Sedimentary environments offshore Norway--Palaeozoic to Recent: proceedings of the Norwegian Petroleum Society Conference, 3-5 May 1999, Bergen, Norway

Norwegian Petroleum Society (NPF), Special Publication No. 10 Sedimentary Environments Offshore Norway Palaeozoic to R...

Author: Norsk petroleumsforening. Conference | Ole J. Martinsen | Tom Dreyer

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Norwegian Petroleum Society (NPF), Special Publication No. 10

Sedimentary Environments Offshore Norway Palaeozoic to Recent Proceedings of the Norwegian Petroleum Society Conference, 3-5 May 1999, Bergen, Norway

Further titles in the series:

1. R.M. Larsen, H. Brekke, B.T. Larsen and E. Talleraas (Editors) STRUCTURAL AND TECTONIC MODELLING AND ITS APPLICATION TO PETROLEUM GEOLOGY- Proceedings of Norwegian Petroleum Society Workshop, 18-20 October 1989, Stavanger, Norway 2. T.O. Vorren, E. Bergsager, Q.A. DahI-Stamnes, E. Holter, B. Johansen, E. Lie and T.B. Lund (Editors) ARCTIC GEOLOGY AND PETROLEUM POTENTIAL- Proceedings of the Norwegian Petroleum Society Conference, 15-17 August 1990, Tromso, Norway 3. A.G. Dore et al. (Editors) BASIN MODELLING" ADVANCES AND APPLICATIONS- Proceedings of the Norwegian Petroleum Society Conference, 13-15 March 1991, Stavanger, Norway 4. S. Hanslien (Editor) PETROLEUM EXPLORATION AND EXPLOITATION IN NORWAYProceedings of the Norwegian Petroleum Society Conference, 9-11 December 1991, Stavanger, Norway

5. R.J. Steel, V.L. Felt, E.P. Johannesson and C. Mathieu (Editors) SEQUENCE STRATIGRAPHY ON THE NORTHWEST EUROPEAN MARGIN Proceedings of the Norwegian Petroleum Society Conference, 1-3 February, 1993, Stavanger, Norway 6. A.G. Dore and R. Sinding-Larsen (Editors) QUANTIFICATION AND PREDICTION OF HYDROCARBON RESOURCESProceedings of the Norwegian Petroleum Society Conference, 6-8 December 1993, Stavanger, Norway 7. P. Moller-Pedersen and A.G. Koestler (Editors) HYDROCARBON SEALS- Importance for Exploration and Production

8. F.M. Gradstein, K.O. Sandvik and N.J. Milton (Editors) SEQUENCE STRATIGRAPHY- Concepts and Applications Proceedings of the Norwegian Petroleum Society Conference, 6-8 September 1995, Stavanger, Norway 9. K. Ofstad, J.E. Kittilsen and P. Alexander-Marrack (Editors) IMPROVING THE EXPLORATION PROCESS BY LEARNING FROM THE PAST Proceedings of the Norwegian Petroleum Society Conference, September 1998, Haugesund, Norway

Norwegian Petroleum Society (NPF), Special Publication No. 10

Sedimentary nvironments Offshore Norway P a l a e o z o i c to R e c e n t Proceedings of the Norwegian Petroleum Society Conference, 3-5 May 1999, Bergen, Norway

Edited

by

Ole J. Martinsen Norsk Hydro Research Centre, R O. Box 7190, N-5020 Bergen, Norway

and

Tom

Dreyer

Norsk Hydro Research Centre, R O. Box 7190, N-5020 Bergen, Norway

2001 ELSEVIER

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9 2001 Elsevier Science B.V. All rights reserved. This work and the individual contributions contained in it are protected under copyright by Elsevier Science B.V., and the following terms and conditions apply to its use: Photocopying Single photocopies of single chapters may be made for personal use as allowed by national copyright laws. Permission of the Publisher and payment of a fee is required for all other photocopying, including multiple or systematic copying, copying for advertising or promotional purposes, resale, and all forms of document delivery. Special rates are available for educational institutions that wish to make photocopies for non-profit educational classroom use. Permissions may be sought directly from Elsevier Science Global Rights Department, PO Box 800, Oxford OX5 1DX, UK; phone: (+44) 1865 843830, fax: (+44) 1865 853333, e-mail: [email protected]. You may also contact Global Rights directly through Elsevier's home page (http://www.elsevier.nl), by selecting 'Obtaining Permissions'. In the USA, users may clear permissions and make payments through the Copyright Clearance Center, Inc., 222 Rosewood Drive, Danvers, MA 01923, USA; phone: (978) 7508400, fax: (978) 7504744, and in the UK through the Copyright Licensing Agency Rapid Clearance Service (CLARCS), 90 Tottenham Court Road, London W1P 0LP, UK; phone: (+44) 171 631 5555, fax: (+44) 171 631 5500. Other countries may have a local reprographic rights agency for payments. Derivative Works Tables of contents may be reproduced for internal circulation, but permission of Elsevier Science is required for resale or distribution of such material. Permission of the Publisher is required for all other derivative works, including compilations and translations. Electronic Storage or Usage Permission of the Publisher is required to store or use electronically any material contained in this work, including any chapter or part of a chapter. Except as outlined above, no part of this work may be reproduced, stored in a retrieval system or transmitted in any form or by any means, electronic, mechanical, photocopying, recording or otherwise, without prior written permission of the Publisher. Address permissions requests to: Elsevier Science Global Rights Department, at the mail, fax and e-mail addresses noted above. Notice No responsibility is assumed by the Publisher for any injury and/or damage to persons or property as a matter of products liability, negligence or otherwise, or from any use or operation of any methods, products, instructions or ideas contained in the material herein. Because of rapid advances in the medical sciences, in particular, independent verification of diagnoses and drugs dosages should be made. First edition 2001

Library of Congress Cataloging-in-Publication Data Sedimentary environments offshore N o r w a y w Palaeozoic to Recent/edited by Ole J. Martinsen, Tom D r e y e r - 1st ed. p. cm - (Norwegian Petroleum Society (NPF) Special Publication; no. 10) Includes bibliographical references and index. I S B N 0-444-50241-6 1. Petroleum-Prospecting-Congresses. I. Martinsen, Ole J. II. Dreyer, Tom. III. Series TN271.P4 I47 2001 622'. 1828-dc21

00-021616

ISBN: 0-444-50241-6 @ The paper used in this publication meets the requirements of A N S I / N I S O Z39.48-1992 (Permanence of Paper). Printed in The Netherlands

Dedication to Arne Dalland (1945-1998) Arne Dalland, who was a member of the organising committee for the Norwegian Petroleum Society Conference "Sedimentary Environments Offshore Norway Palaeozoic to Recent", died in 1998. Arne was a pioneer in Norwegian onshore and offshore sedimentology, and he was the first sedimentology graduate from the University of Bergen. He worked under the supervision of professor Anders Kvale, carrying out work on the Mesozoic section at AndOya, northern Norway (Dalland, 1975). Later on, he became a senior lecturer at the University of Bergen and worked extensively in Spitsbergen, before joining Statoil in 1983, where he worked until 1998. During Arne's spell at the University of Bergen, both the editors of this volume had the privilege of taking G 103-"Historical Geology" where Arne Dalland lectured on the offshore sedimentology. His lectures, given with Arne's well-known quiet and calm expression but clear and in-depth knowledge, inspired us in our careers as sedimentologists. There is a close tie between the theme of the conference and the conference proceedings published in this volume, and Arne Dalland's extensive work. Arne had a strong focus on detailed sedimentological and stratigraphic work for understanding palaeogeography. He was a leader and forerunner in formalising offshore stratigraphy leading to the publication of Dalland et al. (1988) on offshore Mid- and North Norway stratigraphy. The innovative use of names (such as boat and fish names) that relate to the all-important fishing industry, rather than traditional place names, in this part of Norway is remarkable. Ame Dalland was also a forerunner in using new methodology on offshore data. He initiated studies on using Sm/Nd isotope stratigraphy for reservoir and provenance studies. Arne Dalland was a highly respected and well-known geologist. He had his heart in the science, always being concerned with acquiring the highest quality analyses and results. We find it appropriate to dedicate this volume to Arne for his inspirational ideas and work on the sedimentary environments offshore Norway. References Dalland, A., 1975. The Mesozoic rocks of AndCya, Northern Norway. Norg. Geol. Unders. Skr., 316: 271-287. Dalland, A., Worsley, D. and Ofstad, K., 1988. A lithostratigraphic scheme for the Mesozoic and Cenozoic succession offshore mid- and northern Norway. Norw. Pet. Dir. Bull., 4:65 pp.

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VII

Preface and Acknowledgements The aim of the Norwegian Petroleum Society Conference "Sedimentary Environments Offshore Norway - - Palaeozoic to Recent" held in Bergen from May 3-5, 1999, was to show a representative selection of case studies which illustrated how the sedimentary environments in presently offshore areas had changed through time. Whether we as the organising committee and editors have been successful in this, can only be judged by the readers of the book. Naturally, it is impossible to produce an all-inclusive book with examples from all stratigraphic levels and various offshore basins. However, we believe that this volume contains a selection of papers that give examples of the offshore development over time particularly in the Mid-Norway and North Sea areas. Furthermore, several analogue outcrop examples to the offshore cases are included. Moreover, the papers in this volume record the issues and stratigraphy currently in focus. We are particularly happy to include several papers related to the Holocene development of the Norwegian margin, because understanding this part of the succession is important for hazard evaluation and burial and uplift history of the offshore areas. Readers of the volume should combine these proceedings with the abstract volume (Martinsen and Dreyer, 1999), where extended abstracts are available of all presentations made at the conference. Neither the conference organisation nor the proceedings volume could have been a success without the inspirational help of an effective organising committee and a highly competent group of referees. We would like to thank the other members of the organising committee for providing innumerable ideas and suggestions: Tom Bugge Arne Dalland (deceased) Lars Magnus F~ilt

Roy Gabrielsen Karin Haugna~ss (secretary) William Helland-Hansen

Johan Petter Nystuen Rodmar Ravnfis

We aimed to have an international group of referees judge the quality of the submitted papers in order to secure an international standard of the volume. The referees were very effective in providing comments and feedback and in returning manuscripts on time. Thus, we extend our thanks to the referees for helping us significantly in the preparation of this volume: Morten Bergan Arnold Bouma Tom Bugge Reidulv B0e Mike Charnock Ed Clifton John Collinson Steve Corfield Bob Dalrymple Tony Dor6 Lars Magnus F~ilt Atle Folkestad Roald Fa~rseth Ashton Embry Steve Flint

Bill Galloway Mike Gardner Rob Gawthorpe John Gjelberg William Helland-Hansen Erik Johannessen Ragnar Knarud Dale Leckie Trond Lien Finn Livbjerg Gunn Mangerud Tor Nedkvitne Wojtek Nemec Johan Petter Nystuen Snorre Olaussen

Torben Olsen Cai Puigdefabregas Rodmar Ravnfis Phillip Ringrose Alf Ryseth Ian Sharp Ron Steel Jim Steidtmann Finn Surlyk Kristian S0egaard Mike Talbot Tore Vorren Roger Walker Brian Zaitlin

Most importantly, we would like to thank the authors for being very co-operative in delivering and returning manuscripts on time. Without effective authors, there would be no publications.

