g. Fig. 4b, c; cf. figures in Sayers et al. 2001). The landward end of both modelled profiles is characterized by the rapid oceanward thinning of the likely Gawler Craton crystalline crust (‘GC’), with Moho shallowing from around 16 –18 km to approximately 12 km over a distance of about 60 km. This thinning is accommodated by faulting, erosion and onlap of the deeper (?Jurassic and Early Cretaceous) sedimentary section, and by ductile deformation of the crystalline crust. The upper sedimentary synrift package (‘T’, Fig. 4b) is equivalent to the Tiger supersequence of Totterdell et al. (2000). This has a seismic velocity of 4500 m s21, with an inferred density of 2.48  103 kg m23. It has an elevated magnetic susceptibility (k) of 0.01 SI and Qa of 2, with a reversed polarity component. The lower Blue Whale supersequence (‘B1.W1’ and ‘B1.W2’) has an average sonic velocity of approximately 4000 m s21 and an inferred density of 2.5  103 kg m23. It is also reversely magnetized (k ¼ 0.02 SI, Qa ¼2). The magnetic properties of the Blue Whale supersequence indicate probable input of volcanic material in this section.

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Anomalous low-density, lower crust (‘granite?’) interpreted by Sayers et al. (2001) has been reinterpreted here as a thinned wedge of lower continental crust (‘GC’: vp ¼ 6600 m s21, r ¼ 2.80 103 kg m23). Beneath this, a second, dense, magnetized lower crustal layer (7000 m s21, r ¼ 2.95  103 – 3.25  103 kg m23; 0.01 SI) has been interpreted. This layer is believed to be serpentinized peridotites of the upper mantle, altered by fluids within brittle faults cutting down to this level prior to breakup. This alteration increases in intensity into the COT. Such features are well documented on other nonvolcanic margins (Reston et al. 2001). The ‘magnetic quiet zone’ (MQZ: Talwani et al. 1979) is associated with this raft of horizontally layered, weakly magnetized continental crust. This interpretation is consistent with that of Sayers et al. (2001). The prominent feature on the GAB margin is a basement ridge at the outboard edge of the zone of maximum crustal thinning (‘R’, Fig. 4b, 4c). This ridge is seismically and structurally complex, and has an associated high-amplitude magnetic anomaly. Forward modelling indicates that the ridge is probably a combination of highly magnetized and variably dense serpentinized peridotites and mafic intrusions or extrusions. Such features can be explained by a combination of mantle exhumation, diapirism and limited partial melting during breakup (cf. Sayers et al. 2001). The magnetic anomaly associated with the ridge has previously been identified as sea-floor spreading anomaly A34 with the underlying crust interpreted as oceanic leading to previous interpretations of the breakup age at about 95 Ma (e.g. Tikku & Cande 1999). However, as with Sayers et al. (2001), the interpretation of this paper is that this crust is of continental origin, albeit heavily modified by structural and volcanic processes. Interpretation of the age of the onlapping sediments ponded in depressions landward of the ridge indicates that the ridge was probably emplaced in the late Cenomanian, during the last part of the long normal polarity periods of chron 34. Thus, while the previous identification of the age of this anomaly is probably correct, we concur with Sayers et al. (2001) that it is incorrect to describe it as a sea-floor spreading anomaly. Between the outboard flank of the basement ridge and the first emplacement of oceanic crust, the GAB margin encompasses a zone of highly complex structuring. This zone is characterized by a thick, well-layered sequence that is down-faulted across listric faults towards the COB. This sequence is sedimentary in seismic character; however, discontinuous, high-amplitude internal reflectors indicate that volcanics and intrusions are probably

also widespread, particularly closer to the COB. On the basis of seismic character correlation (cf. Totterdell et al. 2000), the rocks in this basin are interpreted to be of pre-Turonian age. The potential field model requires diverse, relatively dense (bodies U1– U3: r ¼ 2.67  103 – 2.75  103 kg m23), with significant magnetizations (k ¼ 0.03 –0.07 SI, Qa ¼ 1.6–2), and of both reversed and normal polarity, to fit the complex magnetic field in this area. These results are consistent with this zone being a mixture of igneous complexes, intruded prerift sedimentary sequences, and rafts of seismically transparent, thinned continental lower crust. This basin is most likely underlain by altered mantle rocks, on the basis of high magnetization and high density, characteristic of serpentinized peridotites (‘SM 1–5’: k ¼ 0.01 – 0.03 SI; r ¼ 2.90  103 –3.05  103 kg m23). A feasible alternative may be extensive mafic intrusions within continental crust. The seismic data and integrated modelling suggest that breakup has taken place on a northdipping fault that probably soles out in a highly magnetized layer, possibly serpentinite, at about 14 km depth. In the zone of clearly recognizable slowspreading oceanic crust, oceanic crustal properties for layer 2A in both GAB lines range from 2.47  103 to 2.55  103 kg m23 (average 2.51) and 0.004–0.05 SI (average 0.026), with Qa values of between 1 and 4. The onset of unambiguous sea-floor spreading is marked by a pulse of reversely magnetized basalts emplaced immediately prior to anomaly A33o time, dating final breakup in the early Campanian at between 83 and 80 Ma, consistent with the interpretation of Sayers et al. (2001). The interpretive forward model for seismic line GA-199/07 is shown in Figure 4C. This model can also be divided into four morphotectonic domains, equivalent to those in GA-199/05. The properties of the bodies in this model are essentially the same as those used for equivalent bodies in the model of GA-199/05, especially in the post-rift section and the underlying mantle. In the zone of highly attenuated continental crust in GA-199/07, the Gawler Craton crystalline basement is magnetized (GC: k ¼ 0.03 SI, Qa ¼ 1; normal polarity), compared to the non-magnetic basement in GA-199/05. This effect is probably owing to localized magnetic intrusions or Banded Iron Formations, both of which are irregularly distributed through the Gawler Craton Proterozoic basement (Direen et al. 2005). Altered mantle (‘SM’: r ¼ 3.05  103 – 3.25  103 kg m23) underlying the interpreted base of continental crust gives rise to significant

COTZ ALONG AUSTRALIA’S SOUTHERN MARGIN

positive normal magnetizations (k ¼ 0.027 –0.04 SI). Slight differences in magnetization and density in GA-199/07 compared to those in GA-199/05 may be attributable to the spatially variable nature of peridotite serpentinization (Hopkinson et al. 2004). The basement ridge (R) in GA-199/07 is somewhat broader than that modelled in GA-199/05, and has less relief on the upper surface, as shown in Sayers et al. (2001). Internally, the ridge has a diffuse seismic reflectivity, with some steeply dipping events interpreted as intruded faults or dykes. The properties of the ridge complex (r ¼ 3.05  103 kg m23, k ¼ 0.02 SI) are consistent with it being comprised of either partially serpentinized peridotites, or fault disrupted gabbros, or both, similar to the modelled solution for GA-199/05. In the upper part of the COTZ outboard of the basement ridge (Fig. 4C: ‘E1, E2’), in GA-199/07, high-amplitude layered reflections are interpreted as mafic volcanics. These display short, discontinuous, high-amplitude, sequences with ‘rough’ topography. The properties of these materials (‘E1, E2’: r ¼ 2.50  103, 2.60  103 kg m23, k ¼ 0.03, 0.04 SI, reversed polarity), although having more variance than similar bodies in GA-199/05, are consistent with the spatially complex nature of the properties of interlayered mafic flows and volcaniclastics (e.g. Zhao 2001). Beneath this sequence lies a layered package that may be equivalent to the Blue Whale and/or Bronze Whaler supersequences (Fig. 4c: ‘U’) that lie inboard of the basement ridge. This sequence is characterized by well-layered, often high- to moderate-amplitude reflectors in normal tilt blocks. These blocks show evidence of growth faulting and half-graben development in the overlying section, indicating synrift movements. Deeper reflectors are often folded at block edges owing to block rotations. The high-amplitude reflectors that dominate several of the normal tilt blocks may be sills, given the higher magnetizations (0.02–0.03 SI) in these blocks. In places these reflectors obscure the original bedding geometries. A more seismically transparent sequence of extended, magnetized (0.03 SI) basement (Fig. 4c: ‘M’: r ¼ 2.88  103 kg m23) underlies the COTZ in GA-199/07, and is probably an intruded raft of continental crust, as is interpreted in line GA-199/05. Beneath this sequence a zone of noisy seismic response with some short, very high-amplitude reflectors (Fig. 4c: ‘S’) may be serpentinized peridotites of the upper mantle (S: r ¼ 3.00  103 kg m23, k ¼ 0.01 SI).

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Discussion There are many strong similarities in the overall architecture of the Naturaliste– Diamantina and Great Australian Bight continental margins. In particular, it is possible in both locations to recognize the following elements: † multiply deformed Proterozoic–early Palaeozoic (minimum 522 Ma: Collins 2003) crystalline cratonic basement that constitutes the basement to infra-rift sequences; † probable onset of rifting and deposition of synrift basin sequences commencing in the middle to Late Jurassic; † an inner zone of attenuated, multiply-rifted continental crust; † a basement ridge, or zone, comprising exhumed serpentinized peridotites, and other mafic magmatic– volcanic rocks, beneath the faulted, landward edge of the deep ocean basin; † COTZ that extends from the basement ridge to the COB. These similarities in style of rifting and subsequent breakup are likely to reflect overall similarity in the applied stress fields, lithospheric rheology and magma budgets during breakup, similar to those of analogues in Iberia and the Alps (Whitmarsh et al. 2001). Distinctive differences between the studied areas include: † proximity of the Naturaliste –Diamantina margin to two rift branches of distinct ages, whereas the GAB margin is wholly within a single rift arm; † the east –west-trending Naturaliste –Diamantina margin reworks the slightly older Valanginian Perth north –south-trending margin, and is also overprinted by later Albian hotspot activity; in contrast, the GAB margin cuts orthogonally across older Proterozoic structures without significant reworking/reactivation (Direen et al. 2005); † breakup between Naturaliste Plateau and Bruce Rise in the western part of the SRS appears to have been highly asymmetrical (Stagg et al. 2005; this paper), whereas between the GAB and Wilkes Land in the central part of the SRS it was more symmetrical (Colwell et al. 2006). Based on studies of the west Iberian margin, Whitmarsh et al. (2001) concluded that mechanical or thermally induced weaknesses may influence the location and style of mantle deformation with respect to the overlying crust. This results in exhumation of the upper subcontinental mantle at length

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scales of tens of km (c. 60 kilometres in their example). In our study, the COTZ in the Naturaliste –Diamantina margin is quite wide (c. 250 km) compared to the GAB margin (70 km); however, the width of the exhumed upper mantle implied in each margin is similar (c. 55 km in the Naturaliste– Diamantina margin v. c. 80 km in the GAB margin). As the mantle deformation is similar, the difference in overall width of the COTZ between these margins can be attributed to differential crustal deformation in each location. Several explanations are possible for altering the mechanical strength of the crust during nonvolcanic rifting: † orientations of pre-existing lithospheric weaknesses relative to the rift palaeostress field (van Wees & Beekman 2000); † thermal state of the lithosphere, reflecting crustal basement cooling history, e.g. postorogenic collapse (van Wees et al. 1992); † thermal weakening of the lithosphere, reflecting internal heat production in the crust (Sandiford et al. 2003); † thermal weakening of the upper lithosphere as a result of prior volcanic breakup (White & McKenzie 1989); † other, as yet unexplored, effects. Each of these possibilities is explored briefly below. † The Naturaliste –Diamantina and GAB margins are both latitudinal margins that cut orthogonally across earlier tectonic structures. In the case of the Naturaliste –Diamantina margin, the earlier fabrics include both the north –south oriented Neoproterozoic Pinjarra Orogen (Collins 2003), and the ?Hauterivian Perth AP–Cuvier AP– Gascoyne AP margins (Bradshaw et al. 2003). In the GAB, the margin cuts orthogonally across 1550–1440 Ma shear zones of the Fowler and Christie domains (Direen et al. 2005). It is thus unlikely that major old crustal weaknesses contributed to the NDM being significantly weaker than the GAB. If anything, young east –west structures such as those in the Polda Basin (Stagg et al. 1992) should have contributed to weakening of the GAB relative to the Naturaliste –Diamantina margin. † The Naturaliste –Diamantina and GAB margins are both late Mesozoic rifted margins that developed on Proterozoic crust. In the Naturaliste– Diamantina margin, the age of the youngest prerift metamorphic activity is 522 Ma (Collins 2003); in the GAB, the youngest recognized thermal reworking is at 1440 Ma (Swain et al. 2005). Both of these ages indicate

that the Gondwana lithosphere had sufficient time to thermally re-equilibrate before the onset of the late Mesozoic rifting. It is thus unlikely that some widespread, but unrecognized, late Palaeozoic– early Mesozoic thermal event has weakened the Naturaliste –Diamantina crust relative to the GAB margin. † Heat flow measurements on the Australian crust are sparse (Cull 1982). However, despite this, the nearest measurements of heat flow from the granitic–gneissic basement of SW Western Australia are similar to those from southern South Australia, at 50 mW m22, implying similar contributions of basement heat production to the heat flow in both areas. Thus, anomalous heat production by radiogenically-enhanced basement rocks in the Naturaliste –Diamantina margin is unlikely to be a cause of enhanced thermomechanical weakness during breakup. † Crust immediately north of the Naturaliste–Diamantina margin (Perth AP –Cuvier AP margins) is believed to have undergone volcanic breakup some time during the Valanginian–Hauterivian (Symonds et al. 1998). Such breakup is commonly a result of massive transient heat input into the lithosphere, resulting in widespread thermal and volcanic effects (Coffin & Eldholm 1992). In contrast, lithospheric domains adjacent to the GAB margin, such as the Otway Basin (Finlayson et al. 1998) and its conjugate Terre Ade´lie margin (Colwell et al. 2006), show none of the attributes of volcanic breakup and attendant lithospheric weakening. Thus, the thermal impact of Valanginian–Hauterivian volcanic breakup proximal to the Naturaliste–Diamantina margin cannot yet be ruled out as a significant influence on the thermomechanical behaviour of crust at the free boundary during subsequent non-volcanic breakup later in the Cretaceous. Investigating the impacts of this early volcanic breakup on the later non-volcanic breakup may be fruitful for further research – not least into constraining the nature of and timing of the breakup in the Perth AP, which is currently poorly understood (Bradshaw et al. 2003). † Finally, there may be other poorly understood factors that operated to cause differential mechanical strength between the Naturaliste –Diamantina and GAB margins during breakup. At least one possibility is significant differences in crustal rheological profiles (Cloetingh et al. 1995; van Wees & Beekman 2000); another is significant differences in prerift lithospheric thickness (Bertotti et al. 1997). Both of these factors in the Southern Rift System are unconstrained at present. The former may be investigated through comparison of lithospheric

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seismic velocity profiles (Cloetingh & Banda 1992); the latter through comparison of suites of currently exposed metamorphic assemblages exhumed during the breakup (Sandiford 1985).

Implications Current knowledge of the Australian side of the Southern Rift System (SRS) (Finlayson et al. 1998; Sayers et al. 2001; this study) indicates that at least the Australian flank of this system has compartments of non-volcanic rifting all along its length. Infill studies may demonstrate that the entire approximately 4000 km of the Australian side of the SRS is, in fact, a non-volcanic rifted margin. Studies of data from the continental margins of Wilkes Land and Terre Ade´lie (Colwell et al. 2006) also indicate that this part of the Antarctic conjugate margin is largely non-volcanic, consistent with the GAB margin. It is therefore possible, or even likely, that the entire SRS is a non-volcanic rift system. This assessment contrasts with some global estimates of the volumes of volcanic margins, which are considered the ‘normal’ mode of continental breakup. For example, Menzies et al. (2002) included the Naturaliste –Diamantina margin in their class of volcanic rifted margins. Our study highlights the risks of coming to general conclusions about the relative importance of the different tectonic drivers in rifting when large tracts of rifted margins are still sparsely investigated.

Conclusions This study documents the interpretation of three crustal sections from coincident deep-seismic reflection, gravity and magnetic data acquired on Australia’s southern margin. Forward modelling of data from the Naturaliste Plateau through the Diamantina Zone to the Australian– Antarctic Basin show a transition from thinned continental crust beneath the Naturaliste Plateau, through a wide COTZ to probable slow and normal sea-floor spreading oceanic crust. The COTZ is a geologically and geophysically complex zone, but interpretation of all data, including dredge hauls, is consistent with the presence of a mixture of modified continental lower crust, breakup related volcanics and exhumed continental mantle being present in this area. Alternative models for the structure of the COTZ involving either mafic underplating or aborted sea-floor spreading are possible, but are not considered tenable in this area. Final breakup on the Naturaliste –Diamantina margin is interpreted to have reworked a margin

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formed by earlier Valanginian (c. 125 Ma) volcanic breakup in the Perth Abyssal Plain to the north. The final geometry of breakup of the Naturaliste –Diamantina margin matches well with our new results from modelling the GAB margin, some 2000 km to the east. There, a larger dataset shows small variance along strike in the nature of the lower crystalline crust and its deformation behaviour, as well as the morphology of interpreted exhumed mantle complexes. Similarity in the final architecture of these margins has major implications for the nature of rifting in the Southern Rift System, and may point to the entire 4000 km system being non-volcanic in character. Further studies of the margin between the Naturaliste Plateau and the GAB are required to verify this, but preliminary data from the Antarctic conjugate (Colwell et al. 2006) tend to support this interpretation. If correct, this will have a major impact on worldwide estimates of volcanic v. non-volcanic rifted margins. Second-order differences in geometry of the Naturaliste –Diamantina and GAB margins, in particular the final width of the zone of serpentinized exhumed mantle, are unlikely to be a function of strain rate. Instead, they may reflect localized thermal weakening of the Naturaliste– Diamantina upper lithosphere. This may have been caused either by the impact of Barremian hotspot activity in the Kerguelen Plateau or, alternatively, by differences in overall crustal rheology or thickness. Few data currently exist to test the latter hypotheses. It is clear that the Naturaliste–Diamantina region, and probably its conjugate to the west of the Bruce Rise, were affected by an older Valanginian phase of margin development, that was in turn re-heated from about 120 Ma under the influence of the Kerguelen plume. This two-stage regional lithospheric heating may have produced enhanced thermal weakening and thus could account for differences between the Naturaliste– Diamantina and GAB margins.

Appendix Details of smoothing function applied to stacking velocities for line GA-187/01 to generate the velocity field for depth conversion of migrated data. WB was picked from stacked, migrated data. WB to WB þ 1000 ms (1, 5, 50, 5, 1), i.e. almost no smoothing. WB þ 1000 to WB þ 2000 ms (1, 3, 6, 3, 1), i.e. light smoothing. WB þ 2000 to WB þ 3000 ms (1, 3, 4, 3, 1). WB þ 3000 to WB þ 4000 ms (2, 3, 4, 3, 2). WB þ 4000 to end of data (1, 1, 1, 1, 1), i.e. full smoothing.

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M. Hand (CERG) is thanked for discussion of the impact of internal heat production on lithospheric strength. S. Randle (AU) is thanked for assistance with isostatic loading calculations. S. Mezzomo (GA) drafted figures. The manuscript was improved by positive comments from M. Tischer, J. Hopper, G. Karner and an anonymous reviewer. Geoscience Australia authors publish with the permission of the Chief Executive Officer, Geoscience Australia.

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Constraints on the deformation and rupturing of continental lithosphere of the Red Sea: the transition from rifting to drifting J. R. COCHRAN1 & G. D. KARNER2 1

Lamont-Doherty Earth Observatory, Palisades, NY 10964, USA (e-mail: [email protected]) 2 Present address: ExxonMobil Upstream Research Company, Mail Stop URC-URC-S169A, P.O. Box 2189, Houston, TX, 77252-2189, USA Abstract: The Red Sea is an ideal locale for testing differing models and hypotheses for rift evolution and the initiation of sea-floor spreading. The Red Sea is an active rift system that formed by the rifting of Precambrian continental lithosphere beginning in the late Oligocene, leading to breakup and sea-floor spreading by approximately 5 Ma in the southern Red Sea to the south of about 198300 N. In the northern Red Sea, north of approximately 238300 N, organized sea-floor spreading is not observed, although individual volcanoes are located within discrete ‘deeps’ spaced along the axial depression. These have been interpreted as marking the beginning of a transition from amagmatic, mechanical rifting to magmatic, oceanic spreading. Based on seismic reflection and refraction, gravity, magnetic and heat flow data in the northern Red Sea, it has been suggested that rift development occurs via the rotation of large crustal fault blocks that sole into a zone of plastic creep in the lower crust, resulting in a flat Moho and high upper crustal relief. Melt formed within individual rift segments is focused to form small axial volcanoes. That is, the northern Red Sea is on the verge of replacing horizontal translation with focused mantle upwelling and organized sea-floor spreading. In marked contrast, many passive margins (e.g. West Africa, Brazil, and NW and NE Australia) are characterized by the stacking of regional synrift sag packages, the thickness and distribution of which are inconsistent with the minor amounts of brittle deformation mapped from seismic sections. A fundamental implication of this observation is that rifts characterized by large offset fault systems, (i.e. faults that generate synrift accommodation, such as in the Basin and Range province and East Africa) will not proceed to breakup. The challenge is to understand why different portions of the Red Sea show different stages in the development of a spreading centre during continental rifting. Two hypotheses exist: (1) the structural framework deduced in the north simply continues to the south where sea-floor spreading exists and that the two regions are registering a difference in total extension. Thus, the northern Red Sea has experienced insufficient extension to breach the continental lithosphere but, in time, should develop into a spreading centre; or (2) the lithosphere of the northern Red Sea region is rheologically stronger compared with the lithosphere of the southern Red Sea, perhaps as a consequence of the thermal effects of the Afar plume, and the northern Red Sea will never evolve to sea-floor spreading. The existence of large rotated fault blocks, as implied from the inversion of gravity and magnetic anomaly data, favours the latter.

Exactly how the lithosphere ruptures to form a rifted continental margin and a new ocean basin (i.e. the transition from extension to sea-floor spreading), and why some rifts succeed in making this transition while others fail, is poorly known. The Red Sea is an ideal locale for testing differing models and hypotheses for rift evolution and the initiation of sea-floor spreading. The Red Sea is an active rift system that formed by the rifting of Precambrian continental lithosphere beginning in the late Oligocene, leading to breakup and sea-floor spreading by about 5 Ma in the southern Red Sea to the south of approximately 198300 N (Roeser 1975). In the northern Red Sea, north of approximately 238300 N, organized sea-floor spreading is not observed, although individual volcanoes are located within discrete ‘deeps’ spaced along the axial

depression. These have been interpreted as marking the beginning of a transition from amagmatic, mechanical rifting to magmatic, oceanic spreading (Pautot et al. 1986; Martinez & Cochran 1988; Cochran 2005). Between these two regions (in the central Red Sea) there is a transition area where sea-floor spreading cells have nucleated, but have not linked to form a continuous axis (Pautot 1983; Cochran 1983; Bonatti 1985; Bicknell et al. 1986). A fundamental question is whether sea-floor spreading is propagating from south to north or whether the mode of extension in the northern Red Sea has in some way prevented the development of sea-floor spreading in that region. Based on data in the northern Red Sea, Cochran (2005) suggested that rift development occurs via the rotation of large crustal fault blocks that sole

From: KARNER , G. D., MANATSCHAL , G. & PINHEIRO , L. M. (eds) Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 265– 289. DOI: 10.1144/SP282.13 0305-8719/07/$15.00 # The Geological Society of London 2007.

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into a zone of plastic creep in the lower crust, resulting in a flat Moho and high upper crustal relief. Initial broadly distributed extension is eventually replaced by focused extension at the rift axis leading to rapid lithospheric thinning and melt generation. Melt formed within individual rift segments is focused to form small axial volcanoes (Cochran 2005). With continued extension and magmatism, these volcanoes may become cells of sea-floor spreading that grow and coalesce to form a continuous axis, as exists in the southern Red Sea. That is, the northern Red Sea is on the verge of replacing horizontal translation by focused mantle upwelling and organized sea-floor spreading. Given that considerably more extension has occurred in the southern than in the northern Red Sea, as determined by the difference in opening rate between Arabia and Nubia (Roeser 1975; Joffe & Garfunkel 1987; Chu & Gordon 1998; Jestin et al. 1998), then why even question that seafloor spreading will simply propagate northwards into the northern Red Sea? Karner et al. (2003) and Karner (2005) have shown that the West African, Brazilian and NW Australian passive continental margins are characterized by the stacking of regional synrift sag basins, the amplitude and distribution of which are inconsistent with the minor amounts of brittle deformation mapped from seismic sections. The sag basins exhibit none of the diagnostic features expected of synrift extensional systems (e.g. normal faults that control accommodation, rotation of crustal blocks, prominent rift onset unconformities, synrift sediment wedges). While the South Atlantic margins do show a number of large, late-rift stage faults, these faults generally fail to explain the regional development of synrift accommodation across the region and are thus relatively unimportant in the extension process. The fundamental implication of this model is that rifts characterized by large normal fault systems controlling synrift accommodation, such as in the Basin and Range province, East African rift system, North Sea Basins, Rhinegraben and Newark Basin, will not proceed to breakup. By analogy, because of the existence of large fault blocks in the northern Red Sea, the implication is that it should never proceed to sea-floor spreading. The challenge is to understand why different portions of the Red Sea show different stages in the development of a spreading centre during continental rifting. Cochran (2005) based his model on data from the northern Red Sea and assumes that the structural framework deduced in the north continues to the south where sea-floor spreading exists. However, it is entirely conceivable that the two areas, separated by a major tectonic boundary (the ‘Zabargad fracture zone’ Crane & Bonatti 1987), have

developed differently and observations in the northern Red Sea cannot be extrapolated to the south where few data are available for comparison. Thus, the difference between the two regions may be owing to a fundamental difference in lithosphere rheology, explaining why sea-floor spreading developed in the southern, but not the northern, Red Sea. The counter-argument is that the entire Red Sea is developing similarly, but rifting in the northern and central Red Sea simply has experienced insufficient extension to develop into sea-floor spreading at this time. In this paper, we will compare and contrast the two models presented above by using existing geophysical data from the region to understand why different portions of the Red Sea show different stages in the development of a spreading centre within a continental rift. For a combination of logistic and political reasons, most of the available geophysical data are concentrated in the northern Red Sea. The morphology and other geophysical parameters for much of the central and southern Red Sea are only known in general terms with Backer et al.’s (1975) single-beam sonar survey still the primary data source. First, we briefly review the prerift development of the Red Sea area and how the pre-existing structure influenced the form and distribution of rifting and the influence of the Afar plume. Finally, we will discuss the structure of the northern and central Red Sea and develop a model for how possible differences between these two areas may influence the future development of the Red Sea.

Prerift constraints The Red Sea rift propagates through the late Precambrian (Pan-African) Afro-Arabian shield. The basement rocks exposed on the adjacent rift shoulders consist of a variety of granitic, metamorphic, and mafic igneous rocks comprising a number of different terrains separated by suture zones and interpreted as a series of collapsed and accreted island arcs that were assembled between 900 and 600 Ma (e.g. Camp 1984; Stoeser & Camp 1985; Vail 1985; Kroner et al. 1991). Two nearly north– south-trending late Precambrian suture zones within the Pan-African lithosphere, the Baraka and Onib-Hamisana sutures, intersect the Red Sea Rift near 188N and 228N, respectively. At both locations there are pronounced kinks in the overall N308E trend of the rift and a corresponding offset of the rift axis (Dixon et al. 1987; Crane & Bonatti 1987; Bosworth et al. 2005) (Fig. 1). The hinge zone of the Mediterranean margin is located just to the north of the Gulf of Suez (Steckler & ten Brink 1986). A change in Eurasian– African plate motions in the late Santonian resulted

THE TRANSITION FROM RIFTING TO DRIFTING

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Fig. 1. Red Sea location map. Bathymetry and topography are from Smith & Sandwell (1997), contoured at 500 m intervals. The maps shown in Figures 2 and 4–9 are of the northern region and are outlined in red. The location of Figure 11 and 12 of the central and southern Red Sea region are also outlined in red. All images of data grids assume an artificial shadowing from the NE.

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in transpressional deformation and the development of ‘Syrian Arc’ structures along the Mediterranean margin (Moustafa & Khalil 1990; Guiraud & Bosworth 1997, 1999; Guiraud et al. 2005) with more minor deformation occurring farther south in the region of the future Gulf of Suez and northern Red Sea (Bosworth et al. 1999). These structures were reactivated, again in transpression, in the late Eocene (Moustafa & Khalil 1994; Guiraud et al. 2001, 2005). The combination of the Syrian Arc structures and the transition from continental to stronger oceanic lithosphere defined the nature and location of the northern termination of the Red Sea rift. Upon reaching the main Syrian Arc deformation, Red Sea rift extension, which had been confined to a reasonably narrow rift, spread out over a broad region from the Bitter Lakes west to Cairo (Bosworth & McClay 2001; Bosworth et al. 2005). The Manzala Rift continues north from Cairo under the Nile Delta (Bosworth & McClay 2001), but does not appear to penetrate far into oceanic Mediterranean lithosphere. Later in the mid-Miocene, the Dead Sea rift developed in continental lithosphere just landward of the Mediterranean hinge zone (Steckler & ten Brink 1986), more or less cutting off the Gulf of Suez. A series of rifts, apparently related to rifting between Africa and India/Madagascar developed in the late Jurassic and early Cretaceous across the Horn of Africa and extended into Yemen and Somalia (Bott et al. 1992; Redfern & Jones 1995; Ellis et al. 1996; Brannan et al. 1997). Further south, the similar age Anza Rift extends from the Kenyan continental margin into SE Sudan where there are a series of NW–SE-trending Late Jurassic rifts referred to as the Central African rift system (Bosworth 1992). Later in the Cretaceaous (Barremian–Albian) more northern portions of Sudan, as well as the Sirte Basin in Libya and Abu Gharadiq Basin in the western Desert of Egypt, were subjected to extensional deformation (Bayoumi & Lofty 1989; Guiraud & Maurin 1992; Guiraud 1998). These extensional tectonic events did not extend across the future location of the Red Sea rift and do not appear to have had a significant impact on the lithosphere thickness and thermal structure. Heat flow measurements in Egypt at distances of more than about 30 km from the present Red Sea rift are in the range of 35– 55 mW m22 (Morgan et al. 1985; Boulos 1990), consistent with a stable tectonic regime and thick lithosphere. The few published heat flow measurements from the Arabian portion of the shield are also in the range of 35 –45 mW m22 (Gettings 1982; Gettings & Showail 1982; Gettings et al. 1986). Information on the crustal and lithospheric thickness and structure is limited. Knopoff &

Fouda (1975) utilized Raleigh wave dispersion to deduce a crustal thickness of about 37 km beneath Arabia. Their results also show a pronounced lowvelocity layer, the top of which is at a depth of 100–140 km. The US Geological Survey acquired a 1000 km-long seismic refraction profile across Arabia from near Riyadh to the Farasan Islands at 168N in the southern Red Sea (Healy et al. 1982); crustal thickness was estimated to be in the range of 37 –45 km away from the Red Sea rift. The crust appears to thin by about 5 km approaching the rift, the crust consisting of two layers, each layer being about 20 km thick. The average velocities of the two layers are 6.3 and 7.0 km s21. The crust– mantle boundary appears to be transitional over a distance of 2 –8 km and the mantle velocity is in the range of 8.0–8.2 km s21. Makris et al. (1979, 1983) and Rihm et al. (1991) presented three seismic transects extending inland for about 150 km from the coast in the northern Red Sea. The transect on the Arabian side, at 268N, was interpreted as showing a 40 km-thick crust at distances greater than about 100 km from the coast with a gradual decrease in crustal thickness toward the rift (Makris et al. 1983). The two lines on the African side, at 278N and 268N, show a thinner 30–35 km-thick crust away from the rift and a much more abrupt thinning of the crust approaching the Red Sea (Makris et al. 1979; Rihm et al. 1991), although neither of these seismic lines are reversed.

Initiation of rifting The onset of rifting in the Red Sea was preceded by massive volcanism in Ethiopia and southern Yemen (e.g. Mohr 1983). Based on stratigraphic relationships and 40Ar/39Ar dating, the majority of the basalts were erupted over a short (c. 1.5 Ma) time period at approximately 30 Ma (Baker et al. 1996; Hofmann et al. 1997; Coulie et al. 2003), although activity continued to about 25 Ma (Ukstins et al. 2002). Rhyolitic volcanism began at about the same time and continued until about 26.5 Ma (Baker et al. 1996). This short burst of flood volcanism has been attributed to impingement of the Afar plume head on the lithosphere (e.g. Richards et al. 1989). Bosworth et al. (2005) emphasize the observation that the extensive volcanism from 31 to 25 Ma was not associated with significant extension. Although the massive approximately 30 Ma volcanism was centred in Ethiopia and southern Yemen, coeval basalt flows have been reported as far north as about 188N in Sudan. Drury et al. (1994) describe basaltic flows reaching a thickness of 600 m that were extruded onto a well-developed laterite palaeosol in northern Eritrea. Dating of

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these flows using 40Ar/39Ar techniques gave ages of between 32.9 and 28 Ma. Further north, Kenea et al. (2001) describe basalt flows overlain by rhyolite from the Odi and Adar Ribad basins of the Derudeb region of eastern Sudan. 40Ar/39Ar dating of the volcanics shows that they were erupted over a short period from about 31 to 29 Ma. In both regions, stratigraphic relations indicate that extension and uplift related to Red Sea rifting post-dates the volcanism (Drury et al. 1994; Kenea et al. 2001). Late Oligocene volcanism has not been reported from farther north in northern Sudan or Egypt. An episode of dyke intrusion (Blank 1977) dated at 24–21 Ma on the basis of K–Ar and 40Ar/39Ar data is recorded along the entire length of the Red Sea on the Arabian margin (Bartov et al. 1980; Eyal et al. 1981; Feraud et al. 1991; Sebai et al. 1991). This episode was more intense in the southern Red Sea, being accompanied in places by intrusion of plutonic bodies of granitic–gabbroic composition (Pallister 1987; Sebai et al. 1991; Davison et al. 1994). This event extends for a distance of 1700 km along the trend of the Red Sea with no discernable temporal pattern (Sebai et al. 1991; Bosworth et al. 2005). Although episodic volcanism has continued to the present in Afar and in southern Yemen (e.g. Zumbo et al. 1995a, b; Coulie et al. 2003), there has been little significant volcanism within the Red Sea rift since the Early Miocene dyking event (Coleman et al. 1983; Coleman 1984; Coleman & McGuire 1988; Sebai et al. 1991). The Early Miocene dyke event appears to mark the initiation of extensional rifting in the Red Sea. Bosworth et al. (2005) documented the earliest definitive synrift sediments along the length of the Red Sea, which consistently fell in the range of 21– 24 Ma. Fission track data from Egypt (Omar et al. 1989; Omar & Steckler 1995), SW Saudi Arabia (Bohannon et al. 1989) and Eritrea (Abbate et al. 2002; Ghebreab et al. 2002) also consistently give 21 –24 Ma for the onset of significant denudation (and, thus, by inference, rift flank uplift). Omar & Steckler (1995) also suggest a pulse of uplift at about 34 Ma, but there is no dated sedimentary evidence of significant uplift and erosion at this time (Bosworth et al. 1998). Bosworth et al. (2005) suggest that this event may be related to Syrian Arc tectonics.

Morphology and structure of the northern Red Sea The Red Sea consists of narrow continental shelves and a broad ‘main trough’ at a depth of 400–1200 m. The shelf broadens in the central Red Sea, and in the southern Red Sea the shelf consists of coral reefs that completely choke the main trough.

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From 158N to 198300 N in the southern Red Sea, the main trough is cut by an ‘axial trough’ reaching depths of over 2000 m and floored by oceanic basalt formed at a well-developed mid-ocean ridge spreading centre (Roeser 1975; Cochran 1983; Miller et al. 1985; Garfunkel et al. 1987). The spreading centre becomes discontinuous north of 198300 N and, with the possible exception of the Mabahiss Deep pull-apart basin centred at 258200 N, 368100 E, there is no evidence of oceanic crust north of about 238N. The continental margins of the northern Red Sea consist of steep, faulted scarps with 400– 1000 m of relief that truncate the narrow continental shelves (Fig. 2) and appear to mark the edge of the actively extending rift. Rotated basement fault blocks forming elongated hills and ridges are found 2– 10 km inland from the coast. Mid-Miocene reef platforms are consistently found at an elevation of 200–300 m mantling the seaward edges of these basement fault blocks along both the African (Bosworth et al. 1998) and Arabian (Purser & Ho¨tzl 1988) margins. These fault blocks are inactive and have been marked by uplift and non-deposition since the late Miocene (Montentat et al. 1988). MidPleistocene reef terraces that have been uplifted 20– 50 m are consistently found between the exposed basement fault blocks and the coast. This pattern of uplifted reefs and fault blocks documents a progressive narrowing of the actively extending rift as fault blocks become inactive and are transferred to the uplifting rift flanks. The bathymetry of the active, main trough of the northern Red Sea rift consists of a series of terraces stepping down to an axial depression at a depth of 1100–1200 m (Martinez & Cochran 1988; Cochran 2005). The terraced nature of the bathymetry is clear on the bathymetry profiles shown in Figure 3. Sediment deformation within the axial depression is much more intense than in the remainder of the Red Sea (Knott et al. 1966; Guennoc et al. 1988; Martinez & Cochran 1988), implying that tectonic activity and extension is presently concentrated predominantly in the axial depression. However, the slopes separating terraces often appear to be faulted (Cochran 2005). Also, although earthquakes with ML.3 are concentrated in the axial depression, smaller earthquakes are distributed throughout the northern Red Sea (Badawy 2005). These two observations imply that some extension still occurs within the main trough away from the axial depression. Free-air gravity anomaly highs are located on the seaward edges of the terraces on Figure 3. This is a systematic observation throughout the northern Red Sea (Martinez & Cochran 1988; Cochran 2005). Martinez & Cochran (1988) argue that the density anomalies responsible for the

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Fig. 2. Northern Red Sea bathymetry map contoured at 100 m-intervals. Data sources for the bathymetry are discussed by Cochran (2005). Projected shipboard bathymetry, gravity and magnetic data profiles used in the Parker and Werner inversions are shown as blue lines (labelled A– E). Triangles mark location of volcanoes identified from the magnetic anomalies. Triangles are blue when the identification is confirmed by bathymetry and/or gravity data and white when based only on interpretation of the magnetic data.

gravity anomalies are sufficiently large to reflect basement relief. As free-air gravity anomalies in the northern Red Sea form a series of highamplitude highs and lows elongated subparallel to the strike of the rift (Fig. 4), Martinez & Cochran (1988) interpreted them as arising from a series of large rotated fault blocks. Gaulier et al. (1988a) conducted a seismic survey in the Egyptian half of the northern Red sea that included 13 expanding spread profiles (ESPs). One of these was located less than 20 km

from the coast. It showed a 9.1 km-thick crust with the Moho at a depth of 16.7 km. Depth to the Moho was in the range of 13–14.7 km on nine seismic lines located more than 20 km from the coast that gave either a mantle reflection or a phase interpreted as a mantle reflection. On seven of these lines, the Moho was located at 13.0–13.5 km below the sea surface. Two other lines showed a reflection at 10.5 km below the sea surface that could be either the Moho or an intracrustal reflector (Gaulier et al. 1988a). The final ESP was located over the axial

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Profile A +50

500 free-air gravity

mgal

nT

–50

0 magnetics

bathymetry

0 metres 5000

inverted gravity

Werner magnetic depth estimates 0

25 km

Fig. 3 Single channel reflection seismic (SCS) and projected bathymetry, gravity and magnetic profiles across the northern Red Sea. Parker and Werner deconvolution inversion results show the sediment–basement contact from inversion of gravity anomalies (bold line) and magnetic depth estimates (grey circles), respectively. The Werner circles are the individual depth solutions with circle size being proportional to the magnetic susceptibility. Depth estimates that tend to cluster vertically beneath the true location of the causative body with magnetic basement being interpreted as the top of the vertical ‘streaks’ of the depth solutions. Single channel seismic data (SCS), where available, are also shown. SCS section length is located on each profile. Profiles A–E are located in Figure 2. Profile C is across the Conrad Deep, and shows velocities as a function of depth obtained from the expanding spread seismic experiment of Gaulier et al. (1988a, b). Note terraced nature of the bathymetry and the presence of gravity highs on the seaward edge of the terraces. Clustering of Werner solutions on the apexes of the rotated fault blocks gives confidence in the gravity and magnetic inversions.

depression at the location of Conrad Deep, which Cochran et al. (1986) argue has been the site of recent large intrusions. This ESP gave an unusual crustal structure with no definable Moho. The profile shown in Figure 3C is perpendicular to the ESP lines of Gaulier et al. (1988a). The location at which the profile crosses each ESP line is noted. A synthetic cross-section from Gaulier et al. (1988a) developed from those ESP lines is also shown in Figure 3C. The flat Moho determined from the seismic data is very obvious in the profile. The crustal thickness on the ESPs varied from 5.1 to

8.3 km (Gaulier et al. 1988a) with changes in crustal thickness largely corresponding to relief on the top of the crust. The nearly flat Moho requires that the faults bounding these large crustal fault blocks sole out into a horizontal shear zone in the lower crust or at the Moho. The nearly flat Moho also implies that extension has been distributed across the water-covered portion of the rift. Magnetic anomalies throughout the northern Red Sea north of the Mabahiss Deep pull-apart basin (centred near 258200 N, 368100 E, Fig. 2) are all dipolar anomalies (Fig. 5), implying compact,

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Fig. 3. (Continued )

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Fig. 3. (Continued )

localized sources (Cochran 2005). These anomalies have been interpreted as arising from individual volcanoes or intrusions (Cochran et al. 1986; Guennoc et al. 1988; Martinez & Cochran 1988; Cochran 2005). This interpretation has been documented in some cases by observation of a volcanic edifice

associated with the magnetic anomaly (Guennoc et al. 1988; Martinez & Cochran 1988; Cochran 2005). The dipolar anomalies in the northern Red Sea all imply normally magnetized sources, suggesting that the volcanoes all formed within the past 780 ka.

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Fig. 3. (Continued )

Volcanoes in the northern Red Sea are systematically found at two locations; at the top of scarps bounding the axial depression and along the edge of terraces in the main trough, particularly where the terraces intersect accommodation zones (Cochran 2005). In both cases it appears that melt ascended along faults (Cochran et al. 1986; Cochran 2005). Volcanoes along the edge of the axial depression occur in pairs on opposite sides of the axial depression immediately adjacent to a bathymetric ‘deep’ within the axial depression with one such pair within each rift segment (Cochran 2005). In one segment, an additional volcano is located on

the floor of the axial depression within Shaban Deep near 268150 N, 358220 E (Pautot et al. 1984, 1986) (Fig. 2). Martinez & Cochran (1988) and Cochran (2005) interpret the development of volcanoes at the axial depression as the first step in the establishment of a magmatic spreading centre within the northern Red Sea. According to their model, lithospheric extension and thinning has become concentrated under the axial depression over the past several million years (Buck et al. 1988; Martinez & Cochran 1989). They hypothesize that melt developed under a rift segment is focused along the segment,

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275

Fig. 3. (Continued )

and then ascends to the surface, originally along faults to form the volcanoes perched on the edge of the axial depression and then directly to form volcanoes in the axial depression, such as observed at Shaban Deep (Pautot et al. 1984; Maury et al. 1985). These volcanoes develop into isolated cells of sea-floor spreading, as observed in the central Red Sea (Cochran 1983; Pautot 1983; Bonatti 1985; Bicknell et al. 1986). These cells are postulated to propagate and grow until they form a continuous axis. The southern, central and northern Red Sea, therefore, represent different stages of

this process of developing an oceanic spreading centre within a continental rift. If an oceanic axis is developing in the northern Red Sea, as proposed by Martinez & Cochran (1988) and Cochran (2005), the structure of the extended continental crust and the transition to oceanic crust differs significantly from that observed at some other margins including the wellstudied margin west of the Iberian Peninsula and the Exmouth Plateau of NW Australia, the Iberian margin being considered the type-example of a nonvolcanic rifted margin (e.g. Boillot & Froitzheim

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Fig. 4. Northern Red Sea free-air gravity anomaly map contoured at 10 mGal intervals. Data sources for the gravity are discussed by Cochran (2005). Projected shipboard bathymetry, gravity and magnetic data profiles used in the Parker and Werner inversions are shown as blue lines (labelled A– E). Triangles mark location of volcanoes identified from the magnetic anomalies. Triangles are blue when the identification is confirmed by bathymetry and/or gravity data and white when based only on interpretation of the magnetic data. Note the sequence of NNE– SSE linear trending gravity anomalies across the entire region, suggestive of the location of large, rotated fault blocks.

2001; IODP 2001; Whitmarsh et al. 2001). In the Red Sea, thinned continental crust with a thickness of at least 5 km extends to the site of volcanic activity at the axial depression (Gaulier et al. 1988b). There is no evidence of a zone of exposed mantle rocks, as has been observed at the Iberian margin (e.g. Pickup et al. 1996; Chian et al. 1999). This difference appears to be reflected in the observation that the distance from basically unrifted

continental crust to the axial depression in the northern Red Sea is about 80 km, much less than the average width of over 200 km for this region at the Iberian margin (e.g. Dean et al. 2000). In these respects, the northern Red Sea appears to resemble the Goban Spur margin of the UK, where thin continental crust with a thickness of about 7 km directly abuts oceanic crust (Peddy et al. 1989; Horsefield et al. 1993) and the width

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277

Fig. 5. Total intensity magnetic anomaly map of the northern Red Sea contoured at 50 nT intervals. Data sources for the magnetics are discussed by Cochran (2005). Projected shipboard bathymetry, gravity and magnetic data profiles used in the Parker and Werner inversions are shown as blue lines (labelled A– E). Triangles mark location of volcanoes identified from the magnetic anomalies. Triangles are blue when the identification is confirmed by bathymetry and/or gravity data and white when based only on interpretation of the magnetic data. While large, isolated magnetic anomalies exist in the central region of the northern Red Sea, well developed, linear magnetic anomalies as part of a sea-floor spreading sequence do not exist.

of extended continental crust is also about 80 km (Dean et al. 2000).

Discussion Crustal extension in the northern Red Sea The model for the transition from rifting to sea-floor spreading in the Red Sea developed by Martinez &

Cochran (1988) and Cochran (2005) assumes that crustal extension in the northern Red Sea during the rifting stage is accomplished mechanically by rotational faulting of crustal blocks accompanied by ductile flow in the lower crust. The upper surface of the crust is thus postulated to consist of a series of large-relief rotated fault blocks. This assumption is based on two observations. First, in the Gulf of Suez, the presence of large faulted

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blocks has been confirmed by seismic imaging and well data (e.g. Patton et al. 1994; Bosworth 1995; Salah & Al Sharhan 1996; Bosworth & McClay 2001). As the Gulf of Suez was a northern continuation of the Red Sea until the establishment of the Dead Sea transform in the middle Miocene caused extension to be essentially abandoned within the Gulf of Suez, it is assumed that they had a common evolution up to that time. The second observation is the presence of large-amplitude, elongate gravity highs and lows trending subparallel to the trend of the rift (Martinez & Cochran 1988). Gravity highs are systematically located on the seaward edges of bathymetric terraces (Fig. 3) (Cochran 2005) and are considered to be sourced by basement relief (Martinez & Cochran 1988). We suspect that large fault blocks characterize northern Red Sea extension and will investigate this further by inverting the potential field data across the Red Sea (both gravity and magnetic data). Specifically, we will downward continue the Bouguer gravity anomaly a depth of 5 km, which is the average depth below sea level of the basement (Gaulier et al. 1988b), and expand the resulting equivalent mass layer to determine if the required relief is reasonable and compatible with other geological and geophysical constraints. We will also use Werner deconvolution of magnetic data, which will give estimates to the depth of magnetic source. These estimates should help constrain the gravity inversion by identifying areas of steep basement slopes (i.e. the edges of rift blocks). The robustness of the technique is such that no reduction to the magnetic pole is required and the inversion is applicable for either remnant or induced magnetizations. Gravity inversion. Figure 6 shows a Bouguer gravity anomaly map of the northern Red Sea calculated using a density contrast of 1370 kg m23. This density contrast was chosen to eliminate the gravity effect of bathymetric relief, which arises from the density contrast between sea water and sediments (assumed to have an average density of 2400 kg m23 (Martinez & Cochran 1988). We calculated the gravity effect of the sediment–water interface using four terms in the Parker (1973) fast Fourier transform technique. The resulting grid was sampled at the locations of the free-air gravity measurements, these values subtracted point-bypoint from the free-air anomalies, and the resulting point Bouguer anomalies gridded in the same manner as was done with the point free-water gravity anomaly values. The Bouguer anomalies shown in Figure 6 have also had a regional planar trend removed. The Bouguer anomaly has greatly reduced the amplitude of the large negative gravity anomaly over the axial depression (Figs 4 & 6), but

the lineated gravity highs and lows observed in the free-air gravity anomalies are still present as the dominant feature in the Bouguer anomaly map. As the Moho in the northern Red Sea appears to be basically flat (Gaulier et al. 1988b) (Fig. 3C), we will assume that the Bouguer gravity anomalies result from relief on the sediment –basement interface and calculate the amount of relief on that surface required to produce them. We downward continued the Bouguer anomalies from the sea surface to a depth of 5 km, the average depth to top of the crust estimated from the seismic lines of Gaulier et al. (1988b). Since downward continuation is inherently unstable at short wavelengths, we first applied a low-pass filter to the gravity data that removed all wavelengths less than 20 km and passed wavelengths greater than 30 km. The downward continued gravity anomalies were converted to an equivalent mass layer and relief on that surface determined assuming a density contrast of 400 kg m23 across the sediment –crust interface. We then calculated the gravity effect of the resulting basement surface using four terms in the Parker (1973) technique, sampled the resulting grid at the locations of gravity measurements, subtracted these values from the Bouguer anomalies and gridded the resulting residual anomalies in the same manner as had been done for the Bouguer anomalies. The residual anomalies were filtered, downward continued and converted to basement relief in the manner described above. This residual basement relief was added to that determined in the first step to produce the calculated depth to the top of the crust shown in Figure 7. As would be expected, the calculated relief resembles a filtered version of the gravity anomalies with a series of elongated highs and lows. Because of the filtering necessary to stabilize the inversion, the result has a rounded appearance rather than the asymmetric rectilinear shape characteristic of fault blocks (i.e. the apexes of the blocks have been smoothed). The gravity effect of the calculated basement relief is shown in Figure 8. The filtered residuals between the Bouguer gravity anomalies (Fig. 7) and the calculated gravity (Fig. 8) are consistently less than about 2–3 mGal except in the vicinity of Mabahiss Deep (Fig. 9). The low residual values in Figure 9 indicate that the predicted basement relief explains basically all of the gravity ‘energy’. The inferred top of crust is generally at a depth of 2 –8 km below sea level with the relief of about 2–6 km between adjacent highs and lows. The computed basement surface nearly reaches the sea surface near 268160 N, 348470 E where a bathymetric peak interpreted as a volcano (Martinez & Cochran 1988) comes to within 300 m of the surface on a bathymetric profile. The only location where the

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Fig. 6. Northern Red Sea Bouguer gravity anomaly map contoured at 10 mGal intervals and assuming a Bouguer reduction density of 1370 kg m23 (topographic density of 2400 kg m23 and water density of 1030 kg m23). Projected shipboard bathymetry, gravity and magnetic data profiles used in the Parker and Werner inversions are shown as blue lines (labelled A –E).

computed top of the crust comes above the sea surface is near 268440 N, 358260 E, where it extends a few hundred metres above sea level. This is the location of a large gravity high (Figs 4 & 6) located in an area of sparse data (Cochran 2005), so it is unclear how well constrained this anomaly is by the data. Although our 5 km average basement depth should perhaps have been slightly deeper, the gravity inversion produces basement relief that is reasonable and consistent with available seismic results supporting the conclusion of Martinez &

Cochran (1988) that the northern Red sea is underlain by a series of large crustal fault blocks. The Werner deconvolution method. Magnetic data are often used to estimate the depth of magnetic basement in sedimentary basins and passive continental margins (e.g. Grant & West 1965; Klitgord & Behrendt 1979). This is possible because, in general, the sediments within a basin have relatively weak magnetic susceptibilities compared with those of crystalline–volcanic basement at the base of the

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Fig. 7. Predicted, relative sediment– basement contact obtained by inverting the Bouguer gravity of the northern Red Sea region contoured at 1 km intervals. Maximum range in basement relief is 2 –8 km below sea level (eastern edge of region between Profiles D and E) with an average relief of about 2– 6 km between adjacent highs and lows. A constant sediment–basement density contrast of 400 kg m23 was assumed. Projected shipboard bathymetry, gravity and magnetic data profiles used in the Parker and Werner inversions are shown as blue lines (labelled A–E).

sediment column. Werner (1953) recognized that the usual approach to magnetic interpretation, that of analysing discrete anomalies using such parameters as maxima, minima, inflection points or other intrinsic properties of the observed magnetic anomaly profile, was complicated because of the interference of adjacent anomalies and the effect of noise (e.g. diurnal variations, non-two-dimensionality and induced v. remnant magnetism). For this reason, our Werner deconvolution algorithm uses simple

models for the source and a quadratic form for the source/noise interference to determine the magnetization properties of the causative bodies. The basic assumption of the approach is that all magnetic anomalies are the result of either a sequence of dykes or an interface between juxtaposed half-spaces of different magnetic susceptibility. In particular, the strike length and depth of the causative body are assumed to be infinite while the width of the body is assumed to be either: (1) finite, representing a

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Fig. 8. Calculated gravity effect of the predicted basement relief shown in Figure 7. This is used to ascertain how much energy is being ‘captured’ by the predicted relief. A constant sediment–basement density contrast of 400 kg m23 was assumed. Projected shipboard bathymetry, gravity and magnetic data profiles used in the Parker and Werner inversions are shown as blue lines (labelled A–E).

dyke; or (2) approaching zero such that it represents an interface between two regions of differing magnetic susceptibility. The robustness of the technique is such that no reduction to the pole is required and works effectively with both induced and remnant magnetizations. We have used five dip lines for gravity and magnetic inversion analysis (Fig. 3). To interpret Werner estimates requires knowledge of the role of the data window. For a given pass and data spacing, 11 points are used to estimate a depth, magnetization intensity and magnetization dip. As the window is made wider, deeper magnetic

sources can be potentially recognized. Depth estimates for a ‘true’ shallow source will usually be biased to deeper depths for increasing window lengths. If the deconvolution is successful in defining a ‘real’ magnetic body, then depth estimates should define either the edges of the causative body or the depth range of an interface or the upper boundary of a dyke. Shown in Figure 3A– E are the results of the Werner depth estimates to magnetic source compared with the inversion of the free-air gravity anomalies described above. In general, it is not possible to see any structure below the evaporites

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Fig. 9. Filtered residual gravity anomaly obtained by subtracting the gravity effect of the predicted basement relief from the Bouguer gravity anomaly. Contour interval is 10 mGal. Projected shipboard bathymetry, gravity and magnetic data profiles used in the Parker and Werner inversions are shown as blue lines (labelled A –E).

as imaged by the single-channel seismic data. Magnetic estimates are of two types: (1) volcanic intrusions (e.g. Fig. 3D); or (2) the apex of tilted fault blocks, consistent with rift structure across the northern Red Sea based on crustal refraction velocities determined from expanding spread seismic profiles (Fig. 3C) (Gaulier et al. 1988a, b). There is excellent agreement between the smoothed, sediment –basement relief and geometry, and the depth to the fault block apex using the magnetic data. We conclude from the gravity and magnetic inversion that the northern Red Sea consists primarily of extended continental crust characterized by a series of large rotated crustal fault blocks.

When will sea-floor spreading develop in the northern Red Sea, if ever? Martinez & Cochran (1988) and Cochran (2005) have proposed a model for the development of an oceanic spreading centre within a continental rift based on observations in the northern Red Sea. They suggest that rift development occurs through rotation of large crustal fault blocks along listric faults that sole out into a zone of plastic creep, resulting in a flat Moho and high upper crustal relief, as observed in the northern Red Sea (Fig. 3). This initial, broadly distributed extension eventually becomes centred at the rift axis leading

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Fig. 10. Satellite free-air gravity map of the central Red Sea using the 1  1 min world grid from Smith & Sandwell (1997; version 15.1). Contour interval is 10 mGal.

to rapid lithospheric thinning and melt generation (Buck et al. 1988). Melt formed within individual rift segments is focused to a location within the segment to form small axial volcanoes at the axis (Cochran 2005). With continued extension and magmatism, these volcanoes become cells of seafloor spreading as observed in the central Red Sea (Degens & Ross 1969; Backer et al. 1975; Bonatti 1985; Bicknell et al. 1986). The deeps then propagate along axis and coalesce to form a continuous axis, as is observed in the southern Red Sea

(Roeser 1975; Cochran 1983). According to this model, the northern Red Sea is on the verge of replacing horizontal translation with focused mantle upwelling and organized sea-floor spreading. A corollary of this hypothesis is that the tectonic processes that have shaped the northern Red Sea have been active throughout the entire Red Sea and that the tectonic framework deduced in the north continues to the south. Thus, it is expected that the large rotated fault blocks found in the Gulf of Suez and the main trough of the northern

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Red Sea are also present in the main trough of the central and southern Red Sea. There is presently insufficient data from the central and southern Red Sea to construct free-air or Bouguer gravity maps based on surface ship data. However, freeair anomaly maps derived from satellite altimetry data (Figs 10 & 11) do not show the prominent linear rift-parallel gravity highs and lows that characterize the northern Red Sea. This does not appear to be due to differing resolution in the surface and altimetry data because the linear gravity anomalies characteristic of the northern Red Sea are evident in satellite altimetry data (Cochran 2005). In addition, a reflector interpreted

as main trough basement by Izzeldin (1987) on a 48-channel seismic reflection line across the Red Sea between 208N and 218N does not show the presence of high-relief fault blocks. In marked contrast to the structure observed in the northern Red Sea, Karner et al. (2003) and Karner (2005) have shown that the West African and Brazilian passive continental margins are characterized by the stacking of regional synrift sag basins, the amplitude and distribution of which are inconsistent with the minor amounts of brittle deformation mapped from seismic sections. The sag basins exhibit none of the diagnostic features expected of synrift extensional systems

Fig. 11. Satellite free-air gravity map of the southern Red Sea using the 1  1 min world grid from Smith & Sandwell (1997; version 15.1). Contour interval is 10 mGal.

THE TRANSITION FROM RIFTING TO DRIFTING

(e.g. normal faults that control accommodation, rotation of crustal blocks, prominent rift onset unconformities, synrift sediment wedges). Large post-rift subsidence in regions characterized by relatively minor synrift faulting is most easily explained by depth-dependent extension that is partitioned vertically across a zone of decoupling separating a relatively non-deforming upper crust (upper plate) from a ductile-deforming lower crust and lithospheric mantle (lower plate; Karner et al. 2003). A fundamental implication of this model is that rifts characterized by large offset fault systems, that is faults that generate synrift accommodation such as in the Basin and Range province, East African rift system, North Sea Basins, Rhinegraben and Newark Basin, will not proceed to breakup. Thus, according to this model, because of the existence of large fault blocks in the northern Red Sea, it should not proceed to sea-floor spreading. A possible reason for different responses to lithospheric extension in the northern and southern Red Sea is alteration and weakening of the lithosphere as a result of the impact of the Afar plume head at about 30 Ma. Although the massive burst of volcanism that accompanied impingement of the plume head on the lithosphere is concentrated in Ethiopia and southern Yemen, significant coeval volcanism is reported as far north as 188N in the Red Sea hills of Eritrea and Sudan (Drury et al. 1994; Kenea et al. 2001). Drury et al. (1994) describe basalt flows reaching a preserved thickness of 600 m that were erupted 32.9 and 28 Ma in northern Eritrea. Kenea et al. (2001) describe exposures of identically aged basalt flows with an aggregate thickness of at least a few hundred metres from more than 200 km further north in Sudan. They argue that flood basalts were erupted across an area approximately 1000 km in diameter between about 31 and 28 Ma, prior to onset of any known extension or plateau uplift. The ellipse drawn by Kenea et al. (2001) to illustrate the extent of the Afar flood basalt province extends to almost 208 in the Red Sea and corresponds to the portion of the Red Sea with a well-developed mid-ocean ridge (Roeser 1975; Cochran 1983). Further evidence that the lithosphere in the southern Red Sea may have been weakened and made more susceptible to stretching and melting is given by along-strike variations in the character of the 24 –21 Ma dyke event that occurred along the entire length of the rift in Arabia (Blank 1977; Feraudet al. 1991) and into the Sinai (Bartov et al. 1980), which marks the beginning of extension and synrift sediment deposition (Bosworth et al. 2005). North of Jeddah, the dykes crop out discontinuously and their presence and continuity along strike is determined primarily from linear,

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high-amplitude magnetic anomalies (Blank 1977; Eyal et al. 1981). South of Jeddah, the dyke swarm becomes much more intense (Coleman et al. 1983; Coleman & McGuire 1988), crops out over wide areas, and is accompanied by large, intrusive igneous complexes in at least two locations, Al Lith near 218N (Pallister 1987) and Tihama Asir near 178N (Coleman et al. 1979, 1983). The two igneous complexes give an 40Ar/39Ar age range that coincides with those of the dykes (Feraud et al. 1991; Sebai et al. 1991). Courtillot et al. (1999) have argued that there is a causal link between the impingement of plume heads on the lithosphere and continental breakup to form new ocean basins. In the case of the Red Sea, it is possible that Afar plume head affected the lithosphere north to about 21– 228N, so that when extension began, 5– 7 Ma later, the region of weakened lithosphere responded in a different manner than did the unmodified lithosphere of the northern Red Sea.

Conclusions The northern Red Sea has been undergoing continental rifting and extension since 24– 21 Ma. The crust within the rift is now 5 to 8 km thick with a nearly flat Moho (Gaulier et al. 1988b). From our gravity and magnetic analysis, sediment thicknesses range from 2–8 km with an average of approximately 5 km. Martinez & Cochran (1988) have hypothesized that large-amplitude, linear gravity highs and lows observed in the main trough of the northern Red Sea result from relief on a series of large rotated crustal fault blocks making up the upper crust. We have also shown that this hypothesis is consistent with the inversion of the gravity and magnetic data. There is a transition in the northern Red Sea from continental rifting in the north to sea-floor spreading in the south. As the opening rate is twice as great in the south, this is widely interpreted to reflect the northward propagation of organized sea-floor spreading (e.g. Courtillot 1980, 1982; Cochran 1983, 2005; Bonatti 1985). However, gravity anomalies derived from satellite altimetry data do not show large, linear, rift-parallel gravity anomalies in the main trough of the central and southern Red Sea where oceanic spreading centres have developed. This raises the possibility that continental extension in the northern and southern portions of the Red Sea has been accommodated in different manners, with large rotated fault blocks formed in the northern Red Sea and regional synrift sag basins developed in the south. This difference in the crustal architecture, combined with observations of the structure of continental

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margins and failed continental rifts in other parts of the world (e.g. Karner et al. 2003; Karner 2005), allows speculation that depth-dependent extension leading to development of large sag basins is necessary for continental rifting to proceed to continental lithosphere breaching and the development of an oceanic spreading centre. If this hypothesis is correct, then it is possible that an oceanic spreading centre will not develop in the northern Red Sea. A possible reason for the difference in the response of the lithosphere to extension in the northern and southern Red Sea is the thermal and rheological weakening and alteration of the lithosphere resulting from the impact of the Afar plume head. This work was supported by National Science Foundation grants OCE-9819563, OCE-0548812 (J.R. Cochran) and OCE-0425411 (G.D. Karner). The GMT software package (Wessel & Smith 1998) was used extensively in the preparation of figures. LDEO publication 7010.

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R ICHARDS , M. A., D UNCAN , R. A. & C OURTILLOT , V. 1989. Flood basalts and hot-spot tracks: Plume heads and tails. Science, 246, 103–107. R IHM , R., M AKRIS , J. & M O¨ LLER , L. 1991. Seismic surveys in the Northern Red Sea: asymmetric crustal structure. Tectonophysics, 198, 279– 295. R OESER , H. A. 1975. A detailed magnetic survey of the southern Red Sea. Geologische Jahrbuch, 13, 131– 153. S ALAH , M. G. & A L S HARHAN , A. S. 1996. Structural influence on hydrocarbon entrapment in the northwestern Red Sea. AAPG Bulletin, 80, 101–118. S EBAI , A., Z UMBO , V., F ERAUD , G., B ERTRAND , H., H USSAIN , A. G., G IANNERININ , G. & C AMPREDON , R. 1991. 40Ar/39Ar dating of alkaline and tholeiitic magmatism of Saudi Arabia related to the early Red Sea rifting. Earth and Planetary Science Letters, 104, 473–487. S MITH , W. H. F. & S ANDWELL , D. T. 1997. Global sea floor topography from satellite altimetry and ship depth soundings. Science, 277, 1956– 1962. S TECKLER , M. S. & TEN B RINK , U. S. 1986. Lithospheric strength variations as a control on new plate boundaries: examples from the northern Red Sea region. Earth and Planetary Science Letters, 79, 120– 132. S TOESER , D. B. & C AMP , V. E. 1985. Pan-African microplate accretion of the Arabian Shield. Geological Society America Bulletin, 96, 817– 826. U KSTINS , I. A., R ENNE , P. R. W OLFENDEN , E., B AKER , J., A YALEW , D. & M ENZIES , M. A. 2002. Matching conjugate volcanic margins; 40Ar/39Ar chronostratigaphy of pre- and syn-rift bimodal flood volcanism in Ethiopia and Yemen. Earth and Planetary Science Letters, 198, 289–306. V AIL , J. R. 1985. Pan-African (late Precambrian) tectonic terrains and the reconstruction of the Arabian-Nubian shield. Geology, 13, 839–842. W ERNER , S. 1953. Interpretation of aeromagnetic anomalies as sheet-like bodies, Sveriges Geologista Undersok Series C, Arsbok, 43, N: 06, Stockholm, Sweden. W ESSEL , P. & S MITH , W. H. F. 1998. New improved version of Generic Mapping Tools released. Eos, Transactions of the American Geophysical Union, 79, 579. W HITMARSH , R. B., M ANATSCHAL , G. & M INSHULL , T. A. 2001. Evolution of magma-poor continental margins from rifting to seafloor spreading. Nature, 413, 150–154. Z UMBO , V., F ERAUD , G., B ERTRAND , H. & C HAZOT , G. 1995a. 40Ar/39Ar chronology of tertiary magmatic activity in southern Yemen during the early Red Sea–Aden rifting. Journal of Volcanology and Geothermal Research, 65, 265– 279. Z UMBO , V., F ERAUD , G., V ELLUTINI , P., P IGUET , P. & V INCENT , J. 1995b. First 40Ar/39Ar dating on early Pliocene to Plio-Pleistocene magmatic events of the Afar–Republic of Djibouti. Journal of Volcanology and Geothermal Research, 65, 281 –295.

Observations from the Alpine Tethys and Iberia – Newfoundland margins pertinent to the interpretation of continental breakup ¨ NTENER2, L. L. LAVIER3, T. A. MINSHULL4 & G. MANATSCHAL1, O. MU G. PE´RON-PINVIDIC1 1

CGS-EOST, 1 rue Blessig, F-67084 Strasbourg, France (e-mail: [email protected]) 2

Institute of Geological Sciences, University of Bern, Baltzerstrasse 1, CH-3012 Bern, Switzerland

3

University of Texas Institute for Geophysics, Jackson School of Geosciences, Austin, TX, 78759, USA

4

Southampton Oceanography Centre, European Way, Southampton SO14 3ZH, UK Abstract: Although the Iberia–Newfoundland and Alpine Tethys margins are of different age and ultimately had a different fate, they share remarkable similarities. Both pairs of margins show a change from initially distributed and decoupled extension to later localized, coupled and asymmetric extension that results in thinning of the crust and exhumation of subcontinental mantle. The change in the mode of extension together with the localization of deformation reflects an evolution of the bulk rheology of the extending lithosphere. In this paper we summarize the pertinent geological observations for the Iberia–Newfoundland and Alpine Tethys margins. We describe the stratigraphic evolution, the fault geometry, basin architecture, and magmatic and metamophic evolution of the two pairs of margins from initial rifting to final continental breakup. This description forms a basis for understanding the evolution of the bulk rheology and how the various processes interact during progressive lithospheric extension. For the Iberia– Newfoundland and Alpine Tethys margins initial rifting appears to be controlled by inherited heterogeneities and mechanical localization processes, whereas final rifting and lithospheric rupture is controlled by serpentinization, magmatic and thermal weakening. At other margins, these modes may interact in a different way depending on the prerift conditions and the evolution of the rheology during rifting.