VIII

Preface and Acknowledgements

Finally, we thank the Norwegian Petroleum Society for willingness to organise the conference originally based on a hand-written note. Especially, the effortless work by Karin Haugna~ss is deeply appreciated. We would also like to thank Norsk Hydro for giving us time to complete these proceedings. Ole J. Martinsen Tom Dreyer Bergen, October 2000

Reference Martinsen, O.J. and Dreyer, T. (Editors), 1999. Sedimentary Environments Offshore N o r w a y - Palaeozoic to Recent. Extended Abstracts, Norwegian Petroleum Society/NPF Conference, Bergen, May 3-5, 1999, 258 pp.

IX

List of Contributors

J. A N D S B J E R G

Geological Survey of Denmark and Greenland (GEUS), Thoravej 8, DK-2400 Copenhagen NV, Denmark

S. BACKSTROM

Applied Biostratigraphy, Blekksoppgrenda 41, N-1352 KolsCts, Norway

M.O. BADESCU

Delft University of Technology, Faculty of Applied Earth Sciences, P.O. Box 5028, 2600 GA Delft, The Netherlands

H. B R E K K E

Norwegian Petroleum Directorate, P.O. Box 600, N-4003 Stavanger, Norway

T. BUGGE

Norsk Hydro ASA, N-9480 Harstad, Norway

M. CECCHI

Enterprise Oil Norge Ltd., Lekkeveien 193b, N-4002 Stavanger, Norway

M.A. CHARNOCK

Norsk Hydro Research Centre, P.O. Box 7190, N-5020 Bergen, Norway

S. CORFIELD

Department of Earth Sciences, University of Manchester, Oxford Road, Manchester M13 9PL, UK

T. D R E Y E R

Norsk Hydro Research Centre, P.O. Box 7190, N-5020 Bergen, Norway

K. DYBKJ}ER

Geological Survey of Denmark and Greenland (GEUS), Thoravej 8, DK-2400 Copenhagen NV, Denmark

T. ENOKSEN

Norsk Hydro Research Centre, P.O. Box 7190, N-5020 Bergen, Norway

J.I. FALEIDE

Department of Geology, University of Oslo, N-0316 Oslo, Norway

J.EG. FENTON

Robertson Research International Ltd., Llandudno, North Wales LL30 1SA, UK

A. FOLKESTAD

Statoil, Research and Technology, Department of Reservoir Characterisation, N-4035 Stavanger, Norway

R.H. GABRIELSEN

Geological Institute, University of Bergen, N-5007 Bergen, Norway

G.K. GILLMORE

University College Northampton, School of Environmental Science, Northampton NN2 7AL, UK

J. GJELBERG

Norsk Hydro Research Centre, P.O. Box 7190, N-5020 Bergen, Norway

C. GUARGENA

Enterprise Oil Norge Ltd., Lekkeveien 193b, N-4002 Stavanger, Norway

K.-O. HAGER

Norsk Hydro Research Centre, P.O. Box 7190, N-5020 Bergen, Norway

L. HANSEN

Enterprise Oil Norge Ltd., LOkkeveien 193b, N-4002 Stavanger, Norway

G. HELGESEN

Statoil Stavanger, Forusbeen 50, Stavanger, Norway

EN. JOHANNESSEN

Geological Survey of Denmark and Greenland (GEUS), Thoravej 8, DK-2400 Copenhagen NV, Denmark

M.D. JACKSON

Centre for Petroleum Studies, T.H. Huxley School of Environment, Earth Sciences and Engineering, Imperial College, Prince Consort Road, London SW7 2BP, UK

X

List of Contributors

H.D. JOHNSON

Centre for Petroleum Studies, T.H. Huxley School of Environment, Earth Sciences and Engineering, Imperial College, Prince Consort Road, London SW7 2BP, UK

I. KAAS

Statoil Bergen, Sandslihaugen 30, N-5020 Bergen, Norway

J.M. KJiEREFJORD

Statoil Bergen, Sandslihaugen 30, N-5020 Bergen, Norway

E KJiERNES

Norsk Hydro Exploration, N-0246 Oslo, Norway

T. KJENNERUD

SINTEF Petroleum Research, N-7465 Trondheim, Norway

M. KREINER-MOLLER

Geological Institute, University of Copenhagen, r DK-1350 Copenhagen K, Denmark

I.L. KRISTIANSEN

Norsk Hydro Exploration, N-0246 Oslo, Norway

R. KYRKJEBO

Geological Institute, University of Bergen, N-5007 Bergen, Norway

J.S. LABERG

Department of Geology, University of TromsO, N-9037 TromsO, Norway

M. LARSEN

Geological Survey of Denmark and Greenland (GEUS), Thoravej 8, DK-2400 Copenhagen NV, Denmark

D.A. LEITH

Statoil StjOrdal, Strandveien 4, N-7501 StjOrdal, Norway

S.J. LIPPARD

Department of Geology and Mineral Resource Engineering, NTNU, N-7465 Trondheim, Norway

H. LOSETH

Statoil Research Centre, N-7005 Trondheim, Norway

C. MAGNUS

Norwegian Petroleum Directorate, P.O. Box 600, N-4003 Stavanger, Norway

G. MANGERUD

Norsk Hydro Research Centre, P.O. Box 7190, N-5020 Bergen, Norway

A.W. MARTINIUS

Statoil Research Centre, Arkitekt Ebbellsveg 10, N-7005 Trondheim, Norway Present address: c/o Statoil Venezuela - Sincor Project, N-4035 Stavanger, Norway

O.J. MARTINSEN

Norsk Hydro Research Centre, P.O. Box 7190, N-5020 Bergen, Norway

A. MORK

SINTEF Petroleum Research, N-7465 Trondheim, Norway

M.B.E. M{0RK

SINTEF Petroleum Research, N-7465 Trondheim, Norway

A.H. MUGGERIDGE

Centre for Petroleum Studies, T.H. Huxley School of Environment, Earth Sciences and Engineering, Imperial College, Prince Consort Road, London SW7 2BP, UK

A. NAESS

Statoil Research Centre, Arkitekt Ebbellsveg 10, N-7005 Trondheim, Norway Present address: Statoil StjOrdal, Strandveien 4, N-7501 StjOrdal, Norway

T. NEDKVITNE

Norsk Hydro Exploration, N-0246 Oslo, Norway

L.H. NIELSEN

Geological Survey of Denmark and Greenland (GEUS), Thoravej 8, DK-2400 Copenhagen NV, Denmark

N. NOE-NYGAARD

Geological Institute, University of Copenhagen, Oster Voldgade 10, DK-1350 Copenhagen K, Denmark

S. OLAUSSEN

Norsk Hydro Research Centre, P.O. Box 7190, N-5020 Bergen, Norway

D. OTTESEN

Geological Survey of Norway, N-7491 Trondheim, Norway

D. RHODES

Enterprise Oil Norge Ltd., LOkkeveien 193b, N-4002 Stavanger, Norway

Voldgade 10,

List of Contributors

XI

L. RISE

Geological Survey of Norway, N-7491 Trondheim, Norway

A. ROBERTS

Enterprise Oil Norge Ltd., LOkkeveien 193b, N-4002 Stavanger, Norway

E. ROE

Norsk Hydro Exploration, N-0246 Oslo, Norway

K. ROKOENGEN

Norwegian University of Science and Technology, N-7034 Trondheim, Norway

A. RYSETH

Norsk Hydro Research Centre, P.O. Box 7190, N-5020 Bergen, Norway. Present address: Norsk Hydro Harstad, Storakern 11, Kanebogen, N-9401 Harstad, Norway

J. S/ETTEM

SINTEF Petroleum Research, N-7465 Trondheim, Norway Present address: Sauherad Kommune, N-3812 Akkerhaugen, Norway

I. SHARP

Norsk Hydro Research Centre, P.O. Box 7190, N-5020 Bergen, Norway

H.I. SJULSTAD

Norwegian Petroleum Directorate, P.O. Box 600, N-4003 Stavanger, Norway

M. SMELROR

Geological Survey of Norway, N-7491 Trondheim, Norway

R.J. STEEL

University of Wyoming, Department of Geology and Geophysics, Laramie, WY 82071, USA

L. STEMMERIK

Geological Survey of Denmark and Greenland (GEUS), Thoravej 8, DK-2400 Copenhagen NV, Denmark

F. SURLYK

Geological Institute, University of Copenhagen, Oster Voldgade 10, DK-1350 Copenhagen K, Denmark

K.E. SVELA

Norske Conoco A/S, Tangen 7, N-4070 Randaberg, Norway

B. TVEITEN

Norsk Hydro ASA, E & P International, N-0246 Oslo, Norway

J. UNDERHILL

Department of Geology and Geophysics, University of Edinburgh, Edinburgh EH9 3JW, UK

E. VAGNES

Norsk Hydro Exploration, N-0246 Oslo, Norway

T.O. VORREN

Department of Geology, University of TromsO, N-9037 TromsO, Norway

H.M. WEISS

SINTEF Petroleum Research, N-7465 Trondheim, Norway

R.W. WILLIAMS

Norwegian Petroleum Directorate, P.O. Box 600, N-4003 Stavanger, Norway

S. YOSHIDA

Centre for Petroleum Studies, T.H. Huxley School of Environment, Earth Sciences and Engineering, Imperial College, Prince Consort Road, London SW7 2BP, UK Present address: Surface Processes and Modern Environments Research Group, Department of Geology, Royal Holloway, University of London, Egham, Surrey TW20 OEX, UK

This Page Intentionally Left Blank

XIII

Contents Dedication to Arne Dalland (1945-1998) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Preface and Acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . List of Contributors . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

V VII IX

I. Introductory Papers Sedimentary environments offshore Norway O.J. Martinsen and T. Dreyer

Palaeozoic to Recent: an introduction . . . . . .