In the last decade, studies of rifted margins have benefited from an increasing quantity of highquality data and significant developments in numerical modelling. The ability to use more realistic boundary conditions in numerical modelling not only enables us to investigate the complex interaction between processes controlling rifting but also leads to a demand for more and better datasets. Although in the last years a lot of high-quality geophysical data have become available from distal passive continental margins, direct observations on the subseismic scale from this domain are still the exception. One way forward is to take a comparative approach using direct observations and unlimited sampling on analogous structures preserved on land. Based on this approach, Manatschal & Bernoulli (1998, 1999), Boillot & Froitzheim (2001), Whitmarsh et al. (2001a) and Wilson et al. (2001) were able to demonstrate that marinederived hypotheses of margin development proposed from drilling and seismic surveys can be supported and further developed using observations from the ancient margins exposed in the Alps.

In this paper we present pertinent observations from the Alpine Tethys (Fig. 1) and Iberia – Newfoundland margins (Fig. 2). The aim of this paper is to demonstrate how observations can help to constrain the tectonic and rheological evolution of these magma-poor rifted margins. More generally, this study attempts to discuss how extensional processes evolve during continental rifting and how they interact with sedimentary, hydrothermal and magmatic processes. More detailed observations and explanations, in particular concerning the Alpine examples, are given in Manatschal (2004). A more general description of the physical boundary conditions of the modes of extension discussed can be found in Lavier & Manatschal (2006). In the first part of this paper, the tectonic and palaeogeographic evolution is discussed together with some ideas and models that evolved from the study of the Alpine Tethys and Iberia–Newfoundland margins. Then, a set of fundamental observations of lithospheric extension is presented, which includes geological mapping, structural/sedimentological and petrological observations, marine multichannel

From: KARNER , G. D., MANATSCHAL , G. & PINHEIRO , L. M. (eds) Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 291– 324. DOI: 10.1144/SP282.14 0305-8719/07/$15.00 # The Geological Society of London 2007.

Fig. 1. (a) Palaeogeographic situation of the Alpine Tethys ocean and adjacent margins during Late Cretaceous time (modified after Dal Piaz 1995). Inset map shows the palaeogeographic relation between the Tethys and Atlantic oceanic basins during Late Cretaceous time. Abbreviations: EM, Eastern Mediterranean/Neo-Tethys; MV, Meliata– Vardar Ocean. (b) Tectonic map of the Alps (modified after Polino et al. 1990) showing the distribution of the major palaeogeographic units. Inset map shows the distribution of the palaeogeographic units in Western Europe. (c) Reconstructed palaeogeographic sections across the Alpine Tethys margins; the Grischun Transect modified after Manatschal & Bernoulli (1999); Brianc¸onnais and Dauphine transects after Lemoine et al. (1987); and the Ticino Transect after Bertotti et al. (1993). For the present-day and Late Cretaceous position of the transects see (a) and (b).

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Fig. 3. Conceptual model showing the temporal and spatial evolution relative to a fixed left-hand edge from (a) the prerift situation, to (b) stretching mode to (c) thinning mode to (d) exhumation mode (modified after Lavier & Manatschal 2006). The single sections have been drawn so that the continental crust is balanced approximately. In reality, erosion has to be taken into account, but, given the scale of the model, it can be neglected and in any case during the later stages the continental crust was covered by deep water so no erosion occurred. The contact between the strong and the weak subcontinental mantle corresponds, at least during the initial stage of rifting, to a stable boundary. Therefore, the surface of strong subcontinental mantle is balanced in the four stages, although thermal erosion and serpentinization will affect and change the thickness of the ‘strong’ subcontinental mantle in a final stage of rifting. The top of the asthenosphere, as drawn in the model, corresponds to the 1350 8C isotherm. Consequently, asthenospheric mantle can, by cooling, change into lithospheric mantle. The major uncertainty in the model is related to the spatial and temporal evolution of the thermal structure within the distal margin during final breakup, which is at the moment not resolved by any observational data.

seismic reflection and refraction data, and drilling results. The final sections discuss extensional processes at magma-poor rifted margins, the associated exhumation and magmatic processes, and the rheological evolution of the extending lithosphere deduced from the observations presented in this paper. In order to simplify the description and discussion of the various rift structures representing different domains and stages within the two pairs of margins, all observed sites are located in the conceptual model presented in Figure 3.

From observations to a conceptual model The Alpine Tethys and Iberia– Newfoundland margins: an overview The Alpine Tethys and Iberia–Newfoundland margins result from rifting of a lithosphere previously affected by the Variscan orogeny (Fig. 4b). The Alpine Tethys margins evolved as part of an equatorial Late Triassic –Jurassic rifting system, which extended from the Caribbean to the eastern

Fig. 4. The prerift conditions as constrained from the Alps. (a) Schematic cross-section across the area of future breakup showing the position of the Malenco crust –mantle boundary, the distribution of lithologies in the prerift lithosphere, the position of prerift basins and volcanic systems, and the general rheology at the onset of rifting. (b) Late Triassic restoration of the Mediterranean and central Atlantic realm, and a section through the area of future Jurassic breakup showing the prerift depositional architecture relative to the future Alpine Tethys Ocean (modified from Bosellini 1973 and Laubscher & Bernoulli 1977). (c) The Malenco crust – mantle boundary as exposed in the field and pressure– temperature– time (P –T–t) diagram for the Malenco crust –mantle boundary. The red arrow shows the exhumation path of Permian gabbros welding the crust –mantle boundary and the dashed black line the successive Alpine history (modified after Mu¨ntener & Hermann 2001).

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Mediterranean area (Fig. 1a). The Iberia– Newfoundland margins resulted from Jurassic – Early Cretaceous rifting and opening of the Northern Atlantic (Fig. 2a). Thus, the different segments of the Atlantic –Tethys Ocean system opened at different times and the closure of the Alpine Tethys segment was contemporaneous with seafloor spreading in the Central and North Atlantic. A wealth of data have been collected from the ancient Alpine Tethys and present-day Iberia– Newfoundland margins. These data have enabled a precise description to be made of the architecture and the temporal and spatial evolution of these two pairs of magma-poor rifted margins. At present, these two pairs of margins can be considered as the best-studied examples of so called magmapoor rifted margins. An overview of the Alpine margins including their reconstruction is given by Manatschal (2004), whereas a review of the work carried out over the past three decades at the Iberia margin can be found in Whitmarsh & Wallace (2001). The major results obtained from these margins are: the existence and mapping of exhumed mantle in the ocean –continent transition (OCT) (Decandia & Elter 1972; Boillot et al. 1987; Pickup et al. 1996; Chian et al. 1999; Dean et al. 2000); the discovery and description of detachment systems (Froitzheim & Eberli 1990; Reston et al. 1995; Krawczyck et al. 1996; Manatschal & Nievergelt 1997; Manatschal et al. 2001) and the changing basin architecture and sedimentary evolution across the margin (Lemoine et al. 1987; Froitzheim & Eberli 1990; Wilson et al. 2001).

Rift concepts derived from the Iberia– Newfoundland and Alpine Tethys margins Observations or surveys can resolve either finite geometries at ancient margins or the spatial distribution of tectonic activity in present rifts. In order to understand the temporal and spatial evolution of the whole rifting process, data sampled across different scales have to be compared and examples of margins preserving different stages of the process have to be studied. Therefore, comparative studies played an important role in the understanding of rifted margins. In the 1970s a comparative approach allowed sedimentary facies such as the upper Jurassic Radiolarian Cherts and Majolica Limestones to be traced from the Central Atlantic into the Alpine realm (Bernoulli & Jenkyns 1974). Subsequent studies then compared exposed rift basins or fault systems in the Alps with seismically imaged structures in proximal margins offshore (Graciansky et al. 1979; Bally et al. 1981). These studies resulted in a precise description of the structures

in the proximal parts of the margins (e.g. Lemoine et al. 1986) that supported the pure shear model proposed by McKenzie (1978). In the late 1980s the first drilling of the deep Iberia margin resulted in the unexpected discovery of exhumed mantle. Several modifications of the simple shear model initially proposed by Wernicke (1981) were proposed to explain mantle exhumation (see fig. 3 in Reston et al. 1996). Since then, ODP legs 149, 173 and 210 and geophysical surveys have enabled a much more complete description of the geophysical characteristics of the deep Iberia margin (Whitmarsh & Wallace 2001). Simultaneous with the study of the deep Iberia margin, mapping of detachment faults in the Alps (Froitzheim & Eberli 1990; Florineth & Froitzheim 1994; Manatschal & Nievergelt 1997) and detailed structural (Manatschal 1999); petrological (Mu¨ntener et al. 2000, 2004; Desmurs et al. 2002) and isotopic (Schaltegger et al. 2002) investigations have allowed the geological characteristics of the distal margins to be described. These studies have demonstrated the limitations of the previously proposed pure-shear and simple shear models (McKenzie 1978; Wernicke 1981) and lead to new and more evolved models. A comparison of the Alpine Tethys and Iberia–Newfoundland margins shows that, although they are of different age and ultimately had a different fate, they share many similarities. These similarities permit a comparison of direct observations from the ancient margins with the borehole and geophysical data from the present-day margins. Based on these data, a conceptual model to explain the tectonic evolution of the Iberia–Newfoundland and Alpine Tethys margins was proposed by Whitmarsh et al. (2001a). Modified versions of this conceptual model are presented by Manatschal (2004) and Lavier & Manatschal (2006). Although these models capture only the most general features of the rifting process, they are designed such that they respect all data derived from ODP legs 103, 149, 173 and 210 and all major Alpine observations. Based on time constraints from the stratigraphic record, isotope ages and magnetic anomalies, four stages of development can be described and will be referred to as the ‘prerift situation’, the ‘stretching mode stage’, the ‘thinning mode stage’ and the ‘exhumation mode stage’ (Fig. 3).

Fundamental observations of lithospheric extension Observations constraining the conditions before onset of rifting (Fig. 3a) Crustal thickness estimated from the prerift stratigraphic record. Observations constraining the prerift conditions can be obtained from the prerift

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stratigraphic record, which is poorly known for the Iberia –Newfoundland margins, but excellently preserved and well described for the Alpine Tethys system (Fig. 4). The prerift stratigraphic record in the Alpine Tethys domain is characterized, apart from local exceptions, by a general evolution from continental siliciclastic –sabkha and shallowmarine carbonate dominated depositional systems (Fig. 4b). Carboniferous–Permian clastic sediments and volcanic rocks form the base to the succession. They are commonly interpreted to be deposited in fault-bounded basins forming during post-orogenic extensional collapse of the Variscan orogen (Handy & Zingg 1991; Handy et al. 1999). Triassic depositional environments changed across the future Alpine realm from deeper marine environments to the SE across platform carbonates and sabkha environments into continental siliciclastics systems to the NW (Fig. 4b). The sedimentary sequence is generally thicker in the east and south (Central Austroalpine and Southern Alps; 1–5 km), thins towards the NW (Lower Austroalpine and Brianc¸onnais; less than 500 m) and is locally only a few tens of metres in the future proximal European margin. The thicker Triassic sequences in the eastern and SE parts of the Alpine realm can be explained by its proximity to the Triassic Meliata ocean and the thermal subsidence following the breakup of this ocean (Channell & Kozur 1997; Dercourt et al. 1986; Robertson 2004). For the more western parts, forming the future distal margins of the Alpine Tethys, the stratigraphic record suggests that the prerift continental crust was in isostatic equilibrium after Permian time (Mu¨ntener et al. 2000). Therefore, at the onset of rifting the crustal thickness in the future distal margin is estimated to be in the order of about 30 km (Fig. 4a). Prerift Moho temperature and composition of the lower crust and upper mantle. Constraints on the Moho temperature and composition of the lower crust and upper mantle before onset of rifting can be obtained from an exhumed prerift fossil crust– mantle boundary exposed in the Malenco unit in northern Italy (Mu¨ntener & Hermann 1996; Hermann et al. 1997) (Fig. 4c). This crust –mantle boundary is welded by an Early Permian gabbro, which intruded into lower crustal rocks formed by metapelites and metacarbonates. Granulite facies metamorphism lasted for about 20 Ma (Hermann & Rubatto 2003). Mu¨ntener et al. (2000) demonstrated that after the intrusion of the gabbro in Permian time, the mantle and lower crustal rocks cooled more or less isobarically over 50 Ma. They calculated the temperature and pressure conditions at the crust –mantle boundary for the onset of rifting to be at about 550 8C at 0.9–1 GPa, corresponding to a 33 –36 km-thick crust (Fig. 4a). Based on these data, a rheological model with a

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brittle upper crust, a ductile middle crust, a gabbroic strong lower crust, a brittle upper mantle and a plastic lower mantle can be proposed for the prerift lithosphere (Fig. 4a). Evidence for a widespread lateral continuity of the underplated gabbros beneath the future proximal margins is given by the common occurrence of Permian gabbro xenoliths in Tertiary volcanic rocks in Western Europe (Fe´me´nias et al. 2003). Preexisting Variscan and Permian structures, local prerift volcanic activity and underplated gabbros along the crust– mantle boundary resulted in rheological heterogeneities (Handy 1989), which, as discussed below, may have strongly controlled the distribution of rift structures at the onset of lithospheric extension.

Observations documenting early rifting (Fig. 3b) Stratigraphic record of early rifting. The age of onset of rifting is diffuse and poorly constrained at both the Iberia –Newfoundland and Alpine Tethys margins. In the Iberia –Newfoundland margins interpretations of the onset of rifting are based mainly on geometrical arguments, i.e. determination of geometry related to synsedimentary normal faulting such as onlap onto tilting hanging walls and thickening into footwalls (Driscoll et al. 1995; Wilson et al. 2001). In the Alps, the age of onset of rifting is based on sedimentological, stratigraphic and structural arguments (Eberli 1988; Bertotti et al. 1993; Wilson et al. 2001). The onset of rifting is marked by differential subsidence and associated faulting during Late Triassic time. However, because sedimentation kept pace with differential subsidence, the prerift platform conditions persisted during initial rifting. Also, the most proximal parts of the Adriatic margin may have been affected by post-rift subsidence related to the opening of the Vardar –Meliata Ocean in Triassic time (Fig. 4a). This renders identification of the onset of rifting at the Alpine Tethys margins difficult. Larger rift basins, which are related unambiguously to the rifting leading to the Alpine Tethys Ocean, formed only later, during Early Jurassic time (Hettangian –Sinemurian). Basins such as the Il Motto basin (Fig. 5b) and the Generoso Basin (Fig. 6b) formed simultaneously with large basins on the conjugate proximal European margin (e.g. Bourg d’Oisans Basin: Chevalier et al. 2003) (Fig. 5a). Rifting in the proximal parts of the future margin ceased in the middle –late Early Jurassic, but some of the faults became inactive even earlier and were sealed by Upper Sinemurian sediments (e.g. Il Motto: Eberli 1988; Conti et al. 1994) (Fig. 5b). Later rifting was focused, as discussed below, on the future distal margin.

Fig. 5. (a) Reconstructed section across the future margins during initial rifting and locations of examples of early rift structures in the proximal parts of rifted margins. (b) View of Il Motto showing the break-away of a high-angle normal fault that is sealed by Upper Sinemurian sediments (to the left) and a model for the Il Motto Basin (middle part) based on sedimentary sections across the basin (to the right) (modified after Eberli 1988). For location of the Il Motto Basin see (a). (c) Reflection seismic section across the Jeanne d’Arc Basin (from Tankard et al. 1989). For location see Figure 2a.

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Early rift structures preserved in the future proximal margins. The most prominent rift structures in the proximal margins are fault-bounded rift basins. Such basins have been imaged in seismic reflection profiles from the proximal Iberia margin (Montenat et al. 1988; Murillas et al. 1990) and drilled on the Newfoundland margin (Keen et al. 1987; Driscoll et al. 1995). A typical example is the Jeanne d’Arc Basin (Fig. 5c), which is bounded by a major fault that is well imaged in the upper crust but disappears at middle–lower crustal levels and does not affect the Moho, which is regionally up-warped beneath the extended basin (Keen et al. 1987). Examples of fault-bounded basins are spectacularly exposed in the area of Bourg-d’Oisans near Grenoble in the external French Alps (Lemoine et al. 1986; Chevalier et al. 2003) and at Il Motto in the Ortler nappe along the Swiss– Italian border in the eastern Alps (Conti et al. 1994) (Fig. 5b). The most complete section preserving rift structures of the former proximal Adriatic margin is exposed in the Lombardian and southern Swiss Alps (Bertotti et al. 1993) (Fig. 6a). In this section, the basin architecture is exposed over more than 200 km from the Trento plateau to the east to the Canavese to the west. Bernoulli (1964) demonstrated, based on mapping of abrupt changes in facies and thickness of the Early Jurassic formations across the Lugano –Val Grande fault, the existence of an Early Jurassic rift basin, about 30 km wide and filled with several thousand metres of synrift sediments (e.g. Generoso Basin) (Fig. 6b). Bertotti (1991) mapped the Lugano –Val Grande fault over a horizontal distance of more than 30 km. He demonstrated that this fault cuts across prerift sediments into basement, changing from a localized brittle fault into a ductile mylonitic shear zone (Fig. 6b). The reconstructed geometrical relationship between fault zone and bedding in the overlying synrift sediments led Bertotti (1991) to conclude that the fault had a listric geometry. Based on microstructural investigations, he demonstrated that the fault soled out in greenschist facies conditions, i.e. at about 300– 350 8C and about 10– 12 km depth, assuming a normal temperature gradient of 30 8C km21. The synrift sediments in the rift basins in the proximal margin are characterized by abrupt changes in facies and show significant thickness variations. Basin infill is dominated by turbidites interbedded with hemipelagic sediments. Massflow deposits, such as debris flows or rock falls derived from exposed normal faults, are observed only in a few basins. A particularly instructive section preserving the relationships between the sediment infill and the master fault bounding the basin is exposed at Il Motto (Fig. 5b). The synrift

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sedimentary sequence is characterized by finingand thinning-upwards in the basin and, away from the fault zone, an inversion of the clast stratigraphy in the basin owing to erosion and redeposition of the footwall block (Eberli 1988). Erosion of the upper Triassic sediments in the footwall block and their redeposition in the basin suggests subaerial exposure and erosion of the tip of the tilted footwall block. In the Generoso Basin and the Bourgd’Oisans the highs remained submarine. In both places a change from shallow to deeper marine sedimentation has been documented from the footwall near the break-away of the fault. This observation indicates that both hanging wall and footwall were subsiding, but at different rates relative to sea level, and that rift-shoulder uplift was less than the general subsidence of the extending margin. Thermal and magmatic evolution during early rifting. Constraints on the thermal evolution of the margin during initial rifting can be obtained from cooling ages of basement rocks. Published cooling ages obtained from the footwall and hanging wall of the Lugano –Val Grande fault range from 220 to 180 Ma (Sanders et al. 1996) (Fig. 6c). However, cooling ages cannot be related to exhumation and uplift along the Lugano–Val Grande fault because there was marine sedimentation onto the footwall. Bertotti & Ter Voorde (1994) and Bertotti et al. (1999) interpreted these cooling ages to be related to thermal equilibration of a Ladinian magmatic event described from the eastern southern Alps and the Austroalpine realm (Lu et al. 1997). Their interpretation ignores the fact that Late Triassic –Early Jurassic cooling ages are widespread in the Alpine realm, suggesting that rifting in the proximal margins was associated with cooling of the extending lithosphere. Volcanic activity was widespread during the Triassic across the future European and Adriatic margins. This is well established by tuff horizons and local dolerites within the Triassic sedimentary sequences. However, evidence for volcanic activity during Late Triassic –late Middle Jurassic rifting is not recorded. This observation suggests that rifting related to the opening of the Alpine Tethys ocean can be considered as non-volcanic. Rates of extension during early rifting. The problem of objectively defining the age of onset of rifting generates big error bars on the determination of the duration of rifting and consequently the rates of extension. For the West Iberia margin, rift duration is poorly constrained and commonly overestimated (Wilson et al. 2001). Observations show that rifting is polyphase and that strain rates are difficult to obtain. In the best case, extension rates can be determined for parts of the margins

Fig. 6. (a) Basin architecture and depocentre migration in the Ticino Transect (for location see Fig. 1) (modified from Bertotti et al. 1993). (b) Palinspastic reconstruction of the Monte Generoso Basin and its relation to the Lugano– Val Grande fault bounding the basin (modified after Bertotti 1991). (c) P –T – t evolution from the rocks sampled in the vicinity of the Lugano– Val Grande fault (Sanders et al. 1996).

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or single events. Bertotti et al. (1993) defined for the Ticino Transect average velocities of rifting ranging between 0.46 and 1.28 mm year21. Strain rates were also determined for two faults bounding rift basins in the proximal European margin (i.e. Ornon fault bounding the Bourg d’Oisans Basin: Chevalier et al. 2003; and the Uzer fault: Dromart et al. 1998). Chevalier et al. (2003) showed that the motion along a single fault was discontinuous. Low rates of displacement (0.2–0.4 mm year21) were accommodated by diffuse extensional deformation, whereas high rates (up to 1.8 mm year21) were accommodated along the localized master fault bounding the basin. These authors showed that such high rates were observed to occur over very short time spans (one ammonite zone; less than 0.5 Ma). These studies show that most of the extension has been accommodated during short pulses and that activity is not simultaneous along different fault segments in the margin. This behaviour means that strain rates for the whole margin are difficult to obtain. Available data suggest that overall the extension associated with rifting appears to be slow (,10 mm year21), discontinuous and distributed over tens of millions of years (Bertotti et al. 1993; Chevalier et al. 2003). In contrast, slip rates along single faults were fast, but occurred only over a short time span. These rather slow rates may explain also the cooling ages obtained from the basement underlying the basins in the Ticino transect (Fig. 6c).

Observations documenting advanced rifting (Fig. 3c) Stratigraphic record and early rift structures preserved in the future distal margins. In contrast to proximal margins, distal margins are more difficult to access. Present-day distal margins are at abyssal depth and covered by thick sedimentary units, and fossil examples preserved in collisional orogens are either subducted or strongly overprinted during collision. Therefore the early rift evolution of distal margins is so far only constrained from a few places. The available observations from the Alps show the migration of depocentres towards the area of future breakup, indicating a diachronous evolution of the proximal and distal margins as well as a successive localization of the deformation in the distal margin during early rifting (Lemoine et al. 1986; Eberli 1988; Bertotti et al. 1993; Manatschal & Bernoulli 1998). A similar evolution is also observed in the West Iberia margin, where rifting in the Interior Basin predates rifting in the Iberia Abyssal Plain (Wilson et al. 2001). A reconstruction of the distal Alpine Tethys margins during Toarcian time (c. 175 Ma), about 10 Ma before final

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breakup, is shown in Figure 7. In this reconstruction the distal parts of the future European margin are represented by the Brianc¸onnais domain (see Brianc¸onnais Transect in Fig. 1c). Parts of this domain escaped subduction and were obducted onto more proximal parts of the European margin (Fig. 1b). Distal parts of the Adriatic margin are represented by the Err–Bernina domain in the north and the Canavese –Gozzano domain further to the south (see the Grischun and Ticino transects in Fig. 1c). These units were in the hanging wall of the Tertiary subduction and therefore were overprinted only weakly during Alpine collision. The observed rift structures in the future distal margin are in many respects different from that described previously from the proximal margins. The diachronous evolution of the proximal and distal margins is best recorded by the observation that the early Middle Jurassic fault-bounded basins in the future distal margins formed contemporaneously with post-rift hemipelagic sequences in the proximal margins (see the Grischun Transect in Fig. 1c). Subsidence is indicated by the transition from platform to hemipelagic cherty limestones, which are contemporaneous with the synrift sediments in the proximal basins. In contrast to the proximal margin these sediments show a constant thickness of less than 100 m across the distal margin, clearly indicating that high-angle faulting and related block tilting were subdued in this part of the margin. These lower Jurassic, hemipelagic carbonates – the Agnelli Limestone of Furrer et al. (1985) – are overlain by unsorted, polymictic breccias with variable amounts of a reddish matrix and clasts predominantly derived from the basement. These breccias show a transition into siliciclastic turbidites, interbedded with hemipelagic marls, 200 and 450 m thick (e.g. Saluver Formation of Finger et al. 1982). These sediments are the first direct evidence for significant tectonic activity within the distal margin. The common occurrence of clasts of alkali granite, characteristic of the Bernina nappe (Spillmann & Bu¨chi 1993), together with the arrival of sandstones, indicates uplift and subaerial exposure of parts of the Bernina domain contemporaneous with down-faulting of the Err domain. The age of this event is weakly constrained; it is younger than the hemipelagic sediments dated as Pliensbachian and older than the overlying upper Middle Jurassic Radiolarite Formation (Baumgartner 1987). Uplift of the Bernina domain is, however, indirectly dated as Toarcian (about 180 Ma) by the occurrence of gravity flows within ammonite-bearing hemipelagic limestones in the distal part of the proximal margin exposed in the Ela–Bernina nappe (Eberli 1988) (Fig. 7). In the Brianc¸onnais domain uplift and subaerial exposure is recorded by a prominent erosional

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surface and a karst system within the Triassic and Lower Jurassic shallow-water carbonates that was dated by Lemoine et al. (1986) and Claudel & Dumont (1999) as late Early to Middle Jurassic in age. Constraints for crustal thinning during early rifting in the distal margin are obtained from the lower crustal rocks exposed in the Ivrea zone and the Malenco area in northern Italy (Figs 1c & 7). These lower crustal rocks are separated from upper crustal rocks along the Pogallo fault in the Ivrea zone (Handy 1987; Handy & Zingg 1991) (see Fig. 7) and the Margna fault in the Malenco area (Mu¨ntener & Hermann 2001) (see Fig. 1c). In reconstructed sections, these faults dip beneath the distal Adriatic margin (Handy 1987; Mu¨ntener & Hermann 2001). Cooling ages from the footwall of the Pogallo fault (Handy & Zingg 1991) and the Margna fault (Villa et al. 2000), together with thermo/barometric data (Mu¨ntener & Hermann 2001 and references therein), clearly show exhumation and major thinning of the crust to about 0.4– 0.5 GPa (i.e. about 15 km) during initial rifting in Early Jurassic time. U –Pb dating on titanite from synmylonitic dykes from the Pogallo fault gave an age of 173 + 4 Ma (Toarcian–Allenian), interpreted to result from a discontinuous retrograde metamorphic reaction associated with the crustal thinning that was localized along the Pogallo fault (Mulch et al. 2002). These observations show that thinning of the crust along major mylonitic fault zones in the distal margin (e.g. Pogallo fault; Fig. 7) postdates Sinemurian–Pliensbachian (200 –184 Ma) sealing of fault-bounded basins in the proximal margins (e.g. Il Motto; Fig. 5b), and coincides with migration of the depocentre into the future distal margin (e.g. Ticino Transect; Fig. 6a), and uplift of the Brianc¸onnais and Bernina domains (e.g. Fig. 7). We infer, based on these data, that there is a direct relationship between deep lithospheric thinning and migration of the depocentre into the distal margin. Lemoine et al. (1987) proposed that the contrasting behavior of the two future distal margins was directly related to breakup. In order to explain the uplift of the Brianc¸onnais margin, they proposed a simple shear model for continental breakup with a detachment fault dipping beneath the European margin. Our observations suggest that uplift of the Brianc¸onnais and Bernina domains is more probably related to the thinning of the continental crust along Pogallo type faults (see below for further discussion), which predate the final detachment faults exhuming the mantle (see later discussion) by 10– 15 Ma. Based on these observations and as discussed below, we believe that the Pogallo and Margna faults are very important structures that explain thinning of

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the crust in the future distal margin predating final mantle exhumation.

Observations documenting final stage of rifting, breakup and transition to sea-floor spreading The sedimentary record of final rifting in the Alpine Tethys margins shows that: (1) synrift sediments become younger towards the place of future breakup (Froitzheim & Eberli 1990; Bertotti et al. 1993) (Fig. 6a); (2) final rift basins are associated with downward concave faults leading to a very different architecture of the synrift sediments (Wilson et al. 2001); and (3) the isostatic evolution of the most distal margins is asymmetric (Lemoine et al. 1987). The most prominent structures related to final rifting are detachment faults, which are described and mapped in the Alps (Froitzheim & Eberli 1990; Florineth & Froitzheim 1994; Manatschal & Nievergelt 1997; Manatschal 2004) and seismically imaged as low-angle reflections off Iberia (de Charpal et al. 1978; Boillot et al. 1987; Reston et al. 1995; Krawczyk et al. 1996) (Fig. 8). These structures are interpreted to play an important role in the exhumation of subcontinental mantle during final rifting, leading to the formation of regions that show the characteristics neither of ‘true’ oceanic crust nor of thinned continental crust. Whitmarsh et al. (2001a) introduced the term ‘zone of exhumed continental mantle’ (ZECM) for such regions. In this paper the term ‘ZECM’ is used only for the part of the margin that is floored by subcontinental mantle lithosphere. In contrast, the term ‘OCT’ is used in a more general way for the transition from the distal continental margin to the oldest oceanic crust and implies neither a particular genetic evolution nor a specific composition of the underlying lithosphere. We summarize the major observations related to final rifting and breakup of the lithosphere using two examples. The first is the Err–Platta OCT exposed in the Alps (Fig. 8b), and the second is the southern Iberia Abyssal Plain (Fig. 8c). More detailed descriptions of these two sites are given in Manatschal (2004 and references therein). Err-Platta OCT † Detachment system in the Err –Platta OCT. Remnants of a detachment system preserving primary relations between detachment faults, exhumed mantle rocks, magmatic rocks, and syn- and post-rift sediments are exposed in the Err –Platta nappes in SE Switzerland (Manatschal & Nievergelt 1997) (Fig. 9). For details of the reconstruction of the overall detachment system see Manatschal & Nievergelt

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Fig. 8. (a) Generalized section showing the situation during final rifting based on observations from the Alpine Tethys and Iberia–Newfoundland margins. (b) Tentative palinspastic section showing the observed relationships between the detachment system and associated sediments, exhumed mantle rocks, and intrusive and extrusive magmatic rocks in the Err–Platta ocean–continent transition. (c) Geological interpretation of the depth-migrated Lusigal 12 profile (Krawczyk et al. 1996) showing the distribution of upper and lower crustal rocks, exhumed subcontinental mantle rocks and reflections interpreted as detachment faults. Numbers on top of the section refer to ODP sites.

RIFTED MARGINS AND CONTINENTAL BREAKUP

(1997) and Desmurs et al. (2001). Although the actual transition from continental rocks to exhumed mantle rocks is not preserved in one continuous outcrop, the Err–Platta OCT represents at present the best-documented and most complete section across an OCT of a magma-poor rifted margin. The detachment system reconstructed in the Err–Platta OCT forms one or several breakaways in the continental crust and cuts oceanwards from upper crust directly into subcontinental mantle. The relations between the detachment system and its hanging wall are spectacularly exposed over 6 km in the area of Piz d’Err –Piz Bial in the Err nappe (for more details see Froitzheim & Eberli 1990; Manatschal & Nievergelt 1997). The hanging wall is formed by fault blocks composed of continental basement and prerift sediments, which are tilted, toward the east (continent) and the west (ocean). The fault blocks, a few kilometres to some hundreds of metres across, are truncated systematically at their base by a detachment fault (the Err detachment; Manatschal & Nievergelt 1997) (Fig. 9). Similar blocks, interpreted as extensional allochthons, also overlie mantle rocks in the Parsettens area in the Platta nappe (Fig. 9). Locally, the detachment faults are directly overlain by tectono-sedimentary breccias, pillow basalts (Fig. 9) or post-rift sediments. Mapping of the detachment faults in the Err nappe shows that these structures are corrugated and composed of several fault zones. The geometrical relationships between the different fault zones are described in Manatschal & Nievergelt (1997). † Deformation structures associated with the detachment system. In the continental crust, profiles across the footwall of the detachment faults exhibit a transition from undeformed granite into green cataclasites that are capped along a sharp and well-defined horizon by a characteristic black fault gouge (Manatschal 1999) (Fig. 9). The lack of any crystal plastic deformation in the quartz-rich basement rocks indicates that these faults were active under low-temperature conditions (,300 8C) within the uppermost 10 km in the crust. Based on geochemical, mineralogical and structural investigations, Manatschal (1999) demonstrated the importance of fluid- and reaction-assisted deformation processes for the localization of extension in the brittle upper crust. Of particular importance is the observation that mantlederived elements such as chromium (Cr) and nickel (Ni) are enriched in the fault zone in the continental crust. Manatschal et al. (2000) explained this enrichment by the percolation of

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fluids with high concentrations in Ni and Cr that were derived from a serpentinizing mantle (Frueh-Green et al. 2002). This interpretation leads to the suggestion that serpentinization of the mantle started beneath an extending continental crust, and that serpentinization is intimately related to extensional deformation. This is compatible with the results obtained from the stable isotope investigations of Skelton & Valley (2000) and the modelling results of the rheological evolution during continental breakup by Pe´rez-Gussinye´ & Reston (2001). In the ZECM, a fault zone with a constant topto-the-ocean sense of shear caps the exhumed mantle. This fault zone is formed by foliated serpentinite, serpentinite cataclasite and serpentinite gouges, recording a complex deformation and hydration history under lower amphibolite facies to sea-floor conditions. High-temperature shear zones, formed by peridotite mylonites – ultramylonites, are commonly oblique to the fault zone capping the exhumed mantle, and show a top-to-the-continent sense of shear. The relationship between the ‘cold’ fault zone and the ‘high-temperature’ mylonite shear zones is exposed in an outcrop at Sur al Cant in the Platta nappe (Fig. 9). In this outcrop, the hightemperature top-to-the-continent peridotite mylonites are cut by a top-to-the-ocean fault zone, which also cuts rodingitized basaltic dykes. This is a clear evidence that magmatic activity predates or was coeval with final movement along the top-to-the-ocean fault zone. This observation is fundamental, because it demonstrates that late, ‘cold’ mantle exhumation was associated with magmatic activity. The different transport directions between the ‘high-temperature’ (.700 8C) and later hydrous ‘cold’ deformation structures show a complex relationship between ‘brittle’ and ‘ductile’ deformation processes during continental breakup that is not yet fully understood. Age constraints on the deformation history can be obtained from deformation structures within gabbros that have been dated by U–Pb on zircons at 161 + 1 Ma (Bathonian–Callovian) (Schaltegger et al. 2002). Clasts of gabbros and albitites, also dated at 161 + 1Ma by the same method, are found in pillow and sedimentary breccias overlain by middle–Upper Jurassic radiolarian cherts. This observation suggests that the gabbros were exhumed, probably along detachment faults, to the sea floor directly after their crystallization. Microstructures in these gabbros reveal a deformation history ranging from synmagmatic to brittle conditions. This deformation was acquired during their intrusion into partially serpentinized mantle rocks and their subsequent exhumation at the sea floor.

Fig. 9. Tentative palinspastic section across the Err– Platta OCT showing the observed relationships between the detachment system and its hanging-wall and footwall structures. In the three grey boxes overlying the section, reconstructed geometry of mapped rift-related structures are shown (for legend see Fig. 8). A –E: observed structures in the Err-Platta OCT: (a) pillow basalts overlying exhumed mantle in the Falotta area (Platta nappe); (b) continent-derived breccias separated from exhumed mantle along a sub-horizontal detachment fault in the Parsettens area (Platta nappe); (c) an allochthon truncated by the sub-horizontal Err detachment in the area of Piz Laviner (Err nappe); (d) preserved cross-cutting relationships between a top to the ocean serpentinite gouge zone, a top-to the continent high-temperature mantle shear zone and a dolerite dyke in the area of Al Cant (Platta nappe); and (e) brittle fault zone in the Err nappe related to the Err detachment. The fault zone architecture shows a transition from granite to cataclasites containing syntectonic quartz veins. The fault core is formed by a black gouge. For locations and more explications of the outcrops see Manatschal et al. (2003).

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† Relation between detachment faults and its sedimentary cover. In several places in the Err–Platta OCT synrift sediments overlie the exhumed detachment surface at a low angle (,208) of discordance (Manatschal & Nievergelt 1997). Exposure of the detachment at the sea floor is also supported by the occurrence of clasts of green cataclasites and black gouge derived from the detachment in the synrift sediments (Froitzheim & Eberli 1990). These observations and the lack of any evidence for a major fault-related morphology produced by this fault, despite its large displacement (.10 km), can be explained only if the detachment fault was locally exhumed at the sea floor at a low angle. The observation that sediment sequences formed by turbidites and debris flows, classically interpreted as synrift sediments, are cut by detachment faults but also deposited onto already exhumed detachment faults shows that there are complicated polyphase relationships between synrift sedimentation and detachment faulting. Different types of tectono-sedimentary breccias, referred to as ophicalcite in the Alpine literature, include mainly serpentinite clasts embedded in a fine-grained, microsparitic, calcite matrix. The fabric of these breccias varies considerably between two end-member types. One endmember consists of serpentinite host rock with fractures filled by red limestone and/or white sparry calcite. An intermediate type consists of clast-supported breccias with in situ fragmented serpentinite clasts (Bernoulli & Weissert 1985; ophicalcites I of Lemoine et al. 1987). The other end-member consists of coarse, unsorted, matrix-supported breccias with fragments of serpentinite, gabbro, and continent-derived basement rocks and prerift sediments (ophicalcites II of Lemoine et al. 1987). Geopetal infill of sediment into crevasses and pockets in the mantle rocks indicate that these rocks were exposed at the sea floor. Indeed, the mantle rocks are also stratigraphically overlain by pillow-basalts or radiolarian cherts. Further evidence for the occurrence of mantle rocks at the seafloor is the presence of mass-flow breccias (ophicalcite type II of Lemoine et al. 1987) containing clasts of serpentinite and Triassic dolomites encased in a serpentine arenite or calcite matrix, and the occurrence of serpentinite arenites within the post-rift sediments (Decandia & Elter 1972). † Mantle and magmatic rocks in the ZECM. The mantle rocks in the ZECM are invariably serpentinized peridotites derived from spinel lherzolites and harzburgites into which gabbros and basaltic dykes were intruded. A major and trace element

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study of mantle clinopyroxene across the Platta ZECM (Mu¨ntener et al. 2002) reveals that mantle rocks close to the continent may represent spinel peridotite mixed with (garnet)-pyroxenite layers while those at some distance from the continent are pyroxenite-poor peridotite that equilibrated in the plagioclase stability field. Textural relationships indicate that some plagioclase peridotites in the Platta nappe were formed by melt infiltration and melt-rock reaction (Mu¨ntener et al. 2004). Two-pyroxene and singleorthopyroxene thermobarometric data from the Platta nappe reveal an increase in the equilibration temperature from 850 + 50 8C at 0.8– 1.2 GPa near to the continent to more than 1000 8C further oceanwards (Desmurs 2001). The age of equilibration near the continent is likely to be Permian, as constrained from intrusive relationships with Permian gabbros in the Malenco peridotite (Mu¨ntener et al. 2000). In contrast, the oceanward increase in the equilibration temperature in the ZECM may be related to final continental breakup and onset of sea-floor spreading. The age for the infiltration event is, as suggested by clinopyroxene Nd isotopic data, very likely to be of Jurassic age (Mu¨ntener et al. 2004). However, because infiltration must have occurred under high-temperature conditions, this process must have initiated prior to final exhumation of the infiltrated mantle along cold brittle faults at the sea-floor. Gabbro bodies occupy less than 5% of the total observed serpentinite volume. Desmurs et al. (2002) described smaller spherical bodies, less than 100 m in diameter, and larger sill-like bodies. The smaller bodies are more homogeneous, consist of Mg-gabbro and show a decrease in grain size from the core towards the rim. The larger bodies show a great diversity in composition from primitive olivine-gabbros to highly differentiated Fe–Ti–P-gabbros and diorite, but up to 90% of the body is formed by Mg-gabbro. Massive basalts, pillow lavas, pillow breccias and hyaloclastites crop out in patches of variable thickness and size in the Platta nappe, and their abundance increases oceanward across the reconstructed section (Dietrich 1969). Away from the edge of continental crust, pillow lavas form isolated bodies less than 100 m in diameter and a few tens of metres thick. Oceanwards, the bodies are aligned and appear to be controlled by late, synmagmatic high-angle faults. The basalts stratigraphically overlie serpentinites, ophicalcites, gabbros and associated breccias, indicating that their emplacement post-dates exhumation of the mantle rocks and gabbros at the sea floor. However, basaltic dykes are also truncated by

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detachment faults showing that detachment faulting occurred within a magmatically active system. Based on Mg numbers and Ni contents of equilibrium olivine calculated from primitive basalts and gabbros, Desmurs et al. (2002) demonstrated that few mafic rocks found in the Err–Platta OCT are primary melts. Most represent fractionated compositions ranging from transitional midocean ridge basalts (T-MORB) to normal MORB (N-MORB). The mafic rocks may be explained by low–moderate degrees of melting of a N-MORB type mantle, as indicated by initial Hf isotope data of zircons and Nd data of whole-rock samples (Schaltegger et al. 2002). The source of some basalt, however, is enriched in incompatible elements. This compositional variation seems to correlate, as indicated by a study of Desmurs et al. (2002), with the spatial distribution of the mafic rocks within the OCT in that mafic rocks with T-MORB signatures occur close to the continental margin whereas N-MORB signatures are more frequently found oceanwards. † Age constraints for continental breakup and emplacement of first magmatic rocks. In the Alpine Tethys margins crustal breakup is constrained by the oldest biostratigraphic ages of the sediments sealing both oceanic and continental units. The biostratigraphic ages obtained for these first sediments, which are radiolarian cherts, give Bathonian–Callovian ages (Bill et al. 2001). These ages are similar to radiometric ages dating recrystallization and exhumation of the oldest magmatic rocks in the ZECM. U/Pb zircon ages on gabbros from the ZECM range between 165 and 157 Ma (Lombardo et al. 2002; Schaltegger et al. 2002), and Ar/Ar cooling ages on phlogopite in a pyroxenite of the Totalp serpentinite yielded a cooling age of 160 + 8 Ma (Peters & Stettler 1987). Thus, all data obtained from the ZECM in the Alps converge towards a very short time period, implying that exhumation and cooling of the subcontinental mantle was associated with the onset of magma emplacement and the deposition of radiolarian cherts. This interpretation is compatible with the observation that detachment faulting leading to the exhumation of the mantle was linked directly to emplacement and exhumation of gabbros at the sea floor. The Southern Iberia Abyssal Plain † Evidence for a detachment system in the Southern Iberia Abyssal Plain. In the Southern Iberia Abyssal Plain, several strong intrabasement

reflections have been observed in the region of thinned continental crust and interpreted as fault structures (Beslier et al. 1995; Krawczyk et al. 1996; Whitmarsh et al. 2000) (Fig. 8c). Mapping of these reflections show complex lateral variations from classical high-angle to low-angle top-basement structures (Pe´ronPinvidic et al. 2007). In an east –west-directed section perpendicular to the margin, the L and H reflections form break-aways towards the continent, bound well-developed basins and cut into basement (Figs 8c & 10). Thus, the H and L reflections show clear characteristics of fault structures. In this paper, whenever we refer to these structures as reflectors, we use L and H, and when we refer to them as detachment faults, we use H-Detachment (HD) and LDetachment (LD). Whereas L has been mapped to cut into the mantle defined by velocities higher than 8 km s21, H flattens within a crustal layer showing velocities of between 7 and 8 km s21, is up-warped beneath Hobby High and reaches the top of the basement at Hobby High, where it separates mantle rocks in the footwall (Site 1068) from lower crustal rocks in the hanging wall (sites 900 and 1067) (Fig. 8c). A continentward-dipping reflection, the C reflection, is cut and displaced by HD and LD (Whitmarsh et al. 2000) (Fig. 8c) and must therefore be older than these structures. Only a few kilometres north of Lusigal 12, on seismic line CAM 144, Chian et al. (1999) described a 6.5–7.3 km s21 interface at a depth similar to that of reflection C on Lusigal 12 and interpreted it as the crust –mantle boundary. Based on these data, Whitmarsh et al. (2000) also interpreted C as a contact between continental crust above and serpentinized peridotite below. A simple geometric argument, however, leads to an alternative interpretation of the C reflection. The C reflection is truncated by the HD and has therefore to change across HD from its footwall into its hanging wall going oceanwards. Assuming that the C reflection is a continuous structure, it must reach the top of the basement continentwards of where the HD reaches the top of the basement (Fig. 8c). At ODP sites 900 and 1067 drilling into the hanging wall of HD penetrated lower crustal rocks. Based on this line of evidence, Manatschal et al. (2001) interpreted the C reflection in the area of Site 1065 as an intracrustal reflection (e.g. Fig. 8c). The lower crustal rocks drilled at ODP sites 900 and 1067 show a pervasive brittle overprint at the top, which decreases in intensity down-hole (Manatschal et al. 2001) (Fig. 10).

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Fig. 10. Photographs of the rocks derived from the top of the basement drilled in the OCT in the Iberia Abyssal Plain. Numbers on top of the section refer to ODP sites.

40 Ar/39Ar plagioclase ages were obtained from these rocks: 136 Ma at Site 900 (Fe´raud et al. 1996) and 137 Ma at Site 1067 (Manatschal et al. 2001). These ages, interpreted as cooling ages, show that, at 136 Ma, these basement rocks were not yet at the sea floor, although they are at present overlain by Upper Cretaceous sediments (Fig. 8c). These observations suggest that Hobby High must be capped by a detachment fault in order to explain the cooling ages and the penetrative deformation at the top of the drilled basement sealed by sediments. At Site 1068, on the oceanward side of Hobby High and within the footwall of H, drilling penetrated mantle rocks. These rocks are capped by a serpentinite gouge and overlain by tectono-sedimentary breccias, providing further support for the existence of an exhumed detachment fault at Hobby High. This detachment, the Hobby High Detachment (HHD) (Manatschal et al. 2001), is interpreted to cut from the continental crust at sites 1067 and 900 into mantle at Site 1068 (Fig. 8c). In seismic profile Lusigal 12, HHD is interpreted to break out about 15 km east of Hobby High (Fig. 8c). From there, it can be followed towards Hobby High, where it has been

drilled. Further oceanwards it plunges beneath a high drilled at Site 1069. This high, with a topographic relief of about 1.5 km and 9 km long, is capped by pre-, syn- and post-rift sediments, and is interpreted as a continent-derived extensional allochthon emplaced onto exhumed subcontinental mantle (Fig. 8c). This interpretation requires a large offset along the HHD, which is also compatible with the observation that this fault forms the top of the basement over a distance of more than 20 km. † Zone of exhumed continental mantle (ZECM). In the Southern Iberia Abyssal Plain, the thinned continental crust is bounded oceanwards by an up to 130 km-wide zone of exhumed continental mantle (ZECM). This zone shows distinct geophysical characteristics. Its seismic velocity structure differs from that of the adjacent stretched continental and oceanic crust. A 2– 4 km-thick upper basement layer with a P-wave velocity of 4.5–7.0 km s21 and a high velocity gradient (c. 1 s21) merges into a lower layer 4 km-thick with velocities of approximately 7.6 km s21 and a low velocity gradient (,0.2 s21) (Chian et al. 1999). Moho reflections are weak or absent. The top-basement velocity is lower than that of the adjacent continental

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crust, while the velocity in the lower layer is too high to represent magmatically intruded or underplated continental crust or even oceanic layer 3. The velocity structure can best be explained by mantle serpentinization decreasing with depth. The seismic velocity structure gradually changes oceanwards to typically oceanic approximately 20 km west of the peridotite ridge (Whitmarsh et al. 2001b). North–south-trending, low-amplitude magnetic anomalies in the ZECM indicate that basement magnetization is typically much lower than that of the oceanic basement further to the west (Srivastava et al. 2000; Russell & Whitmarsh 2003). These weak magnetic anomalies have been explained to result either from small magmatic bodies within the ZECM (Russell & Whitmarsh 2003) or from serpentinization within the hinge of downward-concave faults during exhumation of mantle rocks in the ZECM (Sibuet et al. 2007). Morphologically, the acoustic basement may be divided into two regions of north–south-trending basement ridges and deep (c. 9 km) relatively low-relief basement; both narrow northward. ODP cored exhumed mantle rocks at four sites (sites 897, 899, 1068 and 1070) and each time penetrated serpentinite gouge and/or cataclasite in several places overlain by tectonosedimentary breccias (Fig. 10). Because drilling targeted basement highs, sampling is probably biased, but drilling results are compatible with the observed velocity structure, suggesting widespread serpentinization of the mantle in the Iberia Abyssal Plain. Primary-phase chemistry and clinopyroxene trace element compositions obtained from the Iberia margin indicate a heterogeneous mantle depleted by less than 10% partial melting and percolated by mafic melts (Beslier et al. 1996; Abe 2001; He´bert et al. 2001). Serpentinization began, at least locally, before sea-floor exhumation of the peridotite (Skelton & Valley 2000). Despite strong serpentinization, the exhumation of mantle rocks from deep lithospheric levels to the ocean floor has been documented by a sequence of deformation structures under decreasing temperatures. This is shown by peridotite mylonite shear zones, serpentinization, the formation and subsequent cataclastic reworking of serpentinite mylonites, and low-temperature replacement by calcite (Beslier et al. 1996). The serpentinized mantle is capped by tectono-sedimentary breccias reworking adjacent exhumed basement (Manatschal et al. 2001). At Site 1070 (Fig. 10), which is situated above a margin-parallel basement ridge, upper Aptian (112.2–116.9 Ma) sediments overlie subcontinental mantle, which

is intruded by gabbro veins. These gabbro veins are the only rift-related intrusive magmatic rocks drilled in the Iberia margin. Extrusive magmatic rocks were recovered only as clasts within sedimentary breccias in the ZECM (Sawyer et al. 1994). † Dating continental breakup. The age of continental breakup, referred to in this paper as the process that leads to the final mechanical separation of two lithospheric plates that is followed by sea-floor spreading, has been classically determined by the first sea-floor spreading magnetic anomaly and/or a breakup unconformity. Such dating was based on the assumptions that: (1) breakup of continental crust is immediately followed by sea-floor spreading; (2) rifting is of the same age across and along the whole margin; and (3) the sediment architecture in the OCT is simple. The results obtained from the Iberia–Newfoundland margin show that all these assumptions have to be questioned, and that determining the age of continental breakup is difficult (Pe´ron-Pinvidic et al. 2007). The oldest magnetic anomaly in the southern Iberia Abyssal Plain has been interpreted by Whitmarsh & Miles (1995) as M3 and by Srivrastava et al. (2000) as M17. These contrasting interpretations suggest that, at magma-poor margins, magnetic anomalies are not necessarily formed by classical sea-floor spreading processes and consequently are not reliable for dating onset of sea-floor spreading. The same is the case for the sediments. Because only drill sites over highs were able to reach the basement in the OCT, the oldest prerift sediments overlying oceanic as well as continental crust, suitable to date the age of breakup, have not yet been drilled. Dating magmatic rocks emplaced during continental breakup in the OCT was possible only for one albitite clast from a breccia drilled at Site 1070. U/Pb dating on a single zircon yielded an age of 127 + 4 Ma (Beard et al. 2002), which is the age of the overlying magnetic anomaly J. U/Pb ages on zircons from a meta-gabbro dredged at the Deep Galicia margin gave an age of 121.7 + 0.4 Ma, which is the age of the overlying magnetic anomaly M0 (Scha¨rer et al. 1995). In the Alps, the emplacement of gabbros in the ZECM is contemporaneous with the deposition of the first sediments overlying oceanic crust. Therefore, it seems that dating gabbros from the OCT is a reliable way to constrain the age of accretion of the surrounding crust. † Rates of extension during continental breakup. Rates of extension during final rifting can be determined from a kinematic inversion of profile Lusigal 12, presented in Figure 11. In this reconstruction, the thickness of the predetachment

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Fig. 11. (a) Geological interpretation of the depth-migrated Lusigal 12 profile showing the distribution of upper and lower crustal rocks exhumed subcontinental mantle rocks and reflections interpreted as detachment faults. Numbers on top of the section refer to ODP sites. (b) – (d) Evolution of faulting as determined from kinematic inversion of the interpreted Lusigal 12 profile shown in (b): (b) mantle exhumation along downwards-concave fault HHD, the bathymetry is constraint for this stage for Site 1069 (less than 1500 m) (Kuhnt Urquhart 2001; Urquhart 2001); (c) crustal extension accommodated by upwards-concave faults LD, HD and HHD; and (d) situation before onset of extension along LD, HD and HHD (it is important to note that the crust was already thinned to less than 10 km at this stage). For further discussion see the text.

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crust and the former crustal location of the ODP sites (circles labelled with ODP site numbers), as well as amount of thinning, are obtained from the step-by-step kinematic inversion of the profile. For details of the kinematic inversion and a more extended discussion see Manatschal et al. (2001). Within this inverted profile three time markers exist: (1) the Tithonian sediments drilled and dated at sites 901, 1065 and 1069, which are tilted and consequently predate the detachment faults; (2) the 40Ar/39Ar plagioclase ages of 136 Ma at Site 900 (Fe´raud et al. 1996) and 137 Ma at Site 1067 (Manatschal et al. 2001), dating cooling of the Hobby High basement rocks across the 150 8C isotherm and therefore representing a maximum age for first mantle exhumation; and (3) 128 Ma, the age of the M3 magnetic anomaly which is well identified in the Iberia margin (Whitmarsh & Miles 1995, Russell & Whitmarsh 2003). In order to determine the rate of extension one can calculate from the step-by-step kinematic inversion of the profile the total amount of extension that was accommodated along the LD, HD and the HHD in order to exhume the first mantle at the sea floor. The reconstruction shows that the total extension accommodated by LD, HD, and HHD is of the order of 34.7 km (initial, predetachment length is 34.6 km, final post-detachment length is 69.3 km). This estimate of extension is similar to the approximately 40 km of continental extension estimated from crustal thickness variations by Minshull et al. (2001) for the IAM9 profile 50 km to the south. This total amount of extension had to be accommodated between the Tithonian (145 Ma) and 136 Ma, when the first mantle rocks were within the reach of the sea floor. Using these numbers, an average velocity of 3.8 mm year21 is obtained. The velocity of mantle accretion in the ZECM between the first exhumation at Hobby High at about 136 Ma and the first magnetic anomaly M3 (128 Ma) 80 km further oceanwards is approximately 10 mm year21, similar to the half-spreading rate immediately to the west, which is approximately 10–14 mm year21 (Whitmarsh & Miles 1995). These values are of the same order as those observed at ultra slow-spreading ridges (Dick et al. 2003).

Extensional processes in magma-poor rifted margins The very similar strain evolution, stratigraphic record and isostatic response to rifting described from the Alpine Tethys and Iberia –Newfoundland

margins suggest that the underlying processes controlling lithospheric extension in these two pairs of margins are the same or at least very similar. This observation enables us to combine the complementary datasets from the two pairs of margins to describe different types of fault systems and deformation modes, both of which depend on the initial conditions and rheological evolution of the extending lithosphere.

Fault systems in rifted margins The previous descriptions show striking similarities between the observed fault systems in the Alps and seismically imaged structures in the Iberia –Newfoundland margins. Ho¨lker et al. (2002a) modelled synthetic seismic sections using density, velocity and thickness data from the well-exposed detachment structures in the Alps (Ho¨lker et al. 2002b, 2003), and compared the synthetic seismic profiles they obtained with seismic data and drilling results from Iberia. These studies provide some interesting insights into the seismic characteristics of detachments observed in the OCT, but also show the limitations of seismic imaging methods. Ho¨lker et al. (2003) demonstrated that intrabasement detachment structures such as the Err detachment fault (Fig. 9) would be transparent and therefore not imaged in a seismic section. Nevertheless, we believe that some of the reflections observed in the seismic sections in the Iberia– Newfoundland margins have their equivalents in observed fault structures in the Alps. Based on a combination of field observations with seismic reflection data we define three types of fault systems, which are described in the section below (Fig. 3). Upwards-concave faults. Reflections interpreted as upwards-concave faults are typically found in the proximal margins but occur also in the distal margins (e.g. H and L reflection; Fig. 8c). Because most of these reflections are observed on time sections, the downwards increase in seismic velocity tends to accentuate the apparent concaveupwards curvature. However, structures comparable to these reflections are also observed at land. In the Alps, the Ornon fault bounding the Bourg d’Oisans Basin (Chevalier et al. 2003) and the Lugano-Val Grande fault bounding the Generoso Basin (Bertotti 2001; Fig. 6b) were described as upwards-concave (listric) faults. On a map scale, the Lugano– Val Grande fault cuts across the prerift sedimentary cover into the basement over a horizontal distance of more than 30 km soling out in a greenschist facies mylonitic shear zone within a quartz-rich middle crust, i.e. at about 10–15 km depth (Bertotti 1991). Equivalent structures to the H and L reflections are likely to exist in more

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internal parts of the Alps, but were not identified in the Err –Platta OCT.

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reconstruction of this fault system, now preserved in the two conjugate margins, suggests strong riftshoulder uplift (Brianc¸onnais domain) contemporaneous with down-faulting and little upper crustal extension of its hanging wall (distal Adriatic margin) (Fig. 7). Thus, subdued subsidence and upper crustal normal faulting are the surface expression of the fault system that thinned the crust by more than 15 km. Structures comparable to the continentwardsdipping shear zones mapped in the Alps are not observed clearly in the Iberia margin. The only possible candidate is the C reflection. This reflection is cut by HD and the LD (Fig. 8c) and therefore predates final rifting. In contrast to the well-defined intrabasement reflections H and L, the C reflection is less sharp and coherent and has a more complex multicyclic nature (Ho¨lker 2001). The seismic response of ductile shear zones in continental crust has been investigated by Fountain et al. (1984), Hurich et al. (1985) and more particularly for the Ivrea zone by Holliger & Levander (1994). These authors showed that mylonitic shear zones in the continental crust (e.g. Pogallo fault) are able to produce reflections comparable to the C reflection (see also Ho¨lker 2001). If a shear zone is responsible for the C reflection, the reflection may additionally be amplified by a likely juxtaposition of mafic lower crustal rocks (e.g. rocks drilled at sites 1067 and 900) against upper crustal rocks. The continentwards-dipping mylonitic shear zones such as the Pogallo fault are key structures to explain the thinning of the continental crust during an early stage of rifting.

Downwards-concave faults. Downwards-concave reflections are not observed at the Iberia margin, but drilling results from the Iberia Abyssal Plain led to an interpretation of the top-basement reflection as a detachment fault. This structure forms the exhumed part of a downwards-concave fault dipping beneath the Newfoundland margin (Manatschal et al. 2001). The best-studied structure is the HHD, which was drilled during ODP legs 149 and 173. Analogue structures to the HHD are exposed in the Alps. More continentward parts of this fault system, not drilled off Iberia, are preserved in the Err nappe (e.g. Err detachment; Fig. 9). The transition from the continent to the exhumed mantle (e.g. Hobby High) is exposed in the Tasna OCT (Florineth & Froitzheim 1994) and situations corresponding to more oceanward portions (not drilled in Iberia) are found in the Platta nappe. These fault structures are characterized by: (1) a top-to-the-ocean transport direction; (2) a fault zone, some tens to some hundreds of metres across, affected by fluid- and reactionassisted brittle deformation processes and capped by a gouge horizon; (3) depositional contacts with sediments overlying the fault dips at a low angle; and (4) the local occurrence of hanging-wall blocks interpreted as extensional allochthons stranded on both exhumed continental and mantle rocks (Fig. 9). These faults, referred to as ‘topbasement detachment faults’ (Ho¨lker et al. 2003), are therefore interpreted as exhumed segments of no longer active downwards-concave detachment faults that accommodated tens of kilometres of offset and exhumed continental and mantle rocks at the sea floor.

Modes of lithospheric extension observed in magma-poor rifted margins

Continentwards-dipping shear zones. In the Adriatic distal margin, mylonitic shear zones separating lower from upper crustal rocks have been mapped in the Ivrea zone (e.g. the Pogallo fault: Hodges & Fountain 1984; Handy 1987; Handy & Zingg 1991) (Figs 1c & 7) and the Malenco area (e.g. Margna fault: Hermann & Mu¨ntener 1996; Bissig & Hermann 1999; Mu¨ntener & Hermann 2001) (Fig. 1c). P–T –t (pressure –temperature–time) data from the hanging-wall and footwall rocks from these shear zones show that they were active during an early stage of rifting and led to significant thinning of the continental crust by at least 15 km (0.4–0.5 GPa) (Handy & Zingg 1991; Mu¨ntener & Hermann 2001). Although the break-away zones of these faults are not preserved, we suggest in our reconstruction in Figure 7 that they formed the oceanward limit of the Brianc¸onnais domain in the most distal European margin. The

The geological–geophysical reconstruction of the Alpine Tethys and Iberia –Newfoundland margins (Fig. 3) resulted in the identification of three different modes of extension, referred to as ‘stretching mode’, ‘thinning mode’ and ‘exhumation mode’ (Lavier & Manatschal 2006). The description of the different modes relies on the geometry and the strain evolution of extending fault systems and the related basin architecture and isostatic response. The terms ‘stretching’, ‘thinning’ and ‘exhumation’ are not intended to be a quantitative description of the mode of deformation but rather to describe qualitatively how stain is accommodated during different phases of lithospheric extension. In contrast to the more general terms ‘wide rift mode’, ‘narrow rift mode’ and ‘core complex mode’ proposed by Buck (1991); the modes introduced here are based on observations and identify ‘common’ styles of extension observed in magma-poor rifted

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Fig. 12. Cartoons summarizing the sediment architecture, fault geometry and the strain distribution for: (a) stretching mode; (b) thinning mode; and (c) exhumation mode. For further explanations see the text.

RIFTED MARGINS AND CONTINENTAL BREAKUP

margins. The aim of this paper is to describe the geological characteristics of the three modes. Lavier & Manatschal (2006) modelled the single modes. Their results show that the various modes depend on a set of physical conditions such as bulk rheology and Moho temperature at the onset of rifting. Therefore, the modes described below enable us to relate observations to rift processes and, by modelling, to investigate the underlying physics controlling lithospheric extension (Lavier & Manatschal 2006). Stretching mode. The stretching mode is typically associated with fault systems cutting across the brittle upper crust and soling out at middle –lower crustal levels that behave ductilely and decouple deformation in the upper crust and upper mantle (Fig. 12a). Strain distribution in the upper crust is localized along high-angle faults with fault offsets of less than 10 km and producing a total extension, b, of less than 2. Strain distribution in the mantle is for this mode unconstrained by observations. This mode results in classical fault-bounded asymmetric rift basins, up to 6 km deep and 30 km wide, filled commonly by marine deposits (e.g. Bosence 1998). Rift-shoulder uplift is suppressed, as indicated by the subsidence of both hanging wall and footwall. In this mode uplift and erosion of the highs are not observed, exhumation processes are insignificant and thinning of the crust is moderate to minor. Accommodation space increases by vertical movement, the sedimentary bedding lies at a high angle to the fault, and the overall architecture of the basin is determined by sedimentary sequences thickening towards the footwall and onlapping onto the hanging wall (Fig. 12a). Examples of basins formed by this mode of deformation are the Jeanne d’Arc Basin (Fig. 5c), the Generoso Basin (Fig. 6b) and the Bourg d’Oisans Basin (Chevalier et al. 2003). All these basins formed at an early stage of rifting and in the proximal parts of the future margin (Fig. 3b). Thinning mode. The thinning mode is the most enigmatic phase of extension. This mode explains the thinning of the crust to less than 10 km (Lavier & Manatschal 2006). However, in contrast to the other two modes, it is less well constrained by direct geological observations mainly because the fault system and the sedimentary basins related to this mode are overprinted during later breakup and distributed across the two conjugate margins. Therefore, the overall system can not be observed in one piece. Portions of this system are observed in the Alpine margin. Examples of fault systems belonging to this deformation mode are the Pogallo and Margna faults in the Alps (Figs 1c & 7).

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P –T– t data indicate that these faults are very efficient in thinning the crust, and can explain local thinning to less than 10 km (Mu¨ntener et al. 2000). At the Iberia –Newfoundland margin the C ‘reflection’ (Fig. 8c) is in a similar position to the Pogallo and Margna faults in the Alps, but more observations are necessary to interpret this structure. Better examples for the thinning mode can be observed in parts of the margin where thinning was very pronounced but did not lead to continental breakup. Examples are the Galicia Interior Basin (Pe´rez-Gussinye´ et al. 2003), the Porcupine Basin (Reston et al. 2001) and the Rockall Trough (England & Hobbs 1997). In all these examples the continental crust has been thinned to less than 10 km without leading to significant normal faulting and major fault-related morphology and/or tilted blocks in the upper crust. Exhumation mode. The exhumation mode is characterized by downwards-concave faults that pull deeper crustal and mantle rocks from underneath a hanging wall and exhume them in the ZECM. This mode explains why the top of the basement in the ZECM must be considered as an exhumed fault plane. The fact that this fault exhumed deep-seated rocks to the surface and accommodated tens of kilometres of displacement without creating major fault-related topography is best explained by a downwards-concave fault geometry. Such downwards-concave faults have been proposed and referred to as rolling-hinge faults by Buck (1988) and Wernicke & Axen (1988) for the metamorphic core complexes in the Basin and Range in the SW USA. This model explains how large offsets are accommodated along active high-angles faults that rotated near to the surface to inactive lower-angle faults. A similar model has been proposed by Tucholke et al. (1998) to explain mantle exhumation at slow-spreading ridges, leading to so-called megamullions or ‘oceanic core complexes’. Although the overall geometrical concept of downwards-concave faulting is very similar in all three settings, geologically these three settings are characterized by different crustal and lithospheric thickness and a different thermal structure. The basin architecture and the stratal relationships associated with the exhumation mode are completely different from those of classical rift basins (Wilson et al. 2001). The basins form by pulling the footwall from underneath a hanging wall, i.e. the depositional area increases and is floored by an exhumed fault zone that also represents the main detritic sediment source (Fig. 12c). Thus, the relationships between the basal detachment and sediments is analogous to that of a conveyor belt, whereby the exhumed footwall rocks were fractured,

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exposed and redeposited along the same active fault system. This process explains the common occurrence of tectono-sedimentary breccias overlying detachment faults at low angles observed in the Alps and drilled off Iberia. These observations have also major implications for the interpretation of the stratigraphic record within OCT. In contrast to classical rift basins where synrift strata form wedges thickening into the footwall, in these basins the synrift sediments are dispersed and overlie the footwall at a low angle (Fig. 12c).

Mantle exhumation and magmatic processes during continental breakup Mantle exhumation processes in the ZECM. Concave-downwards faults can explain in a relatively simple way the exhumation of subcontinental mantle rocks over a wide area in the ZECM without producing a high-amplitude, fault-bounded topography. Explicit in this interpretation is that the top of the mantle represents a top-basement detachment fault. This is the case in the Alps (Fig. 9) and off Iberia (Fig. 10) where the mantle is capped by a brittle fault zone and overlain by tectonosedimentary breccias. As illustrated in Figure 13, the type of mantle rocks and their distribution in the ZECM strongly depends on the geometry of the exhuming fault system, the depth of the brittle–ductile transition, the number of faults and their overprinting relationships. In Figure 13a two end-member cases are illustrated. In one case (example to the left in Fig. 13a) the fault soles out at the crust –mantle boundary. As a consequence, the exhumed mantle in the ZECM is formed only by the subcontinental mantle that was directly underlying the crust before detachment faulting started. In this example, the overall deformation between crust and mantle is decoupled along the crust –mantle boundary, and the asthenosphere rises passively and thermally overprints the tectonically exhumed mantle. Decoupling between deformation in the crust and the mantle may be explained by serpentinization. At the other extreme (example to the right in Fig. 13a), the fault cuts into the mantle and soles out in a hot and ductile infiltrated lithospheric mantle overlying the rising asthenosphere. In this case the deformation between the crust and the mantle is coupled and, as a consequence, deep mantle levels can be exhumed in the ZECM. The asthenosphere is pulled out underneath the upper plate margin and the ZECM exposes a section across the lithospheric mantle. The results of Desmurs et al. (2001), Mu¨ntener & Piccardo (2003) and Mu¨ntener et al. (2004) demonstrate that the mantle rocks in the Alpine ZECM change from

spinel peridotite mixed with (garnet)-pyroxenite layers next to the continent to pyroxenite-poor peridotite that equilibrated in the plagioclase stability field further oceanwards. These changes in the nature of the mantle rocks in the ZECM are compatible with a detachment fault cutting into the mantle. In contrast, the structural observations clearly indicate that the detachment fault leading to the exhumation of the mantle to the sea floor are late, cold structures that were only active in the serpentine stability field. High-temperature mylonite shear zones that may have been responsible for the exhumation from deeper and hotter mantle portions are cut by colder detachment faults showing an opposite sense of shear (Fig. 9; photo D). Thus, simple monophase detachment models are unable to explain the overall observations related to mantle exhumation. Lavier & Manatschal (2006) proposed a polyphase evolution of detachment faulting in which hot and deep mantle was first exhumed and emplaced underneath a thinned crust during the thinning mode before it was pulled out along a late and cold detachment to the sea floor during the exhumation phase. In order to test this model more data are needed regarding the precise distribution of different mantle rocks in the ZECM and the timing and kinematics of mantle shear zones. In Figure 13b we illustrate how the interaction of detachment faults can explain characteristic features in the ZECM such as the formation of extensional allochthons or peridotite ridges. Propagating faults, i.e. faults that cut in front of the active fault, can explain the formation of allochthons that are soled by detachment faults (e.g. block drilled at Site 1069 imaged in the Lusigal 12 seismic section, Fig. 8c). This block is soled by a strong reflection interpreted as the continuation of the HHD, that separates a block formed by upper continental crust drilled at Site 1069 from the underlying mantle that was drilled at Site 1068 (Fig. 8c). The formation of peridotite ridges can be explained by back-stepping faults, which would explain the formation of the highs as well as the observation that these ridges are, in contrast to allochthons, not soled by strong reflections, i.e. detachment faults. However, such a mechanism would imply that the peridotite ridges formed during mantle exhumation, which is, at least for some examples (sites 897 and 899 and 1277), shown not to be the case (Pe´ron-Pinvidic et al. 2007). Therefore, backstepping detachment faults may form peridotite ridges, but they cannot explain the present day basement topography of the ZECM. Relation between extensional and magmatic processes. Although the Iberia –Newfoundland and the Alpine Tethys margins are commonly referred to as ‘non-volcanic’ rifted margins, several

a

asthenosphere

extensional allochthon

asthenosphere

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1

'lower plate' extrusive B magmatic rocks

a b

subcontinental mantle

MOHO

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thinned continental crust crust-mantle-boundary 'strong' subcontinental mantle

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Asymmetric

Fig. 13. (a) Geometry of detachment faults and its importance for the distribution of magmatic and mantle rocks in the ZECM: the example of a detachment fault soling out at the crust – mantle boundary (to the left) and a detachment cutting into the mantle lithosphere (to the right). (b) Interplay between detachment faults: sequence of propagating detachment faults cutting into the hanging wall leading to the formation of extensional allochthons (to the left) and detachment faults cutting in the footwall leading to the formation of peridotite ridges.

Formation of extensional allochthons

(b) number of detachments

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observations demonstrate that continental breakup within these two pairs of margins was preceded by magmatic activity. Detachment faults cut dolerites (Fig. 9; photo D), and infiltrated mantle peridotites are exhumed together with synrift gabbros along detachment faults at the sea floor. These observations from the Alps and Iberia clearly indicate that detachment faulting interacted with an active magmatic system before continental breakup. Therefore, we believe that the term ‘magma poor’ is more appropriate and should be used instead of ‘non-volcanic’ or ‘non-magmatic’ for the Iberia –Newfoundland and Alpine Tethys margins. At present, it is not yet understood how extensional faulting and emplacement of igneous rocks interact during final rifting, and how magmatic rocks are distributed in the ZECM. Modelling results by Lavier et al. (1999) and Huismans & Beaumont (2002) show that strong thermal/magmatic weakening suppresses asymmetric fault systems and results in localized symmetric systems classically described at mid-ocean ridges. The existence of magma clearly changes the rheological evolution of the margin, but what is the volume necessary to start sea-floor spreading and how do different distribution processes (dyking v. percolation) affect the bulk rheology of the extending lithosphere during continental breakup? The lack of data, in particular from the Newfoundland and European margins, makes it impossible to propose a single model. The evolution of a magmatic system, its relation to detachment faults and the distribution of the magmatic rocks in the ZECM are illustrated in Figure 13a. The major difference between the two models is the degree of asymmetry, the position of the asthenosphere and the distribution of magma in the extending system. In the extreme case where a single downwards-concave fault cuts across the whole lithosphere, all extrusive and shallow intrusive magmatic rocks are emplaced in the ‘upper plate’ margin, while the intrusive magmatic rocks and percolated mantle rocks that are genetically linked to the extrusive rocks are exposed in the ‘lower plate’ margin. Although it is generally accepted that the Iberia–Newfoundland and Alpine Tethys margins are asymmetric, surprisingly little is known about the conjugate margins: in Newfoundland drilling at ODP Site 1276 failed to sample basement rock, and in the Alps most of the European margin has been subducted. Therefore, more studies, in particular on the Newfoundland and Alpine European margins, are needed to constrain the generation and distribution of magma in magma-poor rifted margins, and to understand how magmatism interacts with serpentinization and deformation processes during final breakup.

Conclusions and outlook In this paper we have used observations from the Iberia –Newfoundland and Alpine Tethys margins to describe and discuss the tectonic evolution of an extending lithosphere during magma-poor rifting. The most pertinent observations are: † the polyphase nature of rifting; † the evolution from initially distributed to final localized extension; † the different stratigraphic, isostatic and tectonometamorphic evolution of the distal and proximal future margins during early rifting; † the observed close relationship between fault geometry, basins architecture and isostatic response, enabling us to distinguish between a stretching mode, a thinning mode and an exhumation mode; † the importance of exhumation processes during a final stage of rifting; † the complex relationship between mantle exhumation, detachment faulting and magmatic processes. Models of the tectonic evolution of the Iberia – Newfoundland and Alpine Tethys margins will need to capture these observations. Future studies need to link observations with processes and to discuss the underlying physics controlling lithospheric extension. It will be important to determine the physical conditions for each of the described modes and to understand what controls their transitions into other modes and how these transitions are documented in nature. Of particular importance will be the understanding of the complex interactions of mechanical and magmatic as well as chemical and thermal processes during rifting. It seems likely that each of the processes can change the bulk rheology and consequently control the strain evolution of the lithosphere during rifting. The increasing ability to use more realistic boundary conditions in modelling will also result in a demand for more complete and better datasets. In particular observations at a subseismic scale, constraints on age and palaeobathymetry, and data on tectono-metamorphic, hydrothermal and magmatic processes during final rifting are needed. Such observations are, however, difficult to obtain from rifted margins in situ. Future studies on land not only need to bridge the different observation scales, but also will need to quantify better the observations. The paper benefited from the careful reviews of N. Christie-Blick, A. Robertson and N. Direen that enabled us to improve the manuscript. We thank also A. Bouzeghaia for improving the quality of the figures. Support for this study came for the GDR marges

RIFTED MARGINS AND CONTINENTAL BREAKUP (France) (G. Manatschal and G. Pe´ron-Pinvidic), the Royal Society of London through a Joint Project grant (T. A. Minshull, G. Manatschal and G. Pe´ron-Pinvidic) and the Swiss National Science foundation (grant no. 21-66923 (O. Mu¨ntener and G. Manatschal) and PP002-102809) (O. Mu¨ntener).

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Overview of tectonic settings related to the rifting and opening of Mesozoic ocean basins in the Eastern Tethys: Oman, Himalayas and Eastern Mediterranean regions A. H. F. ROBERTSON Grant Institute of Earth Science, School of GeoSciences, University of Edinburgh, Edinburgh EH9 3JW, UK Abstract: A combination of geophysical studies and deep-sea drilling have in the past suggested that orthogonally rifted margins fall into two end-members: volcanic-rifted margins (e.g. eastern Greenland) and non-volcanic rifted margins (e.g. Iberia–Newfoundland conjugate). This paper explores the rifted margins of the Eastern Tethys stretching from the Eastern Mediterranean, through Oman to the Himalayas. Rifting in these area was typically pulsed, extending over more than 50 Ma. The timing of final continental breakup ranged from Late Permian in the east, in Oman and the Himalayas, to latest Triassic–earliest Jurassic in many parts of the Eastern Mediterranean (e.g. Antalya in SW Turkey; Pindos in Greece). Rifting in the Himalayas and Oman gave rise to a proximal to a distal ramp geometry with scatted seamounts (continental fragments and atolls) located adjacent to the rifted margin. The Eastern Mediterranean was palaeogeographically varied, and was characterized by a number of mainly elongate continental fragments (tens to several hundreds of kilometres long by tens of kilometres wide). These microcontinents subdivided the Eastern Tethys in the Eastern Mediterranean region into several small ocean basins, which rifted at more or less the same time in latest Triassic –earliest Jurassic time. All of the rifted margins of the Eastern Tethys are associated with rift-related volcanic rocks. However, with the exception of the Permian Panjal Traps in the Himalayas, the volumes of magma and corresponding thermal doming were less than for the ideal Volcanic-rifted margin (i.e. eastern Greenland). None of the Eastern Tethyan rifted margins show evidence of features characteristic of Non-volcanic rifted margins (e.g. sea-floor serpentinite exhumation), in contrast to the Iberia– Newfoundland conjugate or the Alps. Most of the Eastern Tethyan rifted margins appear to correspond to an ‘intermediate’ type, characterized by pulsed rifting, limited rift volcanism and a narrow continent – ocean transition zone. Such ‘intermediate-type’ rifted margins may remain to be explored in the modern oceans by deep-sea drilling. There is little evidence to support previous suggestions that the Eastern Tethyan rifts can be considered as back-arc basins above either northward- or southward-dipping subduction zones. Here it is suggested that the Eastern Tethys documents a fundamentally different type of rifting from either the ‘Volcanic-related’ or ‘Non-volcanic’ intracontinental rifts known from the Alps or the North Atlantic region. The dominant controls of rifting are seen as the traction of rising asthenosphere on the base of the lithosphere, related deviatoric tensional stresses, inherited and thermally induced weaknesses in the crust, and slab-pull. Specifically, in the Eastern Tethyan region continental breakup was probably triggered by a combination of long-term asthenosphere flow, slab-pull related to subduction beneath Eurasia and melt-induced crustal weakening associated with pulsed rifting or plume effects. Final continental breakup corresponds to a major (‘Cimmerian’) convergent phase along the opposing Eurasia margin, which further supports the role of plate boundary forces in Eastern Tethyan rifting. The early Mesozoic oceanic basins opened, probably associated with northwestward propagation of a spreading centre through the already weakened periphery of Gondwana, adjacent to less deformable Palaeotethyan oceanic crust. After a lengthy period of passive margin subsidence, locally punctuated by crustal extension and related volcanism, or plume effects, the rifted margins were finally tectonically emplaced during mid-Mesozoic, late Mesozoic or early Cenozoic time in different areas.

Our knowledge of the processes of continental rifting to form ocean basins is still far from complete. Only a handful of rifted margins have yet been studied by academic drilling and, as yet, there is not one complete study of a conjugate margin.

Marine geophysical studies of rifted margins leave many unanswered questions in the absence of direct sampling by drilling. In this respect the record of former rifted margins emplaced on land in orogenic belts around the world provides a

From: KARNER , G. D., MANATSCHAL , G. & PINHEIRO , L. M. (eds) Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 325– 388. DOI: 10.1144/SP282.15 0305-8719/07/$15.00 # The Geological Society of London 2007.

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valuable source of information. This is clearly shown by the Alps, which can be correlated with the rifted margin of the Central North Atlantic (e.g. Manatshal & Bernoulli 1999; Manatshal et al. 2007). A difficulty, however, is that rifted margins in orogenic belts have experienced collisional processes: they are typically deformed, metamorphosed and only sparsely preserved. Possibly for this reason the evidence of rifted margins in orogenic belts has received less attention in recent discussions of rift processes than they deserve. Here, we will take a ‘virtual fieldtrip’ around rifted margins preserved in the Eastern Tethyan region to facilitate comparisons with our knowledge of rifted margins in the Alps and the modern oceans. The discussion draws on the writer’s field experience of these areas, but it should be noted that it is not the intention to present new basic data here. The approach adopted is to focus on the preserved stratigraphic record of the continent– ocean transition zone of the rifted margins of the Eastern Tethyan region. This can be achieved through an integration of several lines of evidence: first, stratigraphic data from surface exposure and wells, mainly based on dating using fossils; secondly, sedimentological evidence, for example to shed light on proximal –distal sedimentary relations; thirdly, geochemical evidence that can help determine the tectonic settings of eruption of rift-related igneous rocks; and, fourthly, structural evidence that allows deformed units to be restored to their original relative positions along rifted margins. Additional information (e.g. from

palaeomagnetism, seismic reflection/refraction and subsidence curves) can be integrated where possible. The purpose of this contribution is, therefore, to provide an overview of rift settings and processes related to the opening of part of the Mesozoic Tethyan Ocean, generally known as the Eastern Tethys. The evidence comes from a wide region encompassing Oman, the Himalayas and the Eastern Mediterranean region (Fig. 1). The Tethys refers to ocean basins located between Africa and Eurasia that have now more or less entirely closed and been replaced by mountain chains (e.g. the Alps and the Himalayas). The Tethys is divisible into the Western Tethys and the Eastern Tethys. The Western Tethys is mainly located in the Western Mediterranean region; it opened during Jurassic time in concert with the opening of the Central Atlantic and closed during the latest Cretaceous –Early Cenozoic, influenced by the convergence of the African and Eurasian plates, creating the Alps and adjacent mountains (Bernoulli & Jenkyns 1974; Ogg et al. 1983; Manatschal et al. 2007). By contrast, the Eastern Tethys encompassed a much larger oceanic realm stretching from the Eastern Alps, through the Eastern Mediterranean, the Middle East, to Central Asia and beyond. It shows a long and complex history related to the opening and partial closure of a super ocean to the east, Panthalassa (e.g. Smith & Briden 1977; Scotese 1991). The Eastern Tethys is subdivided into an older ‘Palaeotethys’, essentially pre-Late Jurassic in age, that is

Fig. 1. Outline map of the Eastern Tethyan region showing the main areas discussed in this paper.

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widely believed to be rooted near the southern margin of Eurasian (e.g. Pontides of Turkey; Caucasus of Black Sea–Caspian region), and a younger ‘Neotethys’ that is generally seen as being located further south (e.g. Taurides of Turkey; Hellenides of Greece: e.g. Robertson & Dixon 1984; S¸engo¨r 1984; Dercourt et al. 1986). There is an on-going largely semantic discussion concerning the use of such terms as ‘Palaeotethys’ and ‘Neotethys’. Some authors have used these terms in different ways over the years, and, unhelpfully, have linked them to their own favoured tectonic models, thus ensuring that the terms are model dependent and inherently untestable. Here, we are concerned only with the younger Mesozoic –Early Cenozoic ‘Neotethys’, and we will refer to different parts of this system, informally, in our discussion of rift settings (e.g. South Tethys; North Tethys, etc.).

Volcanic v. Non-volcanic rifted margins In recent years rifted continental margins have been divided into ‘Non-volcanic’ and ‘Volcanic’ rifted types, based largely on the study of the Central and North Atlantic regions by the Ocean Drilling Program (Whitmarsh et al. 2001). The tectonic architecture of one conjugate Non-volcanic rifted margin setting has been explored, utilizing two legs of drilling on the Iberia margin and, so far, one leg of drilling on the Newfoundland margin (Tucholke et al. 2004, 2007). Additional evidence has come from the drilling of a relatively small number of other settings, notably the Equatorial African transform margin (Mascle et al. 1987), and from back-arc basins including the Western Mediterranean Tyrrhenian Sea (Kastens et al. 1988) and the SW Pacific Woodlark Basin (Taylor et al. 1999a, b; Goodliffe & Taylor 2007). Several other rifted margins were studied mainly using geophysical techniques and have been assigned, tentatively either to Non-volcanic or Volcanic-rifted margin types. Non-volcanic rifted margins include the Great Australia Bight –Wilkies Land (Antarctic) conjugate margin (Direen et al. 2007), whereas Volcanic-rifted margins include the Falklands– Malvinas and the Brazilian margin segment of the South Atlantic. It should be noted that the rift geometry and settings of many rifted margins beneath the world ocean have not been studied by academic drilling and remain largely unknown. Geophysical data alone (e.g. gravity, magnetic and seismic reflection studies) cannot distinguish contrasting margin settings with confidence. For example, in the absence of wide-angle seismic refraction data it is impossible to determine whether inferred seaward-dipping

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seismic reflectors may simply represent rift-related volcanics or instead true marginal oceanic crust. In some areas (e.g. eastern Brazil margin) little faulting is imaged seismically in the upper crust, which might give the false impression that little extension has taken place when in reality much extension was accommodated by exhumation of lower crust –lithospheric mantle. Such strain localization has implications for the interpretation of ancient rifted margins in orogenic belts where only the stratigraphically higher levels of the rifted margins are typically preserved. The theoretical subsidence curves of Volcanic v. Non-volcanic margins differ markedly (Skogseid & Eldholm 1988), but this is of limited value in practice as few if any volcanic rifted margins have been deeply drilled and accurately dated. At present, the available information suggests that contrasting Non-volcanic-rifted and Volcanic-rifted settings may characterize different segments of rifted margins, as inferred around Australia (Direen et al. 2007). However, to what extent are the Non-volcanic v. Volcanic margin types in reality end members of a spectrum of margins that exhibit more or less evidence of volcanism and fault-controlled extension (i.e. ‘intermediate’ rift types)? Some rifts might show a combination of rift volcanism and seaward-dipping reflectors, as suggested for the Baltimore Trough and Carolina Trough (White & McKenzie 1989), although deep-sea drilling is needed to confirm this. The evidence presented in this paper suggests that such ‘intermediate-type’ rifts characterize most of the rifted margins of the Eastern Tethys.

Recognizing rift settings in orogenic belts Before proceeding it is useful to consider what are the key features of both of the end-member rifted margin types, as currently understood, that might be preserved in a mountain belt. The essential features in the stratigraphic record that could allow different tectonic settings to be recognized (e.g. rift; accretionary prism; foreland basin) have been identified and termed tectonic facies (Robertson 1994). In mountain belts generally the proximal (inboard) parts of rifted continental margins are formed on relatively thick continental crust and these are commonly well preserved. By contrast, the distal (outboard) parts of rifted margins are characterized by markedly thinned continental crust and/or marginal oceanic crust, which rarely survive collision and emplacement processes. Where preserved, such distal units typically take the form of imbricate thrust sheets or chaotic units known as me´lange that underlie large emplaced

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ophiolites (e.g. Robertson 2002, 2004). Typically, we are only able to observe the more proximal, stratigraphically higher parts of the former continent– ocean transition. The Alps, where deeper crustal levels are exposed (e.g. intrusive igneous rocks), is an exception. For an ideal Volcanic-rifted margin the basket of features that could allow this setting to be identified in the field include thick proximal –distal units ranging from subaerial flood basalts (massive flows) and subordinate pyroclastic sediments (e.g. air-fall tuffs), to subaqueous pillow lavas, hyaloclastites and volcaniclastic turbidites (Planke et al. 2000; Campbell 2001; Rainbird & Ernst 2001). In some cases such features (e.g. dyke swarms and flood basalts) cover wide areas and could be preserved subaerially after breakup and subsidence (e.g. dyke swarms of the eastern Greenland margin: Karson & Brooks 1999). However, in other areas (e.g. Brazil margin) there is less evidence of a volcanic-rifted margin onshore where mainly basement rocks are exposed. In orogenic belts evidence of magmatism (e.g. sheet flows) could be exposed in foreland areas that underwent less deformation than the main suture zone or escaped later collisional deformation altogether. Domed uplift of at least several kilometres is expected resulting from thermal effects (probably related to magmatic underplating: White & McKenzie 1989) and this would result in an important regional unconformity. On a subsided Volcanic-rifted margin, an ideal (imaginary) deep hole would pass downwards through deep-sea sediments into thick sequences of subaqueous basalts, then subaerially erupted flood basalts, at least hundreds of metres thick, showing a plume-type geochemical signature. Beneath would be a rifted basement and probably some synrift terrigenous sediments. By contrast, an ideal Non-volcanic rifted margin would be characterized by crustal thinning, normal faulting and related footwall uplift affecting a wide area (Ziegler & Cloetingh 2004). Lithosphere extension is generally seen as being accommodated by faulting in the upper crust, although there are exceptions (e.g. eastern Brazil margin: Karner pers. comm. 2005). Lithosphere stretching may lead to rift magmatism (White & McKenzie 1989; Kendall et al. 2005). Rifting is commonly pulsed, lasting over tens of millions of years. Accompanying flexural uplift can lead to deep erosion of a rift shoulder. The structural style varies greatly depending on whether rifting is symmetrical or asymmetrical. Passive margin subsidence ensues after sea-floor spreading begins (Watts 1982; Karner pers. comm. 2005). As extension intensifies prior to continental breakup deep crust and upper mantle are exhumed and covered directly by

mid-ocean ridge (MOR)-type basalts, based on evidence from ODP Site 1277, Leg 120 (Tuckolke et al. 2007; Robertson in press) and the Alps (Manatschal et al. 2007). Thus, an ideal deep-sea core through a Nonvolcanic rift basin directly on the locus of final continental breakup would be expected to pass through a deep-sea terrigenous passive margin sequence, into MOR-type basalt (tens to several hundreds of metres) and then into exhumed peridotite (mantle lithosphere). In principle, it should be relatively simple to differentiate between a distal Volcanic-rifted v. a Non-volcanic rifted margin setting, where preserved in an orogenic belt. However, uncertainties include the following: (1) the ‘ideal’ sections have never been completely drilled and are thus questionable; (2) ‘intermediate’ types of rift may exist, as noted above; (3) the key lower parts of the succession may not be preserved or exposed; and (4) the rift may be asymmetrical with key information missing from the conjugate margin where this is not preserved. These problems can be minimized if the entire rift setting is considered including proximal rift units, which have a higher preservation potential. Also, other evidence can help, notably the geochemistry of the rift-related extrusive rocks, which may help discriminate between, for example, plume- and non-plume-influenced settings.

Nature of the database With the reader in mind who is not a field geologist it is worth spending a moment to outline the sources of information available for syntheses of rift settings in orogenic belts. Our knowledge of the geology of each of the areas discussed (Himalayas, Oman and Eastern Mediterranean), like many others, stems from more than 100 years of research. Each area needed to be geologically mapped, a process that in areas such as the Himalayas remains incomplete, and the units geologically dated mainly using fossils and radiometric techniques. Each unit has to be characterized using a range of, for example, petrological, sedimentological, geochemical, geophysical and other techniques. Structural geology is critical to allow units to be restored to their predeformational state as far as possible. As information increases it becomes possible to produce regional, to global, palaeogeographic maps and plate-tectonic models. Commonly there are inherent assumptions in such plate-tectonic syntheses; for example, as to whether most ophiolites (i.e. emplaced oceanic lithosphere) have formed at divergent spreading centres or at convergent margins (above subduction trenches). The role of rift settings in several alternative plate-tectonic

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models of the Eastern Tethys is outlined towards the end of this paper. Briefly, some of the main contributions that have provided the regional database for this study of rift settings are as follows. For Oman, following early work, it was the search for hydrocarbons that resulted in the first comprehensive mapping and regional study (Glennie et al. 1973), followed by research triggered by international interest in the Oman ophiolite (Coleman 1981; Lippard et al. 1986; Nicolas 1989). Mapping of the country as a whole was completed by a national geological survey under contract to the government of Oman (Le Metour et al. 1995). At the same time, academics from a number of countries, including Britain, France, Japan, Switzerland and the USA, provided much additional, more specialist information, and Omani nationals increasingly contributed. For the Himalayas, much of the relevant information came from numerous expeditions over many decades especially by Italian, French and British groups, with Indian, Pakistani and other Earth scientists playing an increasingly important role. In the Eastern Mediterranean the contributions are even more diverse with the nationals of each of the many countries in the region providing much information, supplemented by long-standing research traditions, especially in France, Germany and the UK. The present summary of rift settings draws on all of the above contributions. Nevertheless, the writer is one of the very few people who has first-hand field experience (over 35 years) of all of the rift settings discussed here (Oman, the Himalayas and the Eastern Mediterranean), and this is the first attempted synthesis of the rift settings from these regions as a whole based on individual fieldwork.

Rifted margins of the Eastern Tethys In slightly more detail, the Eastern Tethyan region extends from the Eastern Alps, through former Yugoslavia, the Balkans, Greece, Turkey, former southern USSR, the Middle East, Afghanistan, to the Himalayas and beyond (Fig. 1). The northern margin of the Eastern Tethys corresponds to the southern margin of Eurasia (Fig. 2). The southern margin of the Eastern Tethys is effectively the northern margin of the supercontinent, Gondwana and satellite microcontinents. During the Mesozoic, Gondwana episodically fragmented and continental units of various scales and locations drifted away, many towards Eurasia (S¸engo¨r et al. 1984; Scotese 1991). These fragments range from India (on a continent scale), to Madagascar (Hanken 1994), the Iranian and Tauride microcontinents (S¸engo¨r 1984), to smaller fragments (e.g.

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Fig. 2. Simplified palaeogeography of the Eastern Tethyan region. Continental fragments rifted from Gondwana and drifted towards Eurasia. Continental breakup took place during the Late Permian in Oman and the Himalayas, but not until the latest Triassic– earliest Jurassic in the Eastern Mediterranean. See the text for explanation.

Seychelles: Plummer & Belle 1995). Many other continental fragments probably remain to be discovered beneath the world ocean and in orogenic belts (Smith 1999). Continent-scale – microcontinent-sized fragments within the oceans were surrounded by rifted margins of various scales and types. Some of these margin units are now preserved, albeit deformed, in mountain belts. Exceptionally, a few of the rifted margins of the Eastern Tethys remain as undeformed deep-water successions. The most notable example is the North African –Levant margin of the Central and Eastern Mediterranean, that remains to the south of the Alpine– Mediterranean orogenic front (e.g. Garfunkel 1998, 2004, 2006). Of such undeformed rifted margins, only the NE margin of Australia (Baillie et al. 1994) has been explored so far by deep-sea drilling (von Rad et al. 1989). Evidence of exotic units in Indonesia suggests that continental fragments were rifted from the Australian continent and drifted northwards, followed by accretion (Pigram & Pannabean 1984). The possible driving mechanisms of the rifting of such continental fragments are discussed towards the end of this paper. Most of the presently available useful information concerning rift settings comes from a few areas, notably: (1) the Oman region; (2) the Himalayas; and (3) the Eastern Mediterranean region. Each of these regions shows evidence of rifting that was essentially orthogonal to the adjacent continent. By contrast, several other areas appear to have been characterized by

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transform-rifted (or oblique-rifted) settings (e.g. western Pakistan; SE Oman; East Africa), but these will not be considered further here. Below, for each of the regions (Oman, the Himalayas and the Eastern Mediterranean) the settings of rifting will be summarized, alternative interpretations will be outlined and a restored margin geometry will be proposed. We will also consider if each of these cases corresponds to either a Volcanic-rifted margin or a Non-volcanic rifted margin end member.

Oman rifted margin Rift-passive margin evolution is excellently exposed in the Oman Mountains, and has probably been studied there in more detail than any other South Tethyan region (Fig. 3). This area exposes the rifted margin of the Arabian continent that was deformed and tectonically emplaced in latest Cretaceous time (Fig. 4). After this emplacement passive margin conditions resumed on the Arabian margin, which has only more recently begun to show the signs of continental collision with the Eurasian margin in the north (i.e. adjacent to the Zagros; Fig. 3). The rift and passive margin units in Oman have thus never experienced ‘full’ continental collision, in contrast to the Himalayas or the Eastern Mediterranean regions discussed later in the paper. The emplacement of the rift/passive margin units onto the Arabian continent margin (Fig. 4) is widely interpreted as being driven by the collision of the Arabian continental margin with a subduction trench (Glennie et al. 1973;

Fig. 3. Sketch map of Oman showing the location of the emplaced rifted margin sediments and volcanics located beneath and to the west of the overriding Semail ophiolite.

Lippard et al. 1986; Robertson 1987). In this model, Tethyan oceanic crust was consumed in a NE-dipping intra-oceanic subduction zone until the trench collided with the margin; the higher levels of the passive margin were then detached and emplaced continentwards. Parts of the leading edge of the passive margin (in the south) were subducted and underwent high pressure– low temperature (HP–LT) blueschist metamorphism, followed by rapid exhumation, by latest Cretaceous time (El-Shazly et al. 1990; Searle & Cox 1999; Searle et al. 2004). Prior to the Upper Cretaceous deformation, the Arabian continental margin in Oman can be restored as an orthogonal rift segment, approximately 800 km long. This was bounded a by a major transform margin in the south (Masirah margin) and by a smaller transform (or obliquely rifted segment) in the north (Dibba zone) (Cooper 1990; Robertson & Searle 1990). The conjugate margin of the Oman rift is assumed lie within Iran to the north. It has been suggested that a continental fragment was rifted from Oman, the SanandajSirjan zone (Fig. 2), and that Neotethyan oceanic crust separated this from the Lut continental block to the north (Glennie 2000). However, insufficient work has been carried out in these areas to allow detailed correlation. Evidence of initial rifting comes from the undeformed continental margin, rather then the overriding, tectonically transported units. The in situ continental margin exhibits a variably complete Late Precambrian–Early Palaeozoic prerift sedimentary sequence that was unconformably overlain by a Late Permian transgressive succession, beginning with thin clastics and passing directly into much thicker shallow-marine carbonates (Glennie et al. 1973; Rabu et al. 1990). The unconformity was exploited as a Late Cretaceous emplacement related thrust in some areas, whereas in others a spectacular angular unconformity is exposed (Mann & Hanna 1990). The underlying Pre-Permian sequence is deformed, cleaved and folded and was subjected to regional low-grade metamorphism in contrast to the overlying Upper Permian and younger sequence. The deformation and low-grade metamorphism are widely explained in terms of a ‘Hercynian’ orogenic event affecting the South Tethyan margin, including Oman, during Late Palaeozoic time (Mann & Hanna 1990; S¸engo¨r 1990). On the other hand, Stampfli et al. (2001) suggest that the absence of preserved sequences between the Late Precambrian–Cambrian? basement and the Late Permian transgression (at least in the Jebel Akhdar; Fig. 3) resulted from rift-related uplift and deep erosion (i.e. rift-shoulder uplift). Rift-related uplift and erosion has indeed taken place, as subsidence curves imply an exhumation

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Fig. 4. Simplified tectonostratigraphy showing the relation of the rifted margins sediments and volcanics (Hawasina Complex) beneath the Semail ophiolite, also the relative positions of the Haybi Complex and the Oman Exotics. Some of the Hawasina Complex was also ‘bulldozed’ ahead of the Semail ophiolite (in the Hamrat-ad-Duru). From Robertson (2004); based on Glennie et al. (1973).

of up to 1 km (Stampfli et al. 2001). However, a purely extensional setting is unlikely to explain the pervasive deformation and low-grade metamorphism of the pre-rift successions in the central Oman Mountains (i.e. Jebel Akhdar; Saih Hatat), and it is likely that the area experienced compressional deformation prior to rifting and continental breakup during the Permian. The overlying Permian rift-related sequence passes upwards into thick platform carbonates of Triassic –Late Cretaceous age, several kilometres thick, forming a regional-scale carbonate shelf sequence that is well exposed in the Jebel Akhdar, Saih Hatat and the Musandam Peninsula (Glennie et al. 1973; Rabu et al. 1990) (Fig. 5a). A foredeeptype sequence (Muti Formation) of Late Cretaceous age at the top the carbonate platform succession records flexural subsidence of the platform related to overthrusting of the margin and ophiolitic units from the Tethyan Ocean (Robertson 1987).

The relatively in situ carbonate platform (Jebel Akhdar) is overthrust by a stack of thrust sheets (Fig. 4). In general, higher thrust sheets originated progressively further out into the Tethyan Ocean. However, out-of-sequence thrusting has played a role in some areas so that an accurate reconstruction cannot always be guaranteed simply by unstacking the thrust sheets (Robertson & Searle 1990). The lowest major thrust sheet is the thickest, most intact, and preserves the Upper Permian and Mesozoic proximal slope successions related to the carbonate platform (Sumeini Group; Fig. 4). This unit was thrust only a relatively short distance onto the in situ carbonate platform (several kilometres), and is traditionally described as being ‘parautochthonous’ (Glennie et al. 1973). Successions begin with Upper Permian shelf-to-slope redeposited carbonates that document a rifted carbonate platform edge setting (Watts & Garrison 1986). Above this, the succession records relatively

Fig. 5. Field photographs of the Oman rifted margin. (a) Mesozoic carbonate platform bordering the Mesozoic Tethys in the Musandam Peninsula (north of the Dibba Zone), Northern Oman Mountains; (b) proximal passive-margin slope sequence in structurally low thrust sheet (Upper Jurassic– mid-Cretaceous interval); Wadi Mi’adin; Central Oman Mountains; (c) distal slope/basin-plain sequence in structurally higher thrust sheets (Upper Jurassic –Lower Cretaceous); Hamrat ad Duru; west of the Central Oman Mountains; (d) Triassic Oman Exotic structurally overlying the Jebel Akhdar carbonate platform succession. Within this exotic (Jebel Misfah) Triassic alkaline lavas and sills are covered by thick Late Triassic shallow-water carbonates. See the text for discussion.

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proximal slope conditions (Fig. 5b) throughout the Mesozoic until Late Cretaceous time (Watts & Garrison 1986) when syntectonic sedimentation appeared (Muti Formation: Robertson 1987). Successions characteristic of the passive margin slope include detached blocks, slump sheets and debris flows, mainly composed of limestones derived from the adjacent carbonate platform. There are also spectacular slope channels infilled with mainly shallow-water carbonates deposited by mass-flow processes. The slope sequences are typically overlain by the Hawasina Complex, made up of a thick pile (several kilometres) of relatively thin thrust sheets, individually tens to several hundreds of metres thick (Figs 4 & 5c). In general, the stratigraphically thicker thrust sheets exhibit relatively proximal slope facies, whereas the stratigraphically thinner thrust sheets preserve more distal deepwater successions (e.g. Hamrat Duru Range, central Oman Mountains; Fig. 3). However, out-of-sequence thrusting complicates the restored geometry in some areas. The most proximal successions are exposed structurally above the Jebel Akhdar carbonate platform, or the Sumeini Group (e.g. Wadi Guweyza, central Mountains). These sediments are similar to the Sumeini Group slope sequences. The slope and basinal units represent preserved parts of a once contiguous depositional system (Bernoulli & Weissert 1987; Cooper 1990; Blechschmidt et al. 2004). In general, the oldest facies preserved in the Hawasina Complex are Early Triassic redeposited limestones, passing upwards into shales and then into thin radiolarian cherts of Middle Triassic age. These are overlain, in turn, by Late Triassic redeposited carbonates and quartzose sandstone turbidites (Zulla and Al Ayn formations). The Early Jurassic is dominated, after a probable hiatus, by shales, radiolarian cherts and subordinate redeposited carbonates and sandstones. The Middle Jurassic is mainly limestone turbidites with abundant redeposited ooids derived from the Jebel Akhdar carbonate shelf (Guwayza Formation). Above this, the Upper Jurassic –Lower Cretaceous interval is characterized by radiolarian chert (Sid’r Formation). Early Cretaceous (Hauterivian– Barremian) time saw a return to the accumulation of redeposited carbonates (Nayid Formation). The successions within the well-ordered thrust sheets of the Hawasina Complex (as exposed in the Hamrat Duru Range; Fig. 3) are no older than Early Triassic. However, an isolated short intact succession of pillow lavas that is depositionally underlain by Upper Permian radiolarian cherts has been discovered in the Hawasina window (De Wever et al. 1990). These Permian volcanics are geochemically transitional between alkali

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basalts and mid-ocean ridge basalt (MORB) (Maury et al. 2001) and, on this basis, could have erupted in either transitional oceanic or seamount settings. A plume-related setting was suggested mainly on the basis of ‘enriched’ chemistry (Lapierre et al. 2004). The Hawasina Complex of mainly deep-water sediments is structurally overlain by a me´lange known as the Haybi Complex (Searle & Malpas 1980) (Fig. 4). This includes up-to-mountain-sized blocks, the Oman Exotics (Glennie et al. 1973) that are critical to the interpretation of the Oman rifted margin. There are two main types of exotic. The first type of exotic includes Permian shallowmarine carbonates, and is represented by the Jebel Qamar exotics in the Northern Oman Mountains (Dibba Zone), and the second is the Ba’id Exotic in the southern Mountains (south of Saih Hatat; Fig. 3). These exotics include thin Permian shallowwater limestones overlain by Triassic limestones. Jebel Qamar is unique, as Late Permian carbonates there are stratigraphically underlain by Lower Palaeozoic siliciclastic sediments (Hudson 1960; Robertson et al. 1990; Pillevuit 1993). These sediments are reportedly overlain locally by a thin unit of shallow-marine sediments of Early Carboniferous age, which are associated with alkaline basalts (Niko et al. 1996). This is significant as it suggests that extension may have been active as early as the Carboniferous as a precursor to Permian rifting, at least in this area. The Late Permian Jebel Qamar shallow-water carbonates are overlain without a break by Late Triassic neritic carbonates. Further south, the huge Ba’id Exotic (Fig. 3) begins with Upper Permian platform carbonates and passes upwards into Middle–Late Triassic pelagic limestones, including slowly deposited (i.e. condensed) facies (Hallstatt facies) and pink ammonite-rich facies (Ammonitico Rosso). In addition, in the Central Mountains (north of Jebel Akhdar; Fig. 3) the small Rustaq Exotic includes a short succession of Upper Permian pelagic carbonates, underlain by basic volcanics of transitional to MOR type (Pillevuit 1993). The second and more numerous type of exotic is dominated by large masses of Upper Triassic shallow-water limestone (Fig. 5d). These carbonates are underlain by basic volcanic rocks and overlain, in some areas, by a thin pelagic carbonate succession. The type example of this exotic is Jebel Misfah, which is located along the NW margin of the Jebel Akhdar carbonate platform (Fig. 3), and includes Jebl Kawr, making up a vast exotic mass that is tens of kilometres long. Voluminous lavas and sills beneath the Upper Triassic neritic carbonate successions are geochemically of within-plate type (Lippard et al. 1986). Overlying shallow-water carbonates, several thousands of

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metres thick, of Late Triassic Carnian –Norian age are covered by ammonite-bearing pelagic carbonate (Ammonitico Rosso) of Early– Middle Jurassic age and then by additional pelagic carbonates of Late Jurassic –Early Cretaceous age (Glennie et al. 1973; Pillevuit et al. 1997). The sediments of the deeper-water Hawasina Complex (Fig. 4) locally include reworked limestones of both Late Permian and Late Triassic ages in different areas. For example, the Late Triassic Misfah Exotic is in structural contact with Late Triassic radiolarites and redeposited limestones, including numerous up-to-house-sized blocks of shallow-water carbonate (e.g. along the south margin of the Jebel Akhdar, near Al Hamra). Elsewhere, isolated detached blocks of Late Triassic neritic, or pelagic, carbonate are locally present within the mid part of the stack of thrust sheets forming the Hawasina Complex (Cooper 1990). The Oman Exotics are in places structurally overlain by the metamorphic sole of the Oman ophiolite, which ranges upwards from greenschistto amphibolite-facies conditions (Glennie et al. 1973; Searle et al. 1980; Lippard et al. 1986) (Fig. 4). The low-grade metamorphic rocks include alkaline volcanics, meta-cherts and metacarbonates. The carbonates include recrystallized equivalents of the Oman Exotics in some areas (e.g. Hawasina window), which is notable as it helps to confirm that these exotics formed in a relatively distal position along the margin or within the South Tethyan Ocean. The various units exposed in different structural units in Oman document rifting related to opening the South Tethys. Several different interpretations have been proposed for the timing and setting of this rifting. The main debatable issues are the palaeogeographical setting of the two different types of Oman exotics and the nature of the crustal basement on which the deep-sea Hawasina Complex formed. A complicating factor is that some of the exotics are located in a high structural position above the distal margin units of the Hawasina Complex (Fig. 4), whereas others, including the huge Misfah Exotics, are located near the base of the thrust stack close to the autochthonous Jebel Akhdar carbonate platform (Fig. 3). There are two main possible explanations of this tectonic setting. First, the Triassic exotics might have formed in originally different palaeogeographic settings on the rifted margin; for example, with the exotics like Jebel Misfah located in a proximal position near the parent platform but other Triassic exotics (e.g. those of the low-grade metamorphic sole) formed more distally. Secondly, the Triassic exotics might all have formed in a relatively distal (oceanward) setting but then have reached their present position in response to structural

complexities during emplacement. Plausibly, the bulldozing of the thinner, more incompetent, deep-sea sediments of the Hawasina Complex took place ahead of the advancing Semail ophiolite nappe leaving the thick relatively competent lavas and carbonates of the Triassic exotics (e.g. Jebel Misfah) in direct contact with the in situ Jebel Akhdar carbonate platform beneath. In reality, both factors are likely. The Triassic exotics formed in a range of palaeographical setting probably in a relatively outboard (oceanic setting), but they sometimes also experienced complex thrusting which in some cases left exotics in an anomalously low structural position (e.g. Jebel Misfah). Based on mapping of the entire Oman Mountains, Glennie et al. (1973) initially proposed two alternative reconstructions of the rifted margin. In the first model, the ocean began to open during Permian time: the Permian exotics formed part of the rifted margin edge (i.e. contiguous with a larger parent carbonate platform), whereas the Triassic exotics developed within the ocean basin at a later stage (i.e. as seamounts). In a second model the ocean opened during the Triassic, following earlier rifting. In this case, the Permian exotics relate to synrift faulting, whereas the Triassic exotics formed at or near a Triassic spreading ridge. Both of these models, either of Permian or Triassic continental breakup, were developed by later workers. Members of the Open University (UK) research group favoured Triassic continental breakup, with the Permian exotics recording rifting (i.e. as synrift facies) and the Triassic exotics being seen as igneous seamounts (i.e. postrift facies) capped by carbonate build-ups within the Tethyan Ocean (Searle & Graham 1982; Lippard et al. 1986). Following the discovery of Late Permian radiolarites associated with basaltic volcanics of alkaline– transitional MOR type, others favoured Permian continental breakup following pulses of rifting since Late Carboniferous time (Blendiger et al. 1990; Stampfli et al. 2001). In this interpretation it was generally accepted that the Triassic exotics (e.g. Jebel Misfah) originated as oceanic seamounts (atolls). However, the Late Permian exotics could be seen, either as part of the rifted margin of a parent Jebel Akhdar carbonate platform (Blendiger 1988; Pillevuit 1993; Pillevuit et al. 1997) or as being formed within the Permian ocean after continental breakup (Bernoulli & Weissert 1987; Be´chennec et al. 1988). Glennie (2000) has also suggested a two-stage rifting model. A narrow ocean basin opened in the Late Permian (‘Neo-Tethys 1’) bordered by Permian platform units on both margins. The Qamar and Ba’id exotics represent parts of the eastern margin of this small ocean basin (‘Kawr Ridge’). The rift axis then ‘jumped’ oceanward

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during the Triassic, and the NE margin of the initial rift remained as part of the subsequent SW margin of a new ocean basin of Triassic age (‘Neo-Tethys’) that opened to the NE. In this interpretation the Upper Triassic exotics (e.g. Jebel Misfah) would relate to Triassic rifting of a more easterly, younger basin, rather than as oceanic seamounts as in most other interpretations. However, in the writer’s view it is doubtful if the localized Permian exotics really represent the easterly margin of a Permian rifted ocean basin, rather than merely fragments of a single westerly Permian rifted margin; these were left stranded when a Permian ocean, initially narrow, later widened with the genesis of Triassic oceanic crust. The question of whether or not the Late Permian exotics with shallow-water carbonates originated along the rifted margin of Arabia or within a Permian ocean is central to an understanding of the rift setting in Oman. Here, it is important to note that (unusually) the Hawasina Complex in the area of the Ba’id Exotic includes Permian deepwater sediments (Blendiger 1988), with carbonate turbidites (reef-derived), radiolarites and minor pelagic limestones (with cephalopods) all being present. These deep-water successions have been restored as parts of a south-facing carbonate slope/basin (Blendiger 1988). The most proximaltype slope facies are located in a structurally high position in the local stack of thrust sheets, whereas the more distal facies are located near to the base of the thrust sheet. This suggests that the relatively proximal facies restore to a position within the rift basin, away from the parent carbonate platform (Jebel Akhdar). However, Blendiger (1988) believed that the proximal facies instead originated in a setting close to the Jebel Akhdar (effectively as part of it) and that they has reached their present structural position is response to outof-sequence thrusting. In this interpretation, the distal facies were first thrust continentwards (towards the SW) over more proximal facies. The thrust sheets were then re-thrust, placing the originally proximal units at the highest level of the thrust stack. These Hawasina thrust sheets structurally overlie the large Bai’id Exotic just to the north, which, as noted above, includes Permian neritic carbonate and Triassic pelagic carbonates. The Ba’id platform unit was assumed to have formed part of the Jebel Akhdar platform during Permian–Mesozoic (Blendiger 1988; Pillevuit 1993). However, structural evidence was not presented to support this interpretation. In the author’s view the Ba’id Exotic is unlikely to represent part of the original Jebel Akhdar carbonate platform for several reasons: First, the Ba’id Exotic is structurally underlain by lithologies including volcanics and radiolarites, typical of the

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distal Haybi Complex in other areas, rather than the proximal slope facies (i.e. the Sumeini Group). Secondly, if the Ba’id Exotic was contiguous with the parent Jebel Akhdar platform it is unclear why a Jurassic–Cretaceous neritic succession was not deposited there as well; structural removal is a possibility, but why? It is suggested here that the Ba’id Exotic represents a small, rifted platform unit that was positioned outboard of the main carbonate platform (Jebel Akhdar) and surrounded by oceanic crust of Upper Permian age. The Permian and younger facies described by Blendiger (1998) would have accumulated on the oceanward side of this exotic platform unit, far from the parent Jebel Akhdar carbonate platform. A similar problem (and possible solution) applies to the Permian Jebel Qamar Exotics in the Northern Oman Mountains (Dibba Zone). A high structural position, and thus a relatively outboard palaeogeographical setting, was inferred by Robertson et al. (1990) (Fig. 6), a setting also favoured by Glennie (2000). However, Pillevuit (1993) re-interpreted these exotics as part of the autochthonous Jebel Akhdar carbonate platform. In this interpretation the exotic material has somehow to be relocated from within the autochthon to a structurally high position above the Hawasina-type rocks in this area during Late Cretaceous emplacement. However, no supporting structural evidence for this was provided. If the Permian exotics formed part of the proximal rifted margin of the Jebel Akhdar carbonate platform how were they exhumed from deep within the regionally thick (many kilometres) Permian–Late Cretaceous shallow-water carbonate succession? The Permian exotics lack any known post-Triassic sequences, which would need to have been structurally removed without trace in this interpretation. The author considers it more likely that the Jebel Qamar Exotics formed in a relatively outboard (oceanward) setting, where shallow-water carbonates were never deposited as the exotics had by then subsided into deep water. In summary, a reconstruction of the rifted Oman margin is shown in Figure 7. The Ba’id and Qamar exotics are seen as rifted continental fragments (detached fault blocks) within Upper Permian oceanic crust (of unknown width). The Rustaq Exotic is seen as representing a small volcanic high within Upper Permian oceanic crust. The Triassic exotics are reconstructed as seamounts (atolls), mainly located in an oceanic setting as in most previous interpretations. In this reconstruction, the Permian Ba’id Exotic subsided and was covered by pelagic carbonate during the Triassic. The Oman Mountains as a whole are seen as the product of near-orthogonal rifting. However, the Northern Oman Mountains is an exception, as transform or at least oblique rifting has been widely

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Fig. 6. Relation of the platform margin to the slope successions, based on evidence from the Dibba Zone, Northern Oman Mountains (from Robertson et al. 1990).

inferred for this area (for references see Robertson & Searle 1990). The unique setting of the Jebel Qamar Exotics in the Northern Mountains (Dibba Zone; Fig. 3) with a Lower Palaeozoic prerift basement and both Permian and Triassic shallow-water

carbonates may relate to a setting of transform, or oblique rifting. This rift setting could have triggered the detachment of a fragment of prerift continental basement, which was then ‘parked’ in a relative outboard setting on the margin as the rift basin opened

Fig. 7. Restored relations of the Oman passive margin sediments to the Triassic Exotics located on oceanic crust outboard of the Oman rifted margin (from Robertson & Searle 1990).

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in the Permian. This could be comparable with the setting of the marginal high of continental crust outboard of the deep-marine Ivory Basin of the Equatorial African margin (Mascle et al. 1987). Also, the Hawasina Complex in the Dibba Zone, unlike most other parts of the Oman rifted margin, remained magmatically active during Triassic time possibly reflecting volcanism along a ‘leaky transform’. Finally, if transform related, this part of the margin may have remained buoyant and undergoing unusually slow thermal subsidence, which would have allowed platform up-building during the Triassic to keep pace with subsidence of the Jebel Qamar exotic carbonate platform (Robertson et al. 1990). The Oman rifted margin can now be compared with the Volcanic-rifted v. Non-volcanic-rifted margin types. The Oman rifted margin is dissimilar to the ideal Volcanic-rifted margin as voluminous rift magmatism is absent, despite the possible plume-like signature recorded in the Permian volcanics (Lapierre et al. 2004). The conjugate margin, inferred to be located in Iran, is deformed and not well documented but does not seem to be of Volcanic-rifted type either (Glennie 2000), although more evidence is needed. However, there is little field evidence to support a correlation with a Non-volcanic rifted margin either. The crust beneath the Hawasina Complex is assumed to be oceanic rather than continental, as otherwise a restored width of several hundreds of kilometres of continental crust needs to have been subducted in latest Cretaceous time, which is unlikely (Robertson & Seale 1990). Fragments of Permian oceanic, or transitional, crust are locally preserved (i.e. in Wadi Hawasina and also as the Rustaq Exotic). The nature of the crust beneath this fragment of earliest formed MOR-basalt is unknown, so it cannot be determined whether or not any (simple shear-type) exhumation affected the Oman margin, as for the Iberia– Newfoundland conjugate. However, there is no known evidence, for example, of contemporaneous (Permian) serpentinite-derived clastics from exhumed asthenosphere within the Permo-Triassic deep-sea sediments of the Hawasina Complex. This contrasts with the Alps, Iberia and Newfoundland rifted margins, where serpentinitederived clastics were shed from exhumed low-angle fault scarps (e.g. Tucholke et al. 2004, 2007). Also, there is no evidence of exhumed deep continental crust in Oman as anticipated if the margin underwent intense stretching and exhumation as along the Iberian rifted margin. If lower continental crust had been exhumed it is likely (but not inevitable) that some of this material would have been incorporated into the thrust stack when the ocean was closing during the latest Cretaceous, but, as yet, this has not been discovered. In summary,

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Oman appears to differ somewhat from either of the ideal rift end members.

Himalayan rifted margin Rifted margin units of Permian and Mesozoic age are well exposed in the Himalayas. Proximal (continentward) settings are well exposed, for example in the Salt Ranges of Pakistan and in the Kashmir further north (Gaetani & Garzanti 1991). More distal settings, located near the rifted platform edge are well exposed in the Zanskar Mountains (Figs 8 & 9a). Continental slope and oceanic units are exposed in a variably deformed and metamorphosed state within the Indus suture zone to the north. This suture zone was formed in response to closure of the South Tethyan Ocean that bordered the Indian continent to the south. Similar rift-related and other Tethyan units are exposed along strike for hundreds of kilometres within the Indus–Tsangpo suture zone, where they are, however, generally less well known. It is generally agreed that, during the Permian and Triassic, India remained south of the equator, still attached to Gondwana (Fig. 2), and that it finally detached and drifted rapidly northwards during the Late Cretaceous (from Cenomanian time, c. 95 Ma, onwards). The progressive closure of the South Tethyan Ocean culminated in diachronous collision of India with Eurasia (e.g. Scotese 1991). Thus, unlike Oman, the effects of this collisional deformation have to be unravelled before the rift setting can be reconstructed. Continental rifting during the Permian, to the south of the Indus–Tsangpo suture zone (e.g. in Kashmir and parts of the Zanskar Mountains), was marked by the deposition of transgressive clastic sediments above a regional low-angle unconformity, and also by the extensive outpouring of tholeiitic flood basalts, known as the Panjal Traps, during Early Permian time. These volcanics reach hundreds of metres in thickness and thin northwards towards the axial rift zone. Rifting apparently began as early as Early Carboniferous time (Vannay & Spring 1993). Stratigraphic relations suggest riftrelated exhumation of more than 2 km (Gaetani & Garzanti 1991). The rift volcanics and sediments are overlain transgressively by a mixed terrigenous–carbonate succession of Late Permian age. Late Triassic –Middle Jurassic time was marked by a predominant accumulation of shallow-marine platformal and proximal slope carbonates. Middle Jurassic time was characterized by a regional discordance associated with condensed sedimentation, faulting and the disintegration of the platform edge. The Cretaceous was then associated with deep-water deposition of shales and pelagic carbonates that were accompanied, during the Early

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Fig. 8. Outline geological map showing the location of the rifted continental margin sediments exposed in the Zanskar Mountains and the Indus–Tsangpo suture zone in Ladakh, northern India (from Robertson & Degnan 1993).

Cretaceous (late Aptian), by further alkaline volcanism along parts of the rifted margin (Garzanti 1993). The emplacement of oceanic units (e.g. Spontang ophiolite) was followed by collision with Eurasia during Paleocene – Eocene time (c. 50 Ma). The timing of emplacement of the ophiolites and other Tethyan oceanic rocks, including the rift-related units discussed here, are reported to have taken place either during the Late Cretaceous as in Oman (Searle et al. 1988) or (less probably) not until the Early Cenozoic (Gaetani & Garzanti 1991). This still contentious aspect is outside the scope of this paper. The development of the deep-water rift and passive margin is best documented in Ladakh (northern India) (Searle et al. 1988). The more proximal successions are represented by the regional Zanskar platform, a mixed carbonate–siliciclastic succession that contrasts with the more carbonate-dominate platform succession, as discussed earlier from Oman

(Jebel Akhdar). The Zankar platform succession is highly deformed by collision-related folding and thrusting (Fig. 9b). Recently, the stack of thrust sheets has been restored (Steck 2003) to reveal a transition over approximately 250 km from an area in the south (Chamba Basin) where basement is directly overlain by rift-related Early Permian volcanics and volaniclastic sediments (Panjal Traps). This has been interpreted as an area of flexural uplift (rift shoulder), together with an area slightly further north (Tandi syncline), where basement rocks are overlain by a relatively thin and poorly dated PermoTriassic and possibly younger succession (Steck 2003). Further north, Palaeozoic rocks are overlain by a thin Upper Palaeozoic (Carboniferous– Permian) succession, overlain in turn by a thick shelf succession (Surchu–Baralacha; Kharnag–Tso Morari), which then passed northwards into a slope succession (Nyimaling–Markha), as exposed within the Indus suture zone.

Fig. 9. Field photographs of the Himalayan and Eastern Mediterranean rifted margins. (a) Mesozoic Zanskar platform, backthrust over the Mesozoic deep-water passive margin sequence (Lamayuru-Karamba unit) during the mid-Cenozoic; Sapi-La area; western Ladakh; (b) from left to right: far distance-Zanskar shelf succession; middle distance, right-Mesozoic deep-water passive-margin slope succession (Lamayuru Complex); the pinnacles are former carbonate-filled channels, now vertical; also the high hill, left of centre (dark), represents an accreted oceanic arc unit (Dras arc complex); Namika-La area, east of Mulbeck; (c) Triassic of Phyllite– Quartzite unit, southern Greece; foreground, Triassic shallow-water limestones and sandstones; beyond the bay, Triassic deltaic sandstones; in the distance, downfaulted (overlying) regional-scale thrust sheet of Mesozoic carbonate platform rocks (Tripolitza unit); the assemblage is restored as a Triassic rift and post-rift capping carbonate platform; Vai area, eastern Crete; (d) northern Syria. Jurassic– Cretaceous carbonate platform succession (Jebel Agraa) with overthrust deep-water passive margin slope and ophiolitic material above (wooded on right); taken from near Raas-Al Bassit looking north. See the text for discussion.

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The distal deep-water successions of the Indus suture zone are represented by deformed greenschist-metamorphosed facies known as the Lamayuru Complex (Bassoullet et al. 1981; Robertson & Degnan 1993; Sinha & Upadhyah 1994) (Fig. 9b, c). The base of the stratigraphic succession of the Lamayuru Complex is tectonic so that the nature of the substratum remains unknown. The oldest preserved units, only exposed locally (e.g. in the Markha Valley, Eastern Ladakh), are Late Permian redeposited carbonates (turbidites) that accumulated in a base-of-slope setting (Stutz 1988). Triassic time was marked by the accumulation of deep-water terrigenous sediments, mainly as deep-water turbidites that contain quartz, mica and lithic grains, ultimately derived from the Precambrian Indian craton to the south, and redeposited carbonates were supplied from the Zanskar shelf. The Early Jurassic sequence was dominated by relatively thin-bedded calciturbidites, rich in ooids derived from the Zanskar carbonate platform. During the early Middle Jurassic (Aalenian) the carbonate platform edge disintegrated in response to a regional pulse of extension, giving rise to exotic blocks of platform carbonate in proximal slope areas (‘Mulbeck exotics’). Further east, numerous submarine channels carried large volumes of shallow-water carbonate detritus from the eroding platform into the basin. The Cretaceous record of the Lamayuru Complex is limited by strong deformation. However, it includes sandstones and pelagic limestones (Bassoullet 1978). The stratigraphically lower part of the Lamayuryu Complex is intruded by swarms of diabase sills of chemically alkaline composition (Searle 1983; Robertson 2000). The Lamayuru Complex in the west, in the vicinity of the Suru River (Fig. 8) is in thrust contact with a generally more distal unit of deepwater sedimentary and volcanic rocks known as the Karamba Formation or Karamba Complex (Sutre 1991; Robertson & Sharp 1998). The preserved record of this key unit begins with the deposition of siliciclastic turbidites, subordinate redeposited limestones and basic volcanics of Middle–Late Triassic age. Early–Middle Jurassic time was characterized by the deposition of intercalated radiolarian sediments, siliciclastic turbidites and redeposited shelf-derived calciturbidites. Radiolarites, dated as Middle Jurassic in age (Danelian & Robertson 1997), are exposed as sedimentary intercalations with alkaline basaltic lavas and volcaniclastic sediments, together making up a succession several hundreds of metres thick. The Karamba Formation is interpreted as representing the stratigraphically higher and relatively distal part of the North Indian passive margin (Fig. 10). In this interpretation the Lamayuru Complex

formed the more proximal part of the margin and passed distally (oceanwards) into the Karamaba Formation, at least in the west (Suru River – Shegol area), where exposure is sufficient to determine primary relationships. Counterparts of the Karamba Formation are likely to have existed along strike to the east and west, although these are fragmentary, or absent, possibly because they were removed at some stage of the multiphase emplacement history that affected this margin. Basin-plain sediments similar to the Hawasina Complex in Oman (Hamrat Duru Group) are not preserved in Ladakh, again probably owing to tectonic removal during emplacement. The Karamba Formation alkaline volcanics are definitely interbedded with Indian margin-derived sediments including quartzose terrigenous turbidites and shales (Danelian & Robertson 1997; Robertson & Sharp 1998). This establishes that the Karamba Formation formed an essential part of the North Indian rifted margin and rules out an earlier suggestion that this unit represents an oceanic seamount that was accreted and obducted onto the North Indian margin during closure of Tethys (Sinha & Misra 1994). If it were an oceanic seamount, coarse terrigenous turbidites would be absent, as for the Triassic exotics in Oman, whereas the volcanics are clearly interbedded with terrigenous sediments. The Karamba Formation is also important as it is indicative of alkaline magmatism in a deep-water rifted margin setting during the Middle Jurassic; i.e. long after the inferred continental breakup of this part of the North Indian margin during Late Permian time. Evidence of more distal, oceanic units is preserved as me´lange that is mainly exposed above the Zanskar shelf succession, beneath the Spontang ophiolite (Reuber et al. 1992; Corfield et al. 1997), and is also known locally in the Indus suture zone further north (Bassoullet 1978; Robertson 1998a). Recent mapping has confirmed the existence of a discrete thrust sheet, termed the Photang thrust sheet, beneath the Spontang ophiolite (Corfield et al. 1997). In one succession (south of Photaskar), Upper Permian chemically evolved alkaline basalts, several hundreds of metres thick, are interbedded and overlain by red silty carbonates. Pillow lavas of possibly Late Triassic age occur above this, and then Cretaceous pelagic limestones. A second succession begins with Upper Permian limestone capped by a metalliferous crust, and is then overlain by radiolarian mudstones, tuffs and varicoloured limestones. These sediments contain ammonites and conodonts of probable Late Triassic age. Additional (undated) lavas occur above this. Elsewhere in the Photaskar thrust sheet, limestones and marls of probable Cretaceous age overlie Triassic massive lavas and pillow lavas. Another smaller

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Fig. 10. Restored relations of the passive margin sediments and igneous rocks on the Zanskar shelf and the slope successions of the Lamayuru Complex, Himalayas (from Robertson & Degnan 1993).

‘exotic’ includes pillow lavas, overlain by Late Triassic pink pelagic limestone (Hallstatt facies). Additional smaller exotics comprise Permian bioclastic limestone, locally overlain by Upper Cretaceous pelagic limestone (Reuber et al. 1992). Elsewhere, the me´lange is associated with debris flows made up of limestones of different ages, shed from volcanic edifices that were capped by carbonate rocks of Permian and Triassic age. In addition, a small but instructive ‘exotic’ occurs within the Indus suture zone near the Lamayuru monastery (Bassoullet et al. 1981; Robertson 1998a). The intact succession within this ‘exotic’ records Upper Permian shallow-marine carbonate deposition followed by tilting, fissuring and subsidence. The rotated fault block was then covered by ammonite-bearing pink pelagic limestone and volcaniclastic sediment during Early Jurassic time. Fe/Mn crusts accumulated, followed by the eruption of alkaline basalts, probably during Late Triassic time. Much of the history of rifting is therefore encoded within this single exotic block. Numerous exotics are also exposed within the Indus suture zone in Eastern Ladakh (e.g. Markha Valley), although these have undergone greenschist facies metamorphism and are more recrystallized (Fuchs 1986). Kilometre-sized exotics there

comprise thick piles of within-plate-type basaltic lavas, overlain by shallow-water limestones of Late Triassic age, containing the large bivalves Megalodonta (author’s unpublished data). The North Indian margin is interpreted to have undergone a series of extensional rift events from Early Carboniferous time (i.e. pulsed rifting), culminating in continental breakup probably during Late Permian time, with thermal subsidence from Late Permian time onwards (Gaetani & Garzanti 1991). The Early Permian outpouring of tholeiitic continental flood basalts (Panjal Traps) was associated with thermal uplift, perhaps related to magmatic underplating (Stampfli et al. 2001). The exotics include Permian shallow-water carbonates (Lamayuru and Photaskar exotics) that are interpreted as remnants of several relatively small carbonate platforms that were positioned within Upper Permian oceanic crust. However, it is unclear whether the basement was composed of either oceanic or continental crust. In addition, Triassic seamounts formed within the South Tethys (e.g. Markha Valley Exotics), much as in Oman. After the inferred Permian continental breakup and onset of thermal subsidence, the North Indian continental margin experienced several subsequent

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pulses of renewed extension and related alkaline volcanism. A strong pulse (or pulses) of extension during Middle Jurassic time resulted in a regional discordance, collapse of the platform edge and the alkaline magmatism within the deep-sea slope succession (Karamba Formation). A pulse of extension is also recorded on the Zanskar shelf in the form of Early Cretaceous alkaline volcanism. These rift pulses can be related to the attempted break-way

of India from Gondwana, which was finally achieved during Late Cretaceous time, followed by passive margin subsidence. In summary, the inferred setting of the North Indian rifted margin is shown in Figure 11a, b. The repeated pulses of extension and alkaline volcanism for more than 100 Ma differs strongly from the ideal rifted margin, which simply subsides after spreading begins.

Fig. 11. Summary of the inferred tectonic setting of the north-Indian passive margin. (a) During Permian–Early Jurassic and (b) during Late Jurassic– Early Cretaceous. See the text for discussion.

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The North Indian margin can also be compared with the ideal division into Volcanic-rifted and Non-volcanic-rifted margin types. On this basis, the North Indian margin is quite similar to the Volcanic-rifted type. The extensive Lower Permian magmatism (Panjal Traps) of the proximal rifted margin can be related to the effects of a mantle plume, resulting in regional-scale alkaline magmatism (Panjal Traps), thermal uplift and deep erosion. However, the volcanism is focused well south of the axial rift, unlike, for example, the east Greenland margin, which suggests that the main role of plume magmatism may have been to thermally weaken the crust rather that control the locus of final breakup. Unfortunately, the conjugate margin, assumed to lie somewhere further north in Tibet, is not available for study and thus whether or not the rifted margin exhibited a asymmetrical geometry is unknown. On the other hand, there is no field evidence of the features that characterize Non-volcanic rifted margins; for example, exhumed deep crust or mantle, or lowangle detachment faulting, although as in Oman

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this could reflect the uncertainties of tectonic emplacement.

Eastern Mediterranean rifted margins The Eastern Mediterranean region exhibits considerably greater diversity of rift-related settings than either Oman or the Himalayas. There is thus considerable potential if the rift settings can be disentangled from the complicated emplacement history. One of the main features of the Eastern Mediterranean region is that several sutures are present, associated with rift-related and ophiolitic rocks. This opens the possibility that several different types of rifted margin could be present so that individual cases need top be considered separately. Role of microcontinents and multiple small ocean basins. In the Eastern Mediterranean region rift-related units are exposed in several generally subparallel belts (or zones; Fig. 12). In many current interpretations each of these belts represents the suture of a Mesozoic small ocean basin (e.g.

Fig. 12. Outline tectonic map of the Eastern Mediterranean region showing the locations of the main ophiolitic sutures in relation to the rift/passive margin units discussed here.