1

Sedimentary environments offshore Norway ~ an overview . . . . . . . . . . . . . . . . . . . . . . . . H. Brekke, H.I. Sjulstad, C. Magnus and R.W. Williams

II. Palaeozoic The alluvial cyclicity in Hornelen Basin (Devonian western Norway) revisited: a multiparameter sedimentary analysis and stratigraphic implications . . . . . . . . . . . . . . . . . . . . . . A. Folkestad and R.J. Steel

39

Upper Permian lowstand fans of the Bredehorn Member, Schuchert Dal Formation, East Greenland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . M. Kreiner-MOller and L. Stemmerik

51

III. Mesozoic Sedimentology and palaeogeography of the Statfjord Formation (Rhaetian-Sinemurian), North Sea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . A. Ryseth

67

Sedimentary facies in the fluvial-dominated Are Formation as seen in the Are 1 member in the Heidrun Field . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . K.E. Svela

87

Sedimentology of the heterolithic and tide-dominated Tilje Formation (Early Jurassic, Halten Terrace, offshore mid-Norway) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . A.W. Martinius, I. Kaas, A. Na~ss, G. Helgesen, J.M. Kja~refjord and D.A. Leith

103

Sequence stratigraphy of the Lower Jurassic Dunlin Group, northern North Sea . . . . . . . . . . . M.A. Charnock, I.L. Kristiansen, A. Ryseth and J.EG. Fenton

145

Divergent development of two neighbouring basins following the Jurassic North Sea Doming event: the Danish Central Graben and the Norwegian-Danish Basin . . . . . . . . . . . . . . . J. Andsbjerg, L.H. Nielsen, EN. Johannessen and K. Dybkjaer

175

An integrated study of the Garn and Melke formations (Middle to Upper Jurassic) of the SmOrbukk area, Halten Terrace, mid-Norway . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . S. Corfield, I. Sharp, K.-O. H~iger, T. Dreyer and J. Underhill

199

Middle Jurassic-Lower Cretaceous transgressive-regressive sequences and facies distribution off northern Nordland and Troms, Norway . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . M. Smelror, A. MOrk, M.B.E. MOrk, H.M. Weiss and H. LOseth

211

XIV

Contents

Outcrop studies of tidal sandstones for reservoir characterization (Lower Cretaceous Vectis Formation, Isle of Wight, southern England) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . S. Yoshida, M.D. Jackson, H.D. Johnson, A.H. Muggeridge and A.W. Martinius

233

Lower Cretaceous (Barremian-Albian) deltaic and shallow marine sandstones in North-East Greenland ~ sedimentology, sequence stratigraphy and regional implications . . . . . . . . M. Larsen, T. Nedkvitne and S. Olaussen

259

The depositional history of the Cretaceous in the northeastern North Sea . . . . . . . . . . . . . . . . T. Bugge, B. Tveiten and S. B~ickstr6m Cretaceous faulting and associated coarse-grained marine gravity flow sedimentation, Traill O, East Greenland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . F. Surlyk and N. Noe-Nygaard

279

293

IV. Mesozoic-Cenozoic Transition Cretaceous-Tertiary palaeo-bathymetry in the northern North Sea; integration of palaeo-water depth estimates obtained by structural restoration and micropalaeontological analysis... R. KyrkjebO, T. Kjennerud, G.K. Gillmore, J.I. Faleide and R.H. Gabrielsen Structural restoration of Cretaceous-Cenozoic (post-rift) palaeobathymetry in the northern North Sea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . T. Kjennerud, J.I. Faleide, R.H. Gabrielsen, G.K. Gillmore, R. KyrkjebO, S.J. Lippard and H. LOseth The reconstruction and analysis of palaeowater depths: a new approach and test of micropalaeontological approaches in the post-rift (Cretaceous to Quaternary) interval of the northern North Sea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . G.K. Gillmore, T. Kjennerud and R. Kyrkjebr

321

347

365

V. Cenozoic Outcrop-based classification of thick-bedded, deep-marine sandstones . . . . . . . . . . . . . . . . . M.O. Badescu

383

Use of integrated 3D seismic technology and sedimentology core analysis to resolve the sedimentary architecture of the Paleocene succession of the North Sea . . . . . . . . . . . . . M. Cecchi, C. Guargena, L. Hansen, D. Rhodes and A. Roberts

407

The Maastrichtian and Danian depositional setting, along the eastern margin of the MOre Basin (Mid-Norwegian Shelf): implications for reservoir development of the Ormen Lange Field . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . J.G. Gjelberg, T. Enoksen, R Kj~emes, G. Mangerud, O.J. Martinsen, E. Roe and E. V~gnes

421

Glacial processes and large-scale morphology on the mid-Norwegian continental shelf . . . . . . D. Ottesen, L. Rise, K. Rokoengen and J. Sa~ttem Late Quaternary sedihaentary processes and environment on the Norwegian-Greenland Sea continental margins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . T.O. Vorren and J.S. Laberg Accretionary, forced regressive shoreface sands of the Holocene-Recent Skagen Odde spit complex, Denmark ~ a possible outcrop analogue to fault-attached shoreface sandstone reservoirs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . L.H. Nielsen and EN. Johannessen

441

451

457

References index . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

473

Subject index . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

483

Sedimentary environments offshore Norway--- Palaeozoic to R e c e n t : an i n t r o d u c t i o n Ole J. Martinsen and Tom Dreyer

Sedimentary environments offshore Norway have evolved through time as a response to changing climatic conditions, basin physiography and tectonic setting. This volume includes examples from various time periods and offshore basins, including some adjacent onshore and offshore analogues from Greenland, Denmark and Great Britain. These examples illustrate characteristics of the sedimentary environments of various time periods, from the Devonian of the onshore Hornelen Basin to the Holocene of the Mid-Norway area, including continental, shallow-water and deep-water depositional settings. Cases range from detailed facies analysis of highly prolific, hydrocarbon-bearing Jurassic reservoir rocks, through to Recent, giant submarine slides, showing the changes in processes and setting that the Norwegian offshore areas have experienced. The examples from analogues in East Greenland are particularly important both to understand the downdip evolution of environments in the MOre and VOring Basins, but also because the onshore areas on the Norwegian mainland largely lack a post-Devonian sedimentary record. This volume presents an account of the sedimentary development of the Norwegian basins offshore Mid-Norway and in the North Sea region, and thus complements earlier volumes dealing with the Arctic areas.

Introduction From May 3-5, 1999, more than 200 geologists assembled in Grieghallen, Bergen to discuss the sedimentary development of the Norwegian offshore areas and their analogues. The conference was organised by the Norwegian Petroleum Society. The participants came from universities, government agencies and oil and consulting companies in Norway, Denmark, Great Britain, Holland, France, Spain, Canada and USA. The presentations included 25 talks, 10 core examples and 37 posters. Particular emphasis was on the posters and the core examples, recognising that it was within these presentations that most of the data were presented. Three keynote addresses were given: by Robert Dalrymple on non- and marginalmarine environments, by H. Edward Clifton on shallow-marine environments, and by Arnold Bouma on deep-water environments. For various reasons, papers from these keynote addresses were not included in the volume, but their abstracts are included in the abstracts volume (Martinsen and Dreyer, 1999).

Sedimentary environments in Norway through time This conference volume includes 23 papers, from both the poster and the oral presentations. They range from detailed analysis of single stratigraphic units to

overview articles. Previous volumes treating the Norwegian offshore geology have dealt with correlation methodology (Collinson, 1988), sequence stratigraphy (Steel et al., 1995; Gradstein et al., 1998) or concentrated on Arctic geology (Vorren et al., 1993). Since the present volume has papers mainly from the Mid-Norway and North Sea Basins, it complements the volume edited by Vorren et al. (1993) on the Arctic areas. Together, these two volumes give a comprehensive account of how both the northern and southern basins evolved through time. Both publications are required reading for geologists working in the offshore areas. In the following, we give a brief review of the papers contained in the present volume and put the information into a time-stratigraphic context and provide some comparison with onshore data from Norway. Palaeozoic

As is the case in the Arctic and in the Barents Sea region (cf. Vorren et al., 1993), Palaeozoic sedimentary successions are poorly known from offshore Norway, although they most likely exist in considerable thickness and may link up with the onshore successions (e.g. Fa~rseth et al., 1995). Onshore, Precambrian strata in Finnmark (e.g. Siedlecka, 1975) and Cambrian-Carboniferous sedimentary successions in the Oslo region (e.g. Olaussen et al., 1994) are well

Sedimentary Environments Offshore Norway - Palaeozoic to Recent edited by O.J. Martinsen and T. Dreyer. NPF Special Publication 10, pp. 1-5, Published by Elsevier Science B.V., Amsterdam. 9 Norwegian Petroleum Society (NPF), 2001.

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known. Furthermore, The onshore Devonian sedimentary record from western Norway (Andersen, 1998; Folkestad and Steel, 2001) is a widely known and most impressive record of basin formation and sedimentation by anyone's standard. However, the existence of pre-Devonian sedimentary record in western Norway is less well known (e.g. Brekke and Solberg, 1987; RavnSs and Fumes, 1995 and references therein). Although the Devonian and older rocks are unlikely to form reservoir rocks offshore Norway, their onshore existence and character are valuable for implications of the offshore development. The overview given by Brekke et al. (2001) on pre-Mesozoic offshore sedimentation is valuable although the database is sparse. Carboniferous reservoir successions are well known from the British offshore areas (e.g. Collinson et al., 1993). The Carboniferous succession offshore Norway is perhaps a speculative reservoir target because of large burial depths in most basins, but as of the present day it is untested and still a possibility where burial depth is not excessive. Permian sedimentary rocks are well known in the Barents Sea (Vorren et al., 1993), but further south, there is little information from offshore areas. In the Oslo Graben, there is a Permian sequence (Olaussen et al., 1994). In fact, the Permian may be the least well-known time period in from Norwegian offshore and onshore areas. Therefore, the data and interpretations given by Kreiner-Mr and Stemmerik (2001) from well-exposed Permian deep-water deposits in East Greenland are valuable, because they provide important ideas for the sedimentary evolution of this time period in adjacent areas offshore Norway. Mesozoic The Mesozoic Era is obviously the most important time period offshore Norway for hydrocarbon resources and implicitly for the occurrence of reservoir and source rocks. Especially, there was a major change of sedimentary environments from continental in the Triassic through mainly coastal plain and nearshore marine to deep water in the Cretaceous. Brekke et al.'s (2001) overview paper covers several aspects of these changing sedimentary environments. The Triassic period was dominated by arid, continental deposition (e.g. Steel and Ryseth, 1990; Steel, 1993), while the latest Triassic to earliest Jurassic saw a progressive change into a wetter climate with the initiation of shallow-marine sedimentation. Ryseth (2001) documents the important changes that took place during the deposition of the Statfjord Formation, the major reservoir-bearing sandstone formation that is important both on the eastern and western margin of the Viking Graben in major fields such as

O.J. Martinsen and T. Dreyer

Snorre, Statfjord and Oseberg. In Mid-Norway, equivalent Lower Jurassic sandstones are also highly important as hydrocarbon reservoirs and Svela (2001) describe the Are Formation in the Heidrun Field. Another important stratigraphic unit in the Mid-Norway area is the Tilje Formation, which forms both a primary and secondary reservoir unit in several Mid-Norway fields. The tidally dominated Tilje Formation marks the change from mainly fluvial deposition in the ,~re Formation to shallow-marine, tidally dominated sedimentation. Martinius et al. (2001) document the facies and sedimentary patterns within the Tilje. Early Jurassic sedimentation in the North Sea Basin was dominated by shallow marine sedimentation recorded by the Dunlin Group. The Dunlin Group has played a subordinate role as a reservoir unit, but Charnock et al. (2001) document the regional depositional patterns and sequence stratigraphy. The Dunlin Group may have an unexplored reservoir potential, and the hydrocarbon resources could be underestimated. The Brent Group, the most prolific oil reservoir unit on the Norwegian shelf, lies on top of the Dunlin Group. It marks a change from margin fed shallow marine systems in Dunlin Group (see Charnock et al., 2001), to an axial system which prograded and retrograded largely along a N-S axis. The Brent Group has been amply documented in earlier publications (see for instance the papers in Steel et al., 1995). While the Early and Middle Jurassic periods were dominated by relative tectonic quiescence, tectonism becomes very important in the Late Jurassic. Extensional faulting controls depositional patterns, and sedimentary environments vary considerably on a local scale. This situation is important both offshore Mid-Norway, in the North Sea and in adjacent areas, and Corfield et al. (2001), Andsbjerg et al. (2001) and Smelror et al. (2001) show examples of the relationships between tectonism and sedimentation during this period. One particularly important part of the Jurassic reservoirs is the reservoir behaviour and characterization. Analogue work is important for this aspect and Yoshida et al. (2001) show an example from tidal sandstones in southern England that compares with the tidally dominated Tilje Formation. Other field analogues for the Jurassic and the lower Cretaceous stratigraphy offshore Norway are found in East Greenland, and both Larsen et al. (2001) and Surlyk and Noe-Nygaard (2001) show important examples. Tectonism and rifting in the Late Jurassic and earliest Cretaceous mark the change from shallow-marine sedimentation to deep-water sedimentation. The example by Larsen et al. (2001) from Greenland (see above) shows this change. In the Norwegian offshore