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Robertson & Dixon 1984; S¸engo¨r 1984; Smith 1993, 2006; Channel & Kozur 1997; Golonka & Ford 2000; Stampfli et al. 2001; Stampfli & Borel 2002; Robertson et al. 2004). Many workers, therefore, reconstruct the Eastern Mediterranean region as a number of microcontinents, with individual rifted margins of different scale and local setting, separated by small oceanic basins (Figs 2 & 12). However, there is no agreement as to how such microcontinents should be reconstructed in space and time, or on the rift processes and settings involved. In addition, an earlier and by no means extinct view is that only one main Mesozoic ocean basin existed in the region (until the Late Cretaceous), and that the subparallel zones of rift-related, ophiolitic and other rocks resulted from long-distance overthrusting during the Late Cretaceous –Early Cenozoic closure of Tethys (Aubouin et al. 1970; Bernoulli & Laubscher 1972; Ricou et al. 1974; Dercourt et al. 1986; Stampfli et al. 1991; Ricou 1996). Terrane dispersal should also be considered as a means of duplicating continental fragments within a single oceanic basin and this is considered further below. In one view Permo-Triassic spreading took place in a single Tethyan ocean between Eurasia and Gondwana, while rift basins developed along the southern margin (e.g. Ionian rift basin; Pindos rift basin). The spreading axis then effectively jumped southwards initiating spreading in the Eastern Mediterranean, but not until the Early Cretaceous (Dercourt et al. 1986). In most of the more recent models, however, spreading is assumed to have started much earlier along the northern margin of Gondwana (i.e. Permian or Triassic; see below). In general, throughout many parts of the Eastern Mediterranean region there is evidence of rifting of Gondwana-related units during Late Palaeozoic time, followed by sea-floor spreading during the Mesozoic. This was followed by progressive closure of Tethys in this region as Africa and Eurasia converged during Late Mesozoic –Early Cenozoic time (Livermore & Smith 1984). By contrast, the opposing, northerly continental margin was characterized by the accretion of continental fragments at least from Late Palaeozoic time onwards, most of which ultimately originated from Gondwana and drifted northwards until they collided with Eurasia (Dercourt et al. 1986, 2000; Stampfli & Borel 2002). In the writer’s opinion, several large continental fragments and additional smaller ones must have existed in the Eastern Mediterranean region. In the south, Gondwana was bordered by a large microcontinent known as Adria or Apulia. The timing of separation of this fragment from Gondwana is still debated (Late Permian, Triassic or

Jurassic–Cretaceous), but by the late Mesozoic an ocean basin was clearly present in the east (Ionian Sea); this basin continues to be subducted northwards until today, proving that oceanic crust exists in this region (e.g. Dercourt et al. 2000). Further north, a large elongate continental fragment throughout former Yugoslavia, Albania and Greece is known as the Pelagonian microcontinent (i.e. the Korabi unit in Albania and the Drina –Ivanjica unit in former Yugoslavia). Further east in the Turkish region the Tauride microcontinent (one or several units) extended from the Aegean Sea to Iran. The Pelagonian and Tauride microcontinental units cannot be traced across the Aegean Sea, and so it is not known whether they were linked in any way. The Tauride microcontinent was itself probably split into several smaller rifted fragments (e.g. eastern and western Tauride). Also, outlying continental units (e.g. Nig˘de-Kirs¸ehir massif, Turkey; Fig. 12) can be considered either as smaller microcontinents or as promontories of larger continental units. Additional microcontinents existed further east, in the region of Iran (e.g. Lut Block), although these will not be discussed here. In general, three main Mesozoic basins are recognized, each with their own rifted margin that will be discussed below. The first is the ‘Ionian Sea’, which separated Gondwana from the Adria (Apulia) microcontinent. The second is the Pindos Ocean in the Balkan area, which separated Adria from the Pelagonian microcontinent to the north. This can be broadly correlated with the ‘Southern Neotethys’ further east, which separated Gondwana from the Tauride microcontinent. The third is the Vardar Ocean in the Balkan region that can be broadly correlated with the ‘Northern Neotethys’ further east and which separated the Tauride microcontinent from Eurasia. Several other oceanic elements constituted marginal basins within the Eurasian margin and these will not be discussed here (e.g. Guevgueli, northern Greece; Ku¨re, northern Turkey: see Robertson et al. 1996). Below, each well-documented rifted margin segment associated with each of these three oceanic basins will be discussed individually and then compared to provide a regional picture. It is worth noting that useful insights into rift settings and processes are possible even if agreement as to the regional palaeogeography remains elusive. Rifted margins of North Africa and the Levant. The most southerly located passive margin in the Eastern Mediterranean region is the rifted margin of North Africa–Levant (Fig. 12). This extends from the Western Mediterranean region (adjacent to Morocco, Algeria and Tunisia), eastwards through the Ionian Sea into the Eastern Mediterranean, to the south of the Sea of Crete (off

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Fig. 13. Morphology and structure of the rifted margin of the Southern Neotethys in the Levant Basin. From Garfunkel (1998).

the coast of Libya) and then into the easternmost Mediterranean (south of Egypt). The margin then takes a nearly right-angle bend to follow the Levant margin, off Israel, Syria and southern Turkey, before bending eastwards around the Arabian margin (northern Syria and Turkey) and then extending southeastwards (through Iran) to link with the Oman margin that was discussed earlier. Limited on-land exposure data, combined with scarce available industry well and seismic data (e.g. Makris et al. 1983; Ben-Avraham et al. 2002), chart a history of several rifting events affecting the North African margin during Late Palaeozoic –Mesozoic time (Guiraud et al. 2001). Important rift events appear to have occurred during the Late Carboniferous and the Triassic, but it is not possible to determine from the evidence of rifting from the North African margin alone (on land) when exactly spreading began in the adjacent South Tethyan oceanic basin: estimates range from Late Permian, to Middle– Late Triassic, to Late Jurassic –Early Cretaceous (see Briand 2000). However, it can be inferred that a limited amount of sea-floor spreading took place in the latest Triassic –earliest Jurassic from evidence from several areas (i.e. Cyprus, Levant; SW Turkey; see below). It has also been argued that extensive Late Jurassic –Early Cretaceous volcanism throughout the region is rift related (Laws & Wilson 1997). In the east, the Levant margin is orientated north–south with respect to the main North African margin, and has often been referred to as a transform-rifted margin (e.g. Robertson & Dixon 1984; Stampfli et al. 2001). However, Garfunkel (1998, 2004) argued that the available evidence is more consistent with near-orthogonal rifting along the Levant margin segment, implying the North

African margin further west (off Egypt) is instead of transform-rifted type. Based on seismic and well evidence, the deep structure of the Levant margin (Fig. 13) does appear to contrast markedly with, for example, the Ivory coast transform-riftedtype margin (Mascle et al. 1987). However, it should be noted that transform-rifted margins have so far been little investigated and may show a range of architectures that remain largely unexplored. There is also the evidence of a pulse of alkaline volcanism of Late Jurassic –Early Cretaceous age as documented in wells and from surface exposures (Garfunkel 1998, 2004); some consider this to be rift related (e.g. Laws & Wilson 1997), others plume related (Garfunkel 1998). In the west, Early Permian–Early Triassic deepwater pelagic facies in western Sicily and eastern Crete include a cosmopolitan warm-water radiolarian fauna, which points to the existence of a deep-water connection between the Eastern Mediterranean and, for example, Oman by this time (Kozur & Krahl 1984; Catalano et al. 1991; Kozur 1993). Triassic rift-related rocks in Crete and in the Peloponnnese are exemplified by the Phyllite –Quartzite unit, as discussed further below (Fig. 9c). We may, therefore, assume that a significant amount of rifting and subsidence took place along the North African margin prior to the Early Permian, probably beginning in the Late Carboniferous based on the presence of deep-water terrigenous turbidites of this age (e.g. in western Sicily). However, there is no evidence of preserved oceanic crust or of related passive margin sequences in these areas (i.e. in Sicily and Crete) that could indicate the formation of oceanic crust during Late Permian time, in contrast to suggestions made by Stampfli et al. (2001, 2003). There are

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also different interpretations of whether the Eastern Mediterranean oceanic basin was connected to the Western Tethys through the Ionian Sea (through Sicily) to the Alpine Tethys (Pidemont –Valais Ocean). Many workers inferred such an oceanic connection, albeit narrow (Dercourt et al. 1986; Robertson & Grasso 1995; Ricou 1996), whereas others interpreted Adria as a promontory of Gondwana with no marine connection (Robertson & Dixon 1984; Stampfli et al. 2001). Much depends on the age and structure of the crust beneath the Ionian Sea. Recent geophysical studies (Catalano et al. 2001) suggest that this crust is likely to be of Late Jurassic –Early Cretaceous (rather than, for example, Triassic) age, and this favours a correlation with the Alpine Tethys and the Central North Atlantic. According to the model of Stampfli et al. (2001), spreading began in the Eastern Mediterranean during Late Permian time leading to the opening of a ‘Southern Neotethys’. Any younger volcanism, notably Early Jurassic alkaline volcanics (Asher Formation) as cored in a deep well (Atlit-1, near Haifa), represent a reactivation of the passive margin. On the other hand, based on a detailed study of the available seismic, well and surface exposure data, Garfunkel (1998) argued that major rifting dates from the Triassic and that the Early Jurassic volcanics (Asher Formation) associated with thick clastic sediments reflect continental breakup during Early Jurassic time to form a small ocean basin directly to the NW. Subsidence curves for the Palmyra rift to the NE appear to support rifting during Permian time, but there is little evidence of continental breakup of the main Levant margin prior to Early Jurassic time. This conclusion is supported by a recent reinterpretation of seismic and well data from the Levant Sea (south of Cyprus) that suggests that this basin only experienced rifting, and that spreading took place further north in the vicinity of Cyprus during the Early Jurassic (Gardosh & Druckman 2006). In summary, the North African margin documents pulsed rifting during Palaeozoic –Early Mesozoic time, followed in the east (Levant) by continental breakup in the latest Triassic –earliest Jurassic to form a small ocean basin in the Eastern Mediterranean (Herodotus Basin area). However, breakup was probably delayed until the Late Jurassic in the west (Ionian Sea). We now go on to consider the evidence of tectonically emplaced units, especially in northern Syria and western Cyprus to evaluate further the timing and setting of rifting of the South Tethyan Ocean bordering Gondwana. Rifted Arabian margin: northern Syria. From northern Syria through Iran to Oman, the northern

margin of Arabia experienced the emplacement of continental margin and oceanic (ophiolitic) units during latest Cretaceous (Maastrichtian) time (Ricou et al. 1974). The driving mechanism of the emplacement is considered to have been trenchmargin collision, as noted previously for Oman (Yılmaz 1993). The emplaced rifted margin units are well documented in the Baer –Bassit area of northern Syria, where they take the form of dismembered thrust sheets and me´lange beneath overriding ophiolites (Delaloye & Wagner 1984; Al-Riyami et al. 2002b) (Fig. 14). The proximal slope of the passive margin (i.e. Arabian carbonate margin) was concealed by thrusting, in contrast to Oman, and is buried somewhere to the north. More distal margin units are exposed, however, above the autochthonous carbonate platform (Fig. 9d). The succession in dismembered thrust sheets and blocks begins with Upper Triassic alkaline volcanics, interbedded with deep-water radiolarian cherts and pelagic limestones, passing upwards into a mainly siliceous sedimentary sequence; there is then an incoming of calcareous facies in the Early Cretaceous including quartzose sandstone turbidites and pelagic carbonates (Delaune-Maye`re 1984; Al-Riyami & Robertson 2002). The distal margin was characterized by one, or several, large volcanic build-ups of alkaline basalt (Tamima unit), interbedded with radiolarite and capped by condensed pelagic limestone of Late Cretaceous (Cenomanian) age. This exotic unit is interpreted as one, or several, seamounts located not far outboard of the passive margin (Al-Riyami & Robertson 2002) (Fig. 15).

Fig. 14. Simplified tectono-stratigraphic relations of rifted margin and ophiolitic units emplaces from the South Tethyan Ocean onto the Arabian margin in northern Syria (Baer–Bassit region).

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Fig. 15. Restored section across the southern margin of the Southern Neotethys in northern Syria (Baer–Bassit region). Note the age and relative locations of the passive margin and exotic units. See the text for explanation.

The Baer –Bassit area is restored as part of the rifted margin of the Arabian continental margin that stretched all the way to Oman. The Triassic sediments and volcanics there record continental breakup in Late Triassic time (or possibly slightly later). There is no evidence of pre-existing Permian deep-water facies. Seamounts formed adjacent to the margin in Late Jurassic –Early Cretaceous time, as indicated by the age of interbedded radiolarites (Al-Riyami et al. 2002a). Some MORtype basalt is present in the metamorphic sole of the Upper Cretaceous Baer –Bassit ophiolite (Al-Riyami et al. 2002b), but this remains undated. Rifted South Tauride margin: SW Turkey and western Cyprus. During Late Mesozoic –Early Cenozoic time, the oceanic crust of a South Tethyan ocean (‘South Neotethys’) was mainly subducted northwards (e.g. Dercourt et al. 1986, 2000; Yılmaz 1993; Robertson 1998b; Robertson et al. 2004). In response, the northerly conjugate margin of the South Tethys was mainly destroyed, possibly as a result of ‘subduction erosion’. Exceptionally, in one well-studied area, the Isparta Angle area of SW Turkey (Fig. 12), the rifted margin stepped northwards on a regional scale creating an oceanic embayment. Today, the Isparta Angle has an inverted V shape. However, this is a relatively recent feature (mid-Cenozoic) resulting from crustal rotations during continental collision (Kissel & Poisson 1986). The Isparta Angle restores effectively as a gulf of the South Tethys that divided the Tauride microcontinent into western and eastern

segments. By contrast, Dilek & Rowland (1993) restored the two limbs of the Isparta Angle as the conjugate margins of a small Cretaceous oceanic basin within the Isparta Angle. However, these two limbs are more likely to represent different parts of the south-facing margin of the south Tethys Ocean (‘Neotethys’). The setting of a regional-scale northward oceanic embayment eventually favoured the preservation of rifted margin units originally several hundreds of kilometres long, known as the Antalya Complex (‘Antalya Nappes’). Deformation and emplacement of the Antalya Complex in latest Cretaceous–Early Cenozoic time was dominated by strike-slip/transpression (Woodcock & Robertson 1982). This setting of oblique convergence in turn favoured the dissection of the former rifted margin into small elongate ‘microterranes’, in which long vertical successions are preserved separated by steep faults. This contrasts with the mainly low-angle thrust sheets as seen elsewhere (e.g. in Oman and the Himalayas). The emplaced rifted margin units are exposed around the periphery of the Isparta Angle (Waldron 1984). However, the exposures are most extensive and complete in the classic SW area, SW of Antalya city (Fig. 16) (see Robertson et al. 2003 for a recent review of the research in this key area). Proximal successions of the rifted margin, generally lacking volcanic rocks, are preserved in thrust sheets directly structurally above the in situ Tauride carbonate platform (Bey Dag˘ları autochthon; Figs 17, column 1 & 18a, b). These units include detached blocks of Late

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Fig. 16. Outline geological map of the Isparta Angle showing the Antalya Complex and other tectonic units, notably the Mesozoic rift-related basinal volcanics and sediments (from Robertson 1993).

Triassic highly fossiliferous reef limestone in more proximal slope settings. Lower-slope deposits include thick-bedded terrigenous turbidites and redeposited limestone (e.g. channelized conglomerates) (Fig. 17, columns 2 and 3). Distal-type successions (Fig. 17, column 4) are dominated by Late Triassic sandstone turbidites, pelagic limestones, calciturbidites and radiolarian sediments, and include pelagic carbonates. In most areas no basement to these thrust sheets is preserved. Further east and structurally higher (Go¨dene unit; Figs 17, column 5 & 18b), thick successions of Upper Triassic (Carnian) pillow lavas, massive lavas and lava breccias are interbedded with minor amounts of deep-water hemipelagic limestone (Halobia Limestone) and radiolarite; detached blocks of shallow-marine coralgal limestone are also locally exposed. The lavas and sediments show evidence of syndepositional normal faulting (Fig. 18c). The extrusive sequence, in places, is depositionally overlain by contrasting sediments, exposed on topographical ‘highs’ and in ‘lows’. The ‘highs’ are depositionally overlain by pink

pelagic carbonate, rich in manganese oxide and ammonites of latest Triassic age (Ammonitico Rosso). The ‘lows’ locally begin with terrigenous turbidites and quartzose conglomerates, in places containing well-rounded clasts. This establishes that terrigenous rocks were exposed locally within the axial rift zone. Above this come Late Triassic hemipelagic Halobia limestones, with carbonate derived from adjacent carbonate platforms, and above this deep-water mainly siliceous and radiolarian sediments of Jurassic–Early Cretaceous age (Robertson & Woodcock 1981) (Fig. 17, column 5). This unit (Go¨dene Zone) is interpreted as volcanic seamounts within the continent–ocean transition zone. Eastwards again, the volcanic –sedimentary successions of the Go¨dene zone are overthrust by, or in high-angle contact with, thick units of Mesozoic platform carbonates (Kemer units or ‘Upper Antalya Nappe’); these sediments directly overlie a basement of Lower Palaeozoic–Middle Triassic prerift–synrift sedimentary rocks (Figs 17, column 6 & 18d). In places, marginal facies that

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Fig. 17. The main sedimentary and volcanic successions exposed on the southwestern part of the Isparta Angle (SW segment). Note the parent carbonate platform (column 1), the rifted margin slope units (columns 2– 4), the Mesozoic volcanic-related units (Go¨dene Zone) (column 5) and the Mesozoic carbonate platform floored by prerift continental basement (column 6; Kemer Zone). The Upper Cretaceous Tekirova ophiolite is exposed further east along the coast (see Fig. 16 for an outline tectonic map).

Fig. 18. Field photographs of rifted margin units in the Antalya Complex (Isparta Angle region, SW Turkey). (a) Mesozoic carbonate platform (Bey Dag˘ları) dipping eastwards beneath emplaced Triassic– Cretaceous deep-water sedimentary units (wooded; Kumluca zone); (b) view from the autochthonous Mesozoic carbonate platform (Bey Dag˘ları) across folded and imbricated Triassic– Cretaceous deep-water sedimentary units (Kumluca Zone), then across Late Triassic deep-water transitional basalts (Go¨dene Zone), with Mesozoic–Triassic– Cretaceous shallow-water carbonates above (C¸albalı Dag˘; Kemer Zone). The contact between the lavas and the limestones here is tectonic; the limestones are restored as exotics constructed on ‘highs’ within the continent– ocean transition zone. Ocean crust (Tekirova Zone) lies over the skyline to the east, beyond which is Antalya Bay; (c) small normal faults within Upper Triassic (Carnian) subaqueous massive lava flows; erupted within the continent– ocean transition zone. The interbedded carbonate (peri-platform) ooze was shed from adjacent exotic carbonate platforms; C ¸ albalı Dag˘; Kemer zone; (d) unbroken Palaeozoic– Upper Cretaceous sedimentary succession (prerift; synrift and post-rift); a rifted fragment of continental crust is capped by shallow-water carbonates (Tahtahlı Dag˘; view west from Antalya Bay). See the text for discussion.

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were derived from the Triassic shallow-water carbonate platforms overlie, or interfinger with, volcanic-sedimentary successions (Robertson & Woodcock 1982). Further east again, the Kemer units are in moderate- to high-angle fault contact with Upper Cretaceous ophiolitic rocks of the Tekirova unit, as exposed along the Mediterranean coast (Juteau 1975) (Fig. 16). The emplacement of these ophiolitic rocks onto the margin has been explained by dominantly strike-slip processes during Late Cretaceous –Early Cenozoic time (Woodcock & Robertson 1982). Taken together, the successions restore as a transition from a proximal to a distal rifted margin from west to east in the SW Antalya area (in present-day co-ordinates; Fig. 19). Initial rift volcanism of Early–Middle Triassic age was mainly tuffaceous and was accompanied by evidence of subsidence and extensional faulting. This was followed by continental breakup and was accompanied by extensive outpouring of volcanic rocks (up to hundreds of metres thick) during Late Triassic (Carnian) time. These volcanic rocks are of alkaline– transitional MOR basalt type and were erupted at the same time as the accumulation of peri-platform ooze (Halobia Limestone) and ribbon radiolarites in deep water (Robertson & Waldron 1990). The presence of locally overlying lenses of turbiditic sandstones and redeposited terrigenous conglomerates (e.g. mass-flow-type conglomerates) is important in establishing that rift-related igneous crust formed along the continental margin rather than out in the ocean (Fig. 19). An origin as emplaced open-ocean igneous seamounts, for example, similar to the Triassic exotics in Oman (which lack interbedded terrigenous sediment) can be ruled out. The Kemer units are interpreted as one, or probably several, rifted continental fragments that were overlain by Late Triassic shallow-water carbonate build-ups. Their setting was similar to the Jebel Qamar Exotics in northern Oman, as discussed earlier. However, the Antalya off-margin platforms continued to be the sites of shallow-water deposition and were not covered by pelagic carbonates until Late Cretaceous time. Several factors may contribute to this contrasting subsidence history. First, the original rifting was probably oblique, which could have favoured transform-type rifting of continental fragments – as, for example, along the Dibba Zone of northern Oman and the modern Gulf of Guinea transform margin (Mascle et al. 1987). Secondly, it is possible that only a small amount of oceanic crust formed during Late Triassic –Early Jurassic time within the adjacent South Tethys and thus tectonic subsidence was limited. Thirdly, the marginal platforms (i.e.

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Kemer units) were located within a semi-isolated re-entrant of the South Tethyan oceanic basin, comparable to the setting, for example of the modern Caribbean (Waldron 1984; Robertson 1993), which may have favoured the persistence of shallow-water carbonate accumulation on marginal highs. The net effect were that tectonic and thermal subsidence was sufficiently limited to allow carbonate platform up-building to be maintained from Late Triassic to mid-Cretaceous time. This was followed by regional submergence and a transition to pelagic carbonate deposition during the Late Cretaceous. The off-margin Kemer units and adjacent ophiolitic units were strongly deformed during latest Cretaceous – Early Cenozoic time by a combination of thrusting–strike-slip–transpressional processes. However, platform deposition (mainly hemipelagic) continued on the parent Bey Dag˘ları carbonate platform into Early Cenozoic time. After regional deformation and emplacement ended (Middle Eocene) the oceanic basin in the Isparta Angle remained partially open to a surviving remnant of the South Tethyan oceanic basin that is, today, the Eastern Mediterranean Sea. In summary, the importance of the Antalya Complex (SW area in particular) for the understanding of rift settings is that the continent–ocean transition is unusually intact and well exposed there, especially volcanics and sediments formed during the transition to spreading oceanic crust of Late Triassic age. Additional important evidence of the northern, rifted passive margin of a South Tethyan Ocean (‘South Neotethys’) comes from the Mesozoic rocks of west Cyprus known as the Mamonia Complex. The Mesozoic allochthonous rocks occur in the west of the island and also in the south (i.e. as the Moni Me´lange: Robertson 1977). It might, therefore, be expected that these rocks relate to the southern margin of the South Tethys. However, this seems unlikely for several reasons. First, the leading edge of the southern margin, represented by the Eratosthenes Seamount, did not begin to collide with southern Cyprus until the PlioQuaternary; during the late Mesozoic this unit was still located well to south, as shown by the results of deep-sea drilling (Leg 160: Robertson 1998b). Secondly, the units exposed in the Mamonia Complex (Mamonia nappes of Lapierre 1972) can be lithologically correlated with the SW area of the Antalya Complex, discussed above (Robertson & Woodcock 1982). Thirdly, the Troodos ophiolite, Cyprus and probably the adjacent Mamonia Complex underwent 90º counter-clockwise rotation during latest Cretaceous –Early Cenozoic time (Moores & Vine 1971; Clube & Robertson 1986; Morris 1996). Thus, the Mamonia rocks restore to an originally northerly position. Taken together,

Fig. 19. Restored relations of the passive margin sediments and igneous rocks on the shelf and the slope of the SW Isparta Angle (i.e. SW segment of the Antalya Complex). Carbonate platform units there are underlain by prerift crust (‘Upper Antalya Nappe’) restored in more outboard settings.

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Fig. 20. Reconstruction of the Upper Triassic tectonic setting of the Mamonia Complex, western Cyprus. Note that the parent carbonate platform is not exposed in Cyprus (from Robertson 1990).

these factors favour a north margin-related origin of the Mamonia Complex, together with the Antalya Complex in the Taurides, southern Turkey. In west Cyprus only the more distal (oceanward) rifted margin successions of an original continent– ocean transition are exposed. Dismembered sedimentary successions, termed the Ayios Photios Group, restore as Upper Triassic deep-water terrigenous turbidites, passing into Jurassic mixed calcareous –siliciclastic and siliceous sequences that accumulated in a deep-water base-of-slope setting (Fig. 20). Early Cretaceous time saw deposition of the thick-bedded quartzose sandstones (Akamas Sandstone) that were probably redeposited from a coastal, aeolian setting by gravity-flow processes. Consequently, the Mamonia sedimentary successions (Ayios Photios Group) restores as the slope/lower slope of a Triassic rifted margin (Robertson & Woodcock 1979) (Fig. 20). An originally more distal (oceanward) unit, known as the Dhiarizos Group, comprises strongly deformed Late Triassic basalts of MOR type (Malpas et al. 1992), interbedded with radiolarites and locally overlain by ammonite-bearing pelagic limestones. Local successions continue upwards into relatively thin-bedded terrigenous and calcareous turbidites, shales and pelagic limestones (Mavrokolympos unit; Swarbrick & Robertson 1980). Consequently, as for the Antaya Complex (Go¨dene Zone), these igneous rocks were formed adjacent to the rifted margin and not within an open ocean far from a continent. In addition, there are also scattered blocks of alkaline volcanics associated with Upper Triassic reefal carbonate build-ups (Lapierre 1972). These lavas are interpreted as seamounts in an oceanic setting (albeit

marginal) comparable to that of the Triassic Oman Exotics (Figs 21 & 22). In summary, the chief importance of the Mamonia igneous rocks is that they provide evidence of Triassic MOR-type basalt with the South Tethys (‘South Neotethys’). This is inconsistent with some tectonic models that suggest that no spreading occurred in this area until the Cretaceous (Dercourt et al. 1986; Dilek & Rowland 1993). Taking the evidence from both the southerly and the northerly rifted margin units, discussed above, the South Tethys in the Eastern Mediterranean region can be restored as an east–west-trending Mesozoic small oceanic basin. Pulsed rifting on a regional scale (Late Carboniferous–Late Permian–Early/ Middle Triassic) was followed by continental breakup during latest Triassic–earliest Jurassic time, associated with extensive alkaline–transitional MOR-type volcanism in different areas. These South Tethyan settings differ from both the ideal Volcanic- and Non-volcanic rifted margin types. There is little evidence of regional-scale voluminous plume-type volcanism with the implied thermal uplift and deep erosion. However, the chemistry of the rift volcanics is similar to some plume-type volcanics (Dixon & Robertson 1999), while perhaps not uniquely fingerprinting a plume origin. In addition, there is little in common with the Non-volcanic type margins, as rift-related volcanics (hundreds of metres thick) are present (e.g. Antalya Complex) and there is no exposed evidence of simple sheartype detachment faulting or deep-crust exhumation. An ‘intermediate’ type of rift setting is again implied with some, but not profuse, rift volcanism and some extension-related faulting.

Fig. 21. Field photographs of rifted margin units in Cyprus, Greece, Albania and Bosnia. (a) Cyprus. Exotic blocks of Upper Triassic reefal limestone associated with Triassic alkaline lava (view west); interpreted as build-ups on small seamounts during initial sea-floor spreading; Petra tou Romiou (Aphrodite’s rock); SW Cyprus; (b) Greece. Several imbricate thrust sheets of Triassic– Early Cenozoic deep-water sedimentary rocks that originated as the rifted passive margin of Adria (Apulia) adjacent to the Pindos Ocean; each subparallel ridge is an imbricate thrust slice; near Kato Klitoria; NW Peloponnese, Greece; (c) Albania. View westward over proximal carbonate slope facies (Krasta unit), with the autochthonous platform (Albanian Alps) behind; more distal units including ophiolite in the foreground (right); (d) Bosnia. Upper Jurassic– Lower Cretaceous deep-water siliciclastic sandstone turbidites and shales deposited along the rifted northern margin of the Dinaride continent, adjacent to the Dinarie Ocean, a westward continuation of the Pindos– Mirdita Ocean in Greece and Albania.

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Fig. 22. Restored section across the southern margin of the Southern Neotethys in western Cyprus. Note the age and relative locations of the passive margin and exotic units. See the text for explanation.

Rifted North Tauride margins. Further north, the northern margin of the Tauride microcontinent was bordered by the North Tethyan Ocean (‘North Neotethys’), which was located between the Tauride microcontinent and Eurasia (Figs 2 & 12). The Tauride ‘microcontinent’ is traditionally divided into an unmetamorphosed unit in the south, known as the Taurides sensu strictu and a metamorphosed unit further north (affected by late Mesozoic –early Cenozoic convergent tectonic), termed the Anatolides (S¸engo¨r & Yılmaz 1981). Regionally, these two units formed the Mesozoic Anatolide –Tauride platform. In most interpretations, the ophiolitic suture to the north of the Anatolide –Tauride platform, known as the Izmir– Ankara– Erzincan suture zone, is seen as accommodating the main oceanic separation between North Africa and Eurasia during Mesozoic –Early Cenozoic time (S¸engo¨r & Yılmaz 1981; Robertson & Dixon 1984; Dercourt et al. 1986, 1990; Go¨ncu¨og˘lu¨ et al. 1996–1997, 2003; Okay et al. 2001; Robertson et al. 2004). Alternatively, this suture was recently interpreted by Stampfli & Borel (2002) as a back-arc basin located above a northward-dipping subduction zone. These authors further suggested that a major Late Palaeozoic– Early Mesozoic ‘Palaeotethyan’ suture separated the Anatolide and Tauride units (i.e. that these were of quite different palaeogeographic origin). However, no such suture has been identified regionally, and the stratigraphy of the Anatolide and Tauride units can be correlated at least from the Early Triassic onwards (S¸engo¨r & Yılmaz 1981; Demirtas¸lı et al. 1984; Go¨ncu¨og˘lu¨ et al. 1996– 1997, 2003; Okay et al. 2001; Robertson et al. 2004). Alternative models are discussed further below, but meanwhile it is assumed that the longestablished interpretation of the northern margin of the Anatolides– Taurides as a rifted passive margin is essentially correct.

There is well-documented evidence of Triassic rifting and continental breakup along the southern side of this northerly ‘Neotethyan’ oceanic basin; i.e. along the northern margin of the Tauride – Anatolide microcontinent (Monod 1977; Demirtas¸lı ¨ zgu¨l 1984; Robertson & Pickett 2000; et al. 1984; O Go¨ncu¨og˘lu¨ et al. 2003). This margin is preserved in several emplaced units over more than 600 km from east to west (Fig. 12). Relatively proximal units are exposed within the Tauride microcontinent (i.e. within the Tauride –Anatolide platform, e.g. Geyik Dag˘), whereas more distal units are preserved within overriding thrust sheets. The more proximal units (relatively autochthonous) document subsidence and both terrigenous and calcareous sedimentation related to rifting during the Triassic; this culminated in the establishment of a subsiding passive margin during Late Triassic – Early Jurassic time. During the Late Cretaceous –Early Cenozoic, the continental margin and ophiolites were thrust southwards onto the Tauride microcontinent as a large pile of thrust sheets (‘Bozkir nappes’; S¸engo¨r & Yılmaz 1981) As in Oman, the more distal parts of the rifted margin were accreted and thrust southwards onto the Tauride microcontinent, where they are exposed, generally beneath overriding ophiolites. However, the more proximal margin units (platform/slope) remained attached to the downgoing plate. As a result, the leading edge of the Tauride margin was subducted and underwent LT –HP (low temperature– high pressure) metamorphism in mid Late Cretaceous time, as well documented in NW Turkey (Okay et al. 1998, 2001). Continentward parts of the subducted platform (i.e. the Anatolide carbonate platform) experienced lower grade metamorphism, and areas even further south (i.e. the Tauride carbonate platform) escaped metamorphism altogether.

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The northern margin of the Tauride–Anatolide microcontinent (southern margin of the North Tethys) is possibly the most complete example of an emplaced continental margin anywhere in the Tethyan region, as different shallow to deep crustal level units are all exposed along strike in different areas of the suture zone (Izmir –Ankara– Erzincan suture zone; Fig. 12), as summarized below. The rifted passive margin subdivides into two contrasting segments. In the west (to the west of the longitude of the Isparta Angle; Fig. 12) continental slope/abyssal units were thrust southwards onto the continental margin (Menderes Massif) as a large pile of thrust sheets, initially in latest Cretaceous time (de Graciansky 1972; Poisson 1977, 1984; Collins & Robertson 1998, 1999; Danelian et al. 2006). These far-travelled thrust sheets, known as the Lycian Nappes (Fig. 23), restore as proximal, to more distal, slope/abyssal plain successions. As in Oman, the higher units in the thrust stack were derived from a more oceanward setting. The more proximal settings comprise fragments of prerift continental basement, overlain by shallow-water carbonates, whereas the more distal

Fig. 23. Simplified tectonostratigraphic relations of rifted margin and ophiolitic units of the Lycian Allochthon (Lycian nappes), NW Turkey, inferred to have been emplaced from the Mesozoic North Tethyan Ocean onto the Tauride microcontinent (Menderes Massif).

units (Ko¨yceg˘iz unit) begin with Triassic MOR-type lavas, overlain by mainly redeposited carbonates of Jurassic–Late Cretaceous age, representing a deep-water continental slope succession. As in many other areas the proximal edge of the rifted platform was concealed by overthrust units. The Lycian thrust sheets restore as a margin that rifted in the Triassic, with the genesis of oceanic crust during latest Triassic– earliest Jurassic time. An oceanward-dipping sedimentary ramp developed above a basement that is mainly not preserved. Higher in the pile of thrust sheets more distal units, including MOR-type basalts, radiolarites and pelagic carbonates are interpreted as accreted abyssal plain material, and are exposed within a me´lange (Lycian Me´lange) that is located beneath an overriding regional-scale ophiolite (Lycian ophiolite) at the highest levels of the thrust stack. The me´lange includes large blocks of Late Triassic shallow-water carbonate, overlain by pelagic sediments (Domuz Dag˘ unit) (Poisson 1977, 1984). These exotics were interpreted either as rifted fault blocks floored by thinned continental crust (Poisson 1977) or as oceanic seamounts underlain by volcanic rocks (Collins & Robertson 1997, 1998). The actual setting may be indeterminate as no basement to the succession is exposed. The palaeogeography of the rifted margin differed considerably further east (east of the longitude of the Isparta Angle; Fig. 12). In this area, another series of thrust sheets encompassing the same age range was emplaced southwards onto the Tauride carbonate platform during latest Cretaceous time (Beys¸ehir –Hoyran –Hadim nappes; Monod 1977; ¨ zgu¨l 1984; Andrew & Robertson 2002). These O units restore as a more complex rift palaeogeography than the Lycian nappes (Fig. 24). The Tauride carbonate platform in this area (Geyik Dag˘) was bordered by a Triassic rift basin; this infilled with Triassic silicic tuffs, terrigenous and carbonate sediments, mostly turbidites and debris flows (Hug˘lu unit). The rift basin was then covered with mainly siliceous and pelagic carbonate sediments of Jurassic–Late Cretaceous age. The siliceous tuffs may reflect melting (anatexis) of underlying siliceous continental basement rocks that are not exposed. The rift was bordered, oceanwards, by a carbonate platform (Boyali Tepe unit), on which shallow-water carbonates accumulated in Late Triassic time. This platform later subsided and was overlain by pelagic Ammonitico Rosso, then by radiolarites and pelagic carbonates during the Jurassic–Early Cretaceous (Fig. 25). Finally, an additional unit is exposed to the SE on the southern margin of the Tauride carbonate platform (locally termed the Bolkar Dag˘), known as the Mersin Me´lange. In essence, this is interpreted as a highly dismembered equivalent of the

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Fig. 24. Restored relations of the passive margin sediments and igneous rocks on the shelf and the slope successions represented by the Lycian nappes, which were rooted within NW Turkey. The rift and passive margin-related successions are exposed in different thrust sheets over a wide region and thus this reconstruction is tentative. From Robertson (2004).

Beys¸ehir –Hoyran –Hadim nappes. It includes Permian shallow-water limestones, Upper Triassic neritic carbonates, Triassic intermediate-silicic volcanics, deep-water terrigenous and calcareous sediments, and younger pelagic sediments and radiolarites (Parlak & Robertson 2004). This unit therefore provides additional evidence of rifting and continental breakup along the northern margin of the Tauride microcontinent. Summarizing, the entire northern margin of the Tauride microcontinent (southern margin of North

Tethyan oceanic basin) represented by the Tauride and Anatolide units, the Lycian nappes, the Beys¸ehir–Hoyran –Hadim nappes and the Mersin Me´lange shows evidence of rifting and continental breakup during Late Triassic –Early Jurassic time. The margin geometry was relatively simple in the west, with a single oceanward-dipping ramp bordered by subsiding rift blocks (Fig. 24). Further east, the Tauride microcontinent was bordered by a rift basin and a subsiding marginal platform, beyond which Mesozoic oceanic crust

Fig. 25. Restored relations of the passive margin sediments and igneous rocks on the shelf and the slope successions, represented by the Beys¸ehir– Hoyran—Hadim nappes, central Turkey, which were rooted within the Inner Tauride Ocean. These allochthonous units underwent complex thrusting during latest Cretaceous and Middle–Late Eocene time and thus cannot be restored simply by simply unstacking the thrust pile. See the text for discussion.

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formed from latest Triassic –earliest Jurassic time onwards (Fig. 25). The extensive record of rift-related volcanism, but with an absence of exhumed continental basement, suggests that the south margin of the North Tethyan Ocean does not correspond to either of the ideal Volcanic-rifted or Non-volcanic rifted margin types, but is again more of an ‘intermediate’ type. South Tethyan margins in the Balkan region. Moving westwards back into the Balkan region, we have already considered the margin of the most southerly oceanic basin (Ionian and Herodotus basins) that separated North Africa from the Adria (Apulia) microcontinent (Fig. 2). This ocean basin is nowadays still subducting northwards along the Aegean arc, extending from western Greece to south of Crete. As noted earlier in the paper, the timing of initial rifting of this oceanic basin is controversial, as its margins remains deeply buried beneath the Sea of Crete and the Ionian Sea. However, breakup may have taken place during the Late Jurassic –Early Cretaceous (Catalano et al. 2001) (Fig. 2). The southerly rifted margin of the Ionian oceanic basin is represented by onshore and offshore areas of North Africa. By contrast, the northerly margin is largely beneath the sea or was removed by ‘subduction erosion’ related to Cenozoic northward subduction. However, parts of the northerly margin are exposed within a stack of thrust sheets in Crete and the Peloponnese, as mentioned earlier in the paper. Northerly rifted margin of Adria. Further north, the next oceanic basin, the Pindos Ocean, was bordered on its southern side by the rifted margin of Adria (Apulia). Its conjugate margin is represented by the southerly margin of the Pelagonian microcontinent (Robertson et al. 1991; Smith & Rassios 2003; Smith 2006). An alternative (but not favoured) interpretation is discussed later in the paper. The southerly margin of the Pindos oceanic basin is well exposed in the Pindos Mountains, stretching from Albania to southern Greece (Dercourt 1964; Aubouin et al. 1970; Fleury 1980). To the west, the in situ ‘foreland’ is represented by a Mesozoic carbonate platform (Gavrovo –Tripolitza platform) that overlies Triassic synrift sediments (Phyllite– Quartzite unit) and an inferred prerift continental basement (only rarely exposed). The Adria (Apulia) foreland was overridden by a wellorganized stack of thrust sheets that developed along the rifted margin to the east during Early Cenozoic time when the Pindos ocean basin finally closed (Degnan & Robertson 1998, 2006). The Mesozoic volcanic– sedimentary rift succession (i.e. Phyllite –Quartzite unit) and the

overlying carbonate platform (Gavrovo–Tripolitza Nappe) are separated by a major Cenozoic detachment. The underlying Triassic rift succession was detached and subducted together with the downgoing plate where it experienced HP –LT metamorphism before being exhumed (Seidel 1978; Seidel et al. 1982; Zulauf et al. 2002). By contrast, the more oceanward part of the rifted margin (Gavrovo–Tripolitza nappe) and the former slope and abyssal plain units (Pindos–Olonos nappe) were retained within the overriding plate at a high structural level and remained unmetamorphosed. The underlying Phyllite–Quartzite unit records pulsed rifting during Late Palaeozoic–Early Mesozoic Triassic time, with extensive extrusion of intermediatecomposition–silicic volcanics, and deposition of mixed tuffaceous–carbonate–terrigenous sediments. Evidence from Crete indicates this volcanism was mainly of Early Triassic age (Krahl et al. 1983; Krahl & Kaufmann 2004), contrasting with a previously assigned Late Permian age (Seidel 1978). Facies evidence suggests that the latest stage of rifting during the Late Triassic was followed by passive margin subsidence and the construction of a north-facing carbonate platform (several kilometres thick) that persisted from Late Triassic to Early Cenozoic time. The proximal slope of the carbonate platform was mainly overthrust and concealed, as in many other areas, but locally appears beneath the overriding thrust sheets in the western Peloponnese, Greece. By contrast, the base-of-slope to abyssal-plain sediments (equivalent to the Hawasina Complex in Oman) are well exposed and form a well-ordered imbricate stack of thrust sheets (Pindos–Olonos nappes) (Fig. 21b). Unlike most of the other Tethyan rifted margins, overriding ophiolites are absent above these distal margin/ oceanic sediments, and largely for this reason the deep-sea sedimentary successions are unusually intact stratigraphically (Neumann 2003). The Pindos–Olonos thrust sheets are stacked in a regular piggy-back fashion, which allows the original proximal– distal settings to be restored with confidence. The Pindos–Olonos succession, as documented in the NW Peloponnese (Degnan & Robertson 1998, 2006), begins with locally preserved tholeiites that are depositionally overlain by deepwater pelagic limestones, containing the thinshelled bivalve, Halobia. The succession continues upwards through a range of deep-water redeposited terrigenous and carbonate sediments, siliceous radiolarian sediments and pelagic carbonates. It then culminates in thick Early Cenozoic terrigenous turbidites (Pindos flysch). In southern Greece these terrigenous sediments were mainly derived from Adria (Apulia) to the west (Piper 2006), whereas counterparts in northern Greece (north of the Gulf of

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Corinth) include ophiolitic detritus (Faupl et al. 2002) that was derived from the overriding plate (Pelagonian microplate) as the Pindos ocean closed. A thin unit of me´lange is entrained beneath the basal thrust of the Pindos–Olonos thrust sheets (e.g. Degnan & Robertson 1994; Pe-Piper & Piper 2004). This includes numerous blocks of intermediatesilicic tholeiitic extrusives and related pelagic sediments, locally dated as Triassic in age, together with terrigenous sediment derived from the foreland (Degnan & Robertson 1998). The me´lange is interpreted as a remnant of the original continent– ocean transition of the Pindos Ocean that was accreted and then emplaced onto the foreland during Early Cenozoic closure of the Pindos Ocean (Degnan & Robertson 2006). The thrust front directly above the foreland in the west includes several large (up to kilometresized) blocks of basaltic volcanics, overlain by Triassic ammonite-bearing pelagic limestones (e.g. Glafkos and Meghdovas units). These ‘exotics’ are interpreted as igneous extrusions capped with stratigraphically condensed pelagic sediments that originated along the rifted margin; these then remained as relatively elevated blocks during the passive margin subsidence phase. Clastic sediments by-passed these discontinuous marginal ‘highs’ to accumulate in deeper-water slope–abyssal plain settings (Fig. 26b), and were much later detached and incorporated into the base of the thrust stack when the Pindos ocean basin finally closed during the Early Cenozoic (Robertson et al. 1991). In an area north of the Gulf Corinth the rifted margin of Adria (Apulia) was bordered by a rifted fragment some tens of kilometres across, known as the Parnassus unit (Celet 1960) (Fig. 12). This unit was, in turn, bordered to the west by several thrust sheets, termed the Vardoussia unit (Ardaens 1978); these can be restored as the westerly, rifted margin of a discrete Parnassus microcontinent (Degnan & Robertson 1998) (Fig. 26a). The Vardoussia succession begins with Middle Triassic, mixed carbonate –clastic sediments, tuffs and volcanics, and passes upwards into redeposited slope carbonates, derived from the Parnassos carbonate platform to the east (Ardaens 1978; Pe-Piper & Piper 2004). The deep-water succession bordering Adria (Pindos–Olonos nappes) can be traced northwards through northern Greece and Albania, as far as the Scutari –Pec´ lineament (Fig. 12). This is interpreted as an important continental-margin transform fault, which segmented the Adria margin into palaeogeographically contrasting areas (e.g. Dercourt et al. 2000; Robertson & Shallo 2000). Beyond this, similar deep-water Mesozoic –Early Cenozoic sediments are exposed in Budva Zone of former

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Yugoslavia, but there is no evidence that this basin advanced beyond the rift phase (i.e. it represents an aulacogen). The extension of the Pindos Ocean is instead located northwards of the Dinaride continent (related to Apulia), where it includes a deformed passive margin succession, known as the Bosnian Flysch (Fig. 21d), which is in structural contact with the Dinaride ophiolite belt (e.g. Robertson & Karamata 1994; Karamata 2006). Discussion of the possible linkage, westwards through Bosnia, Croatia, Hungary and Austria, to the Western Tethys rifted margin, although fruitful, is outside the present scope. In summary, evidence from the emplaced northern margin of Adria shows that, following Late Palaeozoic pulsed rifting, the Pindos Ocean began to open during latest Triassic –Early Jurassic time, associated with the development of a proximal carbonate platform, passing eastwards into a distal deep-water succession. Part of the rifted continent– ocean transition is preserved as accreted blocks within me´lange. Rift-related fault blocks adjacent to the ocean subsided and were covered by pelagic carbonate. The rifted margin was bordered by one much larger rifted fragment (Parnassos), with its own small passive margin (Vardoussia). Southerly rifted margin of the Pelagonian microcontinent. Here, it is assumed that the conjugate margin of Adria (Apulia) is represented by the southerly rifted margin of the Pelagonian microcontinent. Alternative models are discussed in the section following this overview of Balkan rifted margins. The former southerly margin of the Pelagonian microcontinent is now represented by thrust sheets and me´lange; these are exposed structurally above the Pelagonian Zone to the east. These marginal units are most intact in the Othris Mountains of central Greece (Smith et al. 1995), but they are also exposed, typically, as me´lange and dismembered thrust sheets over a wide area (Fig. 12), notably in the Pindos Mountains (Avdella Me´lange: Jones & Robertson 1991, 1994) (Figs 27 & 28) on the large island of Evia (Ferrie`re et al. 1988; Robertson 1991; Danelian & Robertson 2001), and in the Argolis Peninsula (Baumgartner 1985; Clift & Robertson 1990) further south. In the interpretation favoured here these early Mesozoic rift-related units were emplaced northeastwards from the Pindos Ocean. The driving force was the collision of a subduction trench with the Pelagonian microcontinent during Middle–Late Jurassic time (Robertson et al. 1991; Smith & Rassios 2003; Rassios & Moores 2006; Smith 2006). Along the former southern margin of the Pelagonian microcontinent, as exposed in the

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Fig. 26. Restored relations of the passive margin sediments and igneous rocks on the shelf and the continental slope successions related to the SW margin of the Mesozoic Pindos Ocean exposed in western Greece. (a) Central western Greece (north of the Gulf of Corinth) where the rifted margin is inferred to have been bordered by a small microcontinental unit, the Parnassus platform (see Fig. 12), and its passive margin (Vardoussia unit). (b) Rifted margin of Adria (Apulia) facing eastwards into the Pindos ocean, as exposed within the Peloponnese, NW Greece. (a) and (b) are effectively conjugate margins. See the text for discussion.

regional Pelagonian nappe, Late Palaeozoic metamorphic basement rocks are depositionally overlain by Upper Permian non-marine clastic sediments (Mountrakis 1986). Above, the Triassic succession includes bimodal volcanics and mixed carbonate– clastic successions, then thick Jurassic platform carbonates, that are well exposed on the island of Evia. The Pelagonian platform cannot be considered as a single homogeneous entity on a regional scale, but was palaeogeographically varied. Several intraplatform basins can be reconstructed. These were floored by intermediate-composition volcanics and

volcaniclastic sediments, in turn overlain by redeposited platform-derived carbonates, as seen in the Argolis Peninsula (Clift & Robertson 1990; Clift 1995) and probably also in northern Greece (Sharp & Robertson 2006). The proximal–distal deep-water rifted margin can be restored from overriding me´lange and thrust sheets (Fig. 29). Successions are most intact, stratigraphically, in the Othris area where both relatively proximal deep-water slope to base-of-slope successions and more distal rift-related alkaline volcanics (Agrilia lavas) are present (Smith et al.

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Fig. 27. Simplified tectonostratigraphic relations of rifted margin, oceanic and ophiolitic units of the Pindos allochthon in the Pindos Mountains, NW Greece, inferred to have been emplaced from the Pindos Ocean onto the Pelagonian microcontinent. The initial emplacement occurred in Late Jurassic time, followed by rethrusting towards the SW during Early Tertiary time. From Robertson (2004).

1975; Price 1976). A wide range of Triassic rift-related rocks are also well exposed further NW in the Avdella Me´lange of the Pindos Mountains (Jones & Robertson 1991, 1994). Lithologies include terrigenous turbidites and platform-derived carbonates, interspersed with alkaline to transitional type composition volcanics and radiolarian cherts. Shallow-water carbonates are interpreted as small platforms that developed on subsiding marginal fault blocks. The Jurassic was characterized by passive margin subsidence with the deposition of slope carbonates and pelagic sediments (e.g. in the Othris area; Price 1976). The passive margin phase was followed by subduction and tectonic accretion, culminating in the emplacement of the rifted margin and associated ophiolitic units onto the Pelagonian microcontinent during Middle –Late Jurassic time (Fig. 28). After this emplacement shallow-water carbonate platform deposition resumed on the Pelagonian

microcontinent during the Late Jurassic (locally) and the Cretaceous (e.g. on Evia). During the Early Cenozoic continental collision finally sutured the Pindos Ocean, and the Pelagonian microcontinent was mainly detached from its prerift basement and internally imbricated to form several huge thrust sheets that were emplaced to the SW. Summarizing, the inferred northerly conjugate margin of the Pindos Ocean (southern margin of the Pelagonian microcontinent) records rifting during Triassic time associated with widespread volcanism, commonly of alkaline (transitional) type; this was followed by passive margin subsidence during the Early Jurassic. The NE part of the Pindos Ocean (bordering the Pelagonian microcontinent) was deformed during Middle–Late Jurassic time, associated with the emplacement of rift-margin units and ophiolites onto the Pelagonian microcontinent to the NE (present co-ordinates). By contrast, the conjugate Adria (Apulia) margin to the

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Fig. 28. Block diagrams showing the restored relations of the passive margin sediments and igneous rocks on the shelf and the slope successions of the represented by the Avdella Melange, Pindos Mountains, NW Greece. Note the role of accretionary processes. From Jones & Robertson (1994).

south escaped this deformation and remained as a subsiding passive margin until final closure of the Pindos Ocean in Early Cenozoic time. The conjugate rifted margins of the Pindos Ocean do not show evidence of voluminous plumetype regional magmatism as in the Volcanic-rifted margin type, or the simple shear-controlled basement exhumation as seen in Non-volcanic margins. On the other hand, Triassic rift volcanism (up to hundreds of metres thick) is regionally extensive and arguably compatible with a plume origin

(Dixon & Robertson 1999). The rifted Pindos Basin again appears to document an ‘intermediate’ rift setting compared to the two end members. North Tethyan margins in the Balkan region. To the north of the Pelagonian microcontinent the third Mesozoic oceanic basin is known as the Vardar Ocean (Figs 2 & 12). The Vardar Ocean extended westwards from northern Greece, through former Yugoslavia and Hungary to the Eastern Alps. Indeed, it was the Late Cretaceous closure of this

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Fig. 29. Restored relations of the rifted margin of the Pelagonian microcontinent, passing into the Pindos Ocean in central and northern Greece (Othris, Evia and Pindos area). See the text for discussion.

ocean basin that triggered the initial compression of the northern margin of the Jurassic Ligurian– Piedmont Ocean, as seen in the Alps (see Manatschal et al. 2007). The Vardar Ocean is the most difficult of all of the Eastern Tethyan ocean basins to interpret, mainly because of only limited study, regional metamorphism of some units and limited exposure. Here it is assumed that the Vardar Ocean was bordered to the south by the northern rifted margin of the Pelagonian microcontinent and by its conjugate, the southern margin of the Serbo-Macedonian continent. However, alternative interpretations are mentioned in the following section. The Vardar Zone is a composite unit and is characterized by several subunits, namely the Almopias unit in the west, the Paikon unit in the centre and the Peonais unit in the east (Mercier 1968, 1979). Each of these three units is internally complex. Successions exposed in each unit are shown in Figures 30 & 31. The northern margin of the Pelagonian microcontinent (Pelagonian Zone) includes Triassic –Jurassic metamorphosed shallow-water carbonates associated with alkaline meta-basalts that are assumed to have erupted during Triassic rifting (Sharp & Robertson 2006). The Almopias Zone was thrust SW over the Pelagonian Zone during the final Early Cenozoic closure of the Vardar Ocean. No rifted-margin slope sediments are known either in the Pelagonian Zone or the Almopias unit. The Almopias unit includes MOR-type basalts (Sharp & Robertson 1994, 2006) in depositional contact with Triassic radiolarites (Stais & Ferrie`re 1991), and is interpreted as a fragment of Triassic Vardar (‘Almopias’) oceanic crust. The Paikon unit is dominated by a Jurassic volcanic arc that relates to partial closure of the Vardar Ocean (Mercier 1986; Brown & Robertson 2003, 2004). The Peonais unit is characterized by

a Jurassic ophiolite (Guevgueli ophiolite) that is interpreted as a marginal basin above a northwarddipping subduction zone (Be´bien et al. 1987). The Peonnais unit includes Jurassic terrigenous turbidites (Svoula Flysch: Kauffmann et al. 1976), interpreted as part of the deep-water slope sequence of the Serbo-Macedonian continental unit to the north. The southern margin of the SerboMacedonian Zone (rifted northern margin of the Vardar Ocean) exposes poorly dated PermoTriassic alluvial conglomerates, overlain by Triassic bimodal MOR-type tholeiites, in turn passing into a Late Triassic carbonate platform succession (Stais & Ferrie`re 1991). In summary, the Vardar Ocean is well exposed in northern Greece and is marked by a poorly preserved Triassic rifted margin on its southern side (Pelagonian microcontinent) and a better preserved Triassic rifted margin on its northern side (Serbo-Macedonian continent). This includes both proximal non-marine, shallow-marine and deepermarine successions. ‘Permo-Triassic’ rifting was followed by passive margin subsidence by Late Triassic time as the Vardar Ocean opened. The record is too sparse to allow useful comparison with either a Volcanic-type or a Non-volcanic-type rifted margin. However, it is notable the rift volcanics on the northern (Serbo-Macedonian) margin are unusually voluminous and could show a plume influence based on the published geochemistry (Dimitriadis & Asvesta 1993). This concentration of preserved rift volcanism on the northern margin is also suggestive of the rift being asymmetrical. The Serbo-Macedonian Zone to the north has in the past been regarded either as part of the regional south margin of Eurasia or as a discrete microcontinent. Recently, Himmercus & Reichmann (2006) have identified what they consider to be a

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Fig. 30. Regional setting of the Vardar Zone, northern Greece between the Pelagonian microcontinent and the Serbo-Macedonian continent to the north. (a) Outline geological map; (b) simplified cross-section. Triassic margin units are exposed in the Pelagonian Zone and the Almopias Zone. (c) Alternative tectonic models for the Vardar Zone during Middle– Late Jurassic time: (A) thrusting from an ocean far to the north beyond the Serbo-Macedonian continent; (B) subduction to SW below the Pelagonian continent; and (C) subduction to NW below the Serbo-Macedonian continent. (D) is supported by structural and geochemical evidence. See the text for data sources and explanation.

‘mid-Alpine’ suture within the high-grade metamorphic rocks of the Serbo-Macedonian Zone. This suture may possibly be an equivalent of the Izmir–Ankara– Erzincan suture in Turkey, which, as noted in a previous section, is widely interpreted as the main suture of the Northern Tethys Ocean. Further north again, the true Eurasian continental margin is represented by the northern part of the Serbo-Macedonian Zone, the Rhodope massif and the Strandja massif. The southern margin of Eurasia in this region appears to have experienced a long history of active margin processes (Karamata & Jankovic´ 2000; Karamata & Vujnovic´ 2000; Karamata 2006) comparable with the Eurasian

margin of Turkey, further east, as exposed in the Pontides and the Caucasus (e.g. Adamia et al. 1995; Nikishin et al. 2001).

Alternative Tethyan reconstructions In the preceding sections, the Mesozoic tectonic setting of the Eastern Mediterranean has been interpreted in terms of several microcontinents separating three main interlinked small oceanic basins bordered by mainly Triassic rifted margins. However, it is fair to note that there are several quite different interpretations in the literature. In the past there have been various attempts to

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Fig. 31. Summary of successions within and adjacent to the Vardar zone, northern Greece. 1 –2, rifted margin of the Pelagonian continent; 3, Upper Triassic oceanic crust; 4– 5, continental margin-type arc constructed on continental crust of the Serbo-Macedonian zone; 6, back-arc marginal basin, rifted within Serbo-Macedonian continental crust; and 7, Serbo-Macedonian continent. See Figure 30a for the location of units and the text for discussion.

simplify an admittedly complex palaeogeography. Several of the currently debated tectonic reconstructions are shown in Figures 32 & 33. In one school of thought (Fig. 32a) the Mesozoic ophiolites and related rifted margin units were thrust hundreds of kilometres generally southwards from a single, wide Mesozoic Tethys located adjacent to Eurasia (Dercourt et al. 1986, 2000; Stampfli et al. 1991). For Turkey and Cyprus, this implies thrusting of all the ophiolites (e.g. Troodos ophiolite) and the rifted margin units (e.g. Mamonia Complex, western Cyprus; Antalya Complex, SW Turkey) over the Tauride –Anatolide platform from a single Mesozoic Tethys to the north during latest Cretaceous –early Cenozoic time (Ricou et al. 1974). However, this can be excluded as these allochthonous unit would have had to have been thrust through a stratigraphical succession, which, at least locally (e.g. Bey Dag˘ları carbonate platform; Isparta Angle apex; Fig. 12), remained intact during the time in question (Late Mesozoic – Early Cenozoic). In the Balkan region the situation is less clear-cut as no such intact ‘blocking’ successions are preserved. It has been argued for this area that all of the Mesozoic ophiolites and the rifted margin units originated in the northerly Vardar

Ocean or even further north (i.e. an ‘ultra-internal origin’) (Michard et al. 1998), and were then thrust over the Pelagonian Zone into the Pindos Zone, which is considered as a continental rift (Aubouin et al. 1970; S¸engo¨r 1984; Dercourt et al. 1986, 2000; Stampfli et al. 1991). However, the available structural evidence from the Jurassic ophiolites, especially the Vourinos ophiolite, is mainly indicative of emplacement from a Pindos ocean, towards the Pelagonian Zone to the NE (Rassios et al. 1994; Rassios & Moores 2006) (Fig. 34). In additional, structural evidence from several of the rifted margin units, especially Othris, supports emplacement from a Pindos ocean located to the SW (Smith et al. 1979; Robertson et al. 1991) (Fig. 35). Another option is that some of the sutures can be considered as exotic terranes, in which different parts of the same Tethyan oceanic basin and its margins were interleaved by strike-slip. Specifically in the Balkan region, it has been suggested that the Pindos and Vardar sutured units formed in a single Mesozoic ocean basin that was later dissected and displaced to form two subparallel exotic terranes (Smith & Spray 1984). However, in the author’s view the tectonostratigraphy of the Vardar Zone

Fig. 32. Cartoons of alternative tectonic reconstruction of Tethys in the Balkan region. (a) Assumes only one Mesozoic Tethyan ocean existed and that the ophiolites and some of the ophiolitic units were thrust hundreds of kilometres generally southwards to their present positions (Dercourt et al. 2000). (b) Assumes the presence of marginal rifts (i.e. Phyllite –Quartzite unit in Crete and south Peloponnese), but with the Pindos and Vardar zones as small Triassic oceanic basins bordered by rifted margins (Robertson et al. 1996; and this paper). (c) Assumes a north-dipping subduction zone now located in the south Aegean region (i.e. Crete and south Peloponnese), with the Pindos and Vardar zones representing back-arc basins. (d) Postulates southward subduction and marginal basin formation. See the text for discussion.

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Fig. 33. Sketch palaeogeographical maps of the Eastern Mediterranean region for the key Late Triassic period (arranged in chronological order): (a) southward subduction model (S¸engo¨r 1984); (b) northward subduction with continental rifts bordering Gondwana (Dercourt et al. 1986, 2000); (c) northward subduction and marginal basin formation (Stampfli & Borel 2002); (d) northward subduction with continental fragments rifted from Gondwana (Robertson et al. 2004; this paper). The main differences between (b) and (d) is the size and location of continental fragments and the timing of spreading. See the text for discussion.

and the Pindos Zone are sufficiently different to make this unlikely. Also, the tectonic contacts are low-angle thrusts and the required major high-angle strike-slip faults have not been identified. For the ‘Turkish region’ application of the terrane hypothesis would imply that the Northern and Southern Tethyan units were once part of a single ocean that was tectonically duplicated during latest Cretaceous –Early Cenozoic time. This, again, seems unlikely; the tectonostratigraphy of rifted margin units north and south of the Tauride– Anatolide platform differ in important respects, and tectonic contacts are, again, low-angle thrusts (except locally as in the Antalya Complex, SW Turkey). In this tectonic interpretation the Tauride– Anatolide platform needs to be part of the north Gondwana margin from which it was transported laterally, ending up outboard (i.e. northward) of

oceanic units. However, in the east the Arabian margin did not suture until the Miocene, and further west in the easternmost Mediterranean the North African (north Gondwana) margin remains in a syn- to pre-collisional setting. Removal of a regional-scale terrane from North Africa and interleaving with oceanic units is difficult to envision. Finally, the near orthogonal (north–south) Africa– Eurasia convergence pathway (Livermore & Smith 1984; Smith 2006) does not provide any obvious driving mechanism for long-distance east–west terrane displacement. In summary, the author does not believe that the multiple sutures can be explained, either by very long-distance southward thrusting of oceanic units or by terrane displacement and interleaving of units from a single Mesozoic ocean basin. However, this does not exclude an important but more

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Fig. 34. Summary of structural evidence of emplacement directions of allochthonous units in Greece and Albania. This supports emplacement from a Pindos oceanic basin to the SW (in present co-ordinates) during Late Jurassic time. See the text for data sources and discussion.

local role for strike-slip during rifting and later emplacement, as exemplified by the Antalya Complex. It is therefore, assumed that the Eastern Mediterranean region, unlike Oman or the

Himalayas, was characterized by several contemporaneous small oceanic basins interspersed with microcontinents during Mesozoic– Early Cenozoic time.

Fig. 35. Comparison of two tectonic models for the setting of the Triassic rifted-margin units in the NW Peloponnese, Greece. (a) Outline structural map and section showing thrust slices of Triassic– Early Cenozoic deep-water sediments of the Adria (Apulia) rifted passive margin (Pindos–Olonos nappes). (b) Interpretation as a rifted margin of a small Pindos oceanic basin (favoured here). (c) As a back-arc-basin above a north-dipping subduction zone. See the text for discussion.

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Why alternative back-arc settings are unlikely? In the foregoing discussion it has been assumed that several continental fragments rifted from Gondwana, without any local association with subduction (e.g. Robertson & Dixon 1984; Dercourt et al. 1986, 2000; Robertson et al. 1991; Smith 1993, 2006). However, several authors have suggested that some, or even all, of the Eastern Tethyan rifts represent back-arc basins (Pe-Piper 1982; S¸engo¨r 1984; Stampfli et al. 2001) that opened related to subduction beneath a continental borderland. A back-arc setting is quite attractive in principal as it provides a viable mechanism to open peripheral rift basins. Such rifts might be comparable with, for example, the Woodlark back-arc basin, SW Pacific, which developed within a longlived active margin setting (see Goodliffe & Taylor 2007) or the Tyrrhenian Sea, Western Mediterranean Sea, which opened in a regional convergent setting (Kastens et al. 1988). Even if valid, this would not preclude valuable insights on rift processes, as shown by studies of the Woodlark rift basin, which is located in a regional back-arc setting (Goodliffe & Taylor 2007). There is widespread evidence of rifting along the southern margin of Eurasia during Late Palaeozoic –Early Mesozoic time related to generally northward subduction (see Robertson 2002, 2004, 2006b). This includes the Jurassic Guevgueli marginal basin, NE Greece (Be´bien et al. 1987) and the Triassic –Jurassic Ku¨re Basin, northern Turkey (Ustao¨mer & Robertson 1994, 1977). These backarc basins are associated with coeval arc volcanism (e.g. Paikon arc, northern Greece). However, it should be noted that these units are unrelated to the Triassic rift settings discussed in this paper. Several different back-arc-type settings have been proposed for the Eastern Tethyan region, as follows.

Rifting above a southward-dipping subduction zone The subduction zone dipped southwards from a preexisting ocean to the north (Palaeotethys) beneath the northern margin of Gondwana as a result a continental fragment (‘Cimmeria’) rifted and drifted into Tethys, opening one or several back-arc basins (Figs 32d & 33a). This model has been applied to the Turkish area (S¸engo¨r & Yılmaz 1981) and Oman (S¸engor 1984), but it can also be applied to the Balkans and the Himalayas. However, several points oppose this interpretation. First, the requisite magmatic arcs of the appropriate age (e.g. Triassic) appear to be absent. They should

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have drifted away from Arabia as marginal basins opened, but are yet to be identified in accretionary terranes further north (e.g. Triassic Karakaya me´lange terrane, northern Turkey). Secondly, in many areas (e.g. Himalayas, Oman, Baer-Bassit, Cyprus, Pindos) the Triassic rift-related basalts lack a subduction influence as expected for a subduction-related setting (Dixon & Robertson 1993; Robertson 2002). On the other hand, Triassic volcanics of the Balkans (e.g. southern Greece, Albania, Serbia, Eastern Alps, Dolomites) commonly do show a subduction-related chemistry, which requires an explanation. Thirdly, the deep-sea sedimentary covers of the more easterly rift settings mentioned above (Himalayas, Oman, Baer –Bassit, Cyprus, Pindos) are mainly terrigenous (Robertson 1981, 1986). This contrasts strongly with marginal basin settings such as the Woodlark rift basin (e.g. Taylor et al. 1999a, b; Robertson et al. 2001) and the Tyrrhenian Sea (e.g. Marsili Basin: Kastens et al. 1988). Fourthly, multiple marginal basins of similar age (Triassic) would be required to explain the multistranded nature of the Early Mesozoic small ocean basins, as discussed earlier in the paper. The absence of related Early Mesozoic magmatic arcs is thus even more problematic. Finally, the subductionrelated rift model requires additional elements, notably accretionary wedges and fore-arc basins of the appropriate ages (infilled with arc-derived volcaniclastic sediments), but these do not appear to be present in the requisite locations.