Sedimentary environments offshore N o r w a y - Palaeozoic to Recent: an introduction

areas, Cretaceous reservoirs have been difficult to locate as the passage from the Jurassic to Cretaceous time also records a change from a sand-rich to a mudrich depositional setting. Bugge et al. (2001) describe the depositional history in the northern North Sea, an area with some Cretaceous discoveries, such as in Agat, but where the volumes of hydrocarbons so far have been too small for commercial exploitation. Understanding the basin development through the Cretaceous-Cenozoic period is challenging, and severe effort has been put into solving this important question for making predictive models. The three papers by KyrkjebO et al. (2001), Kjennerud et al. (2001) and Gillmore et al. (2001) show different approaches of restoring palaeobathymetry. Such modelling is important because they provide ideas on how basins filled and thus where source and reservoir rocks may be located. Sand-rich deep-water turbidite systems are present in the western Voring Basin (Brekke et al., 2001). These systems relate to supply from East Greenland, showing the importance of the western basin margin of the offshore Mid-Norway area for supply of sands (Larsen et al., 2001; Surlyk and Noe-Nygaard, 2001). The western margin of the VOring and MOre Basins still has a high potential for hydrocarbon exploration. Cenozoic

The Palaeocene of the North Sea and the MidNorway area has proven to be a prolific, oil- and gas-bearing reservoir succession. Sedimentation took place in deep-water fans and related depositional systems (Badescu, 2001; Cecchi et al., 2001; Gjelberg et al., 2001). The change from a relatively mud-rich Cretaceous period to more sandy Palaeocene systems relate to basin margin uplift and tectonism in concert with incipient rifting in the North Atlantic (Martinsen et al., 1999; Brekke et al., 2001). This volume lacks papers dealing with the Eocene-Pliocene periods, and thus we give a short review to complete this introduction. The EoceneMiocene period saw a change from deep-water to shallow-marine conditions in the North Sea (Dalland et al., 1988; Isaksen and Tonstad, 1989; Martinsen et al., 1999). This period is in general poorly documented on the Norwegian shelf in terms of sedimentary history. Many wells, drilled for deeper targets, have no data from much of this stratigraphic succession. The petroleum potential of the EoceneMiocene is highly questionable because of shallow burial depths. Nevertheless, the depositional history has more than academic interest, because the accumulation of these stratigraphic successions, and the overlying Pliocene, caused underlying packages to

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reach burial depths where hydrocarbons could form, migrate and accumulate in reservoir-bearing successions of older age. The Pliocene is also relatively poorly documented in general, despite its great importance for burial of older units, and the fact that it records the highest sedimentation rate on the Norwegian shelf (e.g. Jordt et al., 1995; Rokoengen et al., 1995; Henriksen and Weimer, 1996). The thick Pliocene wedges probably relate to glaciation, high erosional rates and consequent high sedimentation rates (cf. Rokoengen et al., 1995 and references therein). The Pleistocene and Holocene periods are dominated by glacial sedimentation and erosion in the offshore area. Ottesen et al. (2001) show how glacial erosion has created large-scale erosional features and depositional products on the Mid-Norwegian continental shelf. In a time-stratigraphic sense, erosion is probably more important than sedimentation in many areas (e.g. Sejrup et al., 1996; Ottesen et al., 2001). A major component of the Pleistocene and Holocene history is the occurrence of giant submarine slides. Vorren and Laberg (2001) describe several examples and their occurrence. Bugge et al. (1987) and Haflidason et al. (1999) described the Storegga slide, the largest known submarine slide (see also the illustration on the front cover of the book). The sedimentary environments and patterns during this recent period have major implications for installation of hydrocarbon-producing equipment on or above the sea floor, and thus the importance of understanding the Quaternary development cannot be underestimated. In addition, studies of the Recent Skagen Odde complex in northern Denmark show a valuable modern analogue for Jurassic, fault block-related shoreface sands (Nielsen and Johannessen, 2001), which shows the importance of studying the present to understand the ancient. Conclusions

This volume covers many important aspects of sedimentation and sedimentary environments offshore Norway from the Palaeozoic to the Recent. The sedimentary geological setting has changed significantly over this time period as a response to changing tectonic setting, basin physiography and morphology. Further knowledge will be attained through an increasing database with new wells, and in particular a growing seismic database where 3-D data play a vital role. Our prediction is that the most important progress in understanding the development of sedimentary environments through time offshore Norway will be made from studying seismic data. There is a general knowledge on how sedimentary environments changed from core data, and naturally these data have

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to be supplemented by new core data. However, a fully three-dimensional understanding of plan view morphology and cross-sectional architecture can only be attained from high-resolution seismic data used with a sedimentologist's eye. References Andersen, T., 1998. Extensional tectonics in the Caledonides of southern Norway, an overview. Tectonophysics, 285:333-351. Andsbjerg, J., Nielsen, L.H., Johannessen, EN. and Dybkja~r, K., 2001. Divergent development of two neighbouring basins following the Jurassic North Sea doming event: the Danish Central Graben and the Norwegian-Danish Basin. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 175-197 (this volume). Badescu, M.O., 2001. Outcrop-based classification of thick-bedded, deep marine sandstones. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore N o r w a y - Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 383-405 (this volume). Brekke, H. and Solberg, EO., 1987. The geology of Atl0y, Sunnfjord, Western Norway. Nor. Geol. Unders. Bull., 410: 677-690. Brekke, H., Sjulstad, H.I., Magnus, C. and Williams, R., 2001. Sedimentary environments offshore N o r w a y - an overview. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 7-37 (this volume). Bugge, T., Befring, S., Belderson, R.H., Eidvin, T., Jansen, E., Kenyon, N.H., Holtedahl, H., Sejrup, H.E, 1987. A giant threestage submarine slide off Norway. Geo-Mar. Lett., 7: 191-198. Bugge, T., Tveiten, B. and B~ickstr0m, S., 2001. The depositional history of the Cretaceous in the northeastern North Sea. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway - - Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 279-291 (this volume). Cecchi, M., Guargena, C., Hansen, L., Rhodes D. and Roberts, A., 2001. Use of integrated 3D seismic technology and sedimentology core analysis to resolve the sedimentary architecture of the Palaeocene succession of the North Sea. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 407-419 (this volume). Charnock, M.A., Kristiansen, I.L., Ryseth, A. and Fenton, LEG., 2001. Sequence Stratigraphy of the Lower Jurassic Dunlin Group, northern North Sea. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway - - Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 145-174 (this volume). Collinson, J.D. (Editor), 1988. Correlation in Hydrocarbon Exploration. Graham and Trotman, London, 381 pp. Collinson, J.D., Jones, C.M., Blackbourn, G.A., Besly, B.M., Archard, G.M. and McMahon, A.H., 1993. Carboniferous depositional systems of the southern North Sea. In: J.R. Parker (Editor), Petroleum Geology of Northwest Europe: Proceedings of 4th Conference. Geological Society, London, pp. 677-687. Corfield, S., Sharp, I., H~iger, K.-O., Dreyer, T. and Underhill, J., 2001. An integrated study of the Garn and Melke Formations (Middle to Upper Jurassic) of the Sm0rbukk area, Halten Terrace, mid-Norway. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway - - Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 199-210 (this volume).

O.J. Martinsen and T. Dreyer Dalland, A., Worsley, D. and Ofstad, K., 1988. A lithostratigraphic scheme for the Mesozoic and Cenozoic succession offshore midand northern Norway. Norw. Pet. Dir. Bull., 4:65 pp. Fa~rseth, R., Gabrielsen, R.H. and Hurich, C.A., 1995. Influence on basement in structuring of the North Sea Basin. Nor. Geol. Tidsskr., 75: 105-119. Folkestad, A. and Steel, R.J., 2001. The alluvial cyclicity in Hornelen Basin (Devonian western Norway) revisited: a multiparameter sedimentary analysis and stratigraphic implications. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway - - Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 39-50 (this volume). Gillmore, G.K., Kjennerud, T., Kyrkjeb0, R., 2001. The reconstruction and analysis of palaeowater depths: a new approach and test of micropalaeontological approaches in the post-rift (Cretaceous to Quaternary) interval of the northern North Sea. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway - - Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 365-381 (this volume). Gjelberg, J.G., Enoksen, T., Kj~ernes, E, Mangerud, G. et al., 2001. The Maastrichtian and Danian depositional setting, along the eastern margin of the MOre Basin (Mid-Norwegian Shelf): implications for reservoir development of the Ormen Lange Field. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway - - Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 421-440 (this volume). Gradstein, F., Sandvik, K.O. and Milton, N.J. (Editors), 1998. Sequence Stratigraphy: Concepts and Applications. Norwegian Petroleum Society (NPF), Special Publication 8. Elsevier, Amsterdam, 437 pp. Haflidason, H., Gravdal, A., Sejrup, H.P., Bryn, E, Lien, R. and Mienert, J., 1999. TOBI iamgery side-scan sonar and seismic data of the northern escarpment of the Storegga Slide off Mid-Norway: evidence for long-term instability. In: O.J. Martinsen and T. Dreyer, (Editors) 1999. Sedimentary Environments Offshore Norway-Palaeozoic to Recent. Extended Abstracts, Norwegian Petroleum Society/NPF Conference, Bergen, May 3-5, 1999, pp. 205-207. Henriksen, S. and Weimer, E, 1996. High-frequency depositional sequences and stratal stacking patterns in lower Pliocene coastal deltas, Mid-Norwegian continental shelf. Bull. Am. Assoc. Pet. Geol., 80:1867-1895. Isaksen, D. and Tonstad, K., 1989. A revised Cretaceous and Ter tiary lithostratigraphic nomenclature for the Norwegian North Sea. Norw. Pet. Dir. Bull., 5:59 pp. Jordt, H., Faleide, J.I., Bj0rlykke, K. and Ibrahim, M.T., 1995. Cenozoic sequence stratigraphy of the central and northern North Sea Basin: tectonic development, sediment distribution and provenance areas. Mar. Pet. Geol., 12: 845-879. Kjennerud, T., Faleide, J.I., Gabrielsen, R.H., Gillmore, G.K., Kyrkjeb0, R., Lippard, S.J. and L0seth, H., 2001. Structural restoration of Cretaceous-Cenozoic (post-rift) palaeobathymetry in the northern North Sea. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 347-364 (this volume). Kreiner-M011er, M. and Stemmerik, L., 2001. Upper Permian lowstand fans of the Bredehorn Member, Schuchert Dal Formation, East Greenland. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway - - Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 51-65 (this volume). Kyrkjeb0, R., Kjennerud, T., Gillmore, G.K., Faleide, J.I. and Gabrielsen, R.H., 2001. Cretaceous-Tertiary palaeo-bathymetry in the northern North Sea; integration of palaeo-water depth estimates obtained by structural restoration and micropalaeontological