Rifting above a northward-dipping subduction zone In this model a subduction zone dipped northwards beneath the Eurasian margin, continuously from Late Palaeozoic to Early Cenozoic time (Stampfli et al. 2001; Stampfli & Borel 2002; Stampfli Kozur 2006) (Figs 32c & 33c). Northward-dipping subduction is indeed widely inferred (Bernoulli & Laubscher 1972; Robertson & Dixon 1984; Dercourt et al. 1986, 2000). For example, there is evidence of northward subduction beneath the Eurasian margin (e.g. Pontides, northern Turkey), in the form of Late Palaeozoic granites (Ustao¨mer et al. 2005; Okay et al. 2006), and a Triassic volcanic arc (C¸angaldag˘: Ustao¨mer & Robertson 1997). In most of these earlier interpretations the subduction zone was located close to the Eurasian margin, whereas the continental terranes further south (e.g. Pelagonian Zone, Serbo-Macedonian Zone, Anatolide platform) were seen as fragments rifted from Gondwana, as discussed previously. However, Stampfli et al. (2001, 2003) position the Permo-Triassic subduction zone south of these

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continental units. In this interpretation the continental fragments formed part of the southern margin of Eurasia by Late Palaeozoic–Early Mesozoic time (they probably rifted from Gondwana at a much earlier stage, but this is not discussed here). The northward-dipping subduction model (Stampfli et al. 2001, 2003; Stampfli & Borel 2002) has two key testable implications. The first is that a Late Palaeozoic –Early Mesozoic northwarddipping subduction zone was located in the southern Aegean region (Fig. 32c), and between the Anatolides and Taurides in southern Turkey. The second is that northward subduction was compensated for by the opening of a wide, southerly ‘Neotethys’ during Late Permian–Triassic time (Fig. 33c). These implications have been tested by detailed fieldwork in the Turkish and Cyprus areas, and have lead to the conclusion that, first, there is no preserved evidence of a north-dipping subduction zone within or adjacent to the Tauride– Anatolide units in southern Turkey (Robertson et al. 2004). Secondly, unlike Oman, there is no evidence of Permian sea-floor spreading in the Eastern Mediterranean, based on a combination of marine and landbased studies (discussed earlier in this paper). For the critical westerly areas (i.e. Crete, southern Peloponnese, Evia and western Sicily) the main conclusion is that rifting rather than subduction/ accretion characterized the Late Palaeozoic– Triassic history of the south Aegean region (e.g. Crete and Peloponnese; Fig. 35b), where there is no evidence of early Mesozoic ophiolites, metamorphism or contemporaneous arc volcanism (Robertson 2006b). In addition, there is no evidence of sea-floor spreading in the south– Aegean during Permian time (see earlier discussion). Other objections to northward-dipping subduction in a southerly location (south Aegean) are similar to the southward-dipping subduction model, and include the need for the formation of multiple marginal basins (e.g. Pindos and Vardar), yet with an absence of coeval magmatic arcs, fore-arc basins and accretionary prisms. Notably, the northern margin of the Vardar Zone records a rifted margin without a subduction influence or arc-derived volcaniclastic sediments. Also, in Greece, calc-alkaline plutonic rocks within the Pelagonian microcontinental unit are Late Palaeozoic in age, much older that the Triassic rifting (Mountrakis 1986; Kotopouli et al. 2000). Finally, an even more complicated scenario, one involving multiple marginal basins, has been suggested based mostly on igneous geochemical evidence. In this, a Late Palaeozoic–Early Mesozoic subduction zone dipped northwards (Pe-Piper 1982), as in Stampfli et al.’s (2001) interpretation, and an additional Triassic subduction dipped southwards from an inferred Pindos back-arc basin

(Pe-Piper & Piper 2004). This model was introduced to explain the igneous geochemical evidence from the Triassic rift basalts, but runs into the same difficulties as with the other subduction-related model, especially the absence of evolved magmatic arcs and accretionary prisms of the appropriate age and location. The geological evidence therefore does not support formation of the Triassic rift basins on the Eastern Tethys as marginal basins above either northward- or southward-dipping subduction zones. However, an outstanding problem is the tectonic setting of some of the Triassic rift volcanics in the Balkan region (eastern Crete, Peloponnese, Albania, Serbia, Eastern Alps, Dolomites). Geochemical plots show considerable variation in chemical composition within and between the suture zones in Greece, with no simple pattern emerging (e.g. MOR-type basalts in one suture v. subduction influenced in another) (Fig. 36). Many of the basalts in both the Pindos and the Vardar zones show a clearly expressed subduction geochemical signature (e.g. Pe-Piper & Piper 2004). An alternative suggestion to Triassic active subduction is that the subduction influence was inherited from earlier subduction in the region (Robertson & Dixon 1984). A possible explanation

Fig. 36. Summary of the inferred tectonic settings of Triassic rift basalts analysed at Edinburgh University, School of GeoSciences. Samples were analysed over a number of years by S. A. M. Brown, P. D. Clift, P. J. Degnan, J. E. Dixon, G. Jones, A. H. F. Robertson and I. R. Sharp. An additional large dataset was published by Pe-Piper & Piper (2004).

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is that southward subduction beneath the north margin of Gondwana was active during the Late Palaeozoic Hercynian ‘orogeny’ and that the subduction signature of Triassic basalts relates to melting of subduction-contaminated lower crust and upper mantle. These volcanics could represent small melt fractions that were strongly influenced by the composition of the underlying crust and mantle lithosphere (Dixon & Robertson 1993). The south Aegean and Balkan region was characterized by calk-alkaline magmatism (e.g. extensive granite extrusion and andesitic volcanism) during Late Palaeozoic time, related to the Hercynian orogeny (Mountrakis 1986; Kotopouli et al. 2000). This pre-existing magmatism was possibly the source of the subduction signature in some of the Triassic rift-related volcanic rocks rather than contemporaneous Triassic subduction. During the Upper Palaeozoic (Early Carboniferous) subduction was perhaps both northwards and southwards, i.e. beneath both Gondwana and Laurasia (Eurasia). This culminated in complete closure of the ‘Hercynian’ ocean in the west (Dercourt et al. 1986, 2000; Ziegler 1990; Okay et al. 2006), whereas Tethys (Palaeotethys) remained partly open further east. Subsequent subduction was northward-directed beneath Eurasia during the Triassic and the late Mesozoic –Early Cenozoic. A relict south-facing Hercynian slab could thus have remained along the north Gondwana margin and, if so, was reactivated during Triassic rifting, resulting in subduction-influenced volcanism as in some, but not all, areas of the Balkans (e.g. eastern Crete but not western Crete; southern Pindos but not northern Pindos, etc.) (Fig. 36). In summary, the Eastern Tethyan rifts cannot be considered as subduction-related back-arc marginal basins; where present, a subduction influence is seen as being relict from previous, probably ‘Hercynian’, subduction.

Comparisons between the Eastern Tethyan and Alpine – North Atlantic rift settings Taken together, each of the rifted margins summarized here, Oman, the Himalayas and the Eastern Mediterranean, show many features in common, as follows. † Rifting did not take place in the midst of a continent like the Atlantic or the Western Tethys (Alps), but was located along the northern margin of the pre-existing supercontinent, Gondwana, facing outwards into an older ocean (Palaeotethys). † Rifting was pulsed prior to onset of sea-floor spreading. In Oman, the margin experienced pulses of rifting dating from the Carboniferous,

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with initial spreading in the Late Permian. In the East Mediterranean (e.g. Crete) rift events along the North African–Levant margin occurred during Late Carboniferous –Permian and the Early–Middle Triassic, preceeding breakup and spreading during latest Triassic – earliest Jurassic time. Rifting and crustal extension in the Eastern Mediterranean had advanced sufficiently to form a deep-water basin with open-marine circulation during Late Carboniferous –Permian time, but there is no evidence of oceanic crust of this age, unlike in Oman. In the Himalayas, pulses of extension and related alkaline magmatism continued to affect the subsiding passive margin after continental breakup. Renewed volcanism, possibly rift-related, is also present in the Levant. † Rifting was associated with widespread volcanism. In the Himalayas, Early Permian volcanism (Panjal Traps) affected a very wide area of the rifted margin. However, the thickness of volcanics decreases northwards towards the rift axis. Late Permian–Early Triassic rift-related volcanism also affected the south Aegean (e.g. Crete). Triassic rift-related volcanism affected large areas (e.g. Pindos rift in Greece; Antalya rift in SW Turkey). Taken together, the volcanics show a transition from intermediate-silicic composition lavas in proximal settings, to transitional MORB basalts in more oceanic settings. The volcanics appear to have extended right up to the edge of the rifted margin, as documented in SW Turkey (Isparta Angle), and thus the continent– ocean transition zone appears to have been quite narrow (kilometres rather than tens of kilometres). † There is little evidence of the involvement of large-scale regional mantle plumes in the breakup of continents in the Eastern Mediterranean region. Large-scale domal uplifts and associated regional-scale deep erosion are not observed. Unconformities relate more to footwall uplift during pulsed rifting (e.g. Late Triassic of Crete). However, small-scale plume-type influences are consistent with the geochemical evidence. Also, the thick and regionally extensive Permian Panjal traps and associated regional uplift in the Himalayas are suggestive of a regional-scale plume influence. † The rifted margins were palaeogeographically varied, ranging from oceanward-deepening ramps of deep-sea sediments above a volcanic basement (e.g. Lamayuru/Karamba Formation, Himalayas; Lycian nappes, NW Turkey) to more complex settings characterized by large marginal rift basins and off-margin carbonate platforms (e.g. Beys¸ehir–Hoyran nappes, central Turkey). In addition, several individual rifted margins

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were bordered by rifted fragments up to tens of kilometres in size (e.g. Vardoussia, Greece; Kemer units in the Isparta Angle, SW Turkey). † Along the margins of the ocean basins marginal highs originated as small platform units, rifted from the larger, parent platforms. Such fault blocks were capped by cephalopod-bearing pelagic limestones, and include manganese crusts of hydrothermal origin (Rustaq Exotic, Oman; Go¨dene Zone, Antalya Complex, SW Turkey). The rift blocks were isolated within marginal oceanic crust as spreading began. Such blocks include synrift/prerift facies in some areas (e.g. Jebal Qamar exotics, Oman; Kemer units; Antalya). The fault blocks were covered by shallow-water carbonates, then by pelagic carbonates. This facies change was triggered by tectonic subsidence, perhaps accentuated by eustatic sea-level rise. Exceptionally, in SW Turkey (Antalya), shallow-water carbonate deposition kept pace with tectonic subsidence or sea-level change for more than 100 Ma. This setting of reduced subsidence

probably relates to a setting of transform rifting or oblique rifting, and the local palaeogeographic setting of this area as an embayment of the Southern Tethys. † True volcanic seamounts formed further out in the Tethyan Ocean and were overlain by thick, rapidly subsiding shallow-marine carbonate platforms, passing upwards into deeper water pelagic carbonates (e.g. Oman Triassic Exotics).

Western Tethys – North Atlantic Following the (incomplete) drilling of the Iberia – Newfoundland conjugate transect, the rift/breakup history of the Central Northern Atlantic can now be correlated with the Western Tethyan region (Manatchal et al. 2007; Tucholke et al. 2007). Together, these regions provide the type example of a Non-volcanic rifted margin (Fig. 37a). This margin type has several key characteristics. First, mantle lithosphere was exhumed along the opposing rifted margins (i.e. Iberia and Newfoundland in the Central Atlantic; Eurasian and African

Fig. 37. Comparison of ideal (a) ‘Non-volcanic’ versus (b) ‘Volcanic’-rifted margins, based mainly on ODP drilling in the north Atlantic. Much greater thicknesses of volcanic rocks are to be expected on volcanic-rifted margins, but without evidence of exhumed lower continental crust and mantle. See the text for data sources and discussion.

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margins in the Alps). Secondly, the exhumed mantle is directly overlain by MOR-type basaltic volcanics (Tucholke et al. 2004), as discovered at ODP Site 1277 (Robertson 2007). Thirdly, the conjugate margins were, in turn, overlain by deep-water sediments, as the facies vary markedly under the influence of local palaeogeography. For example, the Newfoundland rifted margin was strongly influenced by the input of terrigenous clastic sediments that were locally derived from adjacent continental basement to the west and north (e.g. Avalon platform), whereas the sediments on the Iberia margin are more calcareous and pelagic owing to the trapping of clastic sediments by the rifted Galicia rifted basin to the east. The evolution of the Western Tethys is now quite well understood. The central North Atlantic and the Piedmont –Ligurian Ocean in the Alpine region opened in concert. The Western Tethyan rifts, as exposed in the Alps, are characterized by the exhumation of lower crust –upper mantle and covering by MOR-type basalts (see Manatschal et al. 2007). This rift basin was clearly asymmetrical. The Eurasian margin of the Piedmont –Ligurian Ocean was relatively abrupt (upper plate) and was dominated by terrigenous sediments, whereas the conjugate Adria (Apulia) margin (lower plate) exhibits a greater abundance of pelagic sediments (e.g. radiolarites, pelagic carbonates) reflecting isolation of the rifted margin from the ocean by a wide rifted borderland (e.g. in the Dolomites, northern Italy). On an oceanic scale, the Western Tethys can be seen as a gulf of the central North Atlantic. However, one uncertainly is whether only a single ocean basin existed in this region (Piedmont– Ligurian Ocean), as is likely, or if two subparallel oceanic basins co-existed (Stampfli & Marchant 1997). Despite the similar breakup history, the evolution of the Central Atlantic and Western Tethys eventually diverged. The North Atlantic continued to open, whereas the northerly Eurasian margin of the Western Tethys began to be deformed in latest Cretaceous time. This was related to closure of the most westerly of the Eastern Tethys ocean basins (Vardar Ocean), and was followed by complete closure of the Western Tethys during Early Cenozoic time. Contrasting with the Non-volcanic rifted margin, Volcanic-rifted margins are mainly known from two legs of drilling of the Eastern Greenland margin (e.g. Larsen & Saunders 1998; Fitton et al. 2000) (Fig. 37b). The conjugate volcanic rifted margin (e.g. Vøring plateau) remains less well understood. Key features of volcanic-rifted margins are: first, seaward-dipping seismic reflectors originated as lava flows that were erupted as flood basalts. Secondly, there is a narrow continent– ocean transition zone (several kilometres), with MOR-type

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oceanic crust reaching in close to the rifted margin. Thirdly, the rifted margin may be characterized by swarms of diabase dykes, as exposed in eastern Greenland. Fourthly, the volcanism is associated with regional thermal uplift, which leads to the development of a regional domed unconformity in the stratigraphic record. From this summary it is clear that the Eastern Tethyan rifts differ markedly from either of the ideal ‘Non-volcanic’ or ‘Volcanic’ end members. In no case is there evidence of the exhumation of basement lower crust or upper mantle as in Non-volcanic margins (although preservation is always questionable). Also, rift volcanism is more extensive than on known Non-volcanic margins (although deep rift basins there remain largely unexplored). On the other hand, with the exception of the Himalayas, there is little signs of regionalscale plume-type magmatism and thermal doming. In general, volcanism is sparse along the Tethyan suture zone extending through Asia, as noted by S¸engo¨r et al. (1991). The Eastern Mediterranean rifts in particular appear to record a different setting to either the ideal ‘Non-volcanic-rifted’ or the ‘Volcanic-rifted’ types.

Pure shear versus simple shear models The evidence from drilling in the Iberia – Newfoundland conjugate shows that detachment faulting and exhumation play an important role in rifting, comparable to the Basin-and-Range of Western USA. However, drilling and geological observations of ancient margins cannot distinguish between these two alternative mechanisms and other techniques, mainly geophysical, must be used. Future drilling of the S reflector off Iberia, a possible simple shear detachment fault, is clearly a high priority for the IODP (Integrated Ocean Drilling Program). Some authors have argued that inherited crustal structure plays a key role in controlling the mechanism of rifting; it is suggested that pure shear dominates when rifting is at a high angle to the inherited tectonic fabric (i.e. zones of crustal weakness) but that simple shear results when the rifting is subparallel to the inherited grain (Ziegler & Cloetingh 2004). However, this assumes that the mechanism of rifting, by pure or simple shear, can be determined unambiguously, which is not the case in even the well-documented Iberia –Newfoundland conjugate. Recent modelling studies suggest that depth-uniform lithosphere stretching can replicate the structure and subsidence histories of intracontinental rifts but that depthdependent stretching plays an important role during final continental breakup, with the lower

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crust and mantle lithosphere being extended much more than the upper crust (see Kusznir & Healey 2006; Kuznir & Karner 2007). This situation is well documented, for example, for the east Brazil margin (Karner pers. comm. 2005). A related approach has been to use the stratigraphic record of rifted margins to calculate subsidence profiles and thus attempt to determine the timing and processes of rifting (e.g. Brunet & Cloetingh 2003). The calculated shapes of the subsidence curves differ for pure shear and simple shear rifted margins during the postrift phase, thereby allowing a theoretical distinction using the stratigraphic record (Watts 1982; Ziegler et al. 2003; Karner et al. 2007). The subsidence curves mainly utilize shallow-water successions deposited on the more proximal rifted margin segments that remained relatively undeformed. In contrast, the deep-water units formed within the continent– ocean transition zone are generally allochthonous, deformed or metamorphosed, limiting their value for subsidence analysis. Utilizing, this approach Wooler et al. (1995) calculated subsidence curves for a number of rifted margins of both the Western Tethys (e.g. Apennines) and the Eastern Tethys (from Greece to Oman). They concluded that there was little evidence that rifting and continental breakup were related to the activity of large-scale mantle plumes (‘active rifting’), especially since major domal unconformities appeared to be absent from the areas studied. This situation contrasts with, for example, the Jurassic of the North Sea, a prime candidate for rifting associated with a mantle plume (Ziegler 1990; Underhill & Partington 1993). Wooler et al. (1995) emphasized the important role of rifting during Triassic time. The calculated subsidence curves suggested that extension occurred even in some marginal platforms where rift-related faults are not obvious in the field (e.g. Triassic of the Dolomites). However, such rift faults may be obscured, for example by slope facies; also many (or even most) rift faults were later inverted and developed into thrusts, again obscuring pre-existing rift normal faults. Wooler et al. (1995) argued that a dominantly pure shear rift model could explain the subsidence curves of the Tethyan rift units they studied. Stampfli et al. (2001) carried out a comparable study, utilizing subsidence curves calculated from evidence of rift units in the Himalayas, Oman, Zagros and the Eastern Mediterranean regions. These authors argued that major rift-related subsidence began throughout the Eastern Tethyan region during Carboniferous time, with further widespread rifting during the Permian. They proposed that sea-floor spreading began during Late Permian time throughout the South Tethyan

region, including the Eastern Mediterranean (based on the presence of inflections in the calculated subsidence curves). They also argued that the subsidence curves obtained by them could be modeled using a simple shear rift model (e.g. for Oman). Subsidence curves provide valuable insights on rifting, especially the timing of initial rifting and the role of extension-controlled subsidence, including areas where rift faults are largely absent, not exposed or converted into emplacement thrusts. However, they have a number of limitations. (1) Subsidence curves may not provide a good guide to the timing of final continental breakup, especially where a breakup unconformity is either subdued (e.g. Newfoundland margin) or apparently absent (e.g. South American margin). (2) Subsidence curves are influenced by whether a depth-dependent or a depth-independent stretching model is adopted, which cannot be assumed a priori for any given setting (see Kuznir & Karner 2007). (3) Subsidence curves can only be properly modelled where the geometry of the complete conjugate margin pair is known, showing whether rifting was symmetrical or asymmetrical. Even in the modern ocean (e.g. Iberia –Newfoundland) this information remains incomplete. For the Alps, asymmetrical rifting is well established from the stratigraphy (see Manatschal et al. 2007). However, in the ancient Eastern Tethyan examples discussed here, only one margin is readily available for study (e.g. Oman). In summary, subsidence curves alone do not permit a convincing distinction between either pure shear or simple shear models, and additional constraints are needed, for example, from geophysical evidence.

Rift driving mechanisms Based on a combination of the global data base and modelling, the main controls of rifting have recently been considered as a combination of plate-boundary forces, frictional forces acting on the base of the lithosphere (as asthenosphere convects) and the tensional stress (deviatoric tension) that develops above upwelling asthenosphere (Ziegler & Cloetingh 2004). Modelling studies suggest that continental breakup and the onset of sea-floor spreading may relate to an upwelling, divergent flow field within asthenosphere and continental lithosphere, implying that upwelling flow may be localized beneath rifts and zones of continental breakup (see Kusznir & Healy 2006; Kuznir & Karner 2007). The relation, if any, of such upwelling flow to ‘mantle plumes’ is currently controversial. The main role of plumes may be to facilitate crustal weakening rather than as a primary rift/

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breakup driving mechanism (Ziegler & Cloetingh 2004). For the Eastern Tethyan rifts, more specifically, several driving mechanism can be briefly considered as follows. † Asthenosphere convection exerting forces on Gondwana lithosphere leading to rifting. This is likely to be a key long-term process, but the forces exerted may be insufficient to cause breakup and sea-floor spreading (Smith 1999). The convection cells are assumed to lie beneath the African continent, rather than specifically at the margins of Gondwana. † Plume-related magmatism (i.e. ‘active rifting’: White & McKenzie 1995). The South Tethyan margins were all volcanically active and, thus, could have been influenced by the activity of mantle plumes, especially for the Early Permian Panjal Traps, northern India. There is little evidence for a string of plumes rising along the Gondwana margin, which would be expected if this were the dominant rift process. On the other hand, strain weakening related to rift volcanism, whether on not plume related, is likely to have played a role in final continental breakup (Hopper & Buck 1996). Once rifts crust becomes magmatically active strain weakening may allow continental breakup to proceed (Hopper & Buck 1996). A model of melt-assisted rifting has recently gained support from seismic anisotropy studies using shear-wave splitting in the Northern Ethiopia rift (Kendall et al. 2005). † Slab-pull. This could relate to subduction beneath the Eurasian margin, or elsewhere. There is, indeed, evidence from the Eurasian margin (e.g. Pontides, Caucasus) of northward subduction, HP –LT metamorphism and accretion (e.g. in the central Pontides, northern Turkey) during the Middle–Late Triassic when the Anatolian microcontinent was rifting from North Africa and drifting towards Eurasia (Okay 2000; Robertson et al. 2004). However, it seems unlikely that slab-pull, as a plate boundary force, could by itself initiate continental rifting (Smith 1999; see below). Also, such slabpull would not propagate across an intervening spreading ridge, if present. † Ridge push. This is not considered to be a significant plate boundary force in rifting (Ziegler & Cloetingh 2004). Several other factors may have influenced the rift setting. The first is that the crust around Gondwana was already thinned and weakened by pulses of rifting and related magmatism, potentially dating from Carboniferous time, or even the time of pan-African continental accretion. Theoretical arguments suggest that crust within the continent–

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ocean transition zone is relatively weak compared to full-thickness continental crust or oceanic crust (Vink et al. 1984; Steckler & ten Brink 1986; Lavier & Steckler 1997). Secondly, a spreading axis may have propagated around the Gondwana margin. A spreading centre appears to have formed adjacent to Oman in the Late Permian and then migrated westwards along the continental borderland of Gondwana. The propagating spreading tip reached the Eastern Mediterranean region by Late Triassic –Early Jurassic time, opening the South Tethyan oceanic basins. However, at present, the timing of spreading initiation in the large area between Oman and the Eastern Mediterranean (i.e. Zagros suture) remains poorly constrained and thus it is not known if spreading (i.e. continental unzipping) was progressive or episodic. A migrating spreading centre can be compared with the westward-stepping migration of the Woodlark spreading centre during Late Miocene – Quaternary time (see Goodliffe & Taylor 2007). In this case, the spreading tip migrated through a broad rift zone, within crust that was previously deformed and thermally immature (Taylor et al. 1999a, b; Goodliffe & Taylor 2007). The spreading tip migrated westwards, episodically over several million years, a timescale that might not be currently detectable if applicable to continental breakup in the Eastern Tethyan region. Also, the spreading tip was unstable and jumped tens of kilometres or more northwards or southwards as it migrated westwards, thus, in effect, isolating continental fragments offset by transform faults. The opening of the South Tethyan ocean basins bordering Gondwana can also be compared with the opening of the Arctic Ocean. In this region, an elongate continental fragment, the Lomonosov Ridge, more than 1500 km long by less than 150 km wide (comparable to the scale to the Tauride or Pelagonian microcontinents), rifted from Siberia during the Paleocene, opening the Arctic Ocean along the ultra-slow-spreading Gakkel Ridge (Lawver et al. 1990; Kristoffersen 1990). This rifting took place as a spreading centre migrated from the North Atlantic, potentially exploiting a pre-existing zone of crustal weakness in northern Eurasia. Similarly, the Eastern Tethyan rifts are seen as exploiting pre-existing weakened continental crust, following earlier deformation and coeval rift volcanism.

Working model of Eastern Tethyan rifting The inferred timing of rifting is summarized in Figure 38. The Eastern Tethyan rifting differed fundamentally from the intracontinental breakup of the

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Fig. 38. Summary of the timing of rifting, continental breakup and sea-floor spreading inferred for areas of the Eastern Tethys discussed in this paper. Timescale: Gradstein et al. (2004).

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North Atlantic (e.g. Newfoundland–Iberia conjugate) or Western Tethys (i.e. Alps). There was a strong tendency for the pre-existing Gondwana margin to fragment into microcontinents, which contrasts with the North Atlantic –Western Tethyan (Alpine) region where rifting took place amidst a pre-exiting continent. The Eastern Tethyan breakup was characterized by the detachment of elongate crustal fragments from a continental margin adjacent to a pre-existing wide ocean (Palaeotethys; Fig. 39a, b). The main forces promoting rifting can be considered as, first, the regional-scale mantle convection beneath Gondwana and, secondly, slabpull beneath Eurasia, or elsewhere. These forces individually were capable of initiating rifting but probably not continental breakup. Breakup occurred, not randomly in time, but instead concentrated in the Late Permian in the east (Oman and Himalaya) and in the latest Triassic– earliest

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Jurassic in the Eastern Mediterranean area. When breakup finally occurred it was quite drastic, initiating several subparallel small ocean basins at more or less the same time in the Eastern Mediterranean region. Pulsed rifting affected the Northern Gondwana margin in the Eastern Mediterranean. Extension there was sufficient to promote adiabatic melting and extensive rift volcanism by the Early Triassic, probably facilitating final breakup by the process of strain weakening. The breakup in the Eastern Mediterranean region corresponded to a time of major tectonic activity along the opposing Eurasian margin, where subduction, accretion and arc volcanism were taking place. If a spreading ridge had been subducted during this time, this could have accentuated slab pull on the southern margin and provided the extra force to trigger final continental breakup.

Fig. 39. Summary of the tectonic settings of rifting in the Eastern Tethys from the Eastern Mediterranean to the Himalayas. (a) Spreading adjacent to Gondwana in the Late Permian; (b) the spreading axis propagates into the Eastern Mediterranean region during the Late Triassic–Early Jurassic, isolating several microcontinents within oceanic crust (Neotethys); and (c) setting of microcontinents rifted from Gondwana in the Balkan region (Greece; Albania) during Late Triassic– Early Jurassic time. See the text for explanation.

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Where present, plumes facilitated continental breakup, particularly by further weakening the crust. For the Himalayas, breakup (probably plume-assisted) took place in the Late Permian (Fig. 39a). Late Permian breakup also took place in Oman, but without evidence of a major local plume effect, although the chemistry of the volcanics is compatible with a plume influence. By the latest Triassic –Earliest Jurassic passive margin subsidence was underway in the Eastern Mediterranean region (Fig. 39b). The resulting assemblage of continental fragments and small ocean basins is shown in Figure 39a –c. Final breakup in the Western Mediterranean took place later, coeval with the Jurassic opening of the central North Atlantic. In the Himalayas, the rifted passive margin was later reactivated several times related to breakaway and the northward drift of India towards Eurasia. The Eastern Mediterranean rifted margins were also magmatically reactivated in some areas long after spreading had begun (i.e. in the Levant). The Eastern Tethyan rift first underwent deformation related to convergent-margin tectonics in the Middle–Late Jurassic in the Balkan region and the Late Cretaceous in the ‘Turkish’ region, Oman and the Himalayas.

Conclusions The rifted margins of the Eastern Tethys (Neotethys) are distributed throughout a number of suture zones, from the Eastern Alps to the Himalayas and beyond. These units provide a useful source of information on the anatomy of rifts and the processes of rifting and continental breakup. The uppermost levels of the slope–abyssal plain settings of rifted margins were typically detached during oceanic basin closure and emplaced onto continental margin as stacks of thrust sheets, overlain by ophiolites. Proximal rift/upper slope areas were commonly obscured by overriding thrust sheets, but are occasionally exposed, as in Oman. The leading edge of the rifted margin was typically underthrust or subducted (e.g. in NW Turkey and Saih Hatat, southern Oman Mountains). All of the Eastern Tethyan margins were associated with rift-related volcanism and are thus dissimilar to ideal Non-volcanic rifted margins, as presently known (e.g. Iberia –Newfoundland transect; Western Tethys in the Alps). However, with the possible exception of the Himalayas, they are also dissimilar to Volcanic-rifted margins (e.g. Eastern Greenland), especially as voluminous plume-related magmatism and corresponding

regional updoming are generally absent. The Eastern Mediterranean rifted margins, therefore, exemplify a diversity of rift-related settings, and show that the current classification into ideal ‘Volcanic-rifted’ and ‘Non-volcanic rifted’ end members is oversimplified. Rifting was typically pulsed, over up to 100 Ma (e.g. North African –Levant; Himalayas). The Himalayan margin was reactivated after spreading began, marked by faulting and extension-related alkaline magmatism, focused within the continent– ocean transition zone. The Eastern Mediterranean Levant margin was also magmatically reactivated, with extensive alkaline volcanism long after spreading began in the area. The production of extension-related melts as in the Eastern Mediterranean Triassic rifts is likely to have facilitated final continental breakup (controlled by strain weakening). The extension of a previously weakened continental crust bordering less deformable oceanic crust acted to favour the detachment of elongate crustal slivers, hundreds of kilometres long by tens of kilometres wide. These fragments drifted northwards towards Eurasia as microcontinents, opening up small ocean basins in their wake. The major control on this long-lived northward drift was northward subduction beneath Eurasia. On a smaller scale, rifting and continental breakup resulted in marginal fault-blocks (including prerift/synrift stratigraphy) being detached and isolated within marginal oceanic crust. These ‘highs’ were capped by shallow-water or pelagic facies controlled by the variable effects of subsidence versus eustatic sea-level change. Volcanic seamounts located further out in the oceans were capped by subsiding carbonate platforms (atolls). Final continental breakup appears to have been triggered by a combination of long-term asthenosphere convection and plate boundary forces, especially slab-pull related to subduction beneath Eurasia. Where present, mantle plumes facilitated continental breakup. Rifting processes and products are thus diverse in space and time, and much additional basic information is still needed from the modern oceans and orogenic belts to further test and constrain alternative tectonic models. I thank G. Karner and InterMARGINS for an invitation to participate in the Pontresina Workshop. I also thank G. Karner for discussions. A constructive review of the manuscripts was kindly provided by Prof. P. Ziegler. In addition, the comments of an anonymous reviewer were taken into account when producing the final manuscript.

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S WARBRICK , R. E. & R OBERTSON , A. H. F. 1980. Revised stratigraphy of the Mesozoic rocks of southern Cyprus. Geological Magazine, 117, 547–563. T AYLOR , B. P., G OODLIFFE , A. & M ARTINEZ , F. 1999a. How continents breakup: insights from Papua New Guinea. Journal of Geophysical Research, 104(B4), 7437–7512. T AYLOR , B., H UCHON , P. ET AL . 1999a. Proceedings of the Ocean Drilling Program, Initial Reports, 180. Ocean Drilling Program, College Station, TX. T UCHOLKE , B., S IBUET , J.-C. ET AL . 2004. Proceedings of the Ocean Drilling Program, Initial Reports, 210. Ocean Drilling Program, College Station, TX. T UCHOLKE , B. E., S AWYER , D. S. & S IBUET , J.-C. 2007. Breakup of the Newfoundland–Iberia rift. In: K ARNER , G. D., M ANATSCHAL , G. & P INHEIRO , L. M. (eds) Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 9–45. U NDERHILL , J. R. & P ARTINGTON , M. A. 1993. Jurassic thermal doming and deflation in the North Sea: implications of the sequence stratigraphic evidence. In: P ARKER , J. R. (ed.) Petroleum Geology of NW Europe: Proceedings of the 4th Conference. Geological Society, London, 337–345. U STAO¨ MER , T. & R OBERTSON , A. H. F. 1994. Late Palaeozoic marginal basin and subduction–accretion: the Paleotethyan Ku¨re Complex, Central Pontides, northern Turkey. Journal of the Geological Society, London, 151, 291–305. U STAO¨ MER , T. & R OBERTSON , A. H. F. 1997. Tectonic– sedimentary evolution of the north Tethyan margin in the Central Pontides of northern Turkey. In: R OBINSON , A. G. (ed.) Regional and Petroleum Geology of the Black Sea and Surrounding Region. AAPG Memoirs, 68, 255– 290. U STAO¨ MER , P. A., M UNDIL , R. & R ENNE , P. R. 2005. U/ Pb and Rb/Pb zircon ages for arc-related intrusions in the Bolu Massif (W Pontides, NW Turkey): evidence for Late Precambrian, Cadomian age. Terra Nova, 17, 215 –223. V ANNAY , R. & S PRING , L. 1993. Geochemistry of the continental basalts within the Tethyan Himalaya of Lahul– Spiti and SE Zanskar. In: T RELOAR , P. J. & S EARLE , M. P. (eds) Himalyan Tectonics. Geological Society, London, Special Publications, 74, 237 –249. V INK , G., M ORGAN , J. W. & L ING , Z. W. 1984. Preferential rifting of continents: a source of displaced terranes. Journal of Geophysical Research, 89, 10,072– 10,076. VON R AD , U., T HURROW , J., H AQ , B. U., G RADSTEIN , F., L UDDEN , L. & ODP L EGS 122/123 S HIPBOARD S CIENTIFIC , PARTIES . 1989. Triassic to Cenozoic evolution of the NW Australian continental margin and the birth of the Indian Ocean (preliminary results

of ODP Legs 122 and 123). Geologische Rundschau, 78, 1189– 1210. W ALDRON , J. W. F. 1984. Evolution of carbonate platforms on a margin of the Neotethys ocean: Isparta angle, south-western Turkey. Eclogae Geologicae Helvetiae, 77, 553–581. W ATTS , A. B. 1982. Tectonic subsidence, flexure and global changes of sea level. Nature, 297, 469– 474. W ATTS , K. F. & G ARRISON , R. E. 1986. Sumeini Group, Oman – evolution of a Mesozoic carbonate slope on a South Tethyan continental margin. Sedimentary Geology, 48, 107– 108. W HITE , R. & M C K ENZIE , D. P. 1989. Magmatism at rift zones: the generation of volcanic continental margins. Journal of the Geological Society, London, 149, 841–854. W HITE , R. S. & M C K ENZIE , D. 1995. Mantle plumes and flood basalts. Journal of Geophysical Research, 100, 17,543–17,586. W HITMARSH , R. B., M ANATSCHAL , G. & M INSHULL , T. A. 2001. Evolution of magma-poor continental margins: From final rifting to seafloor spreading. Nature, 413, 150– 154. W OODCOCK , N. H. & R OBERTSON , A. H. F. 1982. Imbricate thrust belt tectonics and sedimentation as a guide to emplacement of part of the Antalya Complex, S.W. Turkey. Journal of the Geological Society, London, 139, 147– 163. W OOLER , D. A., S MITH , A. G. & W HITE , N. 1995. Measuring lithospheric subsidence in Tethyan passive margins. Journal of the Geological Society, London, 149, 517– 537. Y ILMAZ , Y. 1993. New evidence and model on the evolution of the southeast Anatolian orogen. Geological Society of America Bulletin, 105, 251 –271. Z IEGLER , P. A. 1990. Geological Atlas of Western and Central Europe, 2nd edn. Shell International, The Hague. Z IEGLER , P. A. & C LOETINGH , S. 2004. Dynamic processes controlling evolution of rifted basins. EarthScience Reviews, 64, 1– 50. Z IEGLER , P. A., C LOETINGH , S., G UIRAUD , R. & S TAMPFLI , G. M. 2003. Peri-Tethyan platforms: constraints on dynamics of rifting and basin inversion. In: Z IEGLER , P., C AVAZZA , W., R OBERTSON , A. H. F. & C RASQUIN S OLEAU , S. (eds) Peri-Tethys Memoir 5. Peri-Tethyan Rift/Wrench Basins and Passive Margins. Memoirs du Museum National d’Histoire Naturelle, 9–51. Z ULAUF , G., K OWALCZYK , G., K RAHL , J. & S CHWANZ , S. 2002. The tectonometamorphic evolution of high-pressure low-temperature metamorphic rocks of eastern Crete, Greece: constraints from microfabrics, strain, illite crystallinity and paleodifferential stress. Journal of Structural Geology, 24, 1805– 1828.

Continental lithospheric thinning and breakup in response to upwelling divergent mantle flow: application to the Woodlark, Newfoundland and Iberia margins N. J. KUSZNIR1 & G. D. KARNER2,3 1

Department of Earth & Ocean Sciences, University of Liverpool, Liverpool L69 3BX, UK (e-mail: [email protected]) 2

Lamont-Doherty Earth Observatory, Palisades, NY 10964, USA

3

Present address: ExxonMobil Upstream Research Company, Mail Stop URC-URC-S169A, P.O. Box 2189, Houston, TX 77252-2189, USA

Abstract: Depth-uniform stretching is not the dominant deformation process for thinning continental lithosphere leading to breakup; it cannot explain the observed depth-dependent lithosphere stretching and mantle exhumation at rifted continental margins. Depth-dependent lithosphere thinning, in which stretching of the lower crust and lithosphere mantle greatly exceeds that of the upper crust, has been observed at many non-volcanic and volcanic rifted continental margins including conjugate margin pairs. Passive continental margins show a paucity of brittle deformation in the upper crust during continental lithosphere thinning leading to breakup and sea-floor spreading initiation. A new model of rifted continental margin formation has been developed that assumes that deformation and thinning of continental lithosphere leading to breakup occurs in response to an upwelling divergent flow field within continental lithosphere and asthenosphere, and that this deformation evolves into sea-floor spreading. The new model successfully predicts depth-dependent stretching of continental margin lithosphere for both non-volcanic and volcanic margins and mantle exhumation at non-volcanic margins, both of which are observed, but are not explained, by existing depth-uniform continental lithosphere stretching models. The new model provides a balance of extensional strain, supplies an explanation for the paucity of synrift brittle deformation, and offers a simple transition from prebreakup lithosphere thinning to sea-floor spreading. The observed diversity of rifted continental margin structure and width of the ocean–continent transition can be explained by variability in the form of the upwelling divergent flow field. The new upwelling divergent flow model of continental lithosphere thinning leading to continental breakup successfully predicts the observed bathymetry and margin geometry for the most recent segment of sea-floor spreading initiation in the Woodlark Basin in the western Pacific, and the observed bathymetry and free air gravity anomaly for the Newfoundland and Iberian margins.

Continental breakup and the formation of rifted continental margins are an important geodynamic process and a fundamental component of the plate tectonic cycle. The deep-water regions of rifted continental margins are the current exploration frontier for oil and gas, and present major scientific and technical challenges to the hydrocarbon industry. Key questions that need to be answered by academia relate to defining and testing the mechanisms responsible for thinning continental lithosphere prior to continental breakup and sea-floor spreading initiation. On the other hand, the key questions that need to be answered by the hydrocarbon industry when exploring deep-water rifted margin prospects are: (1) where is the location of the ocean –continent transition? (2) What are the structures, compositions and modes of deformation within the ocean –continent transition? Finally, (3) what are

the subsidence and heat-flow history of the ocean –continent transition and the zone of lithospheric thinning? There is a complete parallelism in the themes being asked and investigated by academia and the hydrocarbon industry. In the last decades, two important discoveries have been made at rifted continental margins: lithosphere depth-dependent thinning (Royden & Keen 1980; Driscoll & Karner 1996, 1998; Roberts et al. 1997; Davis & Kusznir 2004); and continental mantle exhumation (Boillot et al. 1987, 1989; Boillot & Winterer 1988; Pinheiro et al. 1992; Sawyer et al. 1994; Pickup et al. 1996; Manatschal & Bernoulli 1999; Whitmarsh et al. 2001). Existing models of continental breakup and sea-floor spreading initiation fail to predict depth-dependent lithosphere stretching and mantle exhumation at non-volcanic margins. Intracontinental rift basins, such as the Viking and

From: KARNER , G. D., MANATSCHAL , G. & PINHEIRO , L. M. (eds) Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 389– 419. DOI: 10.1144/SP282.16 0305-8719/07/$15.00 # The Geological Society of London 2007.

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Central grabens of the North Sea Basin, appear to be satisfactorily explained by depth-uniform lithosphere stretching (McKenzie 1978), where upper crustal stretching associated with faulting is balanced by an equal amount at the whole lithosphere scale measured from post-rift thermal subsidence. In contrast depthdependent lithosphere stretching, in which stretching and thinning of the whole continental crust and lithospheric mantle greatly exceeds that of the upper crust, is observed at both volcanic and non-volcanic rifted margins, and for rifted margin conjugate pairs (Driscoll & Karner 1996, 1998; Roberts et al. 1997; Karner & Driscoll 1999a; Davis & Kusznir 2004; Kusznir et al. 2004). The aim of this paper is to review geological and geophysical observations of depth-dependent lithosphere thinning and mantle exhumation, to present a new model of rifted margin formation, and then apply and test this new model to rifted margin case histories.

Depth-dependent lithosphere stretching and mantle exhumation at rifted continental margins A large number of studies have now demonstrated that extension and thinning of rifted continental

margin lithosphere is partitioned with depth. Depthdependent lithosphere stretching, in which extension and thinning of the whole crust and lithosphere of the rifted continental margin greatly exceeds that of the upper crust, has been observed at many rifted continental margins (Driscoll & Karner 1996, 1998; Roberts et al. 1997; Karner & Driscoll 1999a; Karner et al. 2003, Davis & Kusznir 2004). Extension and thinning of the upper crust may be measured using fault heaves from seismic reflection data, of the whole crust using crustal basement thinning derived from crustal structure utilizing wide angle seismology and gravity studies, and of the whole lithosphere from post-breakup subsidence using flexural backstripping of post-breakup stratigraphic data, or forward modelling of the preserved synrift and post-rift stratigraphic packages of basins and margins (Fig. 1). For example, profiles of the upper crustal b stretching factor and thinning factor (1 2 1/b) for the Goban Spur and South China Sea non-volcanic margins (Fig. 2) show a region of the order of 75–100 km width adjacent to the ocean– continent transition in which stretching of the upper continental crust is substantially less than that of the whole crust and lithosphere. Further towards the continent, stretching estimates for the upper crust, whole crust and lithosphere diminish and tend to converge. Davis & Kusznir (2004)

Measuring lithosphere thinning and extension as a function of depth at rifted continental margins

upper crust – from fault heave summation whole crust – from crustal basement thinning

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Fig. 1. Stretching and thinning of continental lithosphere at rifted margins can be measured for the upper crust from fault heaves, the whole crust from crustal thinning and the whole lithosphere from post-breakup thermal subsidence.

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Fig. 2. Examples of continental breakup depth-dependent lithosphere stretching for (a) Goban Spur and (b) South China Sea non-volcanic rifted margins. b stretching and thinning factors for the upper crust are significantly less than those of the whole crust and lithosphere within 75–150 km of the ocean –continent transition (Davis & Kusznir 2004).

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Fig. 3. Example of depth-dependent lithosphere stretching on a volcanic margin. (a) Upper crustal cross-sections for the northern Vøring and southern Lofoten segments of the Norwegian volcanic rifted continental margins. (b) Breakup thinning factors (1 2 1/b) plotted against distance from the continent –ocean boundary for the whole lithosphere and the upper crust (Kusznir et al. 2004). Stretching and thinning of the upper crust are significantly less than that of the whole lithosphere within 100 km of the ocean– continent transition.

described the evidence for depth-dependent lithosphere stretching at the Goban Spur and South China Sea margins in detail. Depth-dependent lithosphere stretching has also been reported for the South China Sea margins by Clift et al. (2002). Depth-dependent stretching is also observed at volcanic continental margins (Roberts et al. 1997; Davis & Kusznir 2004). Breakup and sea-floor spreading initiation on the Norwegian margin occurred at the end of the Paleocene at approximately 54 Ma (Tsikalas et al. 2001, 2005; Ren et al. 2003) following substantial earlier intra-plate continental rifting in the Triassic, Jurassic and

Cretaceous (Dore´ et al. 1999; D. G. Roberts et al. 1999; Brekke 2000). Profiles of thinning factor (1 2 1/b) for the upper crust and the lithosphere for the southern Lofoten and northern Vøring segments of Norwegian rifted margins are shown in Figure 3 and are adapted from Kusznir et al. (2004). Data analysis to produce these thinning factor profiles is described in detail in Kusznir et al. (2004). Profiles of thinning factors for Lofoten and Vøring margin segments show only a small magnitude of lithosphere extension by faulting (b , 1.1) in the Paleocene and Late Cretaceous immediately preceding breakup at approximately

A NEW MODEL OF LITHOSPHERIC THINNING

54 Ma. In contrast, large lithosphere b stretching factors (.2.5) are required at about 55 Ma to restore the top Lava (top inner lava flow) and top Paleocene reflectors to subaerial depositional environments. The discrepancy between stretching and thinning of the upper crust and whole lithosphere increases towards the continent–ocean transition. Continental lithosphere thinning leading to breakup and sea-floor spreading initiation does not appear to be accompanied by significant stretching and thinning of the upper crust. An extreme example of depth-dependent lithosphere stretching during the formation of a volcanic margin occurs on the Møre segment of the Norwegian rifted continental margin (Fig. 4). The Møre margin requires large lithosphere b stretching factors approaching 4 in order to restore palaeo-water depth indicators to sea level at breakup at 54 Ma, while fault extension in the Paleocene indicates upper crustal stretching that is negligible and less than 1.02. A detailed description and discussion of the analysis techniques used to measure continental margin stretching and thinning are given by Davis & Kusznir (2004). They also consider possible observational explanations for the lack of upper (a)

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crustal extension at rifted continental margins compared with that determined for the whole crust and lithosphere, including subseismic resolution faulting, second generation faulting, non-brittle (aseismic extension) and subaerial erosion. While significant faulting does occur beneath the resolution of seismic reflection imaging, fault scaling relationships suggest that the subseismic missing extension amounts to approximately 35% of the observed extension (Walsh et al. 1991) which is insufficient to explain the observed depthdependent lithosphere stretching. Davis & Kusznir (2004) also considered and discounted other possible explanations including dense underplating, phase transitions, subaerial erosion, mantle –plume interaction, flexural coupling with cooling oceanic lithosphere and lithosphere extension on low-angle detachments. Davis & Kusznir (2004) conclude that depth-dependent lithosphere stretching at rifted margins is a real phenomenon and is not an observational artifact. Xie et al. (2006) have suggested that subduction dynamic topography may contribute to anomalous post-breakup subsidence of the South China Sea rifted margins. However, dynamic topography cannot explain the observed magnitude of

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Distance (km) Fig. 4. Example of depth-dependent lithosphere stretching on a volcanic margin. (a) Upper crustal cross-section for the Møre segment of the Norwegian volcanic margin (Blystad et al. 1995). (b) Stretching factor b for the lithosphere and upper crust. Paleocene upper crustal stretching preceding Eocene–Paleocene boundary breakup on the Møre segment of the Norwegian volcanic margin is negligible compared with whole lithosphere thinning within 150 km of the ocean–continent transition.

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anomalous subsidence and therefore cannot provide an explanation for depth-dependent lithosphere stretching on the South China Sea margins. In addition, depth-dependent lithosphere stretching is observed on rifted margins where a contribution to anomalous post-breakup subsidence from dynamic topography cannot be invoked. Below we present additional evidence for depth-dependent lithosphere stretching on the NW Australian and Congo–Angola rifted continental margins.

Northwest Australian margin Rifting, by definition, deals with the brittle deformation of the upper crust leading to the generation of horst and graben morphologies. A number of key observations, such as normal faulting, divergence of seismic reflectors indicative of differential subsidence and block rotation, and sediment wedge geometries produced by onlap (as opposed to truncation), define both the onset and accommodation created by faulting. Yet, studies of the essentially non-volcanic Exmouth Plateau rifted margin (e.g. Driscoll & Karner 1998; Karner & Driscoll 1999b; Tischer 2006) have clearly shown that these stratigraphic relationships do not characterize the regionally extensive Tithonian– Hauterivian stratigraphic packages, represented by the Barrow Group (yellow package; Fig. 5a), which immediately precede Hauterivian breakup and sea-floor spreading initiation. Figure 5a is a representative seismic section across the southern Exmouth Plateau rifted margin (Tischer 2006) showing the existence of regionally distributed but minor Triassic and Jurassic aged rifting and regional subsidence across the plateau. Barrow Group sediments were deposited as part of a large NW progading deltaic system, the accommodation space being characterized by both a paucity of faulting and regional sagging (Driscoll & Karner 1998). The Barrow Group represents a synextensional depositional package, bracketed by a significant space-forming event in the early Tithonian and the breaching of continental lithosphere and emplacement of oceanic crust in the Gascoyne and Cuvier abyssal plains immediately to the west and SW of the plateau in the Hauterivian (M10; 133.5 – 134.3 Ma: Robb et al. 2005). In the absence of any significant faulting across the Exmouth Plateau of Tithonian–Hauterivian age, the fundamental question becomes: what mechanism was responsible for thinning the crust and lithospheric mantle across the Exmouth Plateau to generate permanent subsidence? Driscoll & Karner (1998) explained the development of deep-water conditions immediately prior to early Cretaceous breakup and the paucity of Tithonian–Hauterivian faulting across the

Exmouth Plateau as the consequence of depthdependent thinning in which only the lower continental crust and lithospheric mantle were involved in the thinning process. Recent refraction data (Fig. 5b) (Tischer 2006) across the Exmouth Plateau confirms that the lower continental crust (20– 23 km; P-wave velocities ranging from 6.5 to 7.5 km s21) is significantly thinner compared with the upper crust (0–20 km; P-wave velocities ranging from 4.5 to 6.1 km s21). In this case, the upper crust consists principally of an extremely thick Palaeozoic sedimentary section, presumably of lower– upper greenschist metamorphic facies in the deeper section because of its significant thickness (10 –15 km: Stagg & Colwell 1994). The ocean –continent transition zone has a width of about 200 km and is flanked on the ocean side by normal oceanic crust approximately 8 km in thickness. If depth-dependent thinning is considered to be a consequence of differential extension, then the abrupt transition in velocities between 250 and 300 km (Fig. 5b) may represent the zone of laterally extended and thinned lower crust and continental mantle pulled out from under the adjacent plateau and emplaced as part of the sea floor (Tischer 2006).

Congo –Angola margin Continental lithosphere breakup and the initiation of sea-floor spreading on the non-volcanic Congo–Angolan margin occurred in the midAptian (c. 115 Ma: Karner & Gamboˆa 2007). Congo and northern Angola seismic reflection sections indicate that the regional subsidence across the West African margin consists of two components: (1) relatively minor Berriasian–early Barremian-agedfault-controlledsubsidence(Fig.6a) (Karner et al. 1997); and (2) significant late Barremian– mid-Aptian regional subsidence, both landward and seaward of the Atlantic hinge zone, culminating in the regional deposition of evaporites (Fig. 6b). The fact that the late Barremian –Aptian prebreakup depositional packages appear to be virtually independent of faulting (or, more appropriately, faulting that controls accommodation space generation) is a proxy for depth-dependent thinning (e.g. Karner et al. 2003). Further, seismic refraction data across the northern Angolan margin (Contrucci et al. 2004) image a continental crust that thins rapidly from 30–34 km in the east to less than 10 km over a distance of 100 km seaward of the Atlantic hinge zone in the same region where seismic reflection data indicate a paucity of faulting (Fig. 6b) (Contrucci et al. 2004; Moulin et al. 2005). As with the Exmouth Plateau, if the crustal thinning is an extension process, then there needs to be a lateral balance of extension. This implies that the continental foundation underlying the pre-salt sag basin may in part be unthinned upper crust in

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Fig. 5. (a) Interpreted multichannel seismic lines GA110-12 and EW0113-2 ( Tischer 2006) across the southern Exmouth Plateau rifted margin showing the major structural elements. Tr, top Triassic; JC, Callovian; JT, Tithonian; KV, Valanginian; KA, Aptian; KS, Santonian; bT, base Tertiary; Pl, early Pliocene. Brittle deformation characterized the late Triassic, and Callovian, locally important in the Exmouth sub-basin. Sag subsidence characterizes the Tithonian–Hauterivian prebreakup subsidence and the Aptian– Present post-breakup subsidence. (b) The derived P-wave velocity model show the stretched continental crust of the western Exmouth Plateau and part of the continent –ocean transition zone (COTZ). The crustal model (thin blue lines) suggests a gradual thinning of the continental crust oceanward. The velocity distribution (thin black lines, contour interval: 0.2 km s21) indicates a distinct layering of the stretched continental crust underlying the plateau and a distinct preferential thinning of the lower continental crust. Layer colours and numbers correspond to a particular P-wave velocity interval. A prominent lateral velocity gradient marks the boundary between continental crust and the COTZ. OBS positions are indicated by blue triangles. Also displayed is the free-air gravity and magnetic anomaly data collected across the margin during R/V Ewing cruise EW0113 (Robb et al. 2005). Tischer (2006) has postulated that magnetic anomalies M4-M10 were created during the exhumation of the lower crust and continental mantle lithosphere rather than by sea-floor spreading.

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Fig. 6. Seismic reflection profiles across the Eastern and Atlantic hinge zones on the Congo margin, and the northern Angolan margin (West Africa) showing minor fault-control on the development of prebreakup accommodation. (a) Early Neocomian rotated and divergent reflectors indicate differential subsidence and block rotation to allow deposition of the earliest synrift depositional packages across the Congo margin. Djeno (Top Djeno: blue line) and Marnes Noires formations sediments have prograded into the basin. The Argilles Vertes Formation also shows evidence of downlap onto the Marnes Noires (Top Marnes Noires: green line). Top and base salt are shown in red. Together, the Djeno, Marnes Noires and the evaporites comprise a prebreakup sag section. Bold vertical lines locate the well control. Seismic data courtesy of HydroCongo. (b) East– west seismic reflection profile located seaward of the Atlantic hinge zone on the northern Angolan margin showing the thick depositional packages of the pre-salt sag basin, regionally developed seaward of the Atlantic hinge of the West African margin, onlapping the hinge zone, and regional but gentle erosional truncation (red arrows). Base salt– top Cuvo is shown by the yellow line. Salt rafts comprise detached and rotated blocks of Albian shallow marine carbonates and sandstones. Top basement is shown by the brown line. Total width of the pre-salt sag basin width is approximately 300 km with a maximum thickness of 7 km and an average of 3 km. The regional accommodation allowing deposition of the pre-salt basin and the overlying evaporites is essentially independent of normal faulting. Modified from Fraser et al. (2005).

A NEW MODEL OF LITHOSPHERIC THINNING

the east and a zone of exposed continental mantle in the west.

Summary The total extension of the continental rifted margin may be determined by horizontally integrating the thinning factor (1 2 1/b) across the continental margin. Total extension is summarized in Figure 7 for a number of rifted continental margins (Kusznir et al. 2005). In all cases, the estimated stretching of the whole crust and lithosphere exceeds that of the upper crust, and the amount of brittle deformation observed is significantly smaller compared with the amplitude and distribution of post-breakup subsidence and crustal thinning. Observed extensional normal faults do not control the regional development of synrift (prebreakup) accommodation, and are therefore not important in the extension process. During continental breakup, continental crust and lithospheric mantle are thinned to the point of being breached, leading to large post-breakup subsidence, but the negligible brittle deformation implies that the upper crust is not involved in the breakup lithosphere thinning process. The depth-dependency of lithosphere deformation associated with the formation of intracontinental rift basins and continental rifted margins may be compared by cross-plotting lithosphere

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thinning factors (1 2 1/b), determined from flexural backstripping of post-rift or post-breakup subsidence and thinning factors for the upper crust, based on fault-heave summation (Fig. 8). For intracontinental rift basins, within observational errors, stretching and thinning of the continental lithosphere is equal to that of the upper crust indicating depth-uniform lithosphere stretching consistent with the McKenzie (1978) model of continental rift basin formation. During intracontinental rifting, lithosphere extension and thinning (Ebinger 1989; Marsden et al. 1990; Lambiase & Bosworth 1995; Withjack et al. 1995; Schlische & Anders 1996; Morley 1999) is characterized by faulting and brittle deformation of the crust balanced by an equal magnitude of distributed plastic deformation within the lithospheric mantle. In contrast, measured lithosphere and upper crustal stretching and thinning factors for rifted continental margins plot in a distinctly different region of the cross-plot where these values are not equal. Thus, the processes involved in lithosphere deformation leading to continental breakup and the initiation of sea-floor spreading appear to be different to those involved in intracontinental rifting. The depth-dependent (or depth-partitioned) deformation of continental margin lithosphere during breakup is not consistent with present models of lithosphere breakup deformation based on pure shear (McKenzie 1978; Royden & Keen 1980). Nor is it consistent

Fig. 7. Comparison of integrated upper crustal, whole crustal and lithosphere extension for the Goban Spur, South China Sea, Exmouth Plateau, Vøring and South Vulcan– Bonaparte rifted margins. Both non-volcanic and volcanic rifted margins show depth-dependent lithosphere stretching and thinning (Davis & Kusznir 2004).

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Geochemistry suggests that in some cases the exhumed mantle has a continental signature implying that it is exhumed continental lithospheric mantle pulled out from under the continent during breakup (Mu¨ntener & Hermann 2001).

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A new model of continental breakup, sea-floor spreading initiation and rifted continental margin formation

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with simple shear extension (Wernicke 1985; Lister et al. 1986); depth-dependent lithosphere stretching is observed on conjugate margins and does not support simple asymmetric lithosphere extensional detachments models of continental breakup (Driscoll & Karner 1998; Davis & Kusznir 2004). Mantle exhumation at rifted continental margins is also not explained by existing models of continental breakup lithosphere thinning. At non-volcanic margins, broad regions of exhumed mantle of the order of 75 –100 km wide have been observed (Fig. 9) separating oceanic crust and thinned continental crust (Boillot et al. 1987, 1989; Boillot & Winterer 1988; Pinheiro et al. 1992; Sawyer et al. 1994; Pickup et al. 1996; Manatschal & Nievergelt 1997; Manatschal & Bernoulli 1999; Whitmarsh et al. 2001; Lavier & Manatschal 2006). Peridotite ridge

What is the mechanism by which lithosphere deformation occurs and is distributed with depth during continental breakup and the formation of continental rifted margins? Sea-floor spreading is characterized by upwelling and divergence of oceanic lithosphere and asthenosphere, and not by pure shear extension and thinning of young oceanic lithosphere. At the time of continent breakup, continental lithospheric thinning is replaced by sea-floor spreading and we assume that there is continuity between the breakup deformation processes and those involved in seafloor spreading. We propose that continental lithosphere thinning leading to breakup is controlled by upwelling and divergence of continental lithosphere and asthenosphere rather than by depth-uniform (pure shear) stretching and thinning. We explore the implications of a fluid-flow model for deformation of continental lithosphere in which upwelling and divergence within continental lithosphere and asthenosphere are parameterized by a vertical upwelling velocity and a horizontal divergence velocity. Such a model explains naturally the depth-dependent extension of the lithosphere observed at many rifted continental margins and builds on earlier models by Kusznir et al. (2005). By default, the new model captures the transition from continental lithosphere thinning and breakup to sea-floor spreading and relates the architectural diversity of observed continental margins to a small number of key parameters. The new model aims to capture the large lithosphere-scale deformation processes responsible for thinning

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continental lithosphere leading to breakup and seafloor spreading initiation; it does not include brittle deformation of the upper crust which, while locally significant, plays only a negligible role in lithosphere deformation leading to breakup. The new model of continental lithosphere deformation leading to breakup and rifted margin formation has been developed using an adaptation of divergent upwelling mantle flow models that have been successfully applied to ocean ridges (PhippsMorgan 1987; Spiegelman & McKenzie 1987; Buck 1991; Spiegelman & Reynolds 1999). Our strategy for modelling rifted continental margin formation is to model the initiation of sea-floor spreading. The pattern of upwelling and divergent fluid flow within continental lithosphere and asthenosphere for our new model of continental breakup deformation is shown in Figure 10a. The flow field is defined kinematically by both a divergent half-rate velocity, Vx, and the upwelling velocity, Vz, and is calculated using an isoviscous streamfunction solution (Bachelor 1967): C ¼ ðAx þ BzÞ þ ðCx þ DzÞ tan1

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The stretching and thinning of continental lithosphere in response to the upwelling divergent flow field is shown in Figure 11. Streamlines (Fig. 11a) have been computed using Vx ¼ 1 cm year21 and Vz ¼ 1 cm year21. The initial geometry of the unstretched continental lithosphere (Fig. 11b) assumes an initial lithosphere thickness of 125 km and crustal thickness of 40 km. The configuration of the lithosphere at time 5 Ma after the initiation of the upwelling divergent flow field is shown in Figure 11c by which time the continental crust has been ruptured exposing continental lithospheric mantle at the surface. By 12.5 Ma after flow initiation (Fig. 11d), the continental lithosphere has also been ruptured bringing asthenosphere to the surface and sea-floor spreading has initiated. Rifted margin lithosphere geometry and temperature within lithosphere and asthenosphere are shown in Figure 11e, f at 25 Ma after a further 12.5 Ma of sea-floor spreading. The new model predicts symmetric depthdependent stretching of rifted margin lithosphere (Fig. 11) from continental lithosphere rupture and sea-floor spreading initiation, and gives ‘upper plate’ behaviour on both conjugate rifted margins, providing an explanation for the ‘Upper Plate Paradox’ as identified by Driscoll & Karner (1998). For Vx ¼ Vz ¼ 1 cm year21, as shown in Figure 11, stretching and thinning of the continental margin lithosphere results in exhumation of the lithospheric mantle oceanwards of the thinned continental crust. This prediction is consistent with the observation of exhumed continental lithospheric mantle on the Iberian non-volcanic margin (e.g. Pickup et al. 1996; Whitmarsh et al. 2001) and in Alpine exposure of the Tethyan margin (Manatschal et al. 2001). Rifted margin subsidence and bathymetry are calculated from the isostatic response to the thinning of the continental crust and the lithosphere temperatures field predicted by the upwelling divergent flow model. Rifted margin lithosphere structure with superimposed bathymetry predicted by the new model is shown in Figure 12a. The model predicts bathymetry at the ocean ridge of approximately 2.5 km deepening towards the ocean – continent transition before shallowing onto the continental margin. Thinning factors for the continental upper crust, whole crust and lithosphere are shown in Figure 12b. The development of bathymetry during prebreakup thinning of the continental lithosphere is shown in Figure 13a –c. The modelling results shown in Figures 11 & 12a, b have been computed using an upwelling divergent flow-field that reaches the lithosphere surface from time zero (Fig. 10a). A more realistic flowfield pattern may be one in which the flow field propagates from base lithosphere to the surface over several millions of years before rupturing the

Fig. 10. A new model of continental breakup, sea-floor spreading initiation and rifted margin formation proposes that continental lithosphere thinning leading to breakup and sea-floor spreading initiation occurs in response to an upwelling divergent flow field within continental lithosphere and asthenosphere. The model assumes that the dominant process controlling breakup lithosphere thinning is upwelling divergent flow rather than depth-uniform stretching and thinning. The upwelling divergent flow field may be defined kinematically by the half divergence rate Vx and upwelling rate Vz (a) A simple model where a fixed pattern of upwelling divergent flow operates to the lithosphere surface at all times with finite Vx and Vz. (b) A more complex model where the pattern of upwelling divergent flow propagates upwards from the base of the lithosphere during continental lithosphere thinning with finite Vz but Vx ¼ 0 at the lithosphere surface prior to breakup followed by finite Vx during sea-floor spreading.

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Fig. 11. The new margin formation uses an analytical stream-function solution to predict the fluid-flow field that is used to advect continental lithosphere material and lithosphere and asthenosphere temperature, leading to continental lithosphere thinning, breakup, and sea-floor spreading initiation. (a) Streamlines for upwelling divergent flow corresponding to Vx ¼ 1 cm year21 and Vx ¼ 1 cm year21 (for model corresponding to Fig. 10a). (b) Initial uniform thickness continental lithosphere and asthenosphere. (c) Predicted deformation and thinning of the continental lithosphere at 5 Ma after flow-field initiation showing rupture of the continental crust; (d) at 12.5 Ma showing rupture and breakup of the continental lithosphere; and (e) at 25 Ma showing sea-floor spreading. (f ) Oceanic and rifted continental margin lithosphere temperature structure at 25 Ma. (For this and all other numerical models described in this paper, Ma denotes elapsed time running in the forward sense.)

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continental lithosphere and initiating sea-floor spreading (Fig. 10b). During the upwards-propagation phase prior to sea-floor spreading initiation we propose that the upwelling divergent flow field takes a form where Vx ¼ 0 at the lithosphere surface consistent with observations of little stretching of the upper continental crust at rifted margins. Only when the upwelling flow field reaches the surface, the continental lithosphere is ruptured and sea-floor spreading commences does Vx become finite. Continental lithosphere margin geometry, temperature structure, isostatic response and thinning factors predicted by the model with an upwards propagation of the upwelling divergent flow field are shown in Figures 12c, d & 14a. In this model the flow field propagates from base lithosphere to the surface in 5 Ma with Vz ¼ 2.5 cm year21 and surface Vx ¼ 0 cm year21 followed by 5 Ma of sea-floor spreading with Vz ¼ 1 cm year21 and

Vx ¼ 1 cm year21. Upwards propagation of the upwelling divergent flow field produces a narrower ocean–continent transition boundary with much less exhumation of the continental lithospheric mantle compared to the model that operates with the flow field to the lithosphere surface at all times. The development of continental lithosphere thinning and bathymetry predicted by the upwelling divergent flow model is compared in Figure 13 for the flowfield pattern that operates to the surface of the lithosphere at all times (Fig. 10a) and for the flow-field pattern that propagates upwards from base lithosphere to surface (Fig. 10b). The model with flow field operating to the lithosphere surface at all times ruptures the continental crust before the continental lithospheric mantle (Fig. 13a–c). In contrast, the upwards-propagating flow-field pattern ruptures the continental lithospheric mantle before the continental crust (Fig. 13d–f).

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The thinning of rifted margin lithosphere and the resulting lithosphere geometry in response to the deformation of the upwelling divergent flow are dependent on the velocity ratio Vz/Vx during early sea-floor spreading. Analytical stream-function

fluid-flow calculations predict that passive sea-floor spreading, assuming iso-viscous lithosphere and asthenosphere, has a velocity ratio of the order of 2/p (Phipps Morgan 1987). However, finite-element fluid-flow calculations using temperature-dependent

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rheology predict passive sea-floor spreading with Vz/Vx  1 (Tymms 2006). Sea-floor spreading at non-volcanic margins mantle is therefore expected to have Vz/Vx of the order of 1. In contrast, mantle convection models of sea-floor spreading initiation (Neilsen & Hopper 2002) suggest that for sea-floor spreading initiation aided by upwards buoyancydriven flow from hot asthenosphere and lithosphere, Vz/Vx is initially greater than 10 for 2– 5 Ma and then decreases to 1 or less. The dependency of rifted margin lithosphere geometry and the temperature-field perturbation within lithosphere and asthenosphere on the early sea-floor spreading velocity ratio Vz/Vx are shown in Figure 14. Figure 14a, b compare the response to a velocity ratio Vz/Vx ¼ 1 and Vz/Vx ¼ 5 during early sea-floor spreading, respectively. For high Vz/Vx during early sea-floor spreading, corresponding to volcanic margin formation, a sharp ocean–continent transition is predicted with no exhumation of continental lithospheric mantle. Corresponding lithosphere cross-sections incorporating subsidence calculated using local isostasy in response to crustal thinning and lithosphere and asthenosphere temperature perturbations are compared in Figure 15. Volcanic and non-volcanic margins are predicted to have fundamentally different margin lithosphere geometries and ocean –continent transitions in addition to the presence or relative absence of volcanic addition. The observed diversity in continental margin structure and ocean–continent transition width can be partly explained by variations in the form of the upwelling divergent flow field leading to continental lithosphere rupture and the velocity ratio Vz/Vx during early sea-floor spreading. The model results shown in Figures 11–15 successfully predict depth-dependent stretching of continental margin lithosphere and, in some cases, mantle exhumation, while maintaining a balance of extensional strain. In the case of the model results shown in Figures 14 & 15, this strain balance is achieved by redistributing lower crust and continental mantle away from the zone of active extension into the hinterland region. In the absence of any contribution to continental lithosphere thinning from depth-uniform pure shear extension, this results in a small amount of thickening of hinterland continental crust and lithosphere (e.g. Figs 14b & 15b). The model applications described above assume that lithosphere thinning leading to continental lithosphere rupture and sea-floor spreading initiation occurs purely in response to an upwelling divergent flow field within continental lithosphere and asthenosphere. Depth-uniform (pure shear) lithosphere stretching and thinning has been ignored. In Figure 16 lithosphere and temperature structure for

rifted continental margin and young ocean basin are compared for models that assume continental lithosphere rupture by depth-uniform pure shear (b ¼ 1) alone (Fig. 16a), continental lithosphere rupture by upwelling divergent flow alone (Fig. 16b), and combined depth-uniform pure shear (b ¼ 2) and upwelling divergent flow. Model predictions show contrasting lithosphere margin geometries and temperature structures. While the observation of depth-dependent lithosphere stretching at rifted margins does not support lithosphere rupture by depth-uniform pure shear alone, a finite contribution to continental lithosphere thinning and breakup by pure shear deformation cannot be ruled out.

Application of the new model of continental breakup and sea-floor spreading initiation to rifted margin case histories The isostatic response to crustal thinning and temperature perturbations within lithosphere and asthenosphere, predicted by the new margin formation model, has been used to predict rifted margin bathymetry that has been compared with observation. Crustal structure, lithosphere temperature and bathymetry derived from the model, together with seismically observed sediment thickness data, have also been used to calculate the free air gravity anomaly. Consequently, the new model may be used to invert observed bathymetry and gravity anomaly data for rifted margins in order to predict continental lithosphere deformation kinematics. The upwelling divergent flow model of continental lithosphere thinning leading to continental breakup has been applied to the formation of the Woodlark Basin in the western Pacific, and to the Newfoundland and Iberian margins. Application of the new model using inverse theory to automatically invert observed bathymetry and gravity data has also been developed and successfully applied to the Goban Spur margin (Healy & Kusznir 2004, 2007).

Application to the Woodlark Basin A continuum of active extensional processes, laterally varying from continental/orogenic rifting to sea-floor spreading, in the western Woodlark Basin –Papuan Peninsula region of Papua New Guinea provides an ideal opportunity to test the new upwelling divergent flow model of continental breakup and sea-floor spreading initiation. A regional unconformity at 8.4 Ma marks the onset of rifting and provides a palaeo-sea-level

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surface for tracking the subsequent subsidence of the northern margin sites from paralic to shelf to the present bathyal water depths (Taylor & Huchon 2002). Seafloor spreading magnetic anomalies indicate that during the last 6 Ma, the formerly contiguous, eastward extensions of the Papuan Peninsula (the Woodlark and Pocklington rises) were separated as a westwards propagating spreading centre opened the Woodlark Basin (Fig. 17a) (Taylor & Huchon 2002). The present spreading tip is at 9.88 S, 151.78 E. Farther west, extension is accommodated by rifting of orogenic forearc/back-arc crust, with associated full and half-graben, metamorphic core complexes, and peralkaline rhyolitic volcanism (Taylor & Huchon 2002). The propagating tip is immediately east of the Moresby seamount region at the present day. Sea-floor spreading rates decrease to the west, with the western Woodlark Basin having faster-spreading characteristics than the eastern basin. A striking feature of the Woodlark Basin is

the obliquity of the present spreading axis relative to older sea-floor fabric and the Brunhes – Matuyama (0.78 Ma) crustal boundary. During the latter part of the Brunhes chron, the entire length of the Woodlark Basin spreading system synchronously reoriented. To the west of the Moresby seamount, ahead of the propagating spreading tip, the extending orogenic crust is characterized by up to 3 km of sag-basin subsidence over a width of 150–200 km, but with little attendant upper crustal faulting, implying depth-dependent lithosphere stretching. Significant, but local, brittle deformation is associated with the late-stage (,0.5 Ma) Moresby fault system. Stacked (and reflected) north– south bathymetric profiles showing the transition from the ocean-ridge axis into the zone of unstretched continental crust for the most recently formed sea-floor spreading segment are shown in Figure 17b. The new upwelling divergent flow model of continental breakup and sea-floor spreading initiation successfully predicts the observed bathymetric profile of young ocean floor and rifted continental margin for the most recently initiated segment of Woodlark Basin sea-floor spreading to the east of Moresby seamount (Fig. 17c). The best-fit bathymetry model requires upwards propagation of the upwelling divergent flow field from base lithosphere to surface for 7 Ma, followed by upwelling divergent flow to surface to 8 Ma, with Vx ¼ 2 cm year21 and Vz ¼ 2 cm year21. Models in which the flow field instantaneously penetrates the lithosphere to its surface, or with Vz/Vx  1, fail to fit the observed bathymetry. The fit of observed and predicted bathymetry is improved by the inclusion of 15 km of depth-uniform pure shear lithosphere stretching and thinning, which is consistent with observed extensional faulting in the newly formed continental margin. Larger amounts of depth-uniform stretching degrade the fit of predicted bathymetry to that observed. Predicted lithosphere structure and crustal thinning for the youngest segment of Woodlark Basin sea-floor spreading initiation is shown in Figure 17d. The predicted profile of thinning factor (1 2 1/b) for upper crust, whole crust and whole lithosphere (Fig. 17e) shows depth-dependent lithosphere stretching, a requirement consistent with subsidence analyses from other rifted continental margins.

Application to the Newfoundland margin The upwelling divergent flow model of continental lithosphere thinning leading to continental breakup has been applied to the formation of the Newfoundland margin. Bathymetry and free air gravity for the Iberia margin are shown in Figure 18a, b. The Newfoundland margin is a non-volcanic margin and underwent multiple rifting beginning in the Triassic

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and Jurassic, and culminating in the late Early Cretaceous (late Valanginian–late Aptian). Oldest oceanic isochrons adjacent to the margin are estimated to have an age of approximately 120 Ma (Mu¨ller et al. 1993), although ocean crust generation is now considered a little younger based on ODP Leg 210 drilling results (Tucholke et al. 2004). Crustal structure for the SCREECH 1 (Studies of Continental Rifting and Extension on the Eastern Canadian sHelf) transect (Fig. 19) (Funck et al. 2003; Hopper et al. 2004) across the northern Newfoundland margin shows a rapid thinning of the continental crust as the ocean–continent transition is approached from the continent direction. A region of very thin crustal basement with a thickness of less than 4 km and width of approximately 75 km separates thicker continental and unequivocal oceanic crust. The composition of this very thin crust is unknown; however, on the evidence of seismic velocities and ODP drilling to the south on the SCREECH 2 profile it is most likely to be thin oceanic crust and exhumed serpentinized mantle. The location of the model application is

superimposed on the bathymetry (Fig. 18a) and corresponds to the SCREECH 1 profile. The new model of rifted margin formation has been used to predict bathymetry and free air gravity for the SCREECH 1 profile. Both predicted bathymetry and free air gravity anomaly incorporate the isostatic response to sediment loading. Sediment thicknesses for the profile have been derived from seismic reflection data (Funck et al. 2003; Hopper et al. 2004). The best fit of present-day and observed and predicted bathymetry and gravity are shown in Figure 18c, d. Predicted rifted margin lithosphere geometry incorporating the isostatic response to crustal thinning, lithosphere and asthenosphere temperature perturbation and sediment loading is shown in Figure 18e. The best-fit model uses an upwelling divergent flowfield pattern that operates to the surface from time zero, a velocity ratio Vz/Vx ¼ 1.5 during continental lithosphere breakup thinning and early sea-floor spreading, and a syn breakup lithosphere b factor of 1. Post-breakup sea-floor spreading has been run for 120 Ma with Vx ¼ 1 cm and Vz/Vx ¼ 1.

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Fig. 17. (a) The Woodlark Basin in the Western Pacific formed by westwards-propagating sea-floor spreading during the last 8 Ma. (b) Stacked reflected bathymetric profiles (A– A0 ) across the most recently propagated segment of sea-floor spreading, which initiated at the beginning of the Brunhes epoch. (c) Bathymetry predicted by the upwelling divergent flow model of continental breakup and sea-floor spreading initiation compared with the stacked observed bathymetry for the Brunhes propagated segment. Predicted lithosphere cross-section (d) and thinning-factor profile (e) from the Woodlark ocean ridge axis into the rifted continental margin.

An oceanic crustal thickness of 7 km has been used in the calculation of the isostatic response to predict bathymetry and free air gravity. The best-fit model successfully predicts the abrupt thinning of the continental crust, and the amplitude and wavelength of the free air gravity anomaly dipole.

Application to the southern Galicia Bank segment of the Iberian margin The new model of rifted margin formation has been applied to the Iberian margin along a profile cutting across the margin on the southern edge of Galicia Bank coincident with a crustal cross-section

A NEW MODEL OF LITHOSPHERIC THINNING

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Fig. 18. Application of the upwelling divergent flow model of continental lithosphere breakup and sea-floor spreading initiation to the northern Newfoundland margin. (a) Bathymetry (IOC, IHO & BODC, 2003). (b) Free air gravity anomaly (Sandwell & Smith 1997). (c) & (d) Comparison of observed and modelled bathymetry and gravity anomaly. (e) Margin lithosphere structure predicted by the model.

Fig. 19. Crustal structure for the northern Newfoundland non-volcanic rifted continental margin determined by reflection and refraction seismology (Funck et al. 2003). The ocean–continent transition consists of a region of very thin crust or exhumed mantle separating continental and oceanic crust.

Fig. 20. Crustal structure for (a) the northern Iberian non-volcanic rifted continental margin (Zelt et al. 2003) and (b) the northern Angolan non-volcanic rifted continental margin (Moulin et al. 2005) determined by reflection and refraction seismology. Both cross-sections show a broad region (100 –200 km wide) of thinned continental crust (thickness ,12.5 km) separating thick landward continental crust from the ocean– continent transition (exhumed mantle?) and oceanic crust. The broad region of thinned crust produces a stepped bathymetry to these margins.

410 N. J. KUSZNIR & G. D. KARNER

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derived from seismic refraction and reflection derived by Zelt et al. (2003) and shown in Figure 20a. Bathymetry and free air gravity for the Iberia margin are shown in Figure 22a & b together with superimposed line location. The Iberia margin, which is conjugate to the Newfoundland margin, formed in the late Early Cretaceous. (a) 20

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Fig. 21. Pausing the upwards propagation of the upwelling divergent flow field during lithosphere thinning produces stepped crustal thinning and bathymetry on the rifted continental margin. (a) Upwards divergent flow field propagates directly from lithosphere base to surface. (b) Flow-field upwards propagation pauses at the mid-crustal level for 5 Ma before continuing to the surface. (c) Flow-field upwards propagation pauses at mid-crustal level for 10 Ma before continuing to the surface.

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Breakup rifting commenced in the late Valanginian (c. 138 Ma) with ocean crust production commencing at approximately the Aptian –Albian boundary (112 Ma; Tucholke et al. 2007). The southern Galicia Bank segment of the Iberia margin shows a broad region of thinned continental crust, approximately 200 km wide, between unthinned continental crust under mainland Iberia and the ocean – continent transition to the west. This broad region of thinned continental crust has a crustal basement thickness of the order of 15 km thickness. The outer Galicia Bank is separated from the Iberia mainland by the Galicia Interior Basin, which locally has a crustal basement thickness of approximately 10 km. To the south of the profile shown in Figure 20a there is a broad domain of exhumed mantle at the ocean –continent transition separating thinned continental crust and ocean crust (Pickup et al. 1996). This region of exhumed mantle may be present on the modelled profile, but is much narrower with a width of the order of 25 km. The stepped thinning of continental crust between unthinned mainland continental crust under mainland Iberia and the ocean –continent transition generates a stepped bathymetry that is typical of many non-volcanic margins. Similar stepped crustal thinning and bathymetry is shown in Figure 20b for the northern Angola profile (Moulin et al. 2005). What process leads to the developed of the stepped thinning of the continental crust seen at these rifted margins? Rifted margin lithosphere geometries, with superimposed isostatic response and bathymetry, are shown in Figure 21 for upwelling divergent flow models of rifted margin formation in which the flow-field pattern propagates upwards from the base of the lithosphere towards the surface but pauses at mid-crustal level for several millions of years before reaching the surface and rupturing the continental lithosphere. The result of pausing the flow field at 20 km depth for 0, 5 and 10 Ma is shown in Figure 21a–c, respectively. Pausing the flow field for 10 Ma prior to it reaching the lithosphere surface and rupturing the lithosphere produces a stepped margin crustal thickness bathymetry with thinned upper crust lying directly on mantle (Fig. 21c). It is suggested that a pause in the upward propagation of the upwelling divergent flow pattern at the base of the brittle upper crust may explain the margin geometries, such as those shown in Figure 21a, b. This paused flow field model has been applied to the Iberia margin profile on the southern edge of Galicia Bank. The best-fit model uses a synbreakup pure shear with b ¼ 1.5 and an upwelling divergent flow field that pauses for 15 Ma at 20 km depth followed by propagation to the surface. Observed and modelled free air gravity profiles are shown in

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Figure 22c & d. The stepped bathymetry of this margin and the superimposed pair of free air gravity dipoles are successfully modelled. The corresponding lithosphere margin structure predicted by the model is shown in Figure 22e and suggests that upper continental crust should lie directly above mantle on this profile on the southern edge of Galicia Bank. Why should the upward propagation of the upwelling divergent flow field pause at the base of the upper crust? Inspection of the offshore Iberia bathymetry shows that, to the south of the modelled profile, the ocean –continent transition is located much nearer to the Iberian coast. It is suggested that the continental lithosphere thinning that led to the formation of the southern Iberian also propagated northwards thinning the lithosphere under the Galicia Interior Basin. However, for reasons

of plate kinematics or lithosphere rheology, the upwelling divergent flow was unable to propagate upwards to the surface within the Galicia Interior Basin and instead either stopped at the level of the mid-crust or migrated westwards where it was able to successfully rupture the continental lithosphere to the west of Galicia Bank. The Galicia Interior Basin was probably partly formed by Jurassic intracontinental rifting (Murillas et al. 1990), however its present-day crustal basement b thinning factors are in excess of 3, which exceeds the upper crustal extension estimated from heave summation on upper crustal faulting (Davies & Kusznir 2004). The Galicia Interior Basin therefore shows lithosphere depth-dependent stretching that may be explained by a thinning of the lower crust and lithospheric mantle under the Galicia Interior Basin by an upwelling divergent flow field which

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A NEW MODEL OF LITHOSPHERIC THINNING

failed to reach the surface and either ceased or migrated westwards.

Discussion and conclusions Many passive continental margins are characterized by a regional distribution and thickness of synrift and post-rift sediment packages that are not consistent with the minor amounts of brittle deformation observed in either seismic sections across the margin (e.g. Exmouth Plateau, NW Australia; Marion Plateau, NE Australia; Grand Banks; Brazilian and West African margins; and West of Shetlands Basins) or from field mapping of synrift stratigraphy (e.g. Brazilian and West African margins). While the geological details and sedimentary facies differ between the various margins, the style of deformation and the regional distribution of accommodation are remarkably similar. The development of significant post-rift accommodation in the same region characterized earlier by minor

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synrift faulting and shallow depositional environments is most easily explained in terms of depthdependent extension. A major problem with depth-dependent lithospheric extension hypotheses for analysing passive margin subsidence and architecture is that extension needs to be horizontally balanced in some way, often assumed to occur in the vicinity of the ocean –continent transition where upper plate extension is counterbalanced by a combination of thinned and ‘rafted’ crustal blocks that expose the continental mantle augmented and masked by emplacement of tholeiitic magmas that comprise landwards-and seawards-dipping reflector packages. That such a balance exists has been difficult to verify. In marked contrast, by analysing lithospheric thinning in terms of a fluid dynamic model for deformation of continental lithosphere in which upwelling and divergence within the continental lithosphere and asthenosphere are parameterized by a vertical upwelling velocity and a horizontal

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divergence velocity, we can explain in a straightforward way depth-dependent extension of the lithosphere and thus the paucity of synrift brittle deformation of the upper crust, the transition from continental lithosphere thinning and the breakup to sea-floor spreading, a balance of extensional strain, and the observed first-order diversity of rifted continental margin structure and width of the ocean –continent transition. The origin of the upwelling divergent flow is possibly driven by asthenosphere flow ahead of laterally propagating sea-floor spreading (Taylor et al. 1999), anomalously hot asthenosphere (plume impingement) and/or convective instabilities in thickened orogenic lithosphere (Gemmer & Houseman 2004). Depth-uniform lithosphere extension is considered subordinate because of the dominance of regional sagging, as opposed to fault-controlled accommodation space generation, prior to breakup.

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Fig. 24. Application of the upwelling divergent flow model of rifted margin formation to the formation of failed breakup basins. (a) Lithosphere structure and bathymetry after flow field upwards propagation for 5 Ma and (b) 12.5 Ma, by which time the flow field has reached mid-crustal level. (c) Upwelling divergent flow continues from 12.5 Ma to 22.5 Ma, but does not propagate upwards, and then ceases. (d) The lithosphere then thermally re-equilibrates and subsides to 100 Ma. Cessation of the upwards propagation of the flow pattern before reaching the lithosphere surface produces thinning of the continental lithospheric mantle and lower crust, but no thinning and stretching of the upper crust. Such failed breakup basins show depth-dependent stretching with large post-rift thermal subsidence, but little upper crustal fault extension. The model is axi-symmetric about distance x ¼ 0 km.

A NEW MODEL OF LITHOSPHERIC THINNING

exhibits large amplitude post-rift thermal subsidence during the upper Miocene corresponding to a lithosphere b stretching factor of approximately 5. In contrast, forward modelling of upper crustal fault extension demonstrates that upper crustal extension indicated by faulting is of the order of b ¼ 1.15. Upper crustal faults are well imaged; the observed discrepancy between upper crustal and whole lithosphere stretching and thinning cannot be explained by poor seismic imaging. Similar estimates of lithosphere and upper crustal stretching were also found by Roberts et al. (1999) for the Nam-Com-Som Basin. The Faroes –Shetland Basin lies along strike of the Møre rifted margin segment of the Norwegian margin that formed at the end of the Paleocene (c. 55 Ma). This basin experienced large-amplitude thermal subsidence since 54 Ma requiring lithosphere thinning factors of 3 or more; however, the Paleocene and Late Cretaceous show negligible upper crustal faulting (Fletcher et al. 2005). While the Faroes –Shetland Basin did experience Late Jurassic and Early Cretaceous intracontinental rifting, b stretching factors for this are unlikely to exceed 1.4 (Hurst 2006) and are insufficient to generate the observed top Paleocene to the present-day thermal subsidence. Figure 24 shows model predictions for the development of lithosphere structure and subsidence in response to an upwards propagating pattern of upwelling divergent flow pattern that propagates from the base of the lithosphere to mid-crustal level but then ceases, and is then followed by lithosphere thermal re-equilibration. The model predicts a broad sag basin with crustal thinning but no fault extension in the upper crust. The Nam-Com-Son and Faroes– Shetland Basins show depth-dependent lithosphere stretching, as predicted by this model. These basins plot with rifted margins rather than with intracontinental rift basins on the cross-plot of the upper crustal versus lithosphere thinning factor shown in Figure 8. We propose that these basins are called failed breakup basins to distinguish them from intracontinental rift basins. We believe that continental lithosphere thinning leading to breakup and sea-floor spreading initiation occurs in response to an upwelling divergent flow field within continental lithosphere and asthenosphere rather by a depth-uniform (pure shear) lithosphere stretching. This process is able to explain depth-dependent lithosphere stretching and thinning observed at rifted continental margins, and also failed breakup basins. Many important questions still remain, however. Does the continental crust rupture before the continental lithospheric mantle during continental breakup, or does the continental lithospheric mantle rupture first? Another important question is whether mantle exhumation

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is caused by unroofing of continental lithospheric mantle or whether it is caused by melt suppression during early sea-floor spreading analogous to mantle exhumation at ultra-slow sea-floor spreading at the Gakkel and South East Indian ridges (Michael et al. 2002; Dick et al. 2003)? It is likely that both processes are operative. A related question is whether oceanic crust nearest to the ocean –continent transition at volcanic margins is underlain by continental lithospheric mantle or by exhumed asthenosphere. The new model of continental breakup and rifted margin formation (SfMargin) has been developed as part of the NERC Ocean Margins iSIMM (integrated Seismic Imaging and Modelling of Margins) project. The iSIMM team comprises researchers from the Universities of Liverpool and Cambridge, Badley Geoscience and Schlumberger, and is funded by a consortium of NERC, DTI, Agip, BP, Amerada Hess, Anadarko, ConocoPhillips, Shell, Statoil, & WesternGeco. iSIMM investigators are N. J. Kusznir, R. S. White, P. A. F. Christie, A. M. Roberts, A. Chappell, J. Eccles, R. Fletcher, D. Healy, N. Hurst, Z. Lunnon, C. J. Parkin, A. W. Roberts, L. K. Smith, R. Spitzer and V. Tymms. The authors gratefully acknowledge discussions with iSIMM team colleagues and constructive reviews by Gianreto Manatschal and Luis Menezes Pinheiro, all of which significantly improved this manuscript.

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Observations from the Basin and Range Province (western United States) pertinent to the interpretation of regional detachment faults N. CHRISTIE-BLICK1, M. H. ANDERS2, S. WILLS2, C. D. WALKER1 & B. RENIK1 1

Department of Earth and Environmental Sciences and Lamont-Doherty Earth Observatory of Columbia University, Palisades, NY 10964, USA (e-mail: [email protected]) 2

American Association for the Advancement of Science, 1200 New York Avenue NW, Washington, DC 20005, USA Abstract: This paper summarizes the results of completed and ongoing research in three areas of the Basin and Range Province of the western United States that casts doubt on the interpretation of specific regional detachment faults and the large extensional strains with which such faults are commonly associated. Given that these examples were influential in the development of ideas about low-angle normal faults, and particularly in making the case for frictional slip at dips of appreciably less than the 308 lock-up angle for m  0.6 (where m is the coefficient of friction), we advocate a critical re-examination of interpreted detachments elsewhere in the Basin and Range Province and in other extensional and passive margin settings. The Sevier Desert ‘detachment’ of west-central Utah is reinterpreted as a Palaeogene unconformity that has been traced to depth west of the northern Sevier Desert basin along an unrelated seismic reflection (most probably a splay of the Cretaceous-age Pavant thrust). The absence of evidence in well cuttings and cores for either brittle deformation (above) or ductile deformation (below) is inconsistent with the existence of a fault with as much as 40 km of displacement. The Pavant thrust and the structurally higher Canyon Range thrust are erosionally truncated at the western margin of the southern Sevier Desert basin, and are not offset by the ‘detachment’ in the manner assumed by those inferring large extension across the basin. The Mormon Peak detachment of SE Nevada is reinterpreted as a series of slide blocks on the basis of detachment characteristics and spatially variable kinematic indicators that are more closely aligned with the modern dip direction than the inferred regional extension direction. A particularly distinctive feature of the detachment is a basal layer of up to several tens of centimetres of polymictic conglomerate that was demonstrably involved in the deformation, with clastic dykes of the same material extending for several metres into overlying rocks in a manner remarkably similar to that observed at rapidly emplaced slide blocks. The Castle Cliff detachment in the nearby Beaver Dam Mountains of SW Utah is similarly regarded as a surficial feature, as originally interpreted, and consistent with its conspicuous absence in seismic reflection profiles from the adjacent sedimentary basin. The middle Miocene Eagle Mountain Formation of eastern California, interpreted on the basis of facies evidence and distinctive clast provenance to have been moved tectonically more than 80 km ESE from a location close to the Jurassic-age Hunter Mountain batholith of the Cottonwood Mountains, is reinterpreted as having accumulated in a fluvial–lacustrine rather than alluvial fan – lacustrine setting, with no bearing on either the amount or direction of tectonic transport. The conglomeratic rocks upon which the provenance argument was based are pervasively channelized, with erosional relief of less than 1 m to as much as 15 m, fining-upwards successions at the same scale and abundant trough cross-stratification – all characteristic features of fluvial sedimentation and not of alluvial fans. The interpretation of the Eagle Mountain Formation as having been deposited within a few kilometres of the Hunter Mountain batholith, which depends strongly on assumptions about the dimensions of alluvial fans, is therefore not required. The result is important because the Eagle Mountain offset has been viewed as representing the strongest evidence for extreme extension in this part of California, and for the existence of detachment faults of regional dimensions.

‘Even one compelling example of a primary LANF [low-angle normal fault] or of LANF slip is sufficient to prove that they may form and slip at low dip, respectively.’ – Axen (2004, 50).

The Basin and Range Province of the western United States has been highly influential in the development of ideas about crustal extension, particularly the role of low-angle normal faults or

From: KARNER , G. D., MANATSCHAL , G. & PINHEIRO , L. M. (eds) Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 421– 441. DOI: 10.1144/SP282.17 0305-8719/07/$15.00 # The Geological Society of London 2007.

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detachments of regional scale (e.g. Armstrong 1972; Crittenden et al. 1980; Wernicke 1981, 1985, 1992; Allmendinger et al. 1983; Davis 1983; Miller et al. 1983; Spencer 1984; Wernicke & Axen 1988; Wernicke et al. 1988; Lister & Davis 1989; Spencer & Chase 1989; Axen et al. 1990, 1993; John & Foster 1993; Livaccari et al. 1993; Axen & Bartley 1997; Brady et al. 2000a; Snow & Wernicke 2000; Livaccari & Geissman 2001; Axen 2004; Carney & Janecke 2005). The detachment concept is now widely applied in extensional and passive margin settings (e.g. Froitzheim & Eberli 1990; Lister et al. 1991; Reston et al. 1996; Driscoll & Karner 1998; Hodges et al. 1998; Osmundsen et al. 1998; Taylor et al. 1999; Boncio et al. 2000; Manatschal et al. 2001; Canales et al. 2004). However, the apparent conflict between generally accepted geological interpretations and rock mechanical and seismological considerations has yet to be resolved satisfactorily (Sibson 1985; Jackson 1987; Jackson & White 1989; Collettini & Sibson 2001; Scholz & Hanks 2004; cf. Axen 1992, 1999, 2004; Scott & Lister 1992; Axen & Selverstone 1994; Wernicke 1995; Rietbrock et al. 1996; Abers et al. 1997; Westaway 1999; Sorel 2000; Collettini & Barchi 2002; Hayman et al. 2003). Beginning in the early 1990s, we became interested in Basin and Range examples for which evidence was regarded by the structural geological community to be the most compelling for slip at dips of appreciably less than 308 – the lowest

plausible frictional lock-up angle for crustally rooted normal faults in the absence of unusual materials (m , 0.6, where m is the coefficient of friction) (Sibson 1985; Collettini & Sibson 2001) – and for the very large extensional strains with which low-angle normal faults are commonly associated (Wernicke et al. 1988; Levy & ChristieBlick 1989; Wernicke 1992; Snow & Wernicke 2000). We reasoned that if progress was to be made in developing a better theoretical understanding, it would be useful to focus on geological examples providing the firmest constraints. This paper summarizes research from three case studies of historical significance and that individually account for many tens of kilometres of current estimates of upper crustal extension in the Basin and Range Province (Fig. 1). While each example has become closely associated with the work of specific investigators over several years, in offering new interpretations of available data we emphasize that our interest is solely in resolving a long-standing paradox. It is not our intent to cast aspersions on colleagues who have struggled hard with the same issues, and contributed much to the present state of knowledge. A re-evaluation of the Sevier Desert detachment of west-central Utah (S in Fig. 1; see also Figs 2 & 3) (McDonald 1976; Allmendinger et al. 1983; Smith & Bruhn 1984; Von Tish et al. 1985; Mitchell & McDonald 1986, 1987; Planke & Smith 1991; Otton 1995; Coogan & DeCelles 1996; Stockli et al. 2001) is now largely published

Fig. 1. Location of three study areas in Basin and Range Province: Sevier Desert (S), west-central Utah (see Fig. 2); Mormon Mountains (M), SE Nevada and Beaver Dam Mountains (B), SW Utah (see Fig. 5); and Eagle Mountain (E), eastern California. Other localities mentioned in text: Cottonwood Mountains (C), Death Valley (D), Resting Spring Range (R) and Nopah Range (N). Physiographical base for the map was modified from Thelin & Pike (1991).

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Fig. 2. Generalized map of Sevier Desert basin, Utah (modified from Wills et al. 2005; see Fig. 1 for location), showing locations of seismic profiles and wells mentioned in text. Seismic profiles (bold lines): COCORP Utah Line 1 (see Fig. 3); PC-8 and V-5 (see Fig. 4). Wells: AE, Argonaut Energy Federal; AMF, ARCO Meadow Federal; CA, Cominco American Federal; GG, Gulf Gronning. The locations of other seismic profiles and wells used by Wills et al. (2005) are included for context, but without labels. Solid, dashed and dotted lines correspond with different seismic datasets. Thrust faults are the Canyon Range thrust in the Canyon Range and the structurally lower Pavant thrust in the Pavant Range (teeth on upper plate). A drafting error in the Wills et al. (2005) original (Pavant thrust) has been corrected.

(Anders & Christie-Blick 1994; Wills & Anders 1999; Anders et al. 2001; Wills et al. 2005; for a discussion of contrasting views, see Allmendinger & Royse 1995; Anders et al. 1995, 1998a; Wills & Anders 1996; Otton 1996; Coogan & DeCelles 1998; Hintze & Davis 2003; DeCelles & Coogan 2006). Here, we summarize the evidence for an alternative interpretation: that the purported detachment is instead a regional unconformity of Palaeogene age (Gradstein et al. 2004) fortuitously aligned down dip in some seismic reflection profiles with a Mesozoic thrust fault. Our research on detachment faults in the Mormon Mountains and Beaver Dam Mountains of SE Nevada and SW Utah (M and B in Fig. 1), and at known slide blocks elsewhere (Anders et al. 2000, 2006; Walker et al. 2007), has dealt mainly with the characteristics of fault zones and with the discrimination of crustally rooted structures from features that are surficial (or rootless). Data in hand are not consistent with the long

accepted interpretation of the Mormon Peak and associated detachments as rooted faults (Wernicke 1981, 1982, 1995; Wernicke et al. 1985, 1988, 1989; Wernicke & Axen 1988; Axen et al. 1990; Axen 1993, 2004), and suggest instead that these structures relate to block-sliding with crustal extension accommodated entirely by high-angle normal faults (Anders et al. 2006; Walker et al. 2007; see also Cook 1960; Tschanz & Pampeyan 1970; Hintze 1986; Carpenter et al. 1989; Carpenter & Carpenter 1994; and Axen & Wernicke 1989 and Axen 2004 for a markedly different opinion). The third example to which we draw attention in this paper relates to the generally accepted geological evidence for more than 400% extension in the Death Valley area of eastern California, between the Cottonwood Mountains and Nopah Range (C and N in Fig. 1) (Stewart 1983; Wernicke et al. 1988, 1993; Snow & Wernicke 1989, 2000; Holm et al. 1992; Snow 1992a; Topping 1993; Brady et al. 2000b; Niemi et al. 2001; for contrasting

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Fig. 3. Part of seismic reflection profile COCORP Utah Line 1, with interpretation of Sevier Desert detachment from Von Tish et al. (1985). See Figure 2 for location. The eastward-dipping panel labeled Olig. (Oligocene) by Von Tish et al. is reinterpreted as Palaeozoic and Neoproterozoic (Canyon Range allochthon), based on a velocity (3.2 km s21) for the basin fill at the Gulf Gronning well (G.G.) that is higher than originally assumed. The east-dipping fault beneath G.G. is inferred in Wills et al. (2005) and in this paper to offset the purported detachment, misaligning an updip portion of the reflection (interpreted as an unconformity) and a downdip portion (interpreted as a Mesozoic thrust fault, most probably a splay of the Pavant thrust).

interpretations of available data, see Stewart 1967, 1986; Wright & Troxel 1967, 1970; Stewart et al. 1968, 1970; Prave & Wright 1986a, b; Corbett 1990; Wernicke et al. 1990; Stevens et al. 1991, 1992; Snow 1992b; Snow & Wernicke 1993; Stone & Stevens 1993; Serpa & Pavlis 1996). Re-examination of the most important individual piece of evidence, a middle Miocene succession of breccia, conglomerate, sandstone, siltstone, limestone and tephra (Eagle Mountain Formation; E in Fig. 1) interpreted by Niemi et al. (2001) to have been deposited in an alluvial fan and lacustrine setting, and to have been transported tectonically more than 80 km from its Cottonwood Mountains source, suggests that the sediments accumulated in a fluvial–lacustrine environment (Renik & Christie-Blick 2004). If this alternative interpretation is correct, these deposits provide no constraint on either the magnitude or direction of extension, and they have no bearing on the interpretation of regional detachment faults.

Sevier Desert detachment Interpretations of the Sevier Desert detachment of west-central Utah (Figs 2 and 3; see also

Appendix 1) are based primarily on 1970s vintage petroleum industry seismic reflection and well data (McDonald 1976; Mitchell & McDonald 1987) and seismic profiles acquired by the Consortium for Continental Reflection Profiling (COCORP; Allmendinger et al. 1983; Von Tish et al. 1985). The low westward dip (118) of the hypothesized detachment, its lateral continuity over 7000 km2, the purported offset of Mesozoic structures by as much as 47 km, the downward termination of high-angle normal faults within the Sevier Desert basin (SDB) and the involvement of sediments as young as Holocene provide seemingly unassailable evidence for large normal offset on a fault that could never have been appreciably more steeply inclined than it is today, and might still be active (Wernicke 1995; Niemi et al. 2004). Difficulties with the detachment hypothesis relate primarily to the absence of anticipated deformation in borehole cuttings and cores, in spite of an extensive study of such materials (Anders & Christie-Blick 1994; Anders et al. 2001), and to details of the subsurface stratigraphic and structural interpretation that are at odds with key elements of published palinspastic reconstructions (Wills et al. 2005). Cuttings from directly above the Palaeozoic– Palaeogene contact where the maximum depth

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never exceeded 1.8 km exhibit no sign of deformation expected of a brittle fault at that depth (Anders & Christie-Blick 1994). Fossil trilobite fragments, recovered in core from 12.8 m below the same contact and in cuttings from as close as 3 m, are undeformed (Anders et al. 2001). A palinspastic reconstruction assuming 47 km of normal slip on the hypothesized detachment places the trilobite-bearing Palaeozoic carbonate rocks at a depth of at least 14 km at the time faulting is inferred to have begun (after the late Oligocene). Given that the region had been volcanically active for at least 10–20 million years (beginning in the Eocene), any reasonable estimate of the geothermal gradient puts the rocks at a temperature (as high as 425 8C) at which mylonites should have developed in carbonate rocks (Anders et al. 2001). No evidence for mylonitization has been observed. The apparent continuity of the Sevier Desert reflection (SDR) from the Palaeozoic–Palaeogene contact beneath the basin to within deformed Neoproterozoic and Palaeozoic rocks of the Cricket Mountains block to the west is restricted to the northern SDB (Figs 2, 3 & 4a) (Wills et al. 2005). In the south, the SDR terminates abruptly at or only a short distance beneath the Cricket Mountains (Fig. 4b). Although seismic attenuation and the

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absence of a sufficiently large contrast in acoustic impedance offer plausible explanations for this observation, a comprehensive re-analysis of seismic reflection and borehole data, including nearly 600 km of previously unavailable profiles, leads to a different conclusion. The feature with which the SDR is approximately aligned in COCORP Utah Line 1, and in virtually all industry and academic dip-oriented profiles from the northern part of the basin, is a Mesozoic thrust fault, most probably a splay of the Cretaceous-age Pavant thrust. That fault, which places Lower Cambrian quartzite atop Upper Cambrian carbonate rocks in the Cominco American Federal well (CA in Fig. 2), rises southwards approximately 2–3 km with respect to a westward projection of the SDR. Even in COCORP Line 1 (Fig. 3) the two features are misaligned by about 0.3 s two-way travel time at a high-angle normal fault system that has been mapped for approximately 80 km along the western side of the basin (cf. Fig. 4a). Our re-evaluation of the subsurface data is inconsistent with a related second pillar of the detachment hypothesis: that as much as 47 km of normal slip is needed to account for the offset of Mesozoic thrust sheets across the SDB (Von Tish et al. 1985; Allmendinger & Royse 1995; DeCelles

Fig. 4. Interpretive line drawings of seismic reflection profiles from northern and southern Sevier Desert basin (modified from figs 6A and 7D of Wills et al. 2005; see Fig. 2 for the location). Grey indicates pre-Cenozoic rocks; white represents Cenozoic basin fill, with stratal geometry shown by fine lines. Both profiles illustrate a persistent pattern of stratal onlap against sub-Cenozoic unconformity of Cricket Mountains block (west) and the existence of basin-bounding high-angle normal faults of relatively small offset. (a) PC-8 (McDonald 1976). Black indicates interstratified basalt and clastic sediments penetrated by Gulf Gronning well 10 km south of section. Domal structures flanked by growth stratigraphy are inferred to be salt-cored. SDR is downdip portion of the Sevier Desert reflection, interpreted as a splay of Pavant thrust. AE is Argonaut Energy Federal well. Vertical scale in kilometres. (b) V-5 (Wills et al. 2005). Listric faults displace panels of subparallel reflections, with stratal growth preferentially at higher stratigraphic levels. Faults interpreted as the Canyon Range thrust (CRT) and Pavant thrust (PT) are erosionally truncated (left). Neither aligns with base of basin to east. WF is Whirlwind Formation (Cambrian). AMF is ARCO Meadow Federal well. Vertical scale in seconds, two-way travel time.

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et al. 1995; Coogan & DeCelles 1996, 1998; Stockli et al. 2001; DeCelles & Coogan 2006). Both the Pavant thrust and the structurally higher Canyon Range thrust are erosionally truncated at the western margin of the southern SDB (Fig. 4b). The SDB is therefore underlain primarily by subPavant rocks, and by the Pavant and Canyon Range allochthons in the north and west as a result of the regional dip of those structures. That a thick succession of Neoproterozoic strata composing the Canyon Range allochthon in the Cricket Mountains projects approximately 60 km NE across the SDB to a tightly folded klippe of the same thrust sheet at the crest of the Canyon Range (Fig. 2) is a measure of the dimensions of the now dissected Mesozoic allochthon rather than of its offset by a Cenozoic detachment fault. Taken together, these observations and others summarized in Appendix 1 lead to a very different interpretation of the SDR: that it is a regional unconformity of Palaeogene age, aligned particularly west of the northern SDB with a Mesozoic thrust fault. The downward termination of high-

angle normal faults and localized stratigraphic growth, also cited as evidence for the existence of a detachment, are attributed to syndepositional deformation of Oligocene salt that today forms residual masses more than 1500 m thick (Argonaut Energy Federal well; AE in Fig. 4a).

Mormon Peak and associated detachments The Mormon Peak, Tule Springs and Castle Cliff detachments (MPD, TSD and CCD, respectively) of SE Nevada and adjacent Utah have been interpreted on the basis of geological mapping to be crustally rooted and of regional extent, and to accommodate 54 + 10 km of Miocene extension (Figs 5 and 6; Appendix 2) (Wernicke 1981, 1982, 1995; Wernicke et al. 1985, 1988, 1989; Wernicke & Axen 1988; Axen et al. 1990; Axen 1993, 2004). Research on these detachments, beginning in the late 1970s, was instrumental in arguing for the importance of low-angle normal faults and for

Fig. 5. (a) Physiographical map of the Mormon Mountains, Tule Springs Hills and Beaver Dam Mountains of SE Nevada and adjacent Utah, with location of range-bounding normal faults (ball on downthrown side), seismic reflection profiles (with shot points) and the eastern portion of the regional geological cross-section A –A0 (modified from Anderson & Barnhard 1993a and Carpenter & Carpenter 1994; see Fig. 1 for the location and Fig. 9 for profile 4-4A). (b) Cross-section A– A0 (simplified from fig. 17a of Axen et al. 1990 and plate 2 of Anderson & Barnhard 1993a; see Fig. 6 for the restored section). Abbreviations: p–C, Precambrian crystalline rocks; P and Mz, Palaeozoic and Mesozoic strata in footwall of Mormon– Tule Springs thrust fault; TS, Precambrian and Palaeozoic rocks in the hanging wall of Tule Springs thrust fault; Cz, Cenozoic strata.

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Fig. 6. Published pre-extension restoration of the regional geological cross section A –A0 from Meadow Valley Mountains to Beaver Dam Mountains (modified from fig. 17c of Axen et al. 1990; see fig. 5b for the present-day section). Vertical and horizontal scales are the same, with elevations corresponding approximately to those of today at the WSW and ENE ends of the section. The dotted line on the left is present topography. The original line length of 49 km is inferred by Axen et al. to have been increased by 54 km in a section oriented 2588 (2558 + 108 is extension direction inferred by Wernicke et al. 1988). The Mormon Peak, Tule Springs and Castle Cliff detachments (indicated by two, three and four ticks, respectively) are notable for their markedly different restored geometry. The portion of the Tule Springs (Mormon) thrust inferred by Axen et al. (1990) to have been reactivated as the Tule Springs detachment is indicated by both thrust (teeth) and detachment (single tick) symbols. Abbreviations: p–C Precambrian crystalline rocks; black, Cambrian clastic rocks; –C-M, Cambrian –Mississippian carbonate rocks; P, Pennsylvanian–Permian strata; Mz, Mesozoic strata. Palinspastic location of block 6365 is compared with that of its footwall at western flank of Mormon Mountains (see Fig. 5a).

high extensional strain in regions not associated with metamorphic core complexes. Initial dips inferred from palinspastic reconstruction are 328 for the CCD to a depth of at least 7 km; 208–288 for the MPD to a depth of at least 6 km; and 38– 158 for the Tule Springs detachment to a depth of 2–5 km, corresponding with a flat in the Mesozoicage Tule Springs thrust (Fig. 6) (Wernicke & Axen 1988; Axen et al. 1990; Axen 1993, 2004). Present dips are 118W for the CCD; ,208E and ,208W, and with a considerable range of azimuth for the MPD; and near horizontal for the TSD in available outcrop. Following earlier work by Carpenter & Carpenter (1994), a challenge to the generally accepted rooted interpretation of these detachment faults has emerged recently from studies of the character and kinematics of deformation associated with the MPD (Appendix 2) (Anders et al. 2006; Walker et al. 2007). As recognized long ago by Wernicke (1982), the MPD is in many places characterized by a 0.1–1m-thick layer of polymictic conglomerate (Fig. 7) that he interpreted as synorogenic gravel overridden by a Mesozoic thrust fault before being carried down in the fault zone of the detachment. More recently, Wernicke re-interpreted the conglomerate as sediments ‘deposited after or near the end of detachment activity in cavern systems that followed the crushed carbonates adjacent to the detachment’ (pers. comm. 2002 to G. J. Axen; p. 61 in Axen 2004). Neither explanation matches available observations. The conglomerate is compositionally distinct from known Cretaceous synorogenic deposits

in this part of Nevada (Carpenter & Carpenter 1994), deposits that are in any case unlikely to have remained unlithified for approximately 50 million years at a depth of up to several kilometres. Although karst is present in these rocks, the grading, flow banding and internal erosional features that characterize the conglomerate are not normally associated with karst infill (James & Choquette 1988). Most important, the conglomerate was demonstrably involved in the deformation, with a lower faulted contact characterized by a thin gouge or cataclastic layer (Fig. 7). As is typical of large gravity slides, a network of clastic dykes is extensively developed above the detachment, but absent below. With one possible exception, no evidence has been found for more than a single emplacement event at any individual locality. A second difficulty for the concept of an extensional allochthon of regional scale is that kinematic indicators for the MPD (slickenlines and minor faults in the basal conglomerate) are of varied orientation, diverge markedly from the published regional extension direction (Fig. 8; 2558+108) (Wernicke et al. 1988), and at many locations correspond approximately with the modern downdip direction of the detachment surface (Anders et al. 2006; Walker et al. 2007; cf. Axen 2004, p. 60). Available evidence is consistent with the rapid surficial sliding of blocks downhill, away from the crest of the Mormon Mountains, as it existed in Miocene time (see Appendix 2) (Anders et al. 2006; Walker et al. 2007). Polymictic conglomerate, similar to that observed at the MPD and recognized as basal layers and dykes within a wide range

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Fig. 7. Cross-section of trench through Mormon Peak detachment beneath block 6365, Mormon Mountains (see Figs 5 and 6 for the location; and Anders et al. 2006 for additional information about this contact). Arrow points to detachment surface. Representative hand specimens on right: a, crackle breccia of Devonian dolomite (Sultan Formation); b, matrix-supported auto-breccia of Devonian dolomite; c, mixture of dolomite and conglomerate; d, polymictic conglomerate with well-rounded particles; e, basal gouge layer; f, brecciated Cambrian dolomite (Bonanza King Formation). The conglomerate illustrated by specimen d is commonly characterized by grading, flow banding and internal erosional features.

Fig. 8. Rose diagram showing the trend of 31 kinematic indicators (mostly slickenlines and grooves) at localities along Mormon Peak detachment in Mormon Mountains, Nevada (data from C. D. Walker). The indicators, which provide no independent measure of sense of displacement, are plotted according to local direction of plunge and in increments of 58. Double-headed arrow indicates inferred extension direction of 2558 + 108 from Wernicke et al. (1988). Azimuths are remarkably dispersed, 29% of them more than 408 from 2558 or 0758. Only two readings (6%) are oriented at more than 408 from the local dip direction of the detachment.

of slides (e.g. Yarnold & Lombard 1989; Shaller 1991; Beutner & Craven 1996; Anders et al. 2000), is thought to be associated with fluidization, which serves to explain the long runouts of such features (Shaller 1991; Anders et al. 2000; Beutner & Gerbi 2005). The Castle Cliff detachment on the western flank of the Beaver Dam Mountains is one of the structures used by Wernicke & Axen (1988) to exemplify their rolling hinge model for extensional unroofing, but its significance is challenged by seismic reflection data and geological mapping that were available in the 1980s (Fig. 9) (Hintze 1986; Carpenter et al. 1989). As interpreted by Wernicke & Axen (1988), the detachment passes beneath now isolated exposures of Cambrian – Mississippian rocks at Castle Cliff and Sheep Horn Knoll (Fig. 5a), as well as a series of smaller blocks and their substrate of Neogene (Gradstein et al. 2004) gravel at the range front, and projects to depth with a dip of 118 (Fig. 5b) (Wernicke et al. 1989; Axen et al. 1990; Axen 2004; cf. Anderson & Barnhard 1993a, b). Surprisingly, no detachment can be discerned in a seismic reflection profile that intersects the range front only 5 km south of the geological cross-section (Figs 5a & 9) (section 4A of Carpenter & Carpenter 1994) or in

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seismic reflection data from anywhere else in the Virgin River depression immediately south and east of the Mormon Mountains and Tule Springs Hills (Bohannon et al. 1993). Instead, profile 4A reveals the presence of a range-bounding normal fault that, after depth conversion and a correction for its oblique orientation, dips at 608 (Fig. 9b). In the absence of any evidence for a transfer fault between section A– A0 and line 4-4A (Fig. 5) (cf. Axen 2004, p. 60), the apparent inconsistency can perhaps be resolved by hypothesizing that the detachment is offset to a depth of at least 4 s (.5.5 km) by the high-angle fault (G. J. Axen pers. comm. 1998). However, that interpretation is at odds with the published cross-section (Fig. 5b) as well as with the rolling hinge model for the detachment (Wernicke & Axen 1988; see Axen & Bartley 1997). A more plausible interpretation is that the Castle Cliff detachment is an expression of landsliding, as

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previously suggested by Cook (1960), Hintze (1986), Carpenter et al. (1989) and Anders et al. (1998b), and has nothing to do with crustal extension. According to that view, the isolated blocks of Palaeozoic rocks have a common origin (they are all rootless), independent of whether their substrate is Precambrian or Neogene; and the sub-Neogene contact at the flank of the Beaver Dam Mountains is in part depositional and in part faulted, but not an expression of the Castle Cliff detachment. The slide block interpretation has been dismissed by Axen & Wernicke (1989) and by Axen (2004) for a host of reasons, none of them definitive or accounting for the specifics outlined here. A key argument relates to the near concordance (,58 – 108; Hintze 1986) between Oligocene–Miocene sedimentary and volcanic rocks and Cretaceous – Palaeogene (?) strata in the northern Beaver Dam Mountains (Axen & Wernicke 1989). It is reasoned

Fig. 9. Part of seismic reflection profile 4-4A (a) with interpretation (b) (modified from fig. 10 of Carpenter & Carpenter 1994). See Figure 5a for the location. Owing to oblique intersection, the normal fault at west flank of Beaver Dam Mountains has an apparent dip of 408; its true dip is 608. The Castle Cliff detachment is either surficial or offset to a depth of at least 4 seconds two-way travel time (.5.5 km for seismic velocity of 2.74 km s21 assumed by Carpenter & Carpenter 1994). Neither interpretation is consistent with that shown in Figure 5b (simplified from Axen et al. 1990). Abbreviations: Ths, Horse Spring Formation (Miocene); Tm, Muddy Creek Formation and younger strata.

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on this basis that an anticline cored by crystalline rocks at the west flank of the range (Wernicke & Axen 1988) must have developed as a result of uplift and exhumation of the footwall of the Neogene extensional system. At issue is: (1) whether the perceived footwall uplift relates to the Castle Cliff detachment or (more likely) to highangle normal faults along the front of the range (compare sections A –A0 and B– B0 in fig. 2 of Wernicke & Axen 1988); and (2) whether the fold itself is an expression of footwall deformation, or is perhaps older. Two observations cast doubt on the purported Neogene age of the fold (plate 2 of Hintze 1986). First, the structure is oblique to and truncated by the range-bounding fault system, not aligned along it. Second, the existence of a panel of strongly overturned Pennsylvanian –Permian strata between the range-front crystalline rocks and the outcrops of Oligocene–Miocene strata is inconsistent with the supposed structural simplicity upon which the argument of Axen & Wernicke (1989) critically depends. For these reasons, we do not think that the fold necessarily has any bearing on the interpretation of the CCD.

Eagle Mountain Formation Upper crustal extension across the Death Valley region of eastern California (D in fig. 1) has long accounted for as much as two-thirds of the approximately 250 –300 km estimate of WNW motion of the Sierra Nevada away from the Colorado Plateau, and for the fraction of the total that, on the face of it, is best constrained from restorations of Neoproterozoic and Palaeozoic isopachs and facies transitions, Mesozoic thrust faults and folds, and palaeoisothermal surfaces (Stewart 1983; Wernicke et al. 1988, 1993; Snow & Wernicke 1989, 2000; Wernicke 1992). The very large inferred strain has also served to explain as much as 10– 15 km of middle and late Miocene exhumation of the Black Mountains on the east side of Death Valley, placing Neogene volcanic and sedimentary rocks in low-angle normal fault contact with rocks as old as 1.7 Ga as a result of tectonic removal of more than 10 km of Neoproterozoic and Palaeozoic strata that is still preserved in surrounding ranges (Holm et al. 1992; Topping 1993). The preferred mechanism for the inferred deformation is westward migration of a rolling hinge (Holm et al. 1992; Wernicke 1992; Axen & Bartley 1997). In spite of the apparent convergence of disparate structural, stratigraphic, geochronological and thermochronological constraints (e.g. Snow & Wernicke 1989), such reconstructions are subject to significant uncertainties: in structural correlation (of late Palaeozoic– Mesozoic thrust faults and

folds), in the spatial variability of stratigraphic thickness that limits confidence in reconstructing both the location and orientation of isopachs, and in matching imprecisely defined facies transitions (e.g. Stewart 1983, 1986; Prave & Wright 1986a, b; Wernicke et al. 1988, 1990; Corbett 1990; Stevens et al. 1991, 1992; Snow 1992a, b; Snow & Wernicke 1993; Stone & Stevens 1993). It is also necessary to assume some simple pre-extensional configuration for whatever markers are selected (e.g. fig. 3 of Snow & Wernicke 2000). To the extent that there is no way independently to verify what the configuration was, particularly for crustal blocks no more than a few kilometres across, the problem becomes intractable in the absence of an independent constraint. Research by Niemi et al. (2001) on the middle Miocene Eagle Mountain Formation provides a series of piercing points that appear to circumvent these inherent difficulties, at least for the structural elements in which these sediments are preserved (Fig. 1; Appendix 3). The presence in conglomerate at Eagle Mountain and in the central Resting Spring Range (E and R in Fig. 1) of a distinctive clast assemblage, including approximately 180 Ma leucomonzogabbro boulders (,1 m) indistinguishable from rocks found in the Hunter Mountain batholith of the Cottonwood Mountains (C in fig. 1) more than 80 km to the WNW, and sedimentological evidence for deposition at an alluvial fan together suggested an original location no more than 20 km from the source (taking that figure as a reasonable upper bound on fan dimensions). Although Niemi et al. (2001) asserted that the ‘lack of dilution of this detritus by other sources’ at least locally and at some stratigraphic levels by itself implies proximity to the Hunter Mountain batholith, other explanations can be considered (Appendix 3). The critical issue that we set out to test is whether the fan interpretation can be sustained. If deposition took place in a fluvial –lacustrine rather than alluvial fan–lacustrine setting, for example, the distribution of the Eagle Mountain Formation may instead reflect the configuration of the mid-Miocene drainage, with no significance for either the magnitude or direction of crustal extension. Sedimentological studies, to be reported in detail elsewhere, do not support the fan interpretation (Appendix 3; Renik & Christie-Blick 2004). Disorganized to diffusely stratified gravel is found in both alluvial fans and bedload rivers of sufficient gradient and discharge (Blair & McPherson 1994). The key to the distinction of these depositional settings in the Eagle Mountain Formation is stratigraphic architecture. Three features in particular point to deposition in a fluvial environment. The leucomonzogabbro-bearing deposits are characterized by pervasive channelization with erosional

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relief from less than 1 m to as much as 15 m (conglomerate-filled incised valleys), by finingupwards successions at the same scale within an overall upwards-coarsening succession 90 m thick, and by abundant trough cross-stratification (Fig. 10). † Channels Rivers are fundamentally channelized (e.g. Campbell 1976; Miall 1992; Best & Bristow 1993; Marzo & Puigdefa´bregas 1993; Willis 1993a; Collinson 1996). Alluvial fans are constructed primarily by sheet-flooding or sediment gravity flows, and composed of relatively tabular sedimentary bodies (Blair & McPherson 1994). Channels exist at alluvial fans only insofar as the fan surface is modified locally or from time to time as a result of changes in sediment flux, gradient or base level, for example, and ultimately through some combination of climate change, variations in sea or lake level and crustal deformation. The best-developed channels, other than the feeder channel at the fan apex, are commonly associated with headward erosion from fault scarps. At steady state, and in the absence of active faults, fan surfaces beyond the intersection point at the feeder channel terminus are relatively smooth, with

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only superficial gullies related to surface run-off between times of active accumulation. † Fining-upwards successions A tendency for upwards fining in bedload-dominated braided systems results from changing conditions during individual flood events and from the filling and abandonment of one channel in favour of another (e.g. Campbell 1976; Willis 1993b). An increase in discharge tends to be associated with renewed cutting of existing channels or the development of new channels. In contrast, the lateral building of alluvial fans results preferentially in the opposite motif at the same scale (metres to tens of metres): upwards coarsening punctuated by stratigraphic discontinuities (sequence boundaries or flooding surfaces; e.g. Steel et al. 1977; Gawthorpe et al. 1990; Blair & McPherson 1994; Lo´pez-Blanco et al. 2000). † Cross-stratification Bedforms are not well developed in all fluvial systems, particularly those transporting gravel. However, crossstratification is observed in most fluvial deposits because bedforms are common in fluvial channels. Cross-stratification is not an important feature in alluvial fans because the processes of sheet-flooding and sediment gravity flow do not involve large bedforms.

35

Siltstone Sandstone Conglomerate Carbonate Hunter Mountain clasts Erosional contact Current ripples Cross-stratification Wavy stratification Soft-sediment deformation

30 25 20 15 10

400 m

5 f.s.

0m Fining Upward

Silt.

Sand. Congl. NE

~500 m

SW

Fig. 10. Detail of inferred fluvial stratigraphy from middle Miocene Eagle Mountain Formation at Eagle Mountain (data from B. Renik). The column corresponds to the boxed portion of one of a series of 13 closely spaced measured sections depicted in the lower right area, within the lower part of unit Te2 of Niemi et al. (2001). The Eagle Mountain Formation unconformably overlies Cambrian Bonanza King Formation. Facies are mostly fluvial below prominent flooding surface (f.s., top of unit 17 in fig. 2 of Niemi et al. 2001), with valley-filling breccia and minor lacustrine sediment at the base; and mostly lacustrine deposits above that flooding surface. The grain size shown is the overall mean for sandstones and siltstones, and mean of granules and larger clasts for conglomerates (Wentworth scale).

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None of this bears on other arguments for large-scale crustal extension across the Death Valley region. At issue, however, is precisely how much extension, and the mechanisms by which it was achieved. If the Eagle Mountain Formation has been misinterpreted, for the reasons indicated, we are forced to retreat to the somewhat unsatisfactory palinspastic uncertainty that existed prior to 2001.

Discussion A critical re-evaluation of three Basin and Range examples, widely regarded prior to our work as providing a firm observational basis for frictional normal-sense slip along gently inclined faults of regional scale, leads us to interpretations that in each case and in quite different ways resolve the mechanical paradox with which such features are associated. (1) The Sevier Desert ‘detachment’ is reinterpreted as a Palaeogene unconformity that has been traced to depth along an unrelated seismic reflection (most probably a splay of the Pavant thrust). (2) The Mormon Peak detachment, the structure that was surely among the most influential in moving the low-angle normal fault concept away from metamorphic core complexes, is reinterpreted as a series of rootless slide blocks on the basis of detachment characteristics and spatially variable kinematic indicators. The nearby Castle Cliff detachment, a defining example for the rolling hinge model (but see Axen & Bartley 1997), is similarly regarded here as a surficial feature on the basis of its conspicuous absence in seismic reflection profiles west of the Beaver Dam Mountains – returning to a view that was generally accepted prior to 1988. (3) Conglomerates and sandstones of the middle Miocene Eagle Mountain Formation, only recently claimed to provide definitive piercing points for the reconstruction of extreme extension across the Death Valley region, turn out to be fluvial rather than alluvial fan deposits – with no bearing on either the amount or direction of tectonic transport, in spite of their unique provenance. We do not imply that the low-angle normal fault paradigm is falsified on the strength of these examples alone. Nor do we deny the reality of extreme crustal extension. Obviously, in the vicinity of the continent–ocean transition at passive continental margins, continental crust thins to zero. The results summarized here are significant primarily because they are unexpected, and because they suggest that a critical re-examination of other generally accepted evidence is now in order. At stake is not the existence of gently dipping normal faults (and other geological discontinuities).

It is how such readily observable field relations developed as a function of time, particularly in regions such as the Basin and Range Province where deformation and magmatism have been protracted, where crustal faults are in some cases distinguished with difficulty from the effects of surficial sliding, and where sediments have accumulated locally on exhumed fault surfaces. The interpretation of detachment faults from geological data depends critically on palinspastic restoration, and on the assumptions and uncertainties that are inherent in such reconstructions. We single out metamorphic core complexes of eastern California and adjacent Arizona for renewed attention, along with less easily studied locations along modern and ancient continental margins at which subcontinental mantle appears to have been dragged tectonically to the sea floor (e.g. Manatschal et al. 2001, 2007). Also of interest are the megamullions at mid-ocean ridges (Canales et al. 2004), a setting in which the rolling hinge model might be expected to apply owing to the inherent weakness of hot young lithosphere (e.g. Garce´s & Gee), and may not (Expedition Scientific Party 2005). In each case, and in contrast to the Sevier Desert and Mormon Mountains, crustally rooted faults are an important part of the story – and, for that reason, these examples are undoubtedly among the most important not included in this paper. Among issues to be resolved are: the manner in which brittle deformation relates to mylonitization (or not); the degree to which faults are folded or tilted to lower dip, as a result of progressive or unrelated deformation; the role of unusual materials such as talc in reducing frictional coefficients (e.g. Moresby seamount, Woodlark Basin; Floyd et al. 2001); and the significance of fluids in fault zones and more generally in crustal rheology (Axen 1992; Manatschal 1999; Collettini & Barchi 2002; Hayman et al. 2003; Collettini & Holdsworth 2004; Evans et al. 2004; Scholz & Hanks 2004). Wills & Buck (1997) examined the possible role of stress-field rotation in the development of rooted detachment faults, and concluded that applied stresses are not in general sufficient to create low-angle normal faults that would propagate to the Earth’s surface (cf. Westaway 1999). There will be no one-size-fits-all solution. Our problem as a community may be, at least in part, one of attempting to shoehorn unrelated phenomena into an overly simplified conceptual model. Experience over the past decade indicates that no matter how important an example is or was regarded, once questions are raised about its significance, other allegedly better examples come to the fore, along with all manner of secondary hypotheses for preserving the original interpretation. It remains

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to be determined whether the now-reinterpreted examples upon which we have worked will prove to be of general significance or (as some critics have asserted) peripheral to an understanding of detachment fault mechanics. The paradox of lowangle normal faulting is after all a frictional issue even in areas in which mylonitic rocks are also present, and most troublesome at the shallowest crustal levels. As a practical matter, hypotheses can be tested only with reference to specific examples, and if evidence is to be regarded as compelling in the manner implied by the quote with which this paper begins, it ought to withstand scrutiny, not simply lead to conclusions that are known in advance to be correct.

Appendix 1: Arguments for contrasting interpretations of Sevier Desert reflection (SDR) References 1

McDonald (1976); 2Allmendinger et al. (1983); 3Smith & Bruhn (1984); 4Von Tish et al. (1985); 5Mitchell & McDonald (1986); 6Mitchell & McDonald (1987); 7 Planke & Smith (1991); 8Anders & Christie-Blick (1994); 9Hamilton 1994; 10Allmendinger & Royse (1995); 11Anders et al. (1995); 12DeCelles et al. (1995); 13 Otton (1995); 14Coogan & DeCelles (1996); 15Otton (1996); 16Wills & Anders (1996); 17Anders et al. (1998a); 18 Coogan & DeCelles (1998); 19Wills & Anders (1999); 20Anders et al. (2001); 21Stockli et al. (2001); 22 Hintze & Davis (2003); Davis (2003); 23Niemi et al. (2004); 24Wills et al. (2005); 25DeCelles & Coogan (2006).

Detachment interpretation 1 – 4,7

Lateral continuity of SDR inclined at 118 over 7000 km2 beneath Sevier Desert basin (SDB) and (down dip) within deformed Neoproterozoic and Palaeozoic rocks of Mesozoic orogen (Fig. 3). 1 – 4,7 Accounts for downward termination of high-angle normal faults at the SDR (Fig. 3). 1 – 7,10,12,14,18,21 Provides explanation for the SDB and for normal sense offset of Mesozoic structures by ,47 km (detachment may in part reactivate a Mesozoic thrust fault1,5). 2,10 Explains gravity high and culmination in crystalline basement rocks beneath the SDB. 1,2,4,14,18,21,23 Stratigraphic growth consistent with displacement on detachment after 28– 26 Ma; youngest sediments involved in deformation are Holocene. 10,21 Fission track ages from footwall rocks of Canyon Range (19– 15 Ma; apatite) and ARCO Meadow Federal well (13.0 + 1.0 Ma to 10.8 + 0.9 Ma; zircon), and from hanging-wall rocks in Cominco

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American Federal well (8.5 + 2.2 Ma; apatite) are attributed to mid- to late Miocene displacement and tectonic exhumation along the hypothesized detachment. 13,15 Brittle deformation in Canyon Range on eastern margin of the SDB interpreted as outcrop expression of hypothesized detachment.

Unconformity interpretation Absence of anticipated deformation problematical 8

No increase in abundance of microcracks in borehole cuttings of Cenozoic clastic sediments no more than 3 m above hypothesized detachment. 8,20 No evidence for ductile deformation in borehole core of Palaeozoic carbonate rock 12.8 m below same contact or in cuttings from as close as 3 m, in spite of a restored depth of at least 14 km and temperatures as high as 425 8C (initial conditions based upon an offset of ,47 km). 16,19 Deformation in Canyon Range at eastern margin of the SDB relates to mapped Mesozoic fold hinges and faults, and to slide blocks that are demonstrably surficial (the SDR projects above the level of outcrop9).

Difficulties with palinspastic reconstruction and reflection geometry 8,24

Re-evaluation of subsurface data indicates that the SDR is fortuitously aligned down dip with a Mesozoic thrust fault west of the northern SDB (Fig. 4a), and not present/imaged west of the southern SDB (Fig. 4b)6,7. 24 Published estimates of extension across the northern SDB (,47 km4,14,25) and southern SDB (6 km7) are markedly different in cross-sections only 20 km apart. 23 Mesozoic thrust faults are truncated by sub-Cenozoic unconformity at western margin of the southern SDB (Fig. 4b; no evidence for termination against the SDR, as required by published palinspastic reconstructions). 17,24 Widespread onlap without growth at the western margin of the SDB (Fig. 4) contrasts with previous interpretations of stratigraphic growth in seismic data (see below) and with expected stratal geometry for a supradetachment basin. 24 SDR offset by high-angle normal faults with up to 500 m of stratigraphic separation at the eastern margin of the SDB, and up to 750 m at the western margin (precluding a throughgoing detachment or placing constraints on its most recent movement; see below). 24 As much as 1.3 km of relief on the SDR in a north–south direction (50 km wavelength) contrasts with apparently planar geometry of more familiar east– west profiles, and lacks an adequate structural explanation (other than an assertion that they are primary corrugations).

Timing constraints and difficulties 6,17,20,23,24

The SDB is relatively long-lived: pre-Oligocene (oldest strata) to Holocene.

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11,17,23

Dipping reflections at the western margin of the SDB, interpreted as Oligocene (28– 26 Ma) and used to infer time of onset of displacement along detachment1,2,4,14,18,21, are likely within Neoproterozoic – Palaeozoic rocks of the Mesozoic orogen if a more plausible (higher) velocity of 3.2 km s21 is assumed for SDB fill at Gulf Gronning well. 24 A high-angle normal fault cutting the SDR is overlapped by strata as old as ?Miocene (providing an approximate younger bound on timing of hypothesized detachment). 11,19,20,24 Fission track ages of 19– 8.5 Ma are subject to multiple interpretations (see below), with no direct bearing on basin formation that began prior to the Oligocene. 20 No evidence for any lateral progression in ages of volcanic rocks above hypothesized detachment (Eocene– Oligocene, mid-Miocene and Pliocene– Holocene volcanism).

Alternative explanations 8,11,20,22,24

SDR is a regional unconformity of Palaeogene age (consistent with kilometre-scale relief) and, where present west of the SDB, a Mesozoic thrust fault. 8,17,24 Downward termination of high-angle normal faults and localized stratigraphic growth are due to syndepositional deformation of intrabasinal Oligocene salt (Fig. 4a). 11 Culmination in crystalline rocks beneath the SDB attributed to a ramp in a hypothesized intrabasement thrust fault (such faults are known in outcrop within Mesozoic orogen). 11,19,24 Fission track ages from Canyon Range (19–15 Ma) are attributed to exhumation of range as a result of displacement along high-angle normal faults; ages as young as 13.0– 10.8 Ma in ARCO Meadow Federal No. 1 well and 8.5 Ma in Cominco American Federal well may reflect contemporaneous magmatism. 24 Origin of SDB: an erosional piggy-back basin within orogen, tilted westward by out of sequence thrusting, augmented by extension- or slab-related regional subsidence, sediment loading and minor high-angle normal faulting.

Appendix 2: Arguments for contrasting interpretations of Mormon Peak and associated detachments

15 Anderson & Barnhard (1993a); 16Anderson & Barnhard (1993b); 17Axen (1993); 18Carpenter & Carpenter (1994); 19 O’Sullivan et al. (1994); 20Wernicke (1995); 21Axen & Bartley (1997); 22Stockli (1999); 23Axen (2004); 24Anders et al. (2006); Walker et al. (2007).