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analysis. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway - - Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 321-345 (this volume). Larsen, M., Nedkvitne, T. and Olaussen, S., 2001. Lower Cretaceous (Barremian-Albian) deltaic and shallow marine sandstones in North-East Greenland - - sedimentology, sequence stratigraphy and regional implications. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway - - Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 259-278 (this volume). Martinius, A.W., Kaas, I., Na~ss, A., Helgesen, G., Kj~eref]ord, J.M. and Leith, D.A., 2001. Sedimentology of the heterolithic and tide-dominated Tilje Formation (Early Jurassic, Halten Terrace, offshore Mid-Norway). In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 103-144 (this volume). Martinsen, O.J. and Dreyer, T. (Editors), 1999. Sedimentary environments offshore Norway Palaeozoic to Recent. Extended Abstracts, Norwegian Petroleum Society (NPF) Conference, Bergen, May 3-5 1999, 258 pp. Martinsen, O.J., Belen, F., Charnock, M., Mangerud, G. and Nottvedt, A., 1999. Cenozoic development of the Norwegian margin 60-64~ sequences and sedimentary response to variable basin physiography and tectonic setting. In: A.J. Fleet and S.A.R. Boldy (Editors), Petroleum Geology of Northwest Europe: Proceedings of the 5th ConfErence. Geological Society, London, pp. 293-3O4. Nielsen, L.H. and Johannessen, RN., 2001. Accretionary, forced regressive shoreface sands of the Holocene-Recent Skagen Odde spit complex, Denmark a possible outcrop analogue to faultattached shoreface Sandstone reservoirs. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 457-472 (this volume). Olaussen, S., Larsen, B.T. and Steel, R., 1994. The Upper Carboniferous-Permian Oslo Rift: basin fill in relation to tectonic development. In: A.F. Embry, B. Beauchamp and D.J. Glass (Editors), Pangea-Global Environment and Resources. Can. Soc. Pet. Geol. Mem., 17:175-197. Ottesen, D., Rise, L., Rokoengen, K. and Saettem, J., 2001. Glacial processes and large-scale morphology on the mid-Norwegian continental shelf. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway - - Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 441-449 (this volume). Ravnfis, R. and Fumes, H., 1995. The use of geochemical data in determining the provenance and tectonic setting of ancient sedimentary successions: the Kalv~g Melange, western Norwegian Caledonides. In: A.G. Plint (Editor), Sedimentary Facies Analysis A Tribute to the Reseach and Teaching of Harold G. Reading. Int. Assoc. Sediment. Spec. PUN., 22: 237-264. Rokoengen, K., Rise, L., Bryn, R, Frengstad, B., Gustavsen, B., Nygaard, E. and Saettem, J., 1995. Upper Cenozoic stratigraphy on the Mid-Norwegian continental shelf. Nor. Geol. Tidsskr., 75: 88-104. Ryseth, A., 2001. Sedimentology and palaeogeography of the Statfjord Formation (Rhaetian-Sinemurian), North Sea. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Off-

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shore Norway Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 67-85 (this volume). Sejrup, H.R, King, E.L., Aarseth, I., Haflidason, H. and Elverhr A., 1996. Quaternary erosion and depositional processes, western Norwegian fjords, Norwegian channel and North Sea Fan. In: M. De Batist and R Jacobs (Editors), Geology of Siliciclastic Shelf Seas. Geol. Soc. London, Spec. Publ., 117: 187-202. Siedlecka, A., 1975. Late Precambrian stratigraphy and structure of the northeastern margin of the Fennoscandian Shield (East Finnmark-Timan region). Nor. Geol. Unders., 316:313-348. Smelror, M., M0rk, A., M0rk, M.B.E., Weiss, H.M. and L0seth, H., 2001. Middle Jurassic-Lower Cretaceous transgressive-regressive sequences and facies distribution off Northern Nordland and Troms, Norway. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 211-232 (this volume). Steel, R., 1993. Triassic-Jurassic megasequence stratigraphy in the Northern North Sea: rift to post-rift evolution. In: J.R. Parker (Editor), Petroleum Geology of Northwest Europe: Proceedings of the 4 th Conference. Geological Society, London, pp. 299-315. Steel, R., Felt, V.L., Johannessen, E. and Mathieu, C. (Editors), 1995. Sequence Stratigraphy on the Northwest European Margin. Norwegian Petroleum Society (NPF), Special Publication 5. Elsevier, Amsterdam, 608 pp. Steel, R. and Ryseth, A., 1990. The Triassic-Early Jurassic succession in the Northern North Sea: megasequence stratigraphy and intra-Triassic tectonics. In: R.F.R Hardman and J. Brooks (Editors), Tectonic Events Responsible for Britain's Oil and Gas Reserves. Geol. Soc. London, Spec. Publ., 55: 139-168. Surlyk, F. and Noe-Nygaard, N., 2001. Cretaceous faulting and associated coarse-grained marine gravity flow sedimentation, Traill O, East Greenland. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway - - Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 293-319 (this volume). Svela, K.E., 2001. Sedimentary facies in the fluvial-dominated ,~re Formation as seen in the .~re 1 member in the Heidrun Field. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 87-102 (this volume). Vorren, T.O. and Laberg, J.S., 2001. Late Quaternary sedimentary processes and environment on the Norwegian-Greenland Sea continental margins. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 451-456 (this volume). Vorren, T., Bergsager, E., Dahl-Stamnes, O.A., Holter, E., Johansen, B., Lie, E. and Lund, T.B. (Editors), 1993. Arctic Geology and Petroleum Potential. Norwegian Petroleum Society (NPF), Special Publication 2. Elsevier, Amsterdam. 751 pp. Yoshida, S., Jackson, M.D., Johnson, H.D., Muggeridge, A.H. and Martinius, A.W., 2001. Outcrop studies of tidal sandstones for reservoir characterization (Lower Cretaceous Vectis Formation, Isle of Wight, southern England). In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 233-257 (this volume).

Norsk Hydro Research Centre, P.O. Box 7190, N-5020 Bergen, Norway Norsk Hydro Research Centre, P.O. Box 7190, N-5020 Bergen, Norway

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Sedimentary environments offshore Norway--- an overview Harald Brekke, Hans Ivar Sjulstad, Christian Magnus and Robert W. Williams

The evolution of sedimentary environments in the Norwegian continental margin since the Early Carboniferous is directly linked with the evolution of the tectonic framework of the broader region of the northern North Atlantic. After the end of the continent-building Variscan and Uralian Orogenies in the Permo-Carboniferous, the tectono-magmatic history is that of a 200 million year period of general extension and rifting of the continent, ending with a final continental rupture and opening of the northern North Atlantic by seafloor spreading in Early Eocene times. The sedimentary environment of the Norwegian continental margin therefore is a record of an evolution from waning orogenic forelands and incipient rifts in an equatorial climate, through a stage of dominantly continental rifting while drifting from equatorial to temperate climates, to the present stage of a passive subsiding continental margin in a temperate to arctic climate. In earliest Carboniferous times the general lateral distribution of sedimentary environments reflects the existence of a northern ocean (the Boreal Ocean) and a southern ocean (the Proto-Tethys) separated by an area of emergent land with a central zone of continental rifting. In the south, the Carboniferous was generally a period of marginal marine, fluvial deltaic and alluvial deposits progressively filling up the central and southern North Sea. Purely continental alluvial, fluvial and lacustrine environments prevailed in the Norwegian Sea and East Greenland. In the Barents Sea, alluvial and fluvial deltaic environments were transgressed at an early stage by marine carbonates and evaporites. The Early Permian period of the North Sea, East Greenland and the Norwegian Sea was a time of continental environments including an early episode of widespread magmatism in the south. In the Barents Sea marine carbonate and evaporite environments prevailed. Middle Permian time was characterised by uplifts and large erosional breaks. Although most prominent in the southern and central areas, such an erosional hiatus is also recorded across much of the Barents Sea. At the end of the Permian period the sea transgressed the low-lying parts of the entire region - - recorded by coarse clastics and evaporites in the south and central area and fine-grained clastics in the Barents Sea. Characteristic of the Triassic period were the numerous marine transgressions and regressions both in the north and south of the region. In the south, the evaporitic environment of the Permian continued but with an increased input of clastics. The northern North Sea, the Norwegian Sea and East Greenland were characterised by marine deposits in the Lower Triassic, followed by continental fluvial and alluvial systems interbedded with marine incursion cycles. A main feature of Triassic times was the shallowing of the Barents Sea by input of large volumes of clastic sediments. A relative sea-level rise, that started in latest Triassic times, caused the Lower and Middle Jurassic of the whole region to become uniformly dominated by shallow marine clastic shelf environments and approximately simultaneous delta oscillations. Early to Middle Jurassic domes and uplifts on regional and semi-regional scales caused a complex pattern of hinterlands, depo-centres and seaways. In the latest Middle Jurassic and through Late Jurassic times, a major sea-level rise considerably deepened the northern and southern seas and finally drowned the central area (between East Greenland and Norway). This caused the widespread accumulation of marine shale with intervals of very rich source rock. Following a period of marked oscillations of the sea level prior to the Aptian, the sea level continued to rise through the Cretaceous period and reached its peak in Late Cretaceous times. Lower Cretaceous deep-water shales and marls accumulated in the basins and rifts of the southern and central parts of the region, while shallow marine and coastal plain deposits dominate on the flanking platforms and in the vast platform of the Barents Sea. The facies pattern of the Lower Cretaceous continues unchanged into the Upper Cretaceous in the central province, while the high sea level gave rise to pelagic limestones in the southern. The central Barents Sea was transgressed with the development of a condensed Upper Cretaceous marine sequence of clastics and carbonate. The volcanism, tectonism and regional uplift preceding the earliest Tertiary continental break-up and subsequent seafloor spreading between Greenland and Norway, effectively ended the carbonate environments in the south, and the whole region became dominated by marine clastic deposits. In the Neogene the stratigraphy is a record of oscillating glaciations. The glaciations and regional uplifts caused deep erosion of the surrounding mainland areas and the Barents Sea shelf in the latest Neogene and the progradation of a huge sediment apron onto the margins of the Norwegian-Greenland Sea.

Introduction

The Norwegian continental shelf is an integral part of the North Atlantic continental margin, and extends

from the central North Sea and well into the Arctic Ocean north of Svalbard. Prior to the opening and seafloor spreading of the Norwegian-Greenland Sea, the present surrounding continental margins were

Sedimentary Environments Offshore Norway - Palaeozoic to Recent edited by O.J. Martinsen and T. Dreyer. NPF Special Publication 10, pp. 7-37, Published by Elsevier Science B.V., Amsterdam. 9 Norwegian Petroleum Society (NPF), 2001.

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joined in one continent, and their sedimentary environments and tectonic history were closely related. The authors have chosen the approach to discuss the sedimentary environments and associated tectonic evolution in a palaeogeographic framework. In this respect, the present study falls in the family of regional compilations like the ones of Ziegler (1987, 1988, 1990), Dor6 (1991, 1992), Dor6 et al. (1999) and Roberts et al. (1999). The scope of the present work is by no means an attempt to make a full revision of the work of Ziegler and others. The aim is to update the areas that are of most relevance to the Norwegian continental margin, i.e. the North Sea, the Norwegian-Greenland Sea, and the western and northwestern Barents Sea, by including some new data and own ideas. Some extra attention is paid to the Carboniferous stratigraphy and tectonic framework, and to the possible configuration of sedimentary basins and emergent land areas at all times in the area of the present Norwegian-Greenland Sea. In the major parts of the area, the Carboniferous is the transition period between convergent plate tectonics and the subsequent long history of intraplate extension and rifting tectonics. Due to overprinting by later events, the complex Carboniferous tectonic framework is subject to much interpretation and assumptions. Some extra details and new information are therefore provided to illustrate the diversity of sedimentary environments of that period. The tectonic framework of the Carboniferous is assumed to have influenced the subsequent tectonic development, including the hinterland/basin configurations through time. Tectonic and stratigraphic framework During the time from the Early Devonian to the Eocene, the region under study developed from a phase of plate conversion and continent growth, through a period of rifting, until subsequent continental rupture (e.g. Anderton et al., 1979; Ziegler, 1988; Glennie and Underhill, 1998; Dor6 et al., 1999). Subsequent to the onset of seafloor spreading, the continental margin has been subject to compression. Through earliest Carboniferous to Late Permian times the region of the present North Atlantic and Barents Sea was a part of the Pangean supercontinent characterised by orogenic accretion around its fringes associated with the Inuitian, Variscan and Uralian Orogenies (Fig. 1). Still, the interior of the region was subject to rifting from the present southern North Sea, through the Norwegian Sea into the central and western Barents Sea in latest Devonian to Middle Carboniferous times. By the Early Triassic onwards, the orogenic events had ended and the supercontinent

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was progressively broken up by successive rifting events along very much the same general trend as in Early Carboniferous time. A compilation of the main tectonic events during the same period shows that, on a broad scale, Carboniferous and Permian times represent a period of very active tectonics all over the area (Fig. 1). The Middle Triassic to late-Early Jurassic seems to have been mainly a period of thermal relaxation. Then intermittent tectonic activity is again seen through Middle and Late Jurassic, Cretaceous and Tertiary times. The timing of events is remarkably coincident across the region, except that the central and easternmost Barents Sea areas decoupled from the tectonic systems to the south and west in Late Cretaceous times (Fig. 1). During Carboniferous to Jurassic times the general lateral distribution of sedimentary environments reflects the existence of a northern ocean (the Boreal Ocean) and a southern ocean (the Proto-Tethys and Tethys) separated by an area with a central zone (the present Norwegian-Greenland Sea) of continental rifting and intermittent shallow seaways. The continuing long history of persistent extension and rifting turned the central zone into a permanent seaway between the northern and southern oceans. A stratigraphic compilation for the whole of the Norwegian continental shelf and adjacent areas from the end of Devonian to Pleistocene times reflects this picture (Fig. 2). At the same time, the northward continental drift caused the climate to change from equatorial through temperate to partly arctic climates. This tectonic and climatic development is reflected by a change in the marine deposits from Palaeozoic carbonates and evaporitic deposits to Mesozoic and Tertiary clastic shelf and basin deposits (Fig. 2). Plate tectonic reconstruction To illustrate the geological development of the region, a set of palaeogeographic maps have been constructed. The palaeogeography is plotted onto appropriate base maps of lithospheric plate tectonic reconstructed configurations. A number of simplifications were introduced in order to arrive at these plate tectonic reconstructions. Firstly, the basic assumption behind the plate tectonic reconstructions is that it is possible to view the tectonic development in terms of three main rifting episodes: the first comprising all rifting from Early Carboniferous to Middle Triassic times, the second comprising all rifting from Middle Jurassic to Late Cretaceous times, and the third the rifting in latest Cretaceous/early Tertiary times (see Fig. 1). Considering the scope of the study, the scale of the maps

Sedimentary environments offshore N o r w a y - - an overview

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Fig. 1. Summary compilation of the tectonic and magmatic evolution of the northern North Atlantic and surrounding mainlands. Blue patches indicate the timing and duration of tectonic events, red 'v's indicate timing of significant magmatic events. Arrows indicate sense of tectonic movements: diverging arrows, extension; converging arrows, compression; lateral arrows, strike-slip movements.

and the lack of detailed knowledge of the rifts of Carboniferous, Permian, Triassic and early Tertiary times, it is deemed sufficient to construct only four different base maps of plate tectonic configurations: at 300 Ma, at 150 Ma, at 70 Ma, and at 53 Ma. These maps illustrate the plate configurations prior to rifting in Carboniferous to Triassic times, prior to rifting in Late Jurassic/Early Cretaceous times, prior to rifting in early Tertiary times, and at the continental break-up in the Early Eocene, respectively. Secondly, it is assumed that the general extension direction between Greenland/North America and northern Europe/Baltic Shield through the whole time span was mainly parallel to the old N W - S E structural grain of the basement (as discussed in Gabrielsen et al., 1999). This is the trend of present prominent lineaments like the Jan Mayen Lineament

and Senja Fracture Zone, documented by several workers to be a controlling factor at least from Late Jurassic times (e.g. Brekke and Riis, 1987; Torske and Prestvik, 1991; Blystad et al., 1995; Dor6 and Lundin, 1996; Dor6 et al., 1999; Brekke, 2000). Here it is assumed that this was the main direction of extension between the regions of present Greenland and Europe also in pre-Jurassic times. The same is assumed for the late Palaeozoic rifting in the Barents Sea shelf (Gudlaugsson et al., 1998). In the rifts of the North Sea, however, detailed models indicate that the direction of rifting changed from NE-SW-directed escape tectonics in Permo-Carboniferous times (Coward, 1993; Corfield et al., 1996; Besly, 1998) to an E - W and/or N W - S E direction in Jurassic-Early Cretaceous times (Faerseth, 1996; Dot6 et al., 1999; Errat et al., 1999). This may imply that the Shetland-

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Sedimentary environments offshore Norway

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British Isles region at times may have moved as an independent block between Europe/Scandinavia and Greenland. Determining the amount of extension for each of the three major rifting episodes involved consideration of a great variety of published models and estimates (e.g. Lippard and Liu, 1992; Skogseid, 1994; Roberts et al., 1995; F~erseth, 1996; Gudlaugsson et al., 1998; Dor6 et al., 1999; Errat et al., 1999; Gabrielsen et al., 1999; Odinsen et al., 2000; Christiansson et al., 2000; Reemst and Cloetingh, 2000). In view of the literature, it was deemed more fruitful to focus on estimating a reasonable order of total extension and to try to decide the relative importance of the three rifting episodes rather than attempting to arrive at exact details. Based on a subjective set of Beta-factors from the literature and distances over which they were to be applied, the authors propose a semi-quantitative model for the total rifting (Fig. 3). The model concentrates on the continental margins of the Norwegian-Greenland Sea, since this occupies the crucial area for the rifting and separation between Scandinavia and Greenland. The model assumes that the total extension of the Carboniferous-Triassic episode was greater than the following Middle Jurassic/earliest Cretaceous episode, as the crustal thinning was probably greater on average and was uniformly distributed over a larger area (e.g. Odinsen et al., 2000; Gabrielsen et al., 1999). The latest Cretaceous/Paleocene rifting episode is suggested to have contributed least to the total crustal extension between Norway and Greenland. This is true if extension estimates are based only on the observed Tertiary subsidence derived from sediment thickness in the outer parts of the continental margin (i.e. western parts of the Voring and MOre Basins) (e.g. Skogseid et al., 1992; Skogseid, 1994). However, large Beta-factors are estimated for this episode assuming underplated magmatic bodies of significant thickness emplaced at the base of the crust at that time (e.g. Skogseid et al., 1992, 2000; Skogseid, 1994). Velocity analysis of the crust of the outer continental margin (beneath the MOre and VOring Marginal Highs) support the existence of high-velocity bodies at the base of the crust that fit with such an underplating mechanism (e.g. Olafsson et al., 1992; Mjelde et al., 2001). However, these deep-seated bodies are not dated, and if they have originated at a different time and/or by a different process, we are left with only the observed subsidence as the key to

extension estimates. The crystalline continental crust underneath the eastern parts of the V0ring Marginal High is almost as thick as that underneath the Tr0ndelag Platform close to mainland Norway (Mjelde et al., 2001). This is not in concert with a wide zone of a high degree of crustal attenuation underneath the V0ring Marginal High, which would have to be the site of the main latest Cretaceous/Paleocene rift. The data of Mjelde et al. (2001) indicate an abrupt western termination of the continental crust, implying a very narrow zone for the actual rifting and final continental break-up. If so, even a locally very high Beta-factor for the latest Cretaceous/Paleocene rifting would not add up to a total extension comparable to that of the two preceding rifting episodes. On the assumption that the general extension direction was N W - S E and that the total extension in each rifting episode did not vary greatly from north to south (no relative rotations), the model gave estimates for the amount of extension in the NorwegianGreenland Sea between Scandinavia and Greenland as follows: For the latest Cretaceous/Paleocene rifting: 45km For the Middle Jurassic/earliest Cretaceous rifting: 80 km For the Carboniferous-Triassic rifting: 135 km For the sake of simplification, the same total distances of extension for each rifting episode are used in the plate tectonic reconstructions for the whole study area south of the Barents Sea (Fig. 3). This implies that the estimated extensions in the North Sea are compensated by less extension west of the British Isles (Fig. 3). In the Barents Sea, the post-Triassic extension is considered as insignificant relative to the Carboniferous-Triassic extension (Fig. 3). Thirdly, the plate tectonic reconstructions were made by progressively subtracting the amount of extension of the three main episodes of extension from the Ypresian plate reconstruction of Srivastava and Tapscott (1986). Because the exact locations of the rifts of Carboniferous-Triassic and latest Cretaceousearly Tertiary times are not known, this exercise was done in a rather schematic manner (see Fig. 3). This allowed for the moving of the coastlines and the known rifts and other tectonic structures into their relative past positions. In this picture, the exact locations of the different Palaeozoic and Cenozoic grabens and rifts were not regarded as critical. The plate reconstruction presented here is an attempt to reflect the

Fig. 2. Summary compilation of the lithostratigraphic evolution of the northern North Atlantic and surrounding mainlands. The compilation is built from a selection of four central stratigraphic columns for each of the three sub-regions, the North Sea, the Norwegian-Greenland Sea, and the Barents Sea. The colour legend is given in Fig. 5b.