Rooted detachments of regional scale 3,4,6,8 – 10,12 – 14,17,20,21,23

Mormon Peak, Tule Springs and Castle Cliff detachments (MPD, TSD and CCD) are interpreted to accommodate 54 + 10 km of crustal extension, with displacements of approximately 25 km, up to 7 km and approximately 24 km, respectively. 13,17,23 TSD follows portion of Tule Springs thrust (a Mesozoic thrust de´collement) for more than 10 km in direction of detachment transport. 4,6,8,13,17,23 Palinspastic restoration implies initial dips of 208 –288 to a depth of at least 6 km (MPD); 38 –158 to a depth of 2 –5 km, steepening down dip at greater depths (TSD); and 328 to a depth of at least 7 km (CCD); present dips are up to 208E and 208W, and with a range of azimuths (MPD), approximately 08 (TSD) and 118 (CCD). 4,6,8,10,13,15,16,21,23 Isostatic rebound of footwall results in uplift, tilting, folding (‘rolling hinge’) and in tendency for detachment to become more gently dipping (or dipping both eastward and westward in case of MPD); near concordance of steeply inclined Oligocene– Miocene strata with Cretaceous– Palaeogene (?) beneath the CCD taken to demonstrate a post-Laramide origin (i.e. not related to crustal shortening). 17,23 Normal faults with up to 3 km of stratigraphic separation in the hanging wall of the TSD do not offset footwall stratigraphic contacts. 6,13,17 Relative ages of detachments (MPD . TSD . CCD) are interpreted on the basis of offset of the MPD and TSD by high-angle normal faults that are assumed (with some observational support in Tule Springs Hills) to have been coeval with displacement on structurally lower detachments. 13 MPD post-dates rhyolitic tuff correlated with Kane Springs Wash volcanic suite (c. 15– 12 Ma). 19,22 Apatite fission track ages in crystalline rocks from Beaver Dam Mountains (16 Ma) are attributed to tectonic exhumation by the 22,23CCD (on the face of it, inconsistent with an age younger than the MPD based upon inferred fault geometry).

Rootless localized slide blocks

References

Geometrical difficulties with rooted detachment interpretation

1

24

Cook (1960); 2Tschanz & Pampeyan (1970); 3Wernicke (1981); 4Wernicke (1982); 5Novak (1984); 6Wernicke et al. (1985); 7Hintze (1986); 8Wernicke & Axen (1988); 9 Wernicke et al. (1988); 10Axen & Wernicke (1989); 11Carpenter et al. (1989); 12Wernicke et al. (1989); 13Axen et al. (1990); 14Wernicke et al. (1990);

Each detachment crops out in a series of isolated exposures (original detachment contiguity is a matter of hypothesis rather than observation). 11,18,24 Seismic reflection profile 4-4A, which obliquely intersects western flank of Beaver Dam Mountains, shows that the range is bounded by a high-angle

REGIONAL DETACHMENT FAULTS normal fault (Figs 5a & 9); the CCD is either surficial or offset to a depth of at least 4 seconds two-way travel time (.5.5 km); neither interpretation is consistent with regional detachment illustrated in Figure 5b; the CCD also not imaged in section 5-5A (fig. 11 of Carpenter & Carpenter 1994; see Fig. 5 for the location). The MPD cuts high-angle normal faults in both the hanging wall and footwall, challenging an assumption used to infer relative ages of detachments. The TSD is especially problematic owing to an interpreted initial dip of 38 –158 and a present-day dip of approximately 08 (inclinations that are appreciably less than the frictional lock-up angle). Folding attributed to footwall deformation in the Beaver Dam Mountains is discordant with the rangebounding fault system; Mesozoic origin cannot be excluded. Unclear why MPD was domed, the TSD was folded by a rolling hinge and the CCD was tilted to a lower dip but remained essentially planar (Fig. 5b).

Difficulties related to character and kinematics of deformation 18,24

Polymictic conglomerate is present widely at the MPD as a 0.1–1 m-thick layer intimately involved in deformation, filling a network of clastic dykes above detachment, with grading, flow banding and internal erosional features, and, with one possible exception, no evidence for more than one generation of conglomerate or dyke intrusion; the conglomerate is compositionally distinct from Cretaceous synorogenic deposits, and not likely to have remained unlithified for about 50 million years at a depth of several kilometres; the geometry, textures and structure of conglomerate bodies are not consistent with karst infill. 24 Asymmetrical deformation: rocks above MPD are pervasively brecciated; those below are virtually undeformed. 24 Microfracture density in isolated quartz grains in conglomerate is lower than typically the case in fault zones; microfractures are abundant in quartz grains of Cambrian Tapeats Sandstone beneath MPD, but with no significant variation with distance from detachment (which would be expected if the two were related). 24 Footwall carbonate rocks are brecciated in thin zone (,1 m), with no evidence for other deformation typically observed along faults of large displacement (mylonites, cross-cutting veins, reduced grain size, preferred orientation of calcite twinning). 24,25 Kinematic indicators (slickenlines at detachment surface, and minor faults involving the conglomerate) diverge from published regional extension direction (2558 + 108; Fig. 8), and correspond approximately with modern dip direction of the MPD.

Alternative explanations 1,2,7,11,18,24,25

The MPD and CCD are basal contacts of several rootless slide blocks; rootedness of the TSD,

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also regarded as a gravity slide in early mapping2, is in doubt by association. 24 All of the characteristics of conglomerate layer and dykes at the MPD are duplicated at known slide blocks, and are consistent with fluidization during block emplacement in a single catastrophic event (for each example). 24 Asymmetrical pattern of brecciation at the MPD, a knife-sharp detachment surface, and up to several centimetres of gouge are typical of known slide blocks. 24 Expected ‘toes’4,6 are absent because structures mapped as the MPD and TSD are as old as middle Miocene, and toe regions of interpreted slide blocks are inferred to have been offset by high-angle normal faults, eroded away or buried by younger gravels. 24,25 Varied orientation of kinematic indicators is consistent with existence of more discrete blocks than implied by current detachment terminology in Mormon Mountains. 24 Age of MPD is not well constrained: youngest westtilted volcanic unit at the NW flank of Mormon Mountains yields 40Ar/39Ar age of 13.61 + 0.06 Ma, and overlies a sequence of volcanic rocks as old as 23.27 + 0.04 Ma, with dips increasing with age; nearhorizontal to east-tilted ash-flow tuffs in Meadow Valley Wash, west of Mormon Mountains are 14.49 + 0.02 Ma (M. H. Anders unpubl.), consistent with ash-flow tuff ages of ,14 Ma5 from Kane Springs Wash caldera; our interpretation is that tilting of volcanic rocks along west flank of Mormon Mountains relates to high-angle normal faulting documented from mapping and seismic reflection profiles (C. D. Walker unpublished data). 19 Fission track ages from Beaver Dam Mountains (16 Ma) are also attributed to exhumation of range as a result of displacement along high-angle normal fault.

Appendix 3: Arguments for and significance of contrasting interpretations of the depositional setting of the Eagle Mountain Formation (middle Miocene) References 1

Stewart (1967); 2Wright & Troxel (1967); 3Stewart et al. (1968); 4Stewart et al. (1970); 5Wright & Troxel (1970); 6 Stewart (1983); 7Prave & Wright (1986a); 8Prave & Wright (1986b); 9Stewart (1986); 10Wernicke et al. (1988); 11Snow & Wernicke (1989); 12Corbett (1990); 13 Wernicke et al. (1990); 14Stevens et al. (1991); 15 16 Holm et al. (1992); Snow (1992a); 17 18 Snow (1992b); Stevens et al. (1992); 19Wernicke (1992); 20Snow & Wernicke (1993); 21Stone & Stevens (1993); 22Topping (1993); 23Wernicke et al. (1993); 24 Brady et al. (2000b); 25Snow & Wernicke (2000); 26 Niemi et al. (2001).

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Alluvial fan and lacustrine 26

Conglomerate at Eagle Mountain and central Resting Spring Range contains a distinctive clast assemblage, including approximately 180 Ma leucomonzogabbro (,1 m) indistinguishable from rocks found in the Hunter Mountain batholith of Cottonwood Mountains more than 80 km to WNW; abundance of these clasts at some locations and as much as 50% modal plagioclase in some sandstone beds implies proximity to Cottonwood Mountains source at time of deposition (c. 15– 11 Ma). 24,26 Sedimentary characteristics of succession at Eagle Mountain indicate deposition at an alluvial fan less than 20 km from source (assuming upper limit on fan dimensions): rock avalanche deposits (conglomerate and breccia, characterized by angular clasts, local provenance, lack of bedding, and lack of matrix); non-cohesive debris flows (massive, clast-supported conglomerate, characterized by lateral continuity of individual units, poor sorting, angular clasts, and lack of clay-rich matrix); sheetflood couplets (granular sandstone and conglomerate, characterized by lithologic alternation, parallel bedding, upper flow regime structures and unimodal palaeocurrents); sheetflood sandskirt and tabular fluvial braidplain (sandstone and pebby sandstone, characterized by lateral continuity of bedding, lithological alternation, and planar and trough cross-stratification); lacustrine (siltstone and fine sandstone, characterized by lack of mudcracks or evaporitic horizons and lack of current-related structures; and micritic limestone, characterized by algal laminae and lenses of rippled fine sandstone). ‘Upward progression from rockfall and/or rockslide to debris flow, sheetflood, and sandskirt facies is consistent with a depositional system that evolved from a relatively small drainage area to a larger one, and. . . with upward change in clast derivation from local bedrock to a more distal source.’ 26 Tectonic transport of Cottonwood Mountains source more than 80 km towards 2938 and away from Eagle Mountain implies more than 400% extension of the upper crust since 11–12 Ma. 1,3,4,6,9 – 11,13,15 – 17,19,20,22 – 25 Interpretation consistent with other estimates of extension based upon restoration of Neoproterozoic and Palaeozoic isopachs and facies transitions, Mesozoic thrust faults and folds, palaeoisothermal surfaces, and Miocene rock-avalanche deposits.

Fluvial and lacustrine Overall characteristics of succession revealed by 13 measured sections and mapping of physical surfaces at Eagle Mountain: (1) basal erosion surface with metres to tens of metres of local erosional relief; (2) locally derived monolithological carbonate breccia and sandstone

(140 m), onlapping basal unconformity; (3) lacustrine siltstone and diamictite (c. 10 m); (4) braided fluvial conglomerate, sandstone, siltstone and minor limestone, pervasively channelized, with abundant trough crossstratification, fining-upwards successions up to 1 m to more than 10 m thick, and overall upwards coarsening over 90 m; (5) disorganized to diffusely stratified, heterolithic conglomerate, with least well-stratified facies filling incised valleys up to 14 m deep; and (6) mostly lacustrine sandstone, siltstone and minor limestone up to 140 m thick, characterized by relatively tabular bedding, parallel stratification to low-angle cross-stratification, current and wave ripples, soft-sediment deformation and rare tufa. ‘Rock avalanche deposits’ of Niemi et al. (2001) are reinterpreted as valley fill; ‘noncohesive debris flows’, ‘sheetflood couplets’ and ‘sheetflood sandskirt’ are fluvial; ‘tabular fluvial braidplain deposits’ are lacustrine event layers (delta front). Rounded clasts of leucomonzogabbro are found exclusively within interpreted fluvial deposits. If deposition took place in a fluvial– lacustrine rather than alluvial fan –lacustrine setting, distribution of Eagle Mountain Formation is hypothesized to reflect midMiocene drainage, with no significance for either magnitude or direction of crustal extension. Local abundance of Hunter Mountain detritus may be due to: (1) preferential erosion owing to local topography or climate-related factors; (2) episodic reaming out of valley/channel system; or (3) relative resistance to breakage during transport.

Difficulties with other evidence for extreme extension between Cottonwood Mountains and Nopah Range 2,5,7,8,12,14,18,21

Apparent convergence of disparate structural, stratigraphic, geochronological and thermochronological constraints is subject to circular reasoning. Difficulties include: uncertainty in structural correlation (thrust faults and folds); spatial variability of stratigraphic thickness that limits confidence in both location and orientation of isopachs; imprecisely defined facies transitions; and assumptions needed about the pre-extensional configuration of all markers. Inferences based on restoration of Miocene rock-avalanche deposits22 hinge on whether provenance can be tied uniquely to specific source areas. Our research has been supported by the National Science Foundation (EAR 99-02782), by the donors of the American Chemical Society Petroleum Research Fund (39706-AC), by a Geological Society of America graduate student grant to B. Renik, and by creative piggybacking. S. Wills and B. Renik acknowledge National Science Foundation graduate research fellowships. We thank Maureen Anders for assistance in drafting, G.S. Lister and an anonymous reviewer for helpful suggestions, and the editors for their patience, which we tested at least as much as any hypothesis. Lamont-Doherty Earth Observatory Contribution number 7011.

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Effects of initial weakness on rift architecture ¨ LLER1 & L. MORESI2 S. DYKSTERHUIS1, P. REY1, R. D. MU 1

University of Sydney School of Geosciences, Baxter Building (H11), University of Sydney, NSW 2006, Australia (e-mail: [email protected])

2

Room 301, Building 28, Monash University Clayton Campus, Victoria 3800, Australia Abstract: Much work has been conducted investigating the primary factors controlling rift architecture, using both computational and laboratory methods. Here, we examine the effects and relative importance that different types of initial weaknesses have on extensional lithospheric deformation style. We find that the type of initial weakness included in the model plays a primary role in determining the subsequent rift mode. Models using single localized weak seeds produce symmetric narrow rifts, irrespective of whether or not strain softening is included, while models that include regions of diffuse weaknesses produce a wide rift mode. Models that include an initial weak fault tend to produce more asymmetric rifting leading to a core complex mode. By distributing the strain laterally over large areas, the ductile lower crust tends to oppose the narrow rift mode forced by the initial weakness. These results suggest that initial weaknesses may play a major role in determining the mode of rifting. Our numerical experiments confirm that low-angle faults can form as the result of rotation of initially high-angle faults. While previous studies have suggested that the rheological and thermal profiles of the lithosphere play the most important role in rift mode determination, our results illustrate that initial weaknesses could play a major role in rift mode determination, highlighting the need to make initial weaknesses a primary consideration when modelling the extensional deformation of the lithosphere.

Rift modes and rift architectures have been the subject of many studies using both laboratory (Brun et al. 1994; Brun & Beslier 1996; Michon & Merle 2003) and computational methods (Buck 1991; Lavier et al. 1999; Huismans & Beaumont 2003; Nagel & Buck 2004; Wijns et al. 2005). While most of these studies have focused on the importance of continental rheology and thermal regime, little consideration has been given to the role initial weakness may have in inducing stress and strain localization (Dunbar & Sawyer 1989). We present a quantitative analysis of the effects initial weaknesses have on the mode of continental extension and the geometry of rift architecture. Having tested various forms of localized and randomly distributed weakness, we find that the type of initial weakness embedded in the model represents a primary controlling factor of continental rift architecture.

Continental extension modes and rift architectures Regardless of the origin of the forces driving extension, continental rifts have been classified into three categories (Buck 1991; Ruppel 1995; Corti et al. 2003). In the narrow rift mode, exemplified by the East African System, extensional deformation localizes along typically 100–150 km-wide, continuous or segmented, rifts. In the wide rift

mode, such as the Basin and Range Province, extensional deformation is more diffuse, spreading over an horizontal length-scale several times larger than the thickness of the lithosphere. In contrast to narrow rifts, wide rifts display relatively small gradient in crustal thickness. The Aegean domain exemplifies the core complex mode of continental extension. As with the wide rifts, core complexes display little variation in crustal thicknesses; however, they differ from the wide rift mode in that deeper crustal levels are exhumed to the surface forming a domical metamorphic core wrapped by a low-angle normal fault connected laterally to a crustal-scale detachment fault. The de´collement/detachment system is overlaid by low-grade rocks and detrital sediments (Coney 1980). The consensus is not yet reached on the significance of the three different modes of continental extension. Some authors see in the core complex mode a mere variant of wide rift mode (e.g. Brun 1999) whereas, for others, core complex and wide rift modes represent the result of fixed boundary collapse and free boundary collapse, respectively (Rey, et al., 2001). For some, the three modes may represent various stages of long-lasting continental extension (Olsen & Morgan 1995). Rheology is commonly accepted as a key parameter. Wide rift mode is thought to be promoted by power-law rheology that dominates at higher temperature and low strain rate, whereas plastic deformation, which

From: KARNER , G. D., MANATSCHAL , G. & PINHEIRO , L. M. (eds) Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 443– 455. DOI: 10.1144/SP282.18 0305-8719/07/$15.00 # The Geological Society of London 2007.

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dominates under lower temperature and high strain rate, is thought to favour the development of narrow rifts (Bassi 1991, 1995). Continental extension modes are also debated in terms of rheological stratification of the continental lithosphere. In particular, the depth of the brittle–ductile interface in the crust, the coupling between brittle and ductile layers, as well as the number of mechanically contrasting layers in the lithosphere, are of primary importance in controlling the style of continental extension (Allemand et al. 1989, 1991; Allemand & Brun 1991; Brun & Tron 1993; Benes & Davy 1996; Brun & Beslier 1996; Huismans et al. 2005). Rifting modes are also discussed in terms of competition between gravitational forces and driving forces (i.e. integrated lithospheric strength), somehow related to the relative thicknesses of the brittle and ductile layers (Benes & Davy 1996; Brune & Ellis 1997). There is a lot of overlap between the various interpretations of the three modes of continental extension and overall they all point toward the conclusion that continental geotherm, rheology and strain rate conspire to dictate the rift mode. Many studies have also emphasized the role of localized initial weakness (as thermal anomalies, compositional heterogeneities or both) in localizing the development of narrow rifts (Fletcher & Hallet 1983; Braun & Beaumont 1987; Dunbar & Sawyer 1989; Bassi 1991, 1995) or promoting the development of core complexes (Brun et al. 1994; Brun 1999). We examine here a range of rheological models in which localized and diffuse rheological heterogeneities are superimposed onto coupled and decoupled rheological profiles.

Model description We use the Ellipsis code, a Lagrangian integration point finite-element code capable of tracking time-dependant variables embedded in an Eulerian mesh, allowing for accurate tracking of surfaces and boundaries through time (Moresi et al. 2001, 2002). While the parameters in the model are scaled to achieve maximum computational efficiency (see the Appendix), ‘real-world’ value parameters are listed in Table 1. Our continental lithosphere is 450 km long and 140 km thick, and includes a 35 km-thick continental crust. Extension is driven by a constant velocity condition applied to the right-hand boundary of the model (Fig. 1), giving an initial strain rate of 1  10215 s21 over the horizontal length of the model. A free-slip boundary condition is applied at the top and bottom surfaces. The basal heat flux of 20 mW m22 and a crustal radiogenic heat production (1.2  1026 W m23) was adjusted to achieve a Moho temperature of 550 8C while the 1300 8C isotherm is reached at 140 km. We use the standard rheological

Table 1. List of natural (non-scaled) model parameter values Parameter

Natural value

Density upper crust Density lower crust Density mantle Gravitational acceleration Cohesion Friction coefficient Maximum strain weakening Saturation strain Thermal diffusivity Crustal heat production Specific heat Basal heat flux

2700 kg m23 2800 kg m23 3300 kg m23 10 m s22 20 MPa 0.4 80% 0.5 1  1026 m2 s21 3.2  1026 W m23 103 J kg21K21 20  1023 W m22

profile of Brace & Kohlstedt (1980) for the continental lithosphere in which a Mohr–Coulomb failure law, augmented by a strain-weakening function, is the dominant failure mechanism at low temperature and high strain rate, whereas the Frank-Kaminetskii approximation (Frank-Kamenetskii 1955) of the temperature-dependent Arrhenius (h) viscosity is used at high temperature and low strain rate: hðTÞ ¼ h0 eT0 T

ð1Þ

where T is the temperature, h0 is the viscosity at a temperature of 0 8C and T0 is a constant. Coupling of the crust to the mantle is implemented by altering the values of h0 and T0. Coupled models have values of h0 and T0 chosen such that the entire crust is initially in the brittle regime, while decoupled models have values of h0 and T0 chosen such that the lower crust is initially in the ductile regime (Fig. 1). Strain weakening in the Mohr–Coulomb failure law is implemented via a power-law function f(1): f ð1Þ ¼

1  ð1  aÞð1=10 Þ1n a

1 , 10 1  10

ð2Þ

where 1 is the accumulated plastic strain, 10 is the saturation strain beyond which no further weakening takes place, 1n is an exponent that controls the shape of the function and a is a maximum value of strain weakening beyond which no further weakening occurs. In our models strain weakening is linear (1n ¼ 1) with a maximum of 80% weakening (equivalent to a 1n value of 0.2) taking place at a saturation strain (10) of 0.5 (Fig. 1), similar to the values used in previous work (Huismans & Beaumont 2003; Wijns et al. 2005). Initial weaknesses are implemented in the following methods. † Single weak seeds: square regions of lowviscosity material are included in the lithosphere,

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Vext

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Vext

Fig. 1. (Top) Model geometry showing material layer thicknesses. Upper and lower crust are a wet quartz composition with a density of 2700 and 2800 kg m23, respectively, while the mantle is a dry olivine composition with a density of 3300 kg m23. Extension is driven by a boundary velocity Vext chosen to achieve a strain rate of 1  10215 over the model. The initial temperature field, controlled by a basal heat flux of 20 mW m22 and radiogenic heat production of 1.2  1026 W m23 in the crust, is initially laterally uniform and increases with depth from 0 8C at the surface, 550 8C at the base of the crust and to 1300 8C at the base of the model. Erosion and sedimentation are not included in the model. (Bottom left) Representative yield strength envelopes of a coupled and decoupled model for a strain rate of 1  10215 with initial strength denoted by solid lines and strain weakened denoted by dashed lines. (Bottom right) Strain-softening behaviour showing strength weakening from 100 to 20% after an accumulated strain of 0.5, after which no further weakening occurs. Dashed lines show the effect of the exponential parameter (En) on the curve (see equation 2).

possibly representing regions of higher radiogenic heat production (if anomaly is located in the crust), regions of locally reduced viscosity due to thermal/pressure perturbations or regions of weakness caused by a material inhomogeneity. † Regions of diffuse weakness: random uniformly distributed regions of prestrained, preweakened material are dispersed through the crust corresponding possibly to accretionary continental crust with crustal inhomogeneities or localized regions of higher radiogenic heat production decreasing the strength of the continental crust at discrete points over a large region. † Faults: discrete regions of low-viscosity material cutting the upper and lower crust at an angle of 458 are included representing a fault or a planar weakness such as a suture zone.

Results We test the differing types of weakness on models where the crust is coupled to the mantle (brittle lower crust) representative of older continental

lithosphere, as well as on models where the crust is decoupled from the mantle (ductile lower crust), representative of younger continental lithosphere (Brun 1999).

Single localized weak seeds In these models, a single square of low-viscosity material represents a localized weak seed. Different models were tested by embedding the weak seed at various depths within the lithosphere. Upon extension, strain focuses around the weak seed resulting in the development of a narrow rift (Figs 2–8). Previous studies (Buck 1991) have demonstrated that narrow rift mode preferentially evolves in lithosphere with high integrated strength, common of cool and thick lithosphere. Our modelling indicates that a narrow rift mode can also develop as a result of the inclusion of a single weak seed in the lithosphere. We tested the sensitivity of the model to the magnitude of the weakness and found that a seed 10 times weaker than that of the surrounding material is sufficient to localize strain. In coupled models, conjugate

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Fig. 2. Coupled upper crustal localized single-seed weakness model including strain weakening after extension of 10, 20, 35 and 50%. Coupling is controlled by variation of the T0 parameter in the viscosity approximation (see text). Upper crust is orange, lower crust is brown, mantle is red while the imposed weak seed is purple. (a) shows strain (blue colouring) while (b) shows only material colouring at the same time deformation step.

normal faults originating from the weak seed lead to one single rift zone whose width corresponds to the depth of the seed (Figs 2–5). This narrow rift accommodates ongoing extension. Allemand & Brun (1991) found that conjugate rift faults meet at the base of the brittle layer and therefore the thickness of the brittle layer controls the width of the rift basin. By placing the weak seed at increasing depth we are locally increasing the depth to the effective base of the brittle layer (the top of the weak seed) and subsequently increasing the width of the resulting rift basin. Decoupling the crust from the mantle (Figs 6–8) results in different rift architecture to that of the coupled models, although the narrow rift mode still accommodates extension. Regardless of where the seed is located, periodic but out-of-phase

zones of necking develop in the upper crust and upper mantle. A weak seed placed in the upper crust results in a small rift basin, the width being controlled by the depth to the crustal brittle – ductile transition. Lower crustal flow towards the rift centre feeds back into the upper mantle causing strain localization and necking. The displacement from upper mantle necking is then transmitted by lower flow inducing further faulting of the upper crust at distances corresponding to the depth of the mantle brittle–ductile transition. When a seed is placed in the ductile lower crust or in the upper mantle, two small narrow basins form in association with two conjugate normal faults. At the surface, the distance between the conjugate faults is about the depth of the crustal brittle–ductile transition. The distance between the two smaller basins is

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447

Fig. 3. Coupled lower crustal localized single-seed weakness model including strain weakening after extension of 10, 20, 35 and 50%.

equal to the depth of the brittle–ductile transition in the mantle. This may have implications for using basin structure to obtain information about the rheology of the rift, with the rift basin widths and distances between rift basins having a direct correlation to the depths of the crustal and mantle brittle–ductile transitions or to crustal heterogeneities. Huismans & Beaumont (2003) suggested that rift asymmetries can be caused by preferential localization of extension over a single weak seed being taken up along one conjugate fault of a rift as a result of the inclusion of strain weakening. The model set-up used here is very similar to that of Huismans & Beaumont (2003), however, our single seed models produce symmetric rift basins both with and without the inclusion of strain softening (Figs 4 & 5). This may be that our models retain computational symmetry while those of Huismans & Beaumont (2003) do not. This serves to illustrate the need for further benchmarking of available dynamic modelling codes. Rift basins formed including strain weakening produce more localized rift basins with

more pronounced topography owing to increased focusing of the strain along the conjugate faults of the rift, but the general rift architecture is the same between the two models. This was observed in all models, indicating that while strain weakening plays a role in modulating the topography of the rift basin, it is not of first-order importance in rift mode selection. The topography generated in the models, even models without strain weakening included, does seem to be exaggerated. In some cases rift flank topography exceeds 10 km. Inclusion of erosional and sedimentary processes (Burov & Poliakov 2001) in the model in the future may serve to help account for the over–estimated topography.

Region of diffuse weakness In these models we implemented the initial weaknesses as small, random (uniformly distributed) regions of weaker material confined to the central 350 km of the models. Upon extension, uniformly distributed weak seeds lead to a wider zone of extension

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Fig. 4. Coupled model including strain weakening with single-seed weakness imposed in the upper mantle. Notice rifting is symmetric in this model.

Fig. 5. Coupled model with single-seed weakness imposed in the mantle and no strain weakening. Rifting is still symmetric in this model; however, a shallower rift basin forms compared to models that include strain softening (Fig. 4).

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Fig. 6. Decoupled model results after extension of 10, 20, 35 and 50% including strain weakening with a localized single-seed weakness in the upper crust. Decoupling is controlled by variation of the T0 parameter in the viscosity approximation (see the text). Imposing a single seed weakness in the upper crust results in a W-type architecture of the narrow rift mode (see the text for discussion).

(c. 500 km after 50% extension) made of numerous, discrete and synchronous basins (Figs 9 & 10). In both coupled and decoupled models, each seed initially produces a set of conjugate normal faults. Upon ongoing extension, only a smaller number of normal faults remain active. The natural selection process is such that in the final configuration the basins are bounded by normal faults merging at the brittle–ductile transition. The final configuration in the decoupled model is very similar to that of our decoupled/single-seed models. However, the strain distribution is a significant difference. In the diffuse weakness model, the basins are synchronous and therefore they accommodate similar amount of extension. Overall, the lateral distribution of the stretching factor is rather homogeneous. This clearly contrasts with the wide extensional zone that results from a single weak seed in decoupled models. In the latter, progressively younger basins develop for the initial

weak seed. Consequently, older basins accommodate a larger amount of extension and the stretching factor is heterogeneous over the length of the model. This model may represent the diffuse rift mode of Ruppel (1995), where topographically high-standing features representing the footwalls of normal faults are interspersed with intervening valleys. The distance between rift-bounding faults and between adjacent rifts in the coupled model approximately corresponds to the distance to the depth of the brittle–ductile transition. In decoupled models the distance between rift-bounding faults corresponds roughly to the depth of the crustal brittle–ductile transition, while the distance between adjacent basins corresponds roughly to the depth of the brittle–ductile transition in the mantle. Although the parameters of our models place them within the narrow rift mode field of Buck (1991), the presence of diffuse weakness produces

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Fig. 7. Decoupled model results after extension of 10, 20, 35 and 50% including strain weakening with a localized single-seed weakness in the lower crust.

a wide rift mode via a series of adjacent narrow rifts, while the geometry of the rift fits the broad description of a wide or diffuse rift (Ruppel 1995). It could be argued that the temporal evolution of our wide rift model does not fit the description of a wide rift (Buck 1991), as strain is initially distributed in many regions with a wide rift mode active from the onset of rifting, instead of strain migrating through. Clearly, the definition of the wide rift mode needs to be revisited.

Faults Faults were implemented in the model as a thin strip of weak material inclined at 458, representing a planar feature such as a fault or a suture zone. The coupled fault model (Fig. 11) produces a strongly asymmetric rift and clearly demonstrates that an initially high-angle fault can rotate to become a low-angle detachment. This

result is consistent with the rolling hinge model of Lavier et al. (1999). As the lithosphere extends, rotation of the west-dipping high-angle fault results in a low-angle fault being exposed at the surface which has accommodated substantial movement. After about 10% extension a secondary, synthetic high-angle fault appears dipping east. Both faults continue to rotate with the final rift geometry (50% extension) showing two lowangle faults dipping in opposite directions, generated from one initially westward-dipping highangle fault. While single-seed weaknesses resulted in narrow rift mode, imposing a fault as initial the weakness favoured a core complex mode. In the coupled model (Fig. 11) a metamorphic core complex initiates with lower crust exposed at the surface over the eastern rift arm after around 35% extension. Following further extension, an oceanic core complex forms with the upper mantle lithosphere being

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Fig. 8. Decoupled model results after extension of 10, 20, 35 and 50% including strain weakening with a localized single-seed weakness in the upper mantle.

exposed at the surface after about 50% extension. Decoupling the crust and mantle (Fig. 12) results in lateral diffusion of extension. In a decoupled system the high-angle fault almost acts as a single weakness with a final rift architecture similar to that seen in the decoupled single upper crustal weak seed model (Fig. 6). Previous studies have required weak

rheological profiles, weak lower crust being a particular condition (Buck 1991; Brun et al. 1994; Wijns et al. 2005), as a requirement for the formation of core complexes. Here we have shown the core complex mode can occur in coupled crust as a result of the inclusion of an initial highangle fault.

Fig. 9. Coupled random weakness model results. Weak seeds (blue triangles) are uniformly randomly distributed through the crust between 50 and 400 km. Strain (blue colouring in a) is localized in many basins resulting a wide rift mode.

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Fig. 10. Decoupled model with weak seeds (blue triangles) uniformly randomly distributed through the crust between 50 and 400 km.

Fig. 11. Coupled model including strain weakening with a 458 fault imposed in the crust. Rifting is very asymmetric and a core complex forms after approximately 70% extension with mantle exposure at the surface. A characteristic feature of the Iberian margin is also demonstrated here with the upper continental crust directly overlying the mantle after approximately 35% extension (see the text for further discussion).

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Fig. 12. Decoupled model including strain weakening with a 458 fault imposed in the crust.

Conclusion While most previous work has concentrated on regional strength discontinuities (such as rheological layering, etc.), here we demonstrate that localized strength anomalies can result in localization of rifting and can also force selection of one of the three modes of rifting (narrow, wide and core complex). Our results suggest that the magnitude of these weaknesses does not have to be large, localized strength reductions to 1/10 of that of surrounding materials are enough to localize strain and onset the rifting process. An initial weakness implemented as a single seed will result in a narrow rift mode, a region of diffuse, random noise will result in a wide rift mode, while an initial weakness implemented as a fault with result in a core complex mode. Decoupling the upper crust and mantle by including a ductile lower crust acts to suppress the rifting mode favoured by the initial weakness, acting to diffuse the rifting over larger lateral areas. The inclusion/ omission of strain weakening does not affect the symmetrical evolution of our single seed models, as has

previously been suggested (Huismans & Beaumont 2003); it only serves to enhance the topography and create deeper rift basins owing to strain being preferentially more focused along the bounding faults of the rift. In single-seed coupled models, the spacing of rift-bounding faults is a function of the effective depth of the weak seed. In addition, in single-seed decoupled models where multiple rifts develop the spacing of adjacent rifts is controlled by the depth of the brittle–ductile transition in the mantle. In a coupled configuration of the diffuse weakness models, the distance between rift-bounding faults and between adjacent rifts roughly corresponds to the depth of the brittle –ductile transition. In the decoupled configuration the distance between riftbounding faults corresponds to the depth of the crustal brittle –ductile transition, while the distance between adjacent basins corresponds to the depth of the brittle–ductile in the mantle. The decoupled models show out-of-phase stretching in the upper crust and upper mantle. We suggest that the core complex mode of rifting can be achieved with the

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inclusion of a weak fault. This conclusion is an alternative to previous studies which found that the rheological profile dictates the formation of core complexes (Buck 1991; Wijns et al. 2005). It should be noted that both our and previous studies into factors controlling core complex formation are not mutually exclusive. The results presented here highlight that rift mode selection can be forced by more than one dominant model parameter.

Appendix: Scaling The ellipsis model input parameters are scaled from realworld values into dimensionless values in order to minimize computation time and increase the accuracy of the solution. We use the non-dimensional scaling approach given by: E ¼ N=S

ð3Þ

where E is a dimensionless ellipsis variable, N is the dimensional real-world parameter and S is a dimensional scaling factor. Scaling from real model length dimensions of 450 km wide by 150 km deep is carried out using a length scaling factor (SL) of 1.5  105 m resulting in non-dimensional model geometry of 3 units wide by 1 unit deep. Using a thermal diffusivity scaling factor (Sk) of 1  1026 a time scaling factor (St) can be found using: St ¼ SL 2=Sk

ð4Þ 21

With a viscosity scaling factor (Sh) of 1  10 Pa and a gravity scaling factor (Sg) of 1 m/s2, a density scaling factor (Sr) can be found: Sr ¼

Sg  SL  St Sh

ð5Þ

Using a velocity scale (Su) defined by: Su ¼ SL =St

ð6Þ

a non-dimensional velocity of 67.5 is calculated to achieve a real-world strain rate of 1  10215 s21 over the length of the model. Temperature is scaled between non-dimensional values of 0.17 and 1 corresponding to temperatures of 273 and 1603 K, respectively, using a temperature scale (ST) of 1603 K.

References A LLEMAND , P. & B RUN , J. P. 1991. Width of continental rifts and rheological layering of the lithosphere. Tectonophysics, 188, 63– 69. A LLEMAND , P., B RUN , J. P., D AVY , P. & V ANDENDRIESSCHE , J. 1989. Symmetry and asymmetry of rifts and mechanism of lithosphere thinning.

Bulletin de la Socie´te´ Ge´ologique de France, 5, 445–451. B ASSI , G. 1991. Factors controlling the style of continental rifting – insights from numerical modeling. Earth and Planetary Science Letters, 105, 430– 452. B ASSI , G. 1995. Relative importance of strain rate and rheology for the mode of continental extension. Geophysical Journal International, 122, 195–210. B ENES , V. & D AVY , P. 1996. Modes of continental lithospheric extension: Experimental verification of strain localization processes. Tectonophysics, 254, 69– 87. B RACE , W. F. & K OHLSTEDT , D. L. 1980. Limits on lithospheric stress imposed by laboratory experiments. Journal of Geophysical Research, 85, 6248– 6252. B RAUN , J. & B EAUMONT , C. 1987. Styles of continental rifting: results from dynamical models of lithospheric extension. In: B EAUMONT , C. & T ANKARD , A. J. (eds) Sedimentary Basins and Basin-forming Mechanisms. Canadian Society of Petroleum Geologists, 241–258. B RUN , J. P. 1999. Narrow rifts versus wide rifts: inferences for the mechanics of rifting from laboratory experiments. Philosophical Transactions of the Royal Society of London Ser. A – Mathematical Physical and Engineering Sciences, 357, 695– 710. B RUN , J. P. & B ESLIER , M. O. 1996. Mantle exhumation at passive margins. Earth and Planetary Science Letters, 142, 161– 173. B RUN , J. P. & T RON , V. 1993. Development of the North Viking Graben – inferences from laboratory modeling. Sedimentary Geology, 86, 31–51. B RUN , J. P., S OKOUTIS , D. & V ANDENDRIESSCHE , J. 1994. Analogue modeling of detachment fault systems and core complexes. Geology, 22, 319–322. B RUNE , J. N. & E LLIS , M. A. 1997. Structural features in a brittle– ductile wax model of continental extension. Nature, 387, 67–70. B UCK , W. R. 1991. Modes of continental lithospheric extension. Journal of Geophysical Research – Solid Earth, 96, 20,161– 20,178. B UROV , E. & P OLIAKOV , A. 2001. Erosion and rheology controls on synrift and postrift evolution: Verifying old and new ideas using a fully coupled numerical model. Journal of Geophysical Research – Solid Earth, 106, 16,461–16,481. C ONEY , P. J. 1980. Cordilleran metamorphic core complexes: an overview. In: C RITTENDEN , M. D. J R , C ONEY , P. J. & D AVIS , G. H. (eds) Cordilleran Metamorphic Core Complexes. Geological Society of America, Boulder, CO, 7– 31. C ORTI , G., B ONINI , M., C ONTICELLI , S., I NNOCENTI , F., M ANETTI , P. & S OKOUTIS , D. 2003. Analogue modelling of continental extension: a review focused on the relations between the patterns of deformation and the presence of magma. Earth-Science Reviews, 63, 169–247. D UNBAR , J. A. & S AWYER , D. S. 1989. How preexisting weaknesses control the styles of continental breakup. Journal of Geophysical Research, 94, 7278– 7292. F LETCHER , R. C. & H ALLET , B. 1983. Unstable extension of the lithosphere: A mechanical model for Basin and Range structure. Journal of Geophysical Research, 88, 7457– 7466. F RANK -K AMENETSKII , D. 1955. Diffusion and Heat Exchange in Chemical Kinetics. Princeton University Press, Princeton, NJ.

INITIAL WEAKNESS ON RIFT ARCHITECTURE H UISMANS , R. S. & B EAUMONT , C. 2003. Symmetric and asymmetric lithospheric extension: Relative effects of frictional– plastic and viscous strain softening. Journal of Geophysical Research – Solid Earth, 108. doi:10.1029/2002JB002026. H UISMANS , R. S., B UITER , S. J. H. & B EAUMONT , C. 2005. Effect of plastic– viscous layering and strain softening on mode selection during lithospheric extension. Journal of Geophysical Research – Solid Earth, 110. doi:10.1029/2004JB003114. L AVIER , L. L., B UCK , W. R. & P OLIAKOV , A. N. B. 1999. Self-consistent rolling hinge model for the evolution of large-offset low-angle normal faults. Geology, 27, 1127– 1130. M ICHON , L. & M ERLE , O. 2003. Mode of lithospheric extension: Conceptual models from analogue modeling. Tectonics, 22. doi:10.1029/2002TC001435. M ORESI , L., D UFOUR , F. & M UHLHAUS , H. B. 2002. Mantle convection modeling with viscoelastic/brittle lithosphere: Numerical methodology and plate tectonic modeling. Pure and Applied Geophysics, 159, 2335–2356.

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Incompressible viscous formulations for deformation and yielding of the lithosphere ¨ HLHAUS2, V. LEMIALE1 & D. MAY1 L. MORESI1, H.-B. MU 1

School of Mathematical Sciences, Building 28, Monash University Clayton, Victoria 3800, Australia (e-mail: [email protected]) 2

Department of Earth Sciences, University of Queensland, St Lucia, Queensland 4072, Australia

Abstract: In models of crustal deformation at a scale of a few tens of kilometres, it is appropriate to use a Mohr–Coulomb yield criterion for lithospheric failure based on the idea that frictional slip occurs on whichever one of many randomly oriented planes happens to be favourably oriented with respect to the stress field. The Drucker– Prager yield criterion has very similar characteristics to the Mohr–Coulomb criterion but is more straightforward to implement, particularly in the context of large-scale fluid-mechanical deformation of the coupled system of mantle, lithosphere and crust. As such models become more sophisticated it is important to be able to use whichever failure model is appropriate to a given part of the system. We have therefore developed a way to represent Mohr– Coulomb failure within the mathematical framework used by mantle-convection fluid dynamics codes. The new formulation is based on the conceptual picture of lithospheric failure incorporated in the Anderson model of fault development and reproduces shear band angles predicted by this model. We use an transversely isotropic viscous rheology (a different viscosity for pure shear to that for simple shear) to define a preferred plane for slip to occur given the local stress field. The simple-shear viscosity and the deformation can then be iterated to ensure that the yield criterion is always satisfied. We assume the Boussinesq approximation can be applied in the model – neglecting any effect of dilatancy on the stress field. An additional criterion is required to ensure that deformation occurs along the plane aligned with maximum shear strain rate rather than the perpendicular plane, which is formally equivalent in any symmetric formulation. We compare the behaviour of this formulation to the Drucker–Prager failure model for twodimensional (2D) layered brittle– ductile systems. The pattern of shear bands is found to be quite similar from each model, although some differences are seen in the characteristic spacings for ‘equivalent’ values of the material properties. In 3D similar patterns are seen but are modulated by additional 3D structure.

On a global scale, the continents drift as an integral part of the surface thermal boundary layer of the convecting mantle. They have retained a distinct identity within the mantle flow for billions of years while developing a strong physical and chemical fabric along the way. Motions in the mantle are described by the equations of fluid dynamics for very large deformation. The rheology needed to describe deformation in the lithosphere is strongly non-linear, and near the surface where temperatures are less than approximately 600 8C it becomes necessary to consider the role of elasticity (Watts et al. 1980a). Plate boundaries are strongly correlated with seismic activity (e.g. Barazangi & Dorman 1969), which highlights the importance of considering failure of the lithosphere in modelling at the plate scale. At the mantle–plate scale, it is necessary to consider the fluid convection of the mantle and the history-dependent viscoelastic–brittle behaviour of the continental crust as a single coupled system. The requirements for a geological simulation code

are therefore an ability to track boundaries and interfaces through extremely large deformation, including fluid convection, of non-linear history-dependent materials. The wide range of physical and temporal scales, and the many coupled physical processes, also impose a need for computational efficiency. The methods should also be very flexible in the rheological laws that they can treat. At the scale of the lithospheric thickness, the details of the failure process itself cannot be ignored in modelling the observed deformation. For example, fault orientation, the order of formation and the occurence of reactivation in response to the lithospheric stress field are primary concerns of structural geology. There is a long tradition of modelling brittle–viscoelastic deformation in the lithosphere using methods derived from solid mechanics plasticity formulations (e.g. Poliakov et al. 1993; Braun & Sambridge 1994; Poliakov & Hermann 1994; Willett & Beaumont 1994; Braun & Beaumont 1995; Ellis et al. 1995; Lavier et al. 2000).

From: KARNER , G. D., MANATSCHAL , G. & PINHEIRO , L. M. (eds) Imaging, Mapping and Modelling Continental Lithosphere Extension and Breakup. Geological Society, London, Special Publications, 282, 457– 472. DOI: 10.1144/SP282.19 0305-8719/07/$15.00 # The Geological Society of London 2007.

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Solid mechanics approaches are usually based upon a viscoelastic displacement formulation and concentrate on the need to predict very precisely the onset and early evolution of the emergent structures of material failure. Where the post-failure deformation needs to be followed through considerable strain, these formulations may be less appropriate than a fluid dynamical approach. Fullsack (1995) developed an incompressible viscous model of lithospheric deformation that incorporated a Drucker–Prager yield criterion to address this need, and Lenardic et al. (2000) demonstrated a similar formulation in a study of mantle derived deformation of cratons and mobile belts. In this paper we will discuss the implementation of yield criteria into viscous fluid formulations of geological deformation. This approach has been most widely used in the mantle dynamics community, but increasingly, the mantle models are being used at much higher resolution and more regional scales to model plate-boundary processes.

where sij is the stress tensor, tij is the deviatoric stress, p is pressure, assumed to be positive in compression, g is the acceleration due to gravity in the vertical direction given by i ¼ 3, r0 is material density at a reference temperature, aT is the coefficient of thermal expansivity and T is temperature. The notation xi indicates partial differentiation with respect to the co-ordinate direction i, summation over repeated indices is assumed. We have also assumed that the variation in density only needs to be considered in the buoyancy forces, not in the continuity equation (the Boussinesq approximation: Boussinesq 1903). Substituting an isotropic viscous constitutive law in equation (1)   1 @ui @uj þ ð2Þ tij ¼ 2hDij ; Dij ¼ 2 @xj @xi gives the equation of motion: ð2hDij Þ,i  pi ¼ gr0 ð1  aT TÞdi3

ð3Þ

Physical model In plate-scale geodynamic models, the Earth is commonly treated as an incompressible, viscous fluid in which motions are driven by internal temperature variations. The existence of plates or, equivalently but more importantly, the boundaries between plates occurs because the cool lithosphere has a finite strength and fails when stressed beyond this critical value. Elastic deformation is most important in the low-temperature part of the system and it is commonly assumed to have a minor effect on the system at the plate scale, which is significantly larger than the typical wavelength of elastic flexural response (Watts et al. 1980b; Moresi et al. 2001). In this section we build up a single mathematical framework that covers this range of deformation styles. Our starting point is the viscous fluid flow model of mantle convection, which we can correct for elasticity and brittle deformation (see Table 1).

Viscous deformation model for large-scale flow We make the assumption that the mantle is an extremely viscous material in which accelerations can be neglected (an infinite Prandtl number fluid) so that viscous stresses are always in equilbrium with driving forces. The force term is a gravitational body force due to density changes. We often assume that these arise, for any given material, through temperature effects so that: sij; j ¼ tij; j  pi ¼ gdi3 r0 ð1  aT TÞ

ð1Þ

Rheology of geological materials on geological timescales The viscosity of the mantle and lithosphere at long timescale is known to be a strong function of temperature, pressure, stress, grain size, composition (particularly water content). Karato & Wu (1993), give the following expression for mantle deformation M exp h / t 1N dgrain

E þ pV  RT

ð4Þ

where T is temperature, t is the second invariant of the deviatoric stress tensor, E* is an activation energy, V* an activation volume, R is the gas constant, dgrain is the grain size, N is a stress exponent and M a grain-size exponent. The exponential dependence on temperature produces distinctive strength characteristics in the deep mantle compared with those of the lithosphere (e.g. Solomatov 1995). Some anisotropy of material properties is observed in mantle materials owing to crystallographic realignment during strain, but the effect is not well characterized and may be more important in diagnosing strain than in influencing the deformation itself (see discussion in Mu¨hlhaus et al. 2004).

Lithospheric strength, length scales and localization Brittle deformation is associated with the stress in the lithosphere reaching the maximum that the rocks can support. The principal feature of

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Table 1. Symbols and their definitions Symbol

Units

Description

x u ()i s sI,II,III P t s¯, t¯ D n s tns snn g r0 aT T k Q Cp h hs L dM grain E* V* R f, g w, w0 c, c0 u T0 gp, g˙p Q KE C, Cp B, Bp

m m s21

co-ordinate vector velocity vector notation indicating derivative @()/@xi stress tensor principal stresses; sI , sII , sIII, sI most compressive pressure deviatoric stress second invariant of stress, deviatoric stress symmetric part of velocity gradient tensor unit vector normal to slip plane unit vector aligned with slip direction shear traction on slip plane with normal n normal force on slip plane with normal n gravitational acceleration reference density coefficient of thermal expansion temperature thermal diffusivity heat production rate per unit mass heat capacity at constant pressure viscosity second viscosity in transverse isotropic flow geometrical part of transversely isotropic constitutive tensor grain size activation energy activation volume gas constant yield function, plastic potential defining flow rule friction angles in yield criteria cohesion values in yield criteria angle between failure plane and direction of sI tensile strength plastic strain, strain rate measures (scalars) healing parameter for plastic strain element stiffness matrix finite-element constitutive matrix (geometrical and by material, point) derivative matrix that expresses strain rate D (geometrical and by material point)

Pa Pa Pa Pa pa s21 – – Pa Pa m s22 kg m23 K21 K m2s21 J kg21 s21 J kg21 s21 Pa s Pa s – m J mol21 m3 mol21 m3 mol21 K21 Pa – Pa radians Pa – – – – –

lithospheric failure at this scale is the formation of faults and the fault systems that make up plate boundaries. The length-scales at which fault-forming processes operate are considerably smaller than the length of the faults themselves, and cannot be resolved in models of the entire lithosphere. We therefore seek a continuum model of the lithosphere incorporating fault formation and evolution. The principal difficulties we encounter arise because the faults themselves are visible on the macroscopic scale, even though the underlying formation processes are microscopic and unresolved. The introduction of constitutive models for brittle deformation introduces new nonlinearities into the equation of motion that immediately change the velocity solution qualitatively as well as quantitatively. Faults and shear bands are structures that emerge from the non-linear characteristics

of the equation of motion and which are strongly dependent on the strain history. In seeking a constitutive model for brittle deformation of the lithosphere we look for two things: the model should naturally develop fault-like structures with a morphology compatible with observational constraints, and the effect of those emergent structures should also be compatible with the influence of faults on the stress field at the large scale. The ingredients for a constitutive law are also twofold: a description of the maximum supportable stress written in terms of a limiting condition on one or more components or invariants of the stress tensor; and a flow law which describes how the material deforms once the stress reaches this limit. Each of these ingredients needs to be developed to best fit the important macroscopically observable phenomena, respect the important underlying

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physics and be tractable to mathematical or computational solution. Which of the phenomenological constraints are considered important depends on the context in which the large-scale model is being used. Different modelling approaches will also represent different failure models with differing levels of fidelity. It is always important to be aware of the fact that all failure models are empirical, to be aware of the underlying physical assumptions and to calibrate numerical models carefully. In the next section we will demonstrate how two different, but related, failure models of the lithosphere at the scale of hundreds of kilometres can be developed from the basic assumption that the lithosphere fails by the development of faults with a strength limited by frictional sliding. We also show how the formulation of these models is influenced by the need to model failure in a model that includes the dynamics of mantle convection.

Faults and failure models We must first outline a simplified description of the phenomena we would like to model: in this case a segment of a fault. The simplest picture of a fault is a two-dimensional (2D) discontinuity surface arbitrarily embedded in unfractured rock that we assume deforms by frictional sliding in the plane of the fault (see Barr & Houseman 1996 for discussion). If we consider a small planar element of the fault (see Fig. 1) with a normal, n ¼ (n1, n2, n3), and an orthogonal vector, s, lying in the fault plane in the direction of maximum shear stress. The traction resolved in each of these directions is

given by snn ¼ ni nj sij

ð5Þ

ssn ¼ si nj sij ¼ si nj tij :

ð6Þ

The yield criterion for the fault segment is given by f ¼ 0 in f ¼ tsn  tan wsnn  c  0

tsn ¼ jssn j:

ð7Þ

Where w the friction angle and c is a cohesion. In 3D it is possible to express the magnitude of the maximum shear traction in the fault plane as pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi ð8Þ tsn ¼ sik sil nk nl  snn 2 eliminating the need for explicitly calculating s.

Continuum failure criterion The failure criterion for a fault segment (equation 7) applies to a patch of a pre-existing fault with a known orientation. For simplicity, it is useful to assume that the same failure criterion should hold for the bulk of the lithosphere before macroscopic fault systems develop. In this case it is necessary to express not only how an existing fault of given orientation behaves, but also to specify the orientation of new fault segments at the point of failure. The continuum approximation rests upon the assumption that, at the small scale, in lithosphere which has not been subjected to sufficient stress that it fails, all orientations of incipient faults are present. For any given stress field, we can then assume that there is at least one incipient fault within every element of the material which is oriented in such a way that it fails earlier than any other fault and begins to grow. For a point in the material we consider the plane that contains the maximum (here most compressive) and minimum principal stresses (sI and sIII. Resolving the stress onto a fault oriented at an angle u measured positive counterclockwise to the sI direction yields 1 1 snn ¼ ðsI þ sIII Þ þ ðsI  sIII Þ cosð2uÞ 2 2 1 ssn ¼ ðsI  sIII Þ sinð2uÞ: 2

ð9Þ ð10Þ

Substituting into the yield criterion (equation 7) gives: sI  sIII ¼

Fig. 1. Forces acting on an element of a pre-existing fault in the normal direction and the direction of maximum shear stress.

tan wðsI þ sIII Þ=2 þ c : sinð2uÞ  tan w cosð2uÞ

ð11Þ

In early theories of shear failure and shear banding it was assumed that failure would take place on spontaneously forming rupture planes or shear bands which are oriented such that the stress difference

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on the left-hand side of (11) is a minimum (e.g. Anderson 1951). It is then simple to compute d(sI 2 sIII)/du to find the value of uB that minimizes the differential stress required for failure to be p w uB ¼ +arctanðtan1 wÞ ¼ + þ : ð12Þ 4 2 Replacing u by uB in equation 11 gives sI  sIII sI þ sIII ¼  sin w þ c cos w: 2 2

ð13Þ

In a pristine material we expect faults to form at an angle between 22.58 and 458 to the most compressive principal stress direction. In the Earth (tan w 1) this means that we expect steep faults to form in extensional environments (sI close to vertical) and shallow faults to form in compressive environments (sI close to horizontal). A note of caution should be introduced at this point – the assumption that the macroscopic frictional sliding criterion should also apply at the microscopic scale is a convenient analogy, but there is no accepted model of rock fracture from which it arises. As summarized by Scholtz (2002), such failure criteria are strictly empirical, and are justified by their usefulness in modelling a range of phenomena.

Drucker – Prager criterion as a modified form of Mohr – Coulomb One characteristic of the Mohr–Coulomb failure criterion is the assumption that failure is independent of the value of the intermediate principal stress, sII. This assumption is true regardless of how close the magnitude of sII comes to that of either of the other principal stresses. As a consequence, small changes in the stress field can produce a sudden transition in the orientation of the favoured failure planes. When the yield surface ( f ¼ 0 in equation 7) is plotted with respect to the principal stress directions (Fig. 2, lower), the locations where these transitions occur show up as edges on the surface. This surface has discontinuous derivatives at these points which can pose difficulties in many computational plasticity implementations. The Drucker–Prager yield criterion (Fig. 2, upper) is an attempt to produce a smooth surface with otherwise similar characteristics to the Mohr– Coulomb yield surface. The criterion is given in terms of invariants of the stress tensor by f ¼ 0 in pffiffiffi ð14Þ f ¼ 3t  tan w0 p  c0  0 pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi where t¯ ¼ ðtij tij Þ=2 and tan w0 and c0 are constants analogous to tan w and c in equation (7). Where the scalings are chosen in this case to

Fig. 2. Drucker–Prager and Mohr–Coulomb yield surfaces drawn with respect to the principal– stress axes. The Drucker–Prager surface is often scaled such that it touches the Mohr–Coulomb surface at the outermost of the edges–the circles indicate these points in the plane where the surfaces have been truncated to fit on the page.

ensure the yield envelope never falls inside the Mohr–Coulomb envelope. In this case, the Mohr–Coulomb criterion and the Drucker–Prager criterion coincide for the conditions of a triaxial compression test. This also implies the following relationships: tan w0 ¼

6 sin w 3  sin w

c0 ¼

6c cos w : 3  sin w

ð15Þ

The Drucker–Prager yield criterion is usually simpler to implement in computational models and, in addition, may be a more appropriate choice of yield function for materials where the deformation is not well represented by frictional sliding on failure planes. This includes many granular material such as soils.

Tensile strength The assumption that post-yield deformation occurs through frictional sliding is valid in compression, but in extension it is also possible to reach the tensile strength of the material, i.e. the most extensional principal stress, sIII is equal to the tensile strength, T0 sIII ¼ T0 :

ð16Þ

The tensile strength is a material property independent of c and tan w. The cross-over to the tensile

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failure mode is found by substituting equation (16) into equation (13)    sin w þ 1 T0 ½sin w þ 1 sI , 2c 1 : ð17Þ cos w 2c cos w This failure mode is characterized by cracks opening perpendicular to the direction of the most extensional principal stress, sIII. Our assumption of incompressible flow is not compatible with the volume change associated with opening one or more macroscopic cracks, so we do not attempt to include the effect of tensile failure in the formulation described here. This amounts to the assumption that the tensile strength is high enough that cracks perpendicular to sIII do not open before the resolved normal stress on the Mohr –Coulomb failure plane reaches zero. This latter condition signifies that the assumptions we have made are no longer valid; this must be tested during deformation to ensure consistency of the numerical solution.

Flow rules for continuum failure model A yield criterion provides a limit on the acceptable stress states in a material. A corresponding flow rule is needed to determine the manner in which the material deforms that the yield stress is reached. The incompressible fluid assumption that we made in formulating the mantle convection problem is a strong constraint on the acceptable flow rules. We make the assumption that we can write the total deformation rate as the sum of contributions from viscous and plastic deformation rates. A standard approach in the theory of plasticity is to write Dij ¼ Dvij þ Dpij

Dpij ¼ g_ p

@g : @ sij

ð18Þ

where D p is the plastic part of the stretching, g˙p the ‘equivalent plastic strain rate’ and g is the ‘plastic potential’. The incompressibility constraint is automatically satisfied if g depends only upon on the deviatoric stress and not the pressure. g depends on the stress invariants according to the symmetry of the material and possibly on other state variables such as temperature. The standard choice for g in connection with a yield criterion of type equation (7) or (11) is p w ð19Þ g ¼ tS u ¼ +uB ¼ + þ : 4 2 This choice ensures that, once the yield surface is reached, subsequent deformation is always confined to that surface. However, unless w ¼ 0, the flow law resulting from equation (19) does not satisfy the

incompressibility constraint since tS is a function of the pressure. In the following subsections we describe flow rules that do satisfy the incompressibility constraint. They are not written as a decomposition into viscous and plastic deformation, nor do we write the plastic strain rate in terms of a plastic potential and equivalent plastic strain rate, but these concepts are revisited in the context of strain softening where only the plastic part of the deformation produce damage. Dilatancy influences the initial orientatation of shear bands (Vardoulakis & Sulem 1995) and also plays the role of a history parameter. In our incompressible formulation, shear band orientation is controlled by the internal slip direction, and strain softening is the history variable that drives localization.

Flow rule for Mohr – Coulomb yield criterion An element of material with an embedded frictional failure plane deforms as a nonlinear, transversely anisotropic medium of the type proposed by Mu¨hlhaus et al. (2002a, b) with the weak orientation aligned to the failure plane. Such a material is characterized by two effective viscosities, h and hs, which apply to pure shear deformation and simple shear deformation, respectively. We write the constitutive law for this material as a correction to equation (2): sij ¼ 2hDij þ 2ðhs  hÞLijkl Dkl  pdij

ð20Þ

where Lijkl ¼ ðni nk dlj þ nj nk dil þ ni nl dkj þ nj nl dik Þ=2  2ni nj nk nl

ð21Þ

and n is a normal surface vector defining the orientation of the failure plane. Moresi & Mu¨hlhaus (2006) have shown that the yield criterion (equation 7) is satisfied at any given point by setting ¼ hIþ1 S

tan wð2hDInn  pÞ þ c I 2g_ ns

ð22Þ

in equation (20) where DInn is the shear strain rate normal to the active slip plane, and g˙Ins ¼ nIi DIij sjI is the shear strain rate resolved on the active slip plane. I is an iteration counter; iteration is necessary to take into account the interaction between individual locations that fail and redistribute the stress throughout the system. In each iteration the stress at all points is recalculated based on the strain rate of all points carried forward from the previous iteration. Iteration is continued until changes in the global solution fall below a specified tolerance. Given the strong non-linearity in this system, the

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Localization, failure history and strain softening

Fig. 3. In a developing shear band, the favoured slip orientations a, b are at u to the most compressive principle stress direction, sI. a is more closely aligned to the sense of shear than b. Macroscopic shear bands that form are strongly aligned with whichever direction, a or b is chosen.

convergence of this iteration is not guaranteed. In practice, however, with any degree of inhomogeneity in the initial distribution of strain, localization patterns are established within a few iterations. In a homogeneous material, there is no a priori reason to suppose one or the other of the two possible slip systems given by u ¼ +uB is more likely to become active at the yield point. However, one of the slip systems is more closely aligned with the local sense of shear strain rate (see Fig. 3) – we assume that during subsequent strain, it is this particular shear band that grows preferentially.

Flow rule for Drucker – Prager yield criterion We can write a correction to the viscous constitutive law for the Drucker–Prager yield criterion (14) as follows sij ¼ 2hDij þ 2ðhy  hÞDij  pdij :

ð23Þ

This has the same form as equation (20) but, since the viscosities are all scalars, can be simplified to sij ¼ 2hy Dij  pdij :

ð24Þ

The condition on the stress invariant is satisfied by a correction to the isotropic viscosity with hy given by the limit of iterating the following expression to an appropriate tolerance.

The failure criteria and flow laws that we have derived above describe the instantaneous response of the lithosphere to stresses including those large enough to produce failure. The constitutive behaviour is strongly non-linear once the yield strength is reached and this can promote non-uniformity in the deformation gradients, including the possibility of localization of deformation into very narrow shear bands. Shear banding in our models corresponds to the formation of fault systems in the lithosphere. Localization of deformation may be mediated by structural effects such as the necking or buckling of the lithosphere under extension/compression; by inherent instabilities in the constitutive law that produce spontaneous localization in homogeneously deforming ‘infinite’ samples; and by progressive failure due to feedback from plastic strain into the material properties, e.g. post-failure strain softening. Structural effects are possible in any of models we construct with material boundaries, and spontaneous localization is weak in incompressible viscoplastic models. Strain softening effects are important in our context and introduce additional material properties. The friction coefficient and/or the cohesion for material points that have failed weaken as slip accumulates during yielding. We consider that there is a competition between damage processes that occur due to continued plastic deformation and healing processes which are likely to depend upon the background deformation rate (in the absence of yielding) and through an Arrhenius law to temperature and pressure. We assume that this strain hardening/softening and healing can be described as a function of the following relative strain measure: p Dgrel tns ¼ g_ p  QðT; PÞ gprel  0: Dt h

In equation (26) 0 , Q , 1 is a dimensionless healing parameter, g˙p is the equivalent plastic strain rate defined in equation (18), and tns/h is a measure of the background strain rate. In the context of the anisotropic viscoplastic flow rule for Mohr–Coulomb yielding, p

aIp þ k ¼ hIþ1 pffiffiffi I y  2 3D pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi I in which D ¼ ðDij Dij Þ=2 Dkk ¼ 0.

g_ ¼ tns ð25Þ

ð26Þ



1 1  hs h

 ð27Þ

and for the isotropic viscoplastic rule for the Drucker –Prager yield criterion, an appropriate

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c(0)

c(1)

p

p

Plastic strain — γ rel /γ 0

Fig. 4. Strain-softening function.

definition is "

# 1 1 g_ ¼ ty  : hy h p

ð28Þ

material derivatives, but a fixed mesh is also used to compute fluid velocities. We have implemented this method into the Underworld finite-element code. Underworld uses a standard mesh to discretize the domain into elements. The shape functions interpolate node points in the mesh in the usual fashion and are used to compute derivatives of nodal variables. Material property variations and history variables, such as failure plane orientation and failure history, are stored on integration points that are also material points of the fluid. The problem is formulated through the usual FEM weak form to give an integral equation that can then be decomposed to a series of element integrals and through the usual Galerkin discretization procedure, give element stiffness matrices, K E: ð KE ¼ BT ðxÞCðxÞBðxÞdV ð31Þ VB

The relationship (26) is assumed to hold even if the shear band is not active, i.e. g˙p ¼ 0. The parameter gprel is not a state parameter but a measure to quantify the relative effect of creation and annihilation of inelastic structures (e.g. shear bands) caused by plastic deformations chosen for heuristic reasons. During the actual calculation we iterate to determine the distribution of points that are currently yielding, we consider first whether each particle has failed in a previous time step and test to see if it will fail given the updated stress distribution. The friction coefficient and/or the cohesion for the material points that have failed weaken as relative slip (equation 26) accumulates during yielding. If a material point has failed in the past but changes in the ambient stress field mean that it is no longer yielding, then the history parameter (equation 26) decreases until gprel ¼ 0. The cohesion and/or friction coefficient are softened according to cðaÞ ¼ acð0Þ þ ð1  aÞcð1Þ ð29Þ tanðfðbÞÞ ¼ b tanðfð0Þ þ ð1  bÞ tanðfð1ÞÞ ð30Þ where a ¼ min(1, gprel/gc0) and b ¼ min(1, gprel/gf0 ) (see Fig. 4). For simplicity, we have assumed b ¼ a or b ¼ 0 for all the examples in this paper.

Numerical implementation Equation (27) contains material derivatives that are most reliably implemented using a Lagrangian reference frame (tracking points embedded in the deforming material). However, in fluid convection, the strains are so large that frequent remeshing is required for accurate solution. Moresi et al. (2003) described a method that allows a Lagrangian reference frame to be used for terms involving

we replace the continuous integral by a summation X wp BT p ðxp ÞCp ðxp ÞBp ðxp Þ: ð32Þ KE ¼ p

Here the matrix B consists of the appropriate gradients of interpolation functions which transform nodal point velocity components to strain-rate pseudo-vectors at any points in the element domain. C the constitutive operator corresponding to equations (20) or (24) is composed of two parts C ¼ C iso þ C aniso. In standard finite elements, the positions of the sample points, xp, and the weighting, wp are optimized in advance. In our scheme, the xps correspond precisely to the Lagrangian points embedded in the fluid, and wp must be recalculated at the end of a time step for the new configuration of particles.

Characterization of the formulation In the following examples we attempt to characterize our algorithm with some simple experiments that derive from common geological problems. We consider the extension and compression of a brittle layer that lies above a viscous layer for a range of cohesion and friction angle values. For the sake of generality, we conduct our experiments in an extremely simple geometry – rectangular domains with depth 1 – and choose reference values of all quantities to be unity unless otherwise specified. In each case the deformation is driven by a boundary condition that uniformly stretches or compacts the background mesh (initial size 3.0  1.0 divided into 360  120 elements) in the horizontal direction with velocity of 1.0. The top and bottom boundaries of the mesh are free-slip surfaces (no shear stress, no normal velocity). The

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evolving interface geometry and material history is recorded on the swarm of particles. A layer of lowviscosity readily compressible material is always included in the calculation above the layered test materials to allow the volume change of the domain to be accommodated, and to mimic a stress-free-surface upper boundary condition on the brittle layer. The volume of the test materials is constant during the computation – during extension the average height of the free surface decreases with time due both to the thinning of the ‘crustal’ layer and also due to the thinning of the mantle. The second of these effects is an artifact of the set-up and would need to be corrected before applying the results of the computations to observed subsidence. In geological models the effect of gravity is always important so we have included this effect. The strength of gravitational acceleration is 10.0 and the density of the layers is 1.0. In the first set of models we also introduce a small perturbation in thickness of the brittle layer that concentrates the initial stress distribution. This encourages the growth of a single pair of conjugate shear bands at an orientation that is not influenced by the growth of competing, intersecting shear bands. This arrangement allows us to examine the relationship between the shear band angle (an emergent, macroscopic phenomenon) and the angle of friction of the material. We also introduce a distribution of initial strain on a small random sample (typically 1%) of material points as follows: p ¼ g0 sin2 kx: grel

ð33Þ

The wavenumber k ¼ 2np/L (L the horizontal dimension of the experiment), which ensures that the initial strain close to the ends of the box are small (where the boundary condition imposes an artificial constraint on the local rotation), and also allows us to explore the sensitivity of localization patterns to the initial periodicity of the solution.

Influence of friction angle on shear band orientation In Figure 5, a Mohr –Coulomb brittle layer of thickness 0.3, viscosity 10.0, and cohesion 4 (compression), 15 (extension) lies above a purely viscous layer of viscosity 1.0. The strain-softening parameter for cohesion only, g0p, was 0.1 for all models shown, with the minimum value of the cohesion was 1.0 at this reference strain. A narrow notch of material in the brittle layer was removed (replaced by material of the lower layer) to provide an initial stress concentration that encourages shear band formation.

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In this experiment the layers were only deformed until the pair of shear bands had developed across the entire layer – total strains between 2 and 5%. The orientation of these shear bands was then compared with the ideal orientation of the slip planes (assuming that the orientation of the principal stress directions is not strongly influenced by the shear band formation), which is indicated in the diagram by the light-coloured line. In compression the fit between the macroscopic shear band orientation and the microscopic material point slip orientation is very close. In extension the shear bands form at shallow angles for low values of the cohesion, steeper angles for higher values, but the agreement between the microscopic slip orientation and the shear bands is not quite as good. Figure 6 shows a similar experiment using the Drucker –Prager yield criterion. In this case the shear band angles show very little sensitivity to tan w or the sense of extensional or compressional deformation. The shear band angles are all close to 458, slightly steeper in compression.