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Fig. 3. Method for progressive removal of the effect of stretching during major post-Devonian rifting episodes. The base map is the 24 anomaly time (53 Ma) plate reconstruction of Srivastava and Tapscott (1986). Coloured corridors indicate to scale the amount of lateral extension resulting from each of the three major rifting episodes as estimated in the continental margin of the Norwegian Sea (see inset legend). The plate-reconstructed base maps of 70 Ma, 150 Ma and 300 Ma for plotting of palaeogeography was derived by closing the light orange, the blue, and the brown corridors, respectively, and in that sequence. The red discontinuous lines indicate major tectonic transfer zones parallel to the direction of extension. See text for comments.

overall picture (Fig. 3), indicating that at the beginning of Carboniferous times the distance between Norway and Greenland was in the order of 200-300 km shorter than at the beginning of Eocene times. It should be noted that the position of Ellesmere Island relative to Greenland and Svalbard at all times is based on the plate tectonic reconstructions of Rowley and Lottes (1988). Furthermore, the palaeogeography includes the notion of "Crokerland" which was a substantial land area north of Ellesmere Island, that served as an important sediment source from the Middle Carboniferous to the end of the Middle Jurassic (Embry, 1993) (see Fig. 4). A possible model for the tectonic framework of the region at the beginning of Carboniferous times has been set up based on the general plate tectonic reconstruction (Fig. 4). At the beginning of Carboniferous times the Caledonian orogenic compression had ended, the Ellesmerian Orogeny in the north was in its waning stages (Embry, 1993), and the Variscan orogenic cycle was at work in the south. The central area between the Variscan and Ellesmerian orogenic fronts was evidently subject to a complicated system of extension of unknown details throughout Carbonif-

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Fig. 4. Tectonic framework of Early Carboniferous times, plotted on 300 Ma plate reconstruction. See text for comments. JL -- Jameson Land; JML -- Jan Mayen Lineament; SBL = Scoresby-Bergen lineament; SL -- Scoresby Sound; OFC = •ygarden Fault Complex.

erous times. It seems likely that Spitsbergen was situated in a strike-slip regime between the Ellesmerian front and the central rift system (Fig. 4). This is supported by the Carboniferous sedimentary facies of Svalbard (Steel and Worsley, 1984). Based on field relationships onshore East Greenland (e.g. Surlyk, 1990; Stemmerik et al., 1993; Escher and Pulvertaft, 1995) and seismic mapping on the Norwegian continental margin (e.g. Blystad et al., 1995; Gudlaugsson et al., 1998) it seems likely that the Carboniferous central rift system between Greenland and Norway was dominated by N-S- to NE-SW-trending normal faults and NW-SE-trending transfer faults including the prominent lineaments like the Jan Mayen Lineament (see discussion in Gabrielsen et al., 1999). In the overall structural picture, there are several large-scale features that have led the present authors to speculate on the existence of yet another prominent N W - S E lineament, here informally termed the Scoresby-Bergen lineament. This runs to the south of, and parallel to, the Jan Mayen Lineament, between the B lossville Coast area on East Greenland and the Bergen area in western Norway (Fig. 4). Onshore East Greenland, in the area between Scoresby Sound and Kangerlussuaq, this lineament fits with the

Sedimentary environments offshore Norway

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southern termination of the late Palaeozoic to Middle Jurassic/earliest Cretaceous rift graben system (see Surlyk, 1978). Offshore, the lineament fits with a fracture zone running right up to the Blossville Coast (see Larsen, 1984). Like other oceanic fracture zones in the Norwegian Sea, this fracture zone may extend as a tectonic lineament into the adjacent continental crust. In the B lossville Coast area, this would be masked by the extensive cover of the onshore Tertiary lava plateau. On the Carboniferous plate tectonic reconstruction, the eastward extension of the lineament falls along a lineament of N W - S E elongated magnetic anomalies running from the oceanic/continental crust transition north of the Erlend Platform, through approximately 62~ 2~ to the west coast of Norway north of Bergen at approximately 61 ~ 4~ (the "Marflo Lineament" of Smethurst, 2000). There, it links directly with the northern end of the Oygarden Fault Complex, which shows up as very prominent arcuate magnetic structure (Olesen et al., 1997). The lineament clearly truncates the pronounced N E - S W magnetic trend of the MOre Basin and the MOreTrOndelag Fault Complex indicating a crustal involvement. It is clear on the magnetic data that the southeastern end of the Scoresby-Bergen lineament (if it exists) then connects with the northeastern end of the Midland Valley-Ling Depression rift through the arcuate Oygarden Fault Complex (see Fig. 4). Parallel to the arcuate magnetic expression of the Oygarden Fault Complex is the onshore, arcuate fault system of the Bergen-Sunnhordland Arcs that show evidence of reactivations in the Devonian, PermoCarboniferous, Early Triassic and Late Jurassic/Early Cretaceous (Fossen, 1998; Wennberg et al., 1998) (Fig. 4). Similar dates are documented on other major basement faults along the southwest coast and well into mainland Norway (e.g. Andersen et al., 1999). It is possible that the Scoresby-Bergen lineament constitutes the southern limit of the whole of the Permo-Carboniferous central rift system between East Greenland and Norway, and hence acted as a transfer zone between the Greenland rift system and extensional fault system in the North Sea in front of the Variscan Orogen south of the Midland ValleyLing Depression rift line (Fig. 4). Any difference in extension between the North Sea and the Norwegian Sea areas would then very likely be compensated for in the area adjacent to the Oygarden Fault Complex (i.e. in the Stord Basin). A rapid Early Carboniferous cooling event related to erosional unroofing of the mainland basement is documented along the adjacent coast, substantiating that the area was tectonically active at the time (Eide et al., 1999). This also implies that it is likely that Lower Carboniferous sediments are present at depth offshore western Norway.

The overall extension estimate takes into consideration the idea of a central area of stable unrifted basement blocks between Norway and East Greenland, so that pre-Tertiary extension occurred in rift zones on both sides (e.g. Fig. 5a). This is a way to achieve comparable amounts of extension to the north and south of the Scoresby-Bergen lineament in Carboniferous to Triassic times (Fig. 3). The sedimentary record in the Jurassic and Cretaceous of the East Greenland and Norwegian Sea continental margins strongly indicates that such a stable, shallow to emergent central area persisted until the final continental break-up in earliest Tertiary (discussed later). The Carboniferous

The Visean map may be taken as representative for the Early Carboniferous palaeogeography (Fig. 5). Evidence from Greenland, Svalbard and BjornOya indicates that, in this period, the whole area north of Scoresby Sound drained northwards (Steel and Worsley, 1984; Stemmerik et al., 1993), while there is strong evidence from Britain and the North Sea of a substantial source hinterland to the northwest of the Midland Valley-Ling Depression rift line. (Anderton et al., 1979; Bristow, 1988; Cliff et al., 1991) (Figs. 4 and 5a). This configuration indicates a regional watershed situated in the area between the Scoresby-Bergen lineament and the Midland ValleyLing Depression rift line.

The Early Carboniferous In the central parts of the eastern Barents Sea, the Early Carboniferous was a marine carbonate environment which was established already in the Devonian, while the Pechora area in the southeast comprised the transition from emergent land through fluvial, delta plain, shallow marine clastic into carbonate facies (Ziegler, 1987, 1988; Johansen et al., 1993). The Nordkapp Basin and Bjarmeland Platform area further west (see Fig. 6 for location) was a complicated system of emergent land, alluvial, fluvial and deltaic environments (Gudlaugsson et al., 1998) at the end of the northward draining system in the main N E - S W rift system (Fig. 5a). The important Carboniferous outcrops of Svalbard and BjOrnOya found themselves in a strikeslip transfer setting relative to the main N E - S W rift system, with the most important fault lines being the Palaeo-Hornsund, Hornsund, B illefjorden and Lomfjorden-Aghardbukta Faults (Steel and Worsley, 1984; NOttvedt et al., 1993a). Typical of the Tournaisian and Visean was subsidence along narrow, isolate zones accommodating large alluvial fans building

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Fig. 5. (a) The palaeogeography of Early Visean times plotted on the 300 Ma plate reconstruction. Some key references for compilation: Ziegler, 1988; Embry, 1993; Stemmerik et al., 1993; Corfield et al., 1996; Besly, 1998; Glennie and Underhill, 1998 supplemented with in-house studies. See (b) for legend. SUB = Southern Uplands block; WBB = Wales-Brabant Block. (b) Colour legend for all palaeogeographic maps and lithostratigraphic compilation in Fig. 2. Palaeolatitudes for all maps based on Scotese, 1997.

from graben edges into swamps, lakes and minor flood plains on Svalbard, and rivers and floodplains on Bj~rn~ya.

In the Namurian there was a change to a larger and more continuous sediment system dominated by large braided fans of quartz arenites, now building

Sedimentary environments offshore N o r w a y - - an overview

Fig. 6. Main structural elements of the Norwegian continental margin.

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from the west onto flood basins and floodplains in the southeast (Steel and Worsley, 1984). This change in sediment pattern was probably controlled by uplift west of the Palaeo-Hornsund Fault. At the end of Namurian (early Bashkirian) times the braided fans retreated and were replaced by humid coal-bearing marsh and floodplain environments, probably due to a decreased subsidence rate (see fan retreat on Fig. 2). The Carboniferous sedimentary facies of the central Barents Sea was probably similar to those developed in the Early Carboniferous grabens and basins of BjornOya and Svalbard, but not subject to the same frequent strike-slip inversions (Gudlaugsson et al., 1998). In Greenland the Lower Carboniferous is entirely continental (Fig. 2). In the north it consists of more than 70 cycles of sandstones and shales with minor coalbeds, interpreted to represent a meandering river system (Hfikansson and Stemmerik, 1984). Palaeocurrents indicate transport from the southwest towards the northeast supporting the drainage model (Fig. 5a). In central East Greenland, the Tournaisian and Visean show a development from red fluvial sandstones and siltstones to later yellow fluvial sandstones, grey siltstones and thin coal seams. This is believed to be part of the northward draining alluvial/fluvial system with the provenance area and associated coalescing alluvial fans and braid plains in the west with a major flood plain with northward draining rivers developed laterally to the fans and braid plains (Stemmerik et al., 1993). It is possible that these outcrops represent the western half of a symmetric rift-drainage system, the eastern half being hidden beneath younger strata on the continental shelves of East Greenland and Norway. In the North Sea area, most of the information comes from onshore outcrops in the British Isles and continental Europe, though oil industry offshore well data are now gaining volume. Lower Dinantian (Tournaisian) red beds are widely distributed and indistinguishable from the underlying Devonian Old Red. During Dinantian times the Old Red Continent was broken up by crustal extension (e.g. Leeder, 1988; Coward, 1993). The area was transgressed progressively from the south through Tournaisian times giving a diachronous sequence of shallow marine limestones, clastics and localised evaporites (Anderton et al., 1979). By Early Visean times regionally extensive carbonate facies became established in the southern areas. The crustal extension gave rise to topographic differentiation, resulting in horsts and grabens that directly influenced the distribution of facies: carbonate platforms on the highs and clastic turbidites and shales in the grabens (Grayson and Oldham, 1987). Some of the large highs were emergent land areas (WalesBrabant, Southern Uplands Block) (Fig. 5).