Influence of cohesion Figure 7 illustrates the effect of varying the cohesion of the upper layer. In this experiment the layer was uniform in thickness and shear bands initiated from the initial distribution of random damage throughout the material. In the Mohr–Coulomb case, when the cohesion was low (C(0) ¼ 6.0), the layer failed easily with many interacting shear bands appearing after 1–2% strain. The number of shear bands decreased with increasing cohesion in the upper layer. At C(0) ¼ 10, the failure occurs on very broadly spaced pairs of conjugate shear bands. For values of C(0) . 10, no shear bands formed from the initial damage, which completely healed at strains of 5–10%. Montesi & Zuber (2000) derived a theoretical spacing for the formation of shear bands in a brittle layer. Their model predicts that the spacing is controlled by rate of the loss of strength compared to the initial strength. This is in agreement with the results shown here: the overall strength of the layer varies while the strain weakening rate is kept fixed. The spacing of localization into shear bands increases with cohesion. This observation also confirms the results of Lavier et al. (2000), Wijns et al. (2005) and Huismans et al. (2005), who showed that the formation and density of shear bands is strongly controlled by the relative integrated strength of the brittle and ductile layers. The pattern is similar when the Drucker– Prager yield criterion is used. We note, however, the equivalent spacing of shear bands is offset to lower values of the cohesion when the Drucker–Prager

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Fig. 5. Shear band formation in compression and extension of a Mohr–Coulomb brittle layer of length 3.0, thickness 0.3 with a small notch removed, viscosity h ¼ 10.0, cohesion 5 (compression) and 10 (extension) and tan w varying between 0.001 and 0.5. The compression/extension velocity is 1.0; the lower layer viscosity is 1.0, the overlying layer viscosity is 0.1 and this material is compressible. The strain-softening parameter was 0.1 for all models shown, and the minimum value of the cohesion of 1.0 at this reference strain; tan w did not soften in these models. The ideal orientation of shear bands assuming the principle stress orientations are horizontal/vertical is indicated by the lines on each image. The graph shows the shear band angle, measured from the images, for extension and compression together with the ideal orientation. The error in measuring was +28 after averaging the dips of the conjugate shear bands. The shading represents the second invariant of the stress tensor plotted between 0 (dark) and 10 in the top layer, and 0 (dark) and 1 in the lower layer.

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Fig. 6. The corresponding experiments to those shown in Figure 5 for the Drucker–Prager yield criterion. Shear band angles show very little sensitivity to the friction angle or the sense of extension/compression. In the range of 0.1 , tan w , 0.5, the shear band angle was 43 + 28 to the horizontal in compression and 49 + 2 in extension. The lines superimposed on this diagram are at 458.

Fig. 7. The effect of changing cohesion for a similar system to Figure 5 but here using a constant thickness layer with no notch to concentrate the initial deformation. C(0) varies from 4 to 11 as indicated, and C(1) is 1 in all the cases. The tangent of the friction angle is 0.5 in every model, and gprel is 0.1. The shading indicates the second invariant of the stress tensor plotted between 0 (dark) and 10 (bright) in the top layer, and 0 (dark) and 1 (light) in the lower layer.

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criterion, normalized according to equation (15) is used. This is unsurprising since this choice ensures the Drucker– Prager criterion is never weaker than the Mohr –Coulomb condition, but it does illustrate the need to calibrate any parameterizations for particular geological problems.

rotated to less favourable angles. In this case, the effect is relatively minor, and is, in fact, confined to reactivation of failure directions that were active in the first 1 –2% of extension.

Large deformation model

All of the models presented so far have been 2D. The formulations are, however, equally applicable to 3D problems. We present two examples in 3D: extension models comparable to those in Figures 5 & 7. In 3D, we are presently limited to a maximum resolution of 96  48  48 elements, which is smaller by a factor of more than 2 than the linear resolution of the 2D cases, and is restricted to a 2  1  1 domain. The resolution should be sufficient to resolve a small number of distinct shear bands so we have chosen material parameters which, from the 2D calculations, are expected to produce only one or two shear bands: a strong upper layer (C(0) ¼ 30; h ¼ 20), rapid strain softening (C(1) ¼ 0.01; gp0 ¼ 0.01). As in 2D, we introduced a some initial damage with random magnitude and orientation (which also includes some variation in the third dimension). Figure 9a shows the pattern of failure in the brittle layer after 4 and 12% strain. The initial pattern of failure is highly distributed and broadly reflects the strength of the seeding. This is very similar to the initial development of the 2D models including those shown in Figures 7 & 8. However, at higher strain, this model does not localize as strongly as we observed in 2D. This is largely due to the fact that the random pattern of shear bands that develops initially produces many planar structures which have a limited extent in the cross-strike direction. In order for a single through-going structure to develop, there needs to be a set of failure planes with the same polarity that are well aligned across the domain or, alternatively, one set of structures must grow through others with the opposite polarity. In the 3D experiment shown, the distribution of shear bands is considerably more diffuse than would be observed in 3D. We ran a related model to see how the observations from this experiment apply when we introduce a finite perturbation to encourage localization. We used the same small notch in the crustal thickness that was used when we analysed the effect of changing friction angle (Figs 5 & 6) but only over one quarter of the domain. The results are shown in Figure 9b; they show that shear bands that develop early can very quickly dominate the deformation pattern and form a near-2D distribution of failure. Although a much larger number of models would be needed to understand the generalization of our models to 3D, we can draw two tentative

Figure 8 shows the evolution of the deformation when the extension is comparitively large. We ran a two-layer model at a resolution of 450  150 elements to a total extension of 85% and plotted the upper layer deformation using initially horizontal stripes of passive markers to indicate the total deformation. An initially large number of active shear bands in the shallow part of the brittle layer reduces during the first 1 –2% of extension as strain softening within the shear bands allows them to compete. During the deformation experiment, owing to stress transfer from the lower viscous layer, there are a six active shear band systems that accumulate significant slip (from a visual inspection of the strain markers). At the end of the experiment is possible to see how the long-lived shear bands have rotated during extension. These are the shear bands that have accumulated significant slip and, consequently, are significantly weaker than surrounding material. Unbroken blocks of material bounded by the major shear bands are strongly tilted during extension and undergo some stretching owing to their inherent viscosity but remain otherwise intact indicating that the principle deformation in the system is along the shear bands. On the right-hand side of Figure 8 we have superimposed a fault interpretation of the shear banding. For each snapshot we have drawn currently active faults in thick dark lines and inactive faults with appreciable strain (from the passive markers) or a positive strain-softening factor, gprel, using thinner lines. The faults are numbered according to the order in which they formed (where there are several interacting faults, the earliest fault in the system is considered). A number of interesting behaviours can be seen in the evolution of these faults. Active systems can offset each other and coalesce to produce complex geometrical block configurations (fault system 4). Reactivation of faults or fault segments is common as the configuration evolves (e.g. fault systems 5 and 6). Towards the end of the experiment, the strain associated with individual faults can be traced along the edges of the rotated blocks well beyond the currently slipping region of the fault. A second generation of faults formed at steep angles in the less deformed blocks (e.g. in fault systems 2 and 6) as the main faults become

Three-dimensional model

Fig. 8. Extension of a two-layer model with C(0) ¼ 1, C(1) ¼ 0.01, g ¼ 100, tan w ¼ 0.5, h ¼ 10, r ¼ 1; initial brittle layer geometry 2.0  0.35, and lower layer viscosity 1. The extension velocity was 1. Passive marker layers were included in this case to show accumulated strain. Left: a series of frames from the simulation with the brightness of shear bands representing g˙p and the intensity of the shading representing grel. The background shading shows the second invariant of the strain rate, which varies between 0 (dark) and 1 (light, but largely hidden by shear bands) in the top layer, and between 0 (dark blue) and 10 (bright yellow) in the lower layer. Right: the same frames with the shear bands interpreted as faults– thick lines indicate regions where slip was occurring and thin lines indicate the zones with a record of high strain.

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Fig. 9. Extension of a two-layer model with C(0) ¼ 30, C(1) ¼ 0.01, g ¼ 10, tan w ¼ 0.5, h ¼ 20, r ¼ 1; initial brittle layer geometry 2.0  1.0  0.4, and lower layer viscosity 1. The extension velocity was 1. A single passive marker layer was included in this case to show accumulated strain coloured yellow for values of the strain rate invariant which exceed 0.1, green for lower strain rates. The viscous layer is coloured by strain rate from 0 (dark blue) and 2 (bright yellow). (a) The layer is of uniform thickness everywhere; (b) a small notch has been removed from the base of the crust in the middle of the box on the distant side. At 12% strain for the system as a whole, we have sketched (by outlines) the main identifiable shear bands – black or white depending on whether they dip to the left or right.

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conclusions: (1) the influence of material properties on spacing of shear bands that we observe in 2D is complicated in 3D by the finite across-strike length of the shear bands which form from initially random distributions of damage; and (2) relatively small perturbations in the strength distribution of the lithosphere can overcome this effect and allow the dominant mode observed in 2D to grow.

Discussion and conclusions The algorithms described in this paper address a prominent problem in geodynamics: how to model the brittle deformation of the uppermost lithosphere that occurs when mantle convection deforms the continents at the same time as modelling the underlying fluid convection which drives the surface deformation. The formulation has been developed from a mathematical description of incompressible viscous fluid flow that is inherently capable of modeling thermally driven convection. By locally satisfying the failure criterion for frictional sliding in a viscous material, and assuming deformation initially occurs along the static characteristics, we can generate macroscopic shear bands which are aligned with the static characteristics of the global stress field. A strain-softening model based on the accumulated slip at failed material points is required for strong localization to occur. The incompressible viscous formulation for yielding in the lithosphere has advantages over the classical viscoelastoplastic solid mechanics formulations: the assumption that very large strains are always present underpins the fluid formulation that naturally copes very easily with thermal convection and very large offsets on shear bands; there are a number of optimized implicit solution methods for mantle convection applications that allow rapid solution of the momentum equation in 2D and 3D, even with large spatial gradients of material properties. There are a number of disadvantages: material history tracking and following material interfaces and the free-surface have to be specially catered for; arbitrary plastic flow rules are hard to implement within the context of many of the efficient convection solvers; elastic effects can be included, but the purely elastic limit is poorly represented. The viscous fluid formulations are ideally applied to plate-scale deformation, but it is important that the models should be able to represent deformation accurately down to the basin scale where elastic solid models are very well established. In a simple extension experiment, the shear bands interact to generate geologically plausible patterns of deformation including rotated fault blocks and multiple generations of faults.

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We have outlined in this paper how the Drucker –Prager and Mohr– Coulomb failure criteria may be implemented into the viscous fluid equations. Although these two criteria are designed to give comparable results in compressible elastoplastic formulations, the suppression of volume changes during failure forces shear banding to occur at 458 when the Drucker –Prager yield criterion is translated into a viscous flow law. The anisotropic viscous formulation of the Mohr–Coulomb failure criterion and flow law that we have outlined recognizes a preferred orientation at the micro-scale which reappears in shear bands at the macroscale in broad agreement with the predictions of Anderson’s faulting theory. Although the formulation is slightly more complicated, the anisotropic Mohr–Coulomb model is therefore a better choice of rheology for viscous –plastic models of the lithosphere. An additional advantage of the anisotropic approach is that the stiffness matrices remain well conditioned during plastic deformation. This improves the performance of iterative solvers and the accuracy of penalty-based direct solution methods. Two-dimensional models have give us a very good intuition for the influence of the many material parameters in the mathematical model. In our preliminary 3D models we can still see the patterns familiar from 2D simulations, but, where models are started from randomly oriented damage planes, we see structure in the third dimension that results from shear bands initially forming with finite extent across strike of the extension direction. Small perturbations to the strength in the third dimension help to excite the same modes we see in 2D. We conclude that the 2D models remain a very good guide to the behaviour in 3D but caution that the across-strike partitioning of deformation may make it difficult to observe the dominant length scales across the entire domain. All failure models are phenomenological at one or more levels. This means that they must be carefully calibrated for the range of length scales, timescales and materials for which they are to be used. The geological application of strainsoftening plasticity models is particularly difficult because the pressure, temperature, porefluid and strain-rate conditions prevailing at the time particular structures formed may be quite poorly constrained, and, if structures are visible only through remote sensing techniques, further ambiguity in material properties and interpretations of the observations are present. The authors acknowledge the support of the Australian Research Council Discovery project grant number DP0345157 and number DP0449979 and the use of the Australian Computational Earth System Simulator software suite.

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References A NDERSON , E. M. 1951. The Dynamics of Faulting and Dyke Formation With Applications to Britain. Oliver & Boyd, Edinburgh. B ARAZANGI , M. & D ORMAN , J. 1969. World seismicity maps compiled from essa, coast and geodetic survey, epicenter data 1961– 1967. Bulletin of the Seismological Society of America, 59, 369– 380. B ARR , T. D. & H OUSEMAN , G. A. 1996. Deformation fields around a fault embedded in a non-linear ductile medium. Geophysical Journal International, 125, 473–490. B OUSSINESQ , J. 1903. The´ory analytique de la chaleur mise en harmonie avec la thermodynamique et avec la the´ory de la lumie`re, Volume II. Gauthier-Villars, Paris. B RAUN , J. & B EAUMONT , C. 1995. Three dimensional numerical experiments of strain partitioning at oblique plate boundaries: implications for contrasting tectonic styles in California and South Island, New Zealand. Journal of Geophysical Research, 100, 18,059–18,074. B RAUN , J. & S AMBRIDGE , M. 1994. Dynamical lagrangian remeshing (dlr): A new algorithm for solving large strain deformation problems and its application to fault-propagation folding. Earth and Planetary Sciences Letters, 124, 211– 220. E LLIS , S., F ULLSACK , P. & B EAUMONT , C. 1995. Oblique convergence of the lithosphere driven by basal forcing; implications for length scales of deformation and strain partitioning in orogens. Geophysical Journal International, 120, 22–44. F ULLSACK , P. 1995. An arbitrary Lagrangian-Eulerian formulation for creeping flows and its application in tectonic models. Geophysical Journal International, 120, 1–23. H UISMANS , R. S., B UITER , S. J. H. & B EAUMONT , C. 2005. Effect of plastic–viscous layering and strain softening on mode selection during lithospheric extension. Journal of Geophysical Research, 110, B02406, doi:10.1029/2004JB003114. K ARATO , S. & W U , P. 1993. Rheology of the upper mantle – a synthesis. Science, 260, 771–778. L AVIER , L. L., B UCK , W. R. & P OLIAKOV , A. N. B. 2000. Factors controlling normal fault offset in an ideal brittle layer. Journal of Geophysical Letters, 105, 23,431–23,442. L ENARDIC , A., M ORESI , L. & M U¨ HLHAUS , H. 2000. The role of mobile belts for the longevity of deep cratonic lithosphere: The crumple zone model. Geophysical Research Letters, 27, 1235– 1238. M ONTESI , L. & Z UBER , M. T. 2000. Spacing of faults at the scale of the lithosphere and localization instability: 1. theory. Journal of Geophysical Research, 109, doi:10.1029/2002JB001923. M ORESI , L. & M U¨ HLHAUS , H.-B. 2006. Anisotropic viscous models of large-deformation Mohr–Coulomb failure. Philosophical Magazine, 86, 3287–3305.

M ORESI , L., M U¨ HLHAUS , H.-B. & D UFOUR , F. 2001. Viscoelastic formulation for modelling of plate tectonics. In: M U¨ HLHAUS , H.-B., D YSKIN , A. & P ASTERNAK , E. (eds) Bifurcation and Localization in Soils and Rocks. Balkema, Rotterdam, 337– 344. M ORESI , L., D UFOUR , F. & M U¨ HLHAUS , H. B. 2003. A lagrangian integration point finite element method for large deformation modeling of viscoelastic geomaterials. Journal of Computational Physics, 184, 476–497. M U¨ HLHAUS , H.-B., D UFOUR , F., M ORESI , L. & H OBBS , B. 2002a. A director theory for viscoelastic folding instabilities in multilayered rock. International Journal of Solids and Structures, 39, 3675– 3691. M U¨ HLHAUS , H. B., M ORESI , L., H OBBS , B. & D UFOUR , F. 2002b. Large amplitude folding in finely layered viscoelastic rock structures. Pure and Applied Geophysics, 159, 2311– 2333. M U¨ HLHAUS , H. B., M ORESI , L. & C ADA , M. 2004. Emergent anisotropy and flow alignment in viscous rock. Pure and Applied Geophysics, 161, 2451 – 2463. P OLIAKOV , A. N. B. & H ERMANN , H. J. 1994. Selforganized criticality of plastic shear bands in rocks. Geophysical Research Letters, 21, 2143– 2146. P OLIAKOV , A. N. B., C UNDALL , P., P ODLADCHIKOV , YU . & L YAKHOVSKY , V. 1993. An explicit inertial method for the simulation of viscoelastic flow: An evaluation of elastic effects on diapiric flow in two and three layer models. In: S TONE , D. B. & R UNCORN , S. K. (eds) Flow and Creep in the Solar System: Observations, Modeling and Theory. Kluwer, Dordrecht, 175– 195. S CHOLTZ , C. H. 2002. The Mechanics of Earthquakes and Faulting, 2nd edn. Cambridge University Press, Cambridge. S OLOMATOV , V. S. 1995. Scaling of temperature- and stress-dependent viscosity convection. Physics of Fluids, 7, 266–274. V ARDOULAKIS , I. & S ULEM , J. 1995. Bifurcation Analysis in Geomechanics. Blackie, Glasgow. W ATTS , A. B., B ODINE , J. H. & R IBE , N. M. 1980a. Observations of flexure and the geological evolution of the pacific basin. Nature, 283, 532–537. W ATTS , A. B., B ODINE , J. H. & S TECKLER , M. S. 1980b. Observation of flexure and the state of stress in the oceanic lithosphere. Journal of Geophysical Research, 85, 6369– 6376. W IJNS , C., W EINBERG , R., G ESSNER , K. & M ORESI , L. 2005. Mode of crustal extension determined by rheological layering. Earth and Planetary Science Letters, 236, 120– 134. W ILLETT , S. & B EAUMONT , C. 1994. Insights into the tectonics of the india-asia collision from numerical models of mantle subduction. Nature, 369, 642–645.

Index Note: Page numbers referring to figures are denoted in italics, those referring to tables are denoted in bold. accretionary wedges 369 –70 Adria margin 358 –9, 360, 361– 2, 368 Adria microcontinent 344 Afar plume 265, 268, 285–6 Aegean Sea 370 Airy isostatic compensation 175, 201 alluvial fan deposits 431–2 Almopias unit 363, 364 Alpine Tethys 2, 291– 324 breakup observations 303–12 conceptual model 294–6 detachments 303, 304, 306 isostatic evolution 297, 303 migration of depocentres 301, 303 ocean–continent transition 303–8 palaeogeography 292 prerift 295, 297 sediments 303, 308 tectonic map 292 Alpine–North Atlantic, Eastern Tethys similarities 371– 3 Anatolide–Tauride platform 355 Anderson model, fault theory 457, 471 Antalya Complex 347, 348, 350, 352, 367, 368 Aptian event 14– 15, 18–19, 19, 26, 27, 28, 31– 3, 32– 8, 39– 41 Aptian–Albian boundary 9, 12, 411 Apulia see Adria Ar/Ar cooling data 81, 308, 309, 312 Arabian margin 330, 346– 7 emplacement mechanism 346, 346 seamounts 346, 347 Argo Abyssal Plain 240 Arrhenius law 463 asthenosphere 47, 139, 318, 325 heat transport 86, 163 –4 and mantle lithosphere 129–30, 140 temperature 85, 167, 169, 170, 414 tensional stress 374 upwelling 39, 87, 90, 139, 157, 165, 218, 374, 398 viscosity 141 see also mantle Atlantic hinge zone 394, 396 Australia, Southern Rift System 239– 63 2.5D forward model 244– 53 aborted sea-floor spreading 254, 259 bathymetry 242 depth-dependent stretching 394

gravity field domains 239, 250, 251 inner COTZ 247 magnetic field domains 239, 250, 251, 253 mantle exhumation 259 models 245–6, 250– 3, 252, 253– 5, 254, 259 oceanic crust 250, 255 outer COTZ 247 –50, 248 seismic data 239, 244 –50, 249 study area 240– 2, 241 Avdella Me´lange 361, 362 Ayios Photios Group 353 back-arc basins 178, 325, 327, 368, 369–71 Baer Bassit, Syria 346–7, 347 Ba’id Exotic 333, 335 Balkan region 358, 362–4, 365, 366, 370– 1 basalt 244, 328, 370, 371, 373 type 48, 328, 333, 351, 373 basin architecture 173, 177, 299, 315, 318, 447, 449 basin fill 173, 176, 299, 315, 316 basin formation 2, 54, 57, 173, 191 failed breakup 413, 414–15, 414 inversion 178 kinematic modelling 173–98 lithosphere extension 174–5, 177 synrift sag basins 6, 7, 218 see also back-arc basins; forearc basins; foreland basins; interior basins Basin and Range Province, regional detachment faults 421–41 Bay of Biscay 65, 68, 70, 73 Beaver Dam Mountains 421, 422, 423, 426, 428, 432 anticline 430 normal faults 429, 429, 434– 5 Beysehir– Hoyran–Hadim nappes 357, 357 BIRPS, seismic reflection data 174 Black Sea 178– 81, 186 3D modelling 188 compression 173, 188 depth-dependent stretching model 186, 187 eastern 181, 182, 186– 8 elastic thickness 186 evolution 173–4, 183 fault related extension 188– 91 kinematic modelling 173–98 location map 179 model detachment depth 186 model profile 184

oceanic crust 191 structural section 180 subsidence 178–88, 191 syn- & postrift stratigraphy 188, 189–90, 191 see also western Black Sea Bosnian Flysch 359 boudinage 77, 83, 88, 89, 105– 6, 106, 151 –3 lower crust 93–4, 94 Bouguer gravity anomalies 63, 66– 9, 67, 68, 74 Bourg d’Oisans Basin 315 breakup 3 –4 melt facilitated 378 steady-state breakup flow 201, 203, 204– 5, 207 –9, 207, 210, 213, 214 triggers 325, 377, 378 volcanics 239 breakup unconformity 9, 10–11, 12, 21, 25, 41, 233, 234, 374 brittle failure 4, 6, 47, 54, 56, 79, 82–4, 83, 84, 85, 101, 104, 106, 174, 232, 299, 305, 389, 394, 397, 399, 414, 432, 458–60 brittle– ductile transition 82, 316, 444, 446, 449, 453 Brunhes axial intrusion 229 Brunhes isochron 219, 406 Byerlee’s Law 145 California 421, 423– 4, 430–2 CAM lines 28 Canyon Range thrust 421, 425, 426 Castle Cliff detachment 421, 428, 434 geometrical difficulties 434–5 initial dip 427, 434 restored geometry 427 rolling hinge model 428 –9, 432 seismic reflection 428–9, 429, 432 slide blocks 429 Central African Rift System 268 Cheshire seamount 217, 224, 225, 229 Chevron construction 175 Congo–Angola margin 394–7, 396 Conrad Deep 270– 1, 271 continent –ocean boundary (COB) 79, 240 continent –ocean transition see ocean–continent transition zone Cottonwood Mountains 421, 430, 436

474 coupling 83, 88, 93, 106, 111, 165, 291, 316, 393 brittle–ductile 121, 444 fluid mechanics 188, 201, 457 mechanical 146– 53, 147, 153–4 zones 154 creep 79, 115, 163, 265– 6, 282 crust 40, 164 brittle strength 57, 113 cooling strengthening 134, 135 fault movement 175–6, 390, 390 inherited weakness 291, 325, 373, 375, 378 lower crust removal 95–6 lower crust strength 129, 130, 131, 134, 135 lower crustal flow 93, 94 mechanical extension 175– 6, 277–8 mechanical localization 291 mechanical strength 258 prerift stratigraphy 296– 7 radiogenic heat production 161 rheology 84, 93, 101 thermally weakened 343 thickness variations 56, 157, 165, 167, 174, 192, 194, 296 –7 transitional 65– 6 velocity gradients 56 weak & strong crust models 115–16 yield strength 164, 170 Cyprus 351–3, 354, 355 Dead Sea Rift 268 Dead Sea transform 278 Death Valley 423–4, 430– 2, 436 Deborah number 142– 3, 146, 153 decompression melting 39, 47, 56, 58, 157, 167 decoupling 55, 87, 88, 129, 134, 285, 291, 315, 316 pattern of 83 within crust 87– 9 zones 93, 105–6 Deep Galicia Margin 101–4, 103 Deep Sea Drilling Project (DSDP) 79, 241, 242 density 139– 56, 176, 250 contrasts 66, 139–41, 140, 142, 146–53, 147– 53, 154, 251–2, 278 numerical model 143–6 seismic velocities 139 temperature dependence 117 D’Entrecasteaux Islands 231–2 depth-dependent stretching 77, 83, 87–90, 88, 101, 186, 212, 373– 4, 414 crust 93– 6, 94 decoupling 285 evidence for 394–8 failed breakup basins 413, 413 fluid flow model 398, 398, 404

INDEX lithosphere rheology 104, 106 lower lithosphere 89– 90, 107 mantle exhumation 3–4, 7, 390–8 model 95, 95, 191, 199 sag basins 6, 7, 286 volcanic margins 390, 392, 393 depth uniform stretching 373, 389, 397, 398, 404 detachments 7, 36, 47, 54, 79, 84, 85, 153, 154, 224, 233, 239, 303, 316, 373 associated strain 421 crustal scale 443 deformation age 305 depth 177, 194 geometry 317 interpretation of 81, 432 low angle 112, 393, 450 magmatic system 318 mapping 296 migrating depth model 194 recognizing 92, 92 regional 421– 41 rooted 432, 434 seismic characteristics 77, 81, 312 serpentine 87, 103– 4, 105–6, 106, 107 top basement 313 Dhiarizos Group 353 Diamantina Zone 239, 241, 243, 244, 247– 50, 248 Dinaride ophiolite belt 359 domino model 96, 97, 98 Drucker–Prager yield criterion 116, 457, 458, 461, 461, 463, 465– 8, 467, 471 ductile layer 5, 54, 79, 120, 232 deformation 4, 305, 443 flow law 143, 160 power law 145, 163 shear zones 48 dykes 23, 230, 328, 373, 421 dynamic models 1, 2, 4 –5, 157– 72, 458 Eagle Mountain Formation 421, 422, 430–2 alluvial fan 430 –2, 436 depositional setting 435– 6 fluvial– lacustrine 430 –2, 431, 436 provenance 421, 424, 436 tectonic transport 424, 436 earthquakes 228, 234 Eastern Mediterranean margins 325– 88 emplacement history 343, 344 field photos 339 microcontinents 343–4 palaeogeography 367 settings 353, 378 subsidence curves 374 terrane dispersal 344

Eastern Tethys margins 325–88, 326 Alpine–North Atlantic similarities 371– 3 breakup 377 deep sea drilling 329 magmatic arc 369 magmatism 371, 375, 377, 378 palaeogeography 329 pulsed rifting 375, 376, 377, 378 sediment settings 369 subdivision 326 –7 subduction 369–71, 375, 377 subsidence curves 374 suture zones 370 tectonic settings 370, 375, 377 transform rifts 330 volcanic v. non volcanic 373, 377 working model 375–8 see also Adria margin; Cyprus; Himalayan margin; Levant margin; North Tauride margins; Oman; Pelagonia margin; Pindos Zone; South Tauride margin; Syria; Vardar Zone Egum graben 219, 224, 225, 229, 234 emplaced rift margins 326–9 database 328–9 proximal/distal 327–8 ERABLE 28, 36 erosion 176, 178, 328, 393 Err –Platta 303–8 breakup age constraints 308 detachment system 303–5, 306, 307 exhumed mantle 305, 306, 307– 8 fault blocks 305 magmatic activity 305, 307 melt compositions 308 pillow lavas 307 sediments 305, 307 serpentinized peridotite 307–8 shear zone 305, 306 Euler pole 63, 74, 232, 234 Eurasian Plate 64– 5, 65– 6 evaporites 13, 18, 50, 54, 394, 396 Exmouth Plateau margin 394, 395 extension accommodation 1, 6, 397, 413 balance 7, 413, 414 distribution of 93, 165, 318 heterogeneous 89, 107, 128–9 localized 111, 121, 318 minimum rate of dissipation 113 onshore/offshore systems 5– 6 pressure reduction 82 quantifying 65–6 rheological evolution 85 scale of 128–9, 174 extension discrepancy 90– 104, 191, 192–4, 196, 393, 397, 414 inverse 93, 95

INDEX solutions 92– 9, 92, 94, 94, 106, 393 extension factor 79, 82, 84, 85, 87, 90, 96, 105–6, 106 extension mode 6, 55, 112, 113, 118– 22, 119, 120, 291, 313– 16, 318 conceptual model 294–6, 294, 296 exhumation 313–15, 314, 315–16 model boundary conditions 291 overprinting of 315 selection 123, 124– 5, 124, 127 thinning 313– 15, 314 total internal dissipation 122, 122 transition 122, 123, 123, 318 extension rate 122, 125– 6, 126, 129, 131– 2, 134, 139– 56 melt suppression 57 model 143– 6, 147 –53, 153 Rayleigh Taylor ratio 139 variations 141–2, 154 Faroes– Shetland Basin 414–15 fault blocks 5, 217, 372, 378 ages 101 –3 internal deformation 96 magnetic estimates 282 rotation 6, 77, 97, 221–4, 222, 233, 265, 269, 277 –8, 282, 394, 396 fault heaves 390, 390, 397, 412, 413 faults 113, 312– 13, 315– 16, 328, 374 2D discontinuity surface 460 3D movement 105– 6 antithetic 225– 8, 226–7, 229 associated strain 468 back stepping 316 and basin architecture 313, 318 concave downwards 313, 315, 316 dip 422 extension discrepancy 90–9, 196 failure models 460–2 failure plane 457, 460, 461, 462, 463, 464 fluid circulation 56, 83– 4, 432 folded 432 formation of 459, 463, 468 friction angles 460, 464, 465 friction coefficient 432, 463, 464 frictional sliding 421, 461 inverted 374 large offset 265, 285 listric 90, 299, 312–13 lock up angle 421, 422 low angle 90– 2, 228, 315, 443 low temperature 305 mantle exhumation 274, 316 migrating 102– 4 model assumptions 175– 6 propagating 316, 317

reactivation 65– 6, 178, 468 restoration of 99, 102–4, 192, 313 reverse 39, 178, 181 rotation 96, 218, 443, 450 second generation 83, 96– 9, 98, 99, 105 –6, 393, 468 seismic resolution 192, 327, 393 strain evolution 313 stress field 459– 60 strike slip 219 synrift accommodation 285 tilted 432 transfer 232 transform 234 transtensional system 219 upwards concave see listric faults see also normal faults; polyphase faulting finite element model 124–7, 157, 159–64, 164, 170, 217, 404 assumptions 164 constraints 159– 60 Eulerian 115 Galicia Bank–Flemish Cap 166 heat equation 162–3 melt production 163–4 parameters 162, 167 remeshing 163 rift structure 159, 160–1, 170 strain rate 168 STRCH95 157, 162– 4 subsidence history 161– 2 targets 160–2 timing of extension phases 159 Flemish Cap 13, 36, 40, 47–61, 55, 63– 76, 79, 80, 153, 157–72 basement high origins 54 Cretaceous motion 63– 76 crustal thinning 52– 5 magma production 56– 7, 58 necking zone 52, 55 reconstruction, N America 69–71 rift phases 58, 71 rock compositions 52 rotation 63, 69, 73–4, 73 SCREECH-1 51 sediments 50, 52, 58 seismic data 47, 53, 57 seismic velocity model 51, 52 serpentinization 55–6 structural trends 71 tectonic setting 49– 52 thermal gradients 52 thinning factors 54, 54 see also Galicia Bank–Flemish Cap Flemish Pass basin 63, 74 flexural backstripping 390, 397 flexural isostasy 173, 175, 176, 177, 328, 458 fluid flow model 398, 400, 404–13, 458

475 assumptions 399, 404 bathymetry 399, 402, 404, 406, 411 cohesion angle 464 depth-dependent stretching 399, 404 exhumed mantle 399, 402, 404 flow field pattern 399–402, 400 free air gravity anomaly 404 Galicia Bank 408– 13 inverse theory 404 iso-viscous lithosphere 403 isostatic response 399, 402 margin geometry 402, 404 rheology 404 sea-floor spreading 401, 402, 403 strain balance 404 streamlines 399, 401, 403 temperature structure 401, 402, 404, 405 thinning factors 402 velocity ratio 403–4 Woodlark Basin 404–6 fluvial– lacustrine deposits 430– 2, 431 forearc basins 369–70 foreland basins 178, 183 Fourier transform technique 278 free air gravity anomalies 199, 202 friction, coefficient of 228, 233–4, 432, 463, 464 frictional plasticity 4, 114, 118–21 Galicia Bank 13, 18, 28–9, 36, 40, 50, 55, 57, 63– 76, 79, 80, 101–4, 106, 153 best fit model 411–12 crustal structure 410 dynamic model 157 –72 fluid flow model 408– 13 free air gravity 411, 412, 412 polyphase faulting 101 S reflection 48 seismic reflection 19, 20–1, 22– 3, 25, 27, 32–3, 35 stepped bathymetry 411, 412, 412 Galicia Bank– Flemish Cap 157–72 exhumed mantle 159 extension duration 159, 160, 165 finite element model 166 locus of extension 157, 159, 170 onset of sea-floor spreading 167 rift structure 159–60 subsidence history 161 –2 Galicia Interior Basin 50, 55, 57, 63, 79, 101, 157, 159, 161, 301, 315 depth-dependent stretching 412 exhumed mantle 411 hinge zones 69 isostatic response 414 locus of extension 170 thinning factors 412 velocity section 100, 102

476 Gavrovo– Tripolitza nappe 358 Gawler Craton 255, 256 GEBCO, topographic grid 66 Generoso Basin 297, 299, 300, 315 Geophysical Seismic Incorporated (GSI) 71, 72 geotherm 28, 57, 77, 82, 85, 86–7, 86, 90, 176, 194, 399, 425, 444 Goban Spur 199–215, 276–7, 390 bathymetry 202, 202 depth-dependent stretching 201, 203, 391 exhumed mantle 201, 212– 13, 212 gravity data 202, 202 location maps 202 restored thinning 201–2 seismic data 212 stretching factors 203 Godene zone 348, 349 Gondwana 240, 242 Gorringe Bank 23–5 GP lines 22–3, 25, 28 Grand Banks 13, 17, 18, 29 gravitational instabilities 139– 56, 201 temperature dependence 146, 153 viscosity variations 146 gravity anomalies see Bouguer gravity anomalies; free air gravity anomalies Great Australian Bight margin 239– 63 basement ridge 256, 257 breakup age 256 crust rheology 258– 9 decompression melting 241 forward model domains 256 intrusions 256 magnetic quiet zone 256 mechanical strength 258 modelled structure 255 Moho shallowing 255 and Naturaliste Plateau 257 oceanic crust 256 prerift lithosphere 258– 9 sequence ages 255, 257 serpentenized peridotites 256 thermal reworking 258 Guevgueli ophiolite 363 Gulf of Suez 278 Hawasina Complex 331, 333, 334, 335, 337 Haybi Complex 331, 333, 335 heat flow 2, 66, 140, 141, 176, 258 Hercynian orogeny 330, 371 Himalayan margin 325– 88, 339 collisional deformation 337 exotics 341 exposures 337– 43 flexural uplift 338 me´lange 340–1

INDEX ophiolite emplacement 338 pulsed rifting 341– 2 reactivation 378 sediments 337– 8, 338 spreading onset 375 subsidence 341, 342, 374 tectonic setting 342 volcanism 337, 340–2, 343 hinge zones 63, 66–9, 67, 74 hornblende blocking temperature 21 Hunter Mountain batholith 421, 430, 436 IAM-9 lines 16, 17, 28, 38 Iberia Abyssal Plain 13, 66, 79–82, 80, 106, 157, 301 H detachment 81 magnetic anomalies 82 seismic refraction data 90– 2 structural geology 81 see also Southern Iberia Abyssal Plain Iberia margin 1, 2, 3– 5, 13, 18, 28, 48, 157, 158 bathymetry 12, 406 C reflection 313, 315 free air gravity 406 hinge zone 66– 9 kinematic modelling 192 –6, 195 magmatic history 157 magnetic anomalies 12 patterns of thinning 77 polyphase rifting 18, 299– 301 reconstruction 301 rift duration 157, 299 S reflector 373 stratigraphic cross section 195–6, 195 subsidence history 157 upwelling divergent mantle flow 389– 419, 412 see also Newfoundland–Iberia; West Iberia Margin Iberia Seismic Experiment 28 Iberian Plate 64– 5, 65– 6, 68, 70 igneous intrusions 16, 30, 191, 282 cooling history 28 magnetic anomaly 232 strain accommodation 232 synrift 77, 87, 90 see also dykes; sills Il Motto basin 297, 298, 299 Indus suture zone 337, 338, 338, 340, 341 inherited strain model 111, 113, 129, 130, 131, 134, 136 interior basins, formation 139, 146, 147, 151–3, 153– 4, 164, 165, 170 InterMARGINS Extensional Deformation of the Lithosphere (IMEDL) 5 –6, 63, 174, 192– 6

intrusions see igneous intrusions Ionian Sea 344, 346, 358 ISE lines 4, 19, 20– 1, 35, 157, 192, 193–4 isostasy 7, 176, 233, 313, 399, 402, 404 see also Airy isostatic compensation; flexural isostasy isoviscous corner flow model 201 isoviscous stream function solution 399 Isparta Angle 347, 348, 349, 352 Izmir–Ankara–Erzincan suture zone 355, 356, 364 J anomaly 17, 25, 38, 40, 310 J Anomaly Ridge 9, 17, 23–5, 38, 39, 40 Jeanne d’Arc Basin 298, 299, 315 Jebel Akhdar carbonate platform 333, 334, 335 Jebel Misfah exotic 333, 334 Jebel Qamar exotics 333, 335 Karamba Complex 340, 342 Kemer units 348–51, 349 Kerguelen Plateau 239, 244 Kerguelen plume 242, 255, 259 kinematic modelling 1, 2, 5, 63, 111–12, 145, 173–98, 177, 217 2D 173–4 3D 173–4, 187– 8 crustal thickness 187, 188, 190 elastic thickness 188 extension phases 194 parameters 176– 7, 177, 218 rift stratigraphy 187–8, 189–90, 191, 194 subsidence 187–8, 192 surface processes 176– 8 temperature re-equilibrium 187 Konigsberger ratio 251, 253 Labrador Basin 71, 74 Lagrangian algorithm 145 integration point 444 reference frame 464 thermo-mechanical model 115 Lamayuru Complex 340, 341 Levant margin 345, 345 see also North Africa– Levant margin Levant Sea 346 lithosphere depth-dependent stretching 106 diffuse weakness region 443, 445, 447– 50, 451– 2, 453 elastic thickness 176, 194 extension evolution 77– 110, 105– 6, 169 faults 443, 445, 450–3, 452– 3 frictional forces 374

INDEX heterogeneity 87, 111– 38, 164, 444, 445 incompressible viscous formulations 457– 72 loading 175, 176, 253 and mantle viscosity 140 melt trapping 87 model of inheritance 127 prerift thermal constraints 160 pristine 115 rheological evolution 77, 83, 104, 106, 291 rheology 6, 157– 9, 160, 162, 174, 257, 265, 318, 443, 444, 457 scale of failure 460 single weak seed 443, 445– 7, 446–51, 453 strain softening 111– 38 strength 153, 154, 160, 162, 458–60 temperature field 86, 157 –9, 173, 176 thermal evolution 118, 121, 176, 443 thermal extension models 56 thermal properties 58, 162, 258, 444 thinning model 398–404 weakness 5, 111, 114 –15, 157, 160, 258, 443, 444, 453 Lugano–Val Grande fault 299, 300, 301, 312 Lycian Nappes 356, 356, 357 McKenzie model 174, 175 Madeira– Tore Rise 9, 17, 23– 5, 30, 38, 39, 40 magmatism 1, 169, 191, 316–18, 328 evolution 56–7, 318 sea-floor spreading 318 synrift 77, 157, 169, 170 volumes 169, 318, 325 magnetic anomalies 3, 9, 16, 27, 32–3, 40, 47, 63, 240, 310 see also J anomaly magnetic field 63 declination 65 inducing 251, 281 inversion 65 polarity 252 –3 magnetic remanence 251, 281 magnetic susceptibility 251, 252–3 Mamonia Complex 351 –3, 353 Man Con Son Basin 413 mantle 55, 89, 316 adiabat 82, 85, 86, 167, 191, 377 anomalously cool 57, 169 boundary 86, 297 downwelling 139, 153 fluid dynamics 199, 377, 457, 458, 471 fluid interaction 55, 83 melting 85–90

necking 5, 157 P-wave velocities 56, 56 potential temperature 85, 86, 86, 87, 89, 157, 163– 4, 167– 9, 169, 170 serpentinization 55–6 solidus 85 strength 164, 170, 253, 458 temperatures 89, 458 velocity structure 56, 57, 87 see also plumes; unroofing; upwardly propagating flow; upwelling divergent flow field; upwelling divergent flow model mantle exhumation 5, 7, 9, 16, 17, 27, 30, 33, 38, 40, 48, 55, 56, 65, 77, 87, 89, 118, 134, 148–53, 153–4, 191, 239, 240, 256–8, 303, 317, 328, 372, 373, 389 depth-dependent stretching 390– 8 geochemistry 398 melt suppression 415 model prediction 199 processes 316 –18 see also zone of exhumed continental mantle (ZECM) mantle upwelling 77, 94– 5, 95, 199, 200, 201, 265, 399 see also upwardly propagating flow; upwelling divergent flow field; upwelling divergent flow model Manzala Rift 268 MARGINS see InterMARGINS Margna fault 303, 315 mass wasting 31, 33, 38, 39, 41, 82 melt discrepancy 86, 90 melt lenses 87 melt production 17, 47, 58, 88, 90, 106, 157, 160, 167 predicted 163– 4, 167–9, 169 suppression 57, 106, 213–14 volumes 47, 85, 86, 107 see also decompression melting melt trapping 87 Mersin Me´lange 356– 7 Mesozoic ocean basins 325–88 metamorphism 330 –1, 334, 340, 358, 432 microcontinents 325, 329, 343 –4, 364, 368, 377, 377, 378 microplates 47, 52, 58, 63–5, 64–5, 74 mineral transformations 113– 14, 114 Moho 52, 54, 56, 91– 2, 183, 265–6, 297 Mohr– Coulomb yield criterion 7, 96, 113, 116, 145, 444, 457, 461–3, 461, 465, 471 Moresby fault 217, 221, 225–8, 226–7, 229, 231, 232 dip 217, 228, 233 –4 dykes 230, 232, 234

477 mineralogy 228 motion on 228–9 synrift sediments 229, 230– 1 Moresby seamount 217, 221, 222, 223, 229, 232, 233, 406 kinematic reconstruction 224 synrift sediments 225–8, 226– 7 uplift 224–5 Mormon Mountains 422, 423, 426 Mormon Peak detachment 421, 426–30, 427, 428, 434–5 alternative explanations 427, 435 geometrical difficulties 434–5 initial dip 427, 434 kinematic indicators 427, 428, 432, 435 polymictic conglomerate 427, 428, 435 slide blocks 427– 8, 432 Mulbeck exotics 340 mylonitization 316, 425, 432, 433 Nam Com Son Basin 414– 15 Naturaliste Fracture Zone 244 Naturaliste Plateau 239– 63 crust rheology 258– 9 dredge hauls 243, 243 exhumed mantle 244 fault-controlled basins 247 free air gravity 244, 250 and Great Australian Bight 257 magmatic source 253–5 magnetics 244 origins of 241– 2 prerift lithosphere 258– 9 sedimentary sequence 241, 247 seismic reflection 242–3, 244 spreading anomalies 242 thermal reworking 258 Navier–Coulomb criterion 145 Navier–Stokes equation 163 necking 4, 5, 47, 52, 55, 93–4, 105–6, 106, 118, 129, 157, 164, 177, 446 Neotethys 327, 344, 345, 346 restored section 347, 355 subduction 347, 370 Nevada 421, 426 Newfoundland basin 4, 11, 11, 25, 26, 34, 36, 66 Newfoundland margin 1, 3– 5, 29, 47–61, 158, 406–8 bathymetry 50, 407– 8, 409 crustal structure 409 fluid flow model 406– 8 free air gravity 407 –8, 409 isostatic response 407 magnetic anomaly map 50 multiple rifting 407 oceanic crust 407 SCREECH survey 49, 407 temperature perturbation 407 upwelling divergent flow 389–419, 409

478 Newfoundland– Iberia 9, 12, 13–33, 40, 111, 129–33, 291– 324, 372 Aptian sediments 28 asymmetric shear 132 bathymetric map 10 biostratigraphic age 14–15 breakup 9– 42, 303–12 chaotic resedimentation 38– 9 compression 33–6 conceptual model 294–6 cross-section 159 deep stratigraphy 28– 33 detachment system 304 extension phases 9, 14– 15, 18, 29, 36, 40, 130, 132 first-order model 134 gravity inversion 133 initial plate positions 63–5, 64–5 limited extension 33, 38 location of extension 129, 132 magmatism 23–5, 30 magnetic anomalies 25, 132 mantle exhumation 9, 27, 131 model constraints 133, 135 restored sections 129 –32, 132, 293 rift structures 17– 18, 299, 301 –3 rifting onset 297, 298 seismic profiles 9, 16–17, 32– 5, 133, 293 seismic sequences 29–30, 31– 3, 38 stress state 41 unconformities 14–15 uplift 39 non-volcanic margins 3, 6, 9, 39, 47–9, 56, 57, 77– 110, 112, 131, 169, 170, 199, 201, 239– 40, 325 along strike complexity 240 bathymetry 406, 411 characteristics 77, 372, 378, 404 cross-section 78 evolution 296 extension processes 240, 312– 18 fluid flow streamlines 405 identifying features 328 mantle exhumation 398 OCT 398, 404 synrift gabbros 318 normal faults 5, 29, 40, 54, 217, 221, 222, 232, 233, 313, 394, 397 crustally rooted 422 dip 96, 98, 228 high angle 423 large offset 224 low angle 2, 7, 47, 421, 422, 432, 433 orientation 63 phases of 99 synrift accommodation 265 North Africa–Levant margin 344–6 North American hinge zone 66–9

INDEX North American Plate 64–5, 65–6, 68, 70 North Atlantic, continent reconstruction 161 North Moresby Graben 221, 222, 223, 229, 231, 233 asymmetry 225 bathymetry 225 igneous intrusion 231 magnetization 225 ring dykes 225, 229 seismic sections 226 shallow angle normal fault 225 subsidence 224– 5 synrift sediments 232 North Tauride margins 355– 8 geometry 357– 8 metamorphism 355 ophiolites 355 palaeogeography 356 volcanic v. non volcanic 358 Northern Angolan margin 410, 411 northern Red Sea 269 –77 axial depression 265, 274, 283 bathymetry 269, 270, 271– 5, 279, 279, 280, 281 Bouguer gravity anomaly 278, 279, 280, 282 calculated gravity 278, 281, 282 crustal extension 271, 277– 82, 285 density contrast 278 expanding spread profiles 270– 1 fault block rotation 269, 277–8, 285 free air gravity anomalies 269– 70, 276, 278– 9, 281 intrusions 271, 273 magnetic anomalies 271–3, 271– 5, 277, 278, 280– 1 magnetic susceptibilities 279–80 mantle upwelling 283 Moho 270, 271, 278, 282 reef terraces 269 sea-floor spreading 274–5, 282– 5 sediments 269, 285 shear zone 271 volcanoes 265, 273–5, 276, 277, 283 Norwegian margin 390, 392, 393, 393 oblique rifting see transform rifts ocean–continent transition zone 1, 2, 9 –10, 15– 16, 26–7, 48–9, 79, 239 –63 definition 239–40 width 258, 389, 394, 395, 402, 404, 414 Ocean Drilling Program (ODP) 4, 15– 16, 17, 27–8, 29, 31, 32– 3, 50, 54, 79, 81, 101, 103, 192, 217, 228, 232, 296, 327, 328, 373

oceanic crust 10–11, 30 anomalous 47, 56, 241 Penrose type 9 –10 Oman 325– 88, 332 two-stage opening model 334– 5 breakup 334, 375 exotics 332, 333, 334, 336 geological mapping 329 Northern Mountains 335– 7, 336 ophiolite 334, 378 pillow lavas 333 plume signature 337 prerift sediments 330–1 reconstruction 334–5, 336 rift evolution 330–7 seamounts 334– 5 serpentinite 337 subsidence 331, 374 syntectonic sedimentation 333 tectonic emplacement 330, 330, 331 thermal subsidence 337 thrust sheets 331, 332, 334, 335 transform rifting 335–7 uplift and erosion 330 orogenic belts 157, 160, 164, 170, 294, 325– 8, 330, 371 Orphan Basin 63–76 hinge zones 69 normal fault orientation 63 prerift sediments 71 rift phases 74 structural map 72 Othris Mountains 359–60 Paikon unit 363 Palaeotethys 326– 7, 369, 371 Panjal Traps 325, 337, 343, 371, 375 Papua New Guinea 217– 38, 404–6 Par(a)ovoz code 145 Parnassus microcontinent 359, 360 Pavant thrust 421, 425– 6, 425 Pelagonian microcontinent 344, 359–62, 361, 363, 363, 364, 365, 365 Peonais unit 363 peridotite 9, 10, 13–15, 17, 81, 244, 307–8 see also serpentinization peridotite ridges 21, 27, 29, 31, 33, 35, 40, 50, 79, 310, 316, 317, 360 Photang thrust 340 Photaskar thrust sheet 340 Piedmont– Ligurian Ocean 373 pillow lavas 328, 333, 348 Pindos Mountains 359–60, 361 Pindos Ocean 344, 358, 359, 360, 361–2, 363, 368 Pindos Zone 365–7, 366, 370 Pindos–Olonos nappes 358– 9, 368 plagioclase blocking temperature 21

INDEX plastic deformation 7, 79, 116, 116, 154, 163, 265– 6, 282, 397, 462 plastic– ductile model 160 plastic gain force 123, 125– 7, 129 plastic viscous model 124– 7, 126 boundary conditions 125 equilibrium equations 117 geometry 125 parameters 117 principle of minimum dissipation 111, 121 –4, 123, 124– 7 statistical heterogeneity 127– 8, 128 plasticity 111 frictional 4, 114, 118–21 solid mechanical formulations 457–8 strength 120, 121 theory of 462 plate boundary forces 325, 374, 375, 377 plate tectonic reconstruction 65–7, 68, 70, 73, 73 plumes 9, 16, 25, 39, 40, 86, 87, 239, 285, 325, 343, 371, 393 crustal weakening 375, 378 geochemical signature 328, 333, 337, 378 large scale 374 Pogallo fault 302, 303, 313, 315 Poisson ratios 52 polyphase faulting 77, 93, 96–101, 99, 101, 104, 107, 316 3D 99 model 91 recognition of 99, 106 Pontresina, Switzerland, workshop 1 –7, 63 Porcupine– Seabight Basin 71–3, 74, 84, 315 pore fluid pressure 114, 114 pure shear 4, 54, 66, 83, 90, 107, 111– 12, 123, 174, 186, 201, 397, 398, 404, 457, 462 mode 111, 122, 122, 123, 129 models 67, 296, 373–4, 407 prebreakup 205 regionally distributed 176 stretching factor 199 radiogenic heat production 445 radiometric age dates 21, 81, 305, 308– 10, 312 Rayleigh Taylor instabilities 4 –5, 139, 140, 147– 53 development of 142 growth rate 139, 141, 142– 3 negative 153 thermo-rheological parameters 142 Red Sea 218, 265–89 axial trough 265, 269 bathymetry 267

crustal thickness 268 dyke intrusion 269 extension difference 265 free air gravity map 283 heat flow 268 hinge zone 265 initiation of rifting 268–9 lithosphere 265, 268 location map 267 prerift constraints 265–8 southern 265, 284–6, 284 spreading centre 269, 285 synrift sediments 269 volcanism 265, 268– 9 see also northern Red Sea rheological model 93, 145, 444–54 Arrhenius viscosity 444 basin formation 447, 449 coupled/decoupled 445–53, 446– 53 density 146 Ellipsis code 444, 454 finite element code 444 free slip boundary condition 444 heat flux 444, 445 internal angle of friction 145 Moho temperature 444 Mohr– Coulomb failure 444 parameter values 146, 444, 444 prerift model 233, 297 rift mode 444– 5 scaling 454 strain softening 443–53, 445, 446– 53 weak seed position 446– 7, 449– 51 rheology 6, 127, 154, 160, 163, 443–4 boundaries 48 coupled/decoupled profiles 444, 445 crustal 432 dominant 127 evolution 85, 305, 318 geological timescales 458 heterogeneities 297 lithosphere 174– 5 non linear 457 power law 444 strain dependent 55 viscous 115 rift architecture 111–38, 139, 153, 203, 327 Australia 239 classification 112 end member styles 112 evolution 218 geometrical hardening 118, 121 initial weaknesses 443– 55 parameters 398 primary controls 443 and rift mode 259 strain softening 443–53, 446– 53

479 rift margins 389 3D 113 active 111, 145 amagmatic mechanical rifting 265 asymmetric extension 118–21, 291, 318, 328 conjugate 63–76 distal structure 301–3, 318, 325 driving mechanisms 373, 374– 5, 377 evolution 39, 139– 56 failed 222, 232 intermediate type 325, 327 onset unconformity 21, 233 partial melting 153, 244, 256 passive rifting 6, 111, 145 polyphase rifting 142 –3, 215, 318, 325, 328, 371 proximal structure 318, 326 rheological evolution 305, 318 synrift sag basins 4, 265, 284–5 tectonic emplacement 325–88 velocity 111, 118, 134 volcanism 160, 257, 371 see also emplaced rift margins rift mode 4, 111, 153, 443, 444 and architecture 443–4 asymmetric plug mode 111–13, 122, 122, 123, 129, 134 asymmetry 47– 8, 55, 112, 143, 146, 147– 8, 154, 233, 374, 443, 447, 450 brittle–ductile transition 446, 449 core complex 443, 450–1, 453– 4 fixed boundary collapse 443 free boundary collapse 443 narrow 112, 443, 445– 6, 449–50, 453 selection 443, 453–4 symmetric plug mode 111, 122, 122, 123, 123, 129, 134 symmetry 111–13, 112, 149–53, 153, 328, 374, 443 wide 112, 443, 449–50, 453 Rockall Trough 71, 74, 315 rocks mechanics data 143, 143 strain-dependent strength 113– 14 rolling hinge model 7, 90, 92, 315, 428–9, 432, 434, 450 Rustaq Exotic 333, 335 SCREECH lines 4, 15– 16, 17, 26, 28, 34, 38, 49, 55, 56, 157 sea-floor spreading 16, 39, 49, 58, 94, 106, 265, 398, 404 accretion 27, 30 asymmetric 160 fluid flow model 401, 402, 403 initiation model 389, 399 magma production 56–7, 160 magnetic anomalies 63, 241, 406 mantle convection models 404

480 sea-floor spreading (Continued) modes 218 normal 33, 39, 40, 160 progressive/episodic 375 propagating 375, 377 rates 56– 7, 139, 153, 239, 415 segments 229 spreading centre 218, 219, 265, 375 seamounts 25, 325, 334–5, 340, 346– 8, 372, 378 sediments 5, 41, 50 accumulation rates 30, 31, 38, 77 channelized 431, 436 coarsening upwards 431 cross stratification 431–2, 436 fining upwards 431, 436 post rift 9, 41, 67, 89, 178, 188 prerift 93, 96 synrift 97, 97, 99 see also basin infill seismic reflection 15, 49, 63, 174, 192, 239, 250 decoupling 87–9 stratigraphic markers 13, 21 seismic refraction 10, 40, 192 seismographs, Ocean Bottom Seismographs 52, 56 Semail ophiolite 330, 331 Serbo–Macedonian continent 363, 364 Serbo–Macedonian Zone 363–4, 365 serpentinization 57– 8, 77, 83–4, 84, 104, 105 –6, 106, 240, 291, 305, 316, 337 compressional wave velocities 17 density reduction 55, 57, 87 detachments 87 exothermic 55, 57 low degree 47 magnetic anomalies 16, 65 and mantle strength 84 rheology 57 thickness 107 Sevier Desert basin 421, 422 high angle normal faults 424– 6 map 423 seismic reflection 425 syndepositional deformation 426 Sevier Desert detachment 421, 422– 3, 424 deformation 424–5, 433 dip 424 interpretations 424–6, 433– 4 lateral continuity 424 offset thrust sheets 425 reflection 433– 4 unconformity 421, 432, 433 SfMargin 199– 215 bathymetry 204, 209, 212–13 coupled fluid mechanics 201 forward modelling 199– 200, 200, 212 Goban Spur application 201– 2, 209–10, 210, 211

INDEX gravity anomalies 204, 209, 212– 13 grid search method 209 –12, 210, 211 least squares method 209 –10, 212 mantle exhumation 204, 212–13 margin formation 200– 1 misfit 209–10, 210, 211, 212 parameter values 199– 200, 200, 201, 202 –8 reference frame 210– 12 stretching factor 203–7, 206 thermal advection-diffusion solution 201 velocity ratio 203, 204– 5, 207–8, 207– 8 Shaban Deep 274–5 shear bands 457, 463, 465– 8 3D 468, 470 cross strike length 468, 471 formation 465, 466, 471 friction angle 457, 465, 466–7 orientation 457, 460–1, 462, 463, 466, 468, 471 patterns 457, 468, 469 strain softening 468, 469, 470 shear stress 145, 460, 460 maximum direction 460, 460 multiple plastic shear modes 128 vertical 175 see also pure shear; simple shear shear zones 4, 71, 74, 105– 6, 106, 115, 118, 127, 134, 299, 313 crustal scale 92 high temperature 305, 306, 316 internal angle of friction 113–14 localized 111, 115, 129 mantle 88, 89, 90 seismic response 313 strain rate 462– 3 strength of 96, 113, 313 sills 23 simple shear 4, 82, 83, 92–3, 107, 111– 12, 174, 398, 462 asymmetric 84 models 218, 296, 303, 373–4 viscosity 457 slide blocks 421, 423, 427– 8, 429, 432, 434–5 South China Sea margin 390, 391, 393– 4 South Moresby Graben 221, 222, 231, 233 South Tauride margin 347–53 emplaced units 347 –51 normal faulting 348 oceanic embayment 347 pillow lavas 348 restored successions 351 seamounts 348, 351 sediments 348, 349 subsidence history 351 volcanics 348, 349, 351

South Tethys 337, 345, 353, 358 Southeast Newfoundland Ridge 9, 17, 23– 5, 30, 38, 39, 40 Southern Iberia Abyssal Plain 308–12, 309 basement cooling ages 309 breakup unconformity 310 C reflection 308 detachment system 308–9, 312– 13 exhumed mantle 309– 10, 311, 312 extension rates 310– 12 H detachment 308–9, 311, 312– 13 Hobby High detachment 309, 311, 312–13 intrusive magmatics rocks 310 kinematic profile inversion 310– 12, 311 L detachment 308– 9, 311, 312– 13 magnetic anomalies 310 ODP 308, 309, 310, 312 seismic reflections 308–10 shear zones 310 time markers 312 Spontang ophiolite 340 strain 1, 115, 315, 318, 327, 443 see also inherited strain model strain rate 4– 5, 79, 127, 142, 145, 163, 165, 239, 259, 299– 301, 444 finite element model 168 strength dependent 114, 120 strain softening 4, 111– 38, 125, 129, 134, 377, 378, 443– 53, 445, 446–53, 453, 462, 471 feedback 111, 115, 116, 118–21, 127, 129, 131, 134, 463– 4 mode selection 113 plastic strength 123, 124 threshold 116, 128 stress 257, 459 field rotation 432 mechanical deformation 56 normal 145 principle direction 461, 461, 463 stable regime 39 strain rate 79 stretching factor 73, 79, 199, 209, 390–3, 391– 3, 449 prebreakup 201, 203–7, 204, 207 subduction 217, 330, 369, 375, 377, 378 subsidence 199, 232, 301, 313, 315, 328, 372, 378, 393– 4 curves 174, 327, 330, 374 fault controlled 174 timing of 5, 6, 7, 174, 218, 390, 397 see also thermal subsidence Sumeini Group 333 suture zones 365, 378

INDEX Syria 346–7 Syrian Arc structures 268 Tagus Abyssal Plain 15, 157 Tauride microcontinent 344, 347, 355 Tauride–Anatolide platform 367 Taurides unit 355 Tekirova unit 351 temperature 163 contrasts 141 field evolution 399 profiles 140 see also thermal tensile stress 28, 30, 146– 54 deviatoric 325 in plane 9, 12, 33, 38, 39, 40 strength 461– 2 transtensional regimes 71, 74 upwelling asthenosphere 374 Tethys back-arc settings 369 –71 blocking successions 365 closure 344 emplacement directions 368 multiple sutures 365–8 North Tethys margins 362–4 ocean basins 326 single Mesozoic ocean 365, 366 strike-slip 365, 368 tectonic reconstructions 364– 8, 366 see also Eastern Tethys margins; Neotethys; Palaeotethys thermal advection/diffusion 129, 201, 399 finite-difference solutions 201 history 199 horizontal gradients 141 model 89 Peclet number 121 processes 318 structure 55, 154, 315 subsidence 141, 153, 174, 194, 337, 389, 390, 413, 414, 414 uplift 373 weakening 259, 291, 318, 325 see also geotherm thermo mechanical model 113, 115– 17, 141 equations 145 geometry 116 lithosphere extension 117– 21 lower crustal strength 118, 119, 120, 130, 131, 135 mode selection 118– 21 numerical 115 –17 power law creep parameters 117 velocity 119, 148 thermo-rheological structure, numerical model 143–6, 144, 144

thinning factor 203, 390 –3, 391–3, 397, 398, 399, 402, 402 thrust sheets emplacement 374, 378 imbricate 327 out of sequence 331–3, 334, 335 piggy back 178, 358 transform faults 234 transform rifts 330, 335–7, 372 transition zone see ocean– continent transition zone Tule Springs detachment 427, 427, 434–5 Tule Springs Hills 426 U/Pb dating 305, 308, 310 unconformities 14–15, 21, 33, 233, 373, 374 see also Aptian event underplating 56, 58, 239, 253, 254, 259, 297, 328, 341, 393 United States 421–41 unroofing 82, 84, 85– 90, 88, 93, 95, 105–6, 107, 241, 415 upper plate paradox 218, 399 upwardly propagating flow 200, 201, 203, 208, 208, 411, 411 field 204–6, 209, 210, 213, 214 model 400, 402, 402, 403, 405, 415 upwelling divergent flow field 94, 199–201, 200, 203, 212, 218, 374, 389, 400, 407, 411–13, 411 upwelling divergent flow model 389–419 advection solution 399 bathymetry 402, 403 boundary conditions 399 ceased propagation 415 divergent velocity 398, 399, 400, 413– 14 failed breakup 414 fixed pattern 402, 403 lithosphere structure 214, 407 mantle exhumation 214, 403 origin of 414 paused flow field 411–13, 411 sag basin 414–15 streamlines 401 temperature structure 407 thinning factors 402, 403 upwelling velocity 398, 400, 413– 14 Utah 421, 426 COCORP Line 424, 425 Vardar Ocean 344, 362–3, 364, 365, 365, 373 Vardar Zone 364, 365–7, 365, 366, 370 Vardoussia unit 359 Variscan Orogeny 157, 160, 164, 170, 294

481 velocity 116, 118, 244, 250, 394 half spreading rate 200, 201 horizontal divergent 398, 399, 400, 413– 14 ratio 203, 204–5, 207–8, 207– 8, 209, 211 stacking velocities 250– 1, 253 upwelling 398, 400, 413–14 viscoelastic displacement formulation 458 viscosity 127, 128, 134, 140, 163, 462 dissipation 125 incompressible 457–72 layered 111 plastic strength 124 pressure dependent 115, 163 temperature dependent 117, 153 transition 123–4, 123, 124, 126–7, 128, 139, 141 viscous deformation 79, 116, 462 viscous penalty 123, 125–7, 129 see also isoviscous viscous fluid flow model 458, 459, 462–8, 471 3D 468– 71, 470, 471 active slip plane 462–3, 463 boundary conditions 464– 5 brittle layer 465, 468, 471 cohesion 465–8, 467 continuum failure model 462– 3 density 458, 465 finite element code 464 flow laws 462–3, 471 Galerkin discretization 464 healing parameter 463–4 incompressible fluid 462– 3, 471 Lagrangian reference frame 464 large deformation 468, 469 localization 463– 4 strain softening 462, 463– 4, 464, 465, 466, 468, 471 stress iterations 462– 3, 464 volcanic arcs 363, 369, 370 volcanic margins 6, 77, 259, 325, 327–8, 372, 373, 378 bathymetry 404, 406 fluid flow streamlines 405 ocean–continent transition 373, 404 see also non-volcanic margins Werner deconvolution 278, 279–82 West Iberia Margin 77–110, 80 depth sections 90– 2 extension discrepancy 99– 104 H detachment 90–2 lower crust 95–6 polyphase faulting 99–101 restored section 99–101 rheological evolution 82–4 thermal evolution 82– 4 western Black Sea 2D kinematic model 181– 6 back-arc extension 178

482 western Black Sea (Continued) cross-sections 181, 183 crustal thinning 183–6 magnitude of extension 178, 183 pure shear 181 resolved faults 183 sediment loading 181 seismic data 178– 81 stratigraphy 178, 183– 6, 185 subsidence 178, 181, 183 thrust belts 181–3, 184 Western Tethys 326, 372– 4 Woodlark Basin 217– 38, 219, 224, 231, 232, 369, 406 acoustic imagery data 234 asymmetric margin 217, 224, 232 bathymetry 221, 234, 406, 408 breakup model 229, 231–2 depth-dependent stretching 406

INDEX Euler poles 218– 21, 220, 228, 229 extension rate 224 faults 220, 221, 222, 225– 9 fluid flow model 404–6 heat flow anomaly 217, 232 igneous intrusions 217, 229– 30, 230, 232, 234 magnetic anomalies 234, 406 neovolcanic zone 224 ODP 221, 222 prerift geology 217, 231 proximal margin 221– 5 rheology 217 sediments 217, 221, 224, 232–3, 234 seismic data 221, 222, 234–5 spreading 218– 21, 224, 225– 30, 230, 233

subsidence 233, 406 synrift sediment velocity model 333 thinning factor 406, 408 unconformities 232–3 uplift 233 upwelling divergent flow model 389– 419, 408 xenoliths, geothermal barometry 86, 86 Zanskar Mountains 338 Zanskar platform 338, 339, 340 Zanskar shelf 340, 341 zone of exhumed continental mantle (ZECM) 10, 79, 303, 305, 309–10, 312, 315, 316– 18, 317

This book summarizes our present understanding of the formation of passive continental margins and their ocean–continent transitions. It outlines the geological, geophysical and petrological observations that characterize extensional systems, and how such observations can guide and constrain dynamic and kinematic models of continental lithosphere extension, breakup and the inception of organized sea-floor spreading. The book focuses on imaging, mapping and modelling lithospheric extensional systems, at both the regional scale using dynamic models to the local scale of individual basins using kinematic models, with an emphasis on capturing the extensional history of the Iberia and Newfoundland margins. The results from a number of other extensional regimes are presented to provide comparisons with the North Atlantic studies; these range from the Tethyan realm and the northern Red Sea to the western and southern Australian margins, the Basin and Range Province, and the Woodlark basin of Papua New Guinea. All of these field studies, combined with lessons learnt from the modelling, are used to address fundamental questions about the extreme deformation of continental lithosphere.