H. Brekke et al.

In the Visean, clastic, mainly deltaic sediments are dominant in northern Britain and the central and northern North Sea (north of 54~ (Besly, 1998) (Fig. 5a). In the Midland Valley, lacustrine oil shales developed (Anderton et al., 1979). The main sediment transport was from the north and west, and coarse clastics spread gradually southwards. However, this infilling of clastics was contested by a gradually rising sea level that is recorded by a gradual increase in coal abundance and the onset of Yoredale cycles at the end of the Visean. These cycles consist of prograding delta clastics capped by abandonment phase shallow marine limestones (Ramsbottom, 1979; Anderton et al., 1979). The Middle Carboniferous

In Namurian times the tectonic activity in the North Sea ceased and was replaced by regional subsidence, drowning even carbonate platforms on relict highs (Collinson, 1988). At the same time, clastic input from the north increased dramatically, perhaps due to a change to more humid conditions (Van der Zwaan et al., 1985). This ended the carbonate deposition and initiated a sequence of major delta advances (Millstone Grit) (Figs. 2 and 7). Deltas in Scotland and northern England experienced a greater marine influence and Yoredale facies continued into the Namurian there; the route of the marine transgression is not known (Besly, 1998). As the carbonate deposition ended in the North Sea, it started its major advance in the Barents Sea (N0ttvedt et al., 1993a; Johansen et al., 1993) (Fig. 2). From its more restricted occurrence in the easternmost Barents Sea in the Early Carboniferous, marine carbonate platform facies had spread all over the Barents Sea by Bashkirian-Moscovian times (Fig. 7). The Bashkirian of Svalbard was a return to the deposition in narrow zones of subsidence along the major fault lines, like in the Tournaisian_Visean. En echelon arrangements of these troughs and local inversions indicate strike-slip movements also in Middle Carboniferous times. This may be explained as a consequence of the geometric position of Svalbard/Bj0rnOya relative to the main NE-SW trend of the Carboniferous rifting in the central Barents Sea (Fig. 7). Otherwise, Bashkirian times saw a gradual change to semiarid terrestrial climate and a general rise in sea level. In the basinal areas, the Gipsdalen Group is marked by the incoming of red, alluvial fan deposits and their laterally equivalent fan-delta deposits, evaporites and carbonates. Three separate major depositional systems are described in the Bashkirian (Steel and Worsley, 1984; NOttvedt et al., 1993b).

Sedimentary environments offshore N o r w a y - an overview

(1) In central and east Spitsbergen (Ebbadalen Formation in B illefjorden) alluvial fans and fan-deltas are building out eastwards from the B illefjorden Fault into restricted marine and sabkha environments (dolomites, limestones, evaporites). These facies display rhythmic intercalations organised in upwardscoarsening cycles, believed by Steel and Worsley (1984) to be the result of an intricate interplay of local fault activity and regional transgression. (2) The St. Jonsfjord Trough is characterised by alluvial fans from both east and west, passing into floodplain, shoreline, and open marine environment northwards. (3) The Hornsund Basin saw alluvial fans building out eastwards. The ongoing transgression produced fan-delta systems and interbedded carbonates and repeatedly submerged coastal plain environments. During the Moscovian all of these systems and adjacent basement were transgressed and overlain by shallow marine carbonate platform deposits. Contemporaneous evaporite development in basihal troughs caused the later pronounced salt tectonics (e.g. in the TromsO and Nordkapp Basins) (Stemmerik and Worsley, 1989; Gudlaugsson et al., 1998). The salt facies of the deep basins were probably fringed by carbonate-evaporite facies on the basin flanks (Fig. 7). In northern Greenland, the Middle-Late Carboniferous consists of three fining-upwards mega-cycles of sandstone-carbonate rhythmics, probably resulting from tectonic activity overprinting the regional transgression (Hftkansson and Stemmerik, 1984). Towards the top of each mega-cycle the sandstones become rare being replaced by thick-bedded carbonates with minor shales and evaporites. The sandstonecarbonate rhythmics are capped by a sequence of thick shallow to deep-water marine shelf carbonates. This may be correlated with the transition from clastics to carbonates on Svalbard and Bj~m~ya (Ebbadalen to Nordskioldbreen Formations and Kapp Hanna to Kapp Duner Formations). Further south in central East Greenland, the northward draining alluvial/fluvial system of the Early Carboniferous persisted (Stemmerik et al., 1993). A typical feature of Middle-Late Carboniferous times are the upper Bashkirian lacustrine black shale facies patchily located along the western tectonic boundary of the Carboniferous alluvial fan/braid plain system (i.e. proximal to the clastic source areas) (Stemmerik et al., 1991). In the absence of other evidence, one may expect similar environments in the Mid-Upper Carboniferous of the Norwegian Sea margin.

17

The Late Carboniferous

During Middle-Late Carboniferous times the North Sea was filled in by clastics (Anderton et al., 1979; Corfield et al., 1996; Besly, 1998) (Figs. 2 and 7). During early-middle Westphalian (late Bashkirian) times, shallow water deltaic environments prevailed on a major scale. Low-lying, near-emergent conditions existed over the whole of the Variscan foreland, and sediment provenance changed from a northerly source to a westerly. The early-middle Westphalian period saw two cycles of flooding and regression. In the flooding periods elongate deltas were infilling shallow water-bodies on a basin-wide scale. Marine bands are common while regionally correlatable coals are few. In the intermittent regressive periods much of the area was emergent with some freshwater lakes being infilled by local deltas. Marine bands are rare and major regional coal beds are developed. In middle-late Westphalian (Moscovian) times the earlier sedimentary environments were modified by the northward advance of the Variscan deformation front. This caused uplift and folding resulting in unconformities and widespread deposition of red beds in the North Sea. The red beds were spreading both from the north and as pulses of molasse from the south (Corfield et al., 1996; Dahlgren and Corfu, 2001). Above continental red beds in the extreme east, in the Oslo Graben, a Moscovian marine limestone is identified in a braid delta sequence prograding into a shallow marine basin (Olaussen et al., 1994) (Fig. 2). But the marine fossils can be correlated with the Russian marine stages and indicate that the regional Bashkirian-Moscovian sea-level rise caused a transgression from the north, opening a seaway across the Baltic Shield from the eastern Barents Sea to the Oslo area at the peak of the Moscovian transgression (Fig. 7). The plant fossils and freshwater fauna of the unconformably overlying red-bedded fluvial delta/lacustrine sequence indicate that by Stephanian times the Oslo area was again comparable to the West European/North American realm. The Carboniferous of the northern North Sea

The Carboniferous of the Norwegian sector of the North Sea is practically unknown. Nearby, however, Visean to Namurian fluvial red beds, alluvial/deltaic coal-bearing sequences and lacustrine facies are recorded in the Outer Moray Firth, and there are indications of sediment transport from the Viking Graben. Westphalian red beds are reported from the South Viking Graben in well UK9/13a-22 (Cameron, 1993).

18

H. Brekke et al.

Fig. 9. Well correlation of pre-Triassic deposits in the Utsira High. See Figs. 8, 10 and 11 for well locations. Fig. 7. The palaeogeography of Bashkirian-Moscovian times plotted on the 150 Ma plate reconstruction. Some key references for compilation: Hfikansson and Stemmerik (1984); Steel and Worsley (1984); Ziegler (1988); Stemmerik et al. (1991, 1993); Embry (1993); Olaussen et al. (1994); Corfield et al. (1996); Besly (1998); supplemented by in-house studies. See Fig. 5b for legend.

Regional seismic data in eastern parts of the northern North Sea, supplemented by some well data, however, strongly indicate the existence of preLate Permian strata of possible Carboniferous age across the Stord Basin and Utsira High (Fig. 8). Data on the Utsira High show a clear pre-Late Permian half-graben development. The lower parts of the wells 25 / 10-2 and 25/10-4 penetrated pre-Zechstein (Upper Permian) clastics (Fig. 9). In both wells, the Zechstein Group marginal facies overlie a polymict, fluvial/alluvial conglomerate, interpreted to be an Upper Rotliegende equivalent. This conglom-

erate rests on an angular unconformity cutting a tilted sequence of massive, medium to coarse, well sorted sandstone to the east (25/10-4). To the west, the conglomerate apparently overlies gneissic basement in the well 25/10-2; this is thought to be juxtaposition by faulting. This fits exactly with the well positions in the geometry of the half-graben revealed by the seismic data (Fig. 10). A strong reflector at what is believed to be the top of the sandstone sequence, is tentatively interpreted to be a coal horizon at the base of a shaly interval (Fig. 10). Alternatively, the strong reflector may be a volcanic horizon or a sill. By lithostratigraphic and seismic correlation, it seems very likely that the red conglomerates and sandstone in well 25/12-1 are equivalents to those of 25/10-2 and-4 (Fig. 9). This implies the existence of the eroded remnants of two parallel half-grabens of possible Carboniferous age within the Utsira High basement (Fig. 11).

Fig. 8. Regional geoseismic profile, based on seismic line CNST-82-08, showing possible distribution of pre-Permian strata in the Stord Basin and local grabens in the Utsira High of the northeastern North Sea. See Fig. 6 for line location.

19

Sedimentary environments offshore N o r w a y - an overview

Fig. 10. Detailed geoseismic profile, based on seismic line CNST-82-08, of a Palaeozoic half-graben in western Utsira High (see Fig. 8 for location). Note the unconformable relationship between the sandstone (light yellow) and possible shales and coal-beds (green with black stringers) of possible Carboniferous age, and the overlyingRotliegende conglomerate. The juxtaposition of the Upper Permian strata and the basement, and the unconformable overstepping nature of the Triassic, indicate that the boundary fault moved through the Late Permian, but had stopped by Triassic times (Fig. 10). The angular relationship between the Upper Rotliegende conglomerate and the underlying strata shows that the half-graben fill was tilted prior to Late Permian times. Faerseth (1996) proposes a Devonian age for the lower sandstone unit. However, the good porosities and non-metamorphic state of the sandstone does not compare at all with the lowmetamorphic state of the Devonian of the Norwegian mainland nearby. An Early Permian age also seems unlikely as the sandstones are devoid of any traces of volcanic material, which elsewhere is the trade mark of the Lower Rotliegende equivalents. Hence, the present authors prefer a Carboniferous age for these clastics, although it may also be argued that they may be correlated with the non-metamorphosed Devonian sandstones offshore Scotland (e.g. Downie, 1998). The Permian

Permian tectonics were characterised by the final stages of the Variscan Orogeny in the south and the

evolving Uralian Orogeny in the northeast (Fig. 1) (e.g. Ziegler, 1988; Glennie and Underhill, 1998). In Late Permian/Early Triassic times there was a widespread tectonic phase all across the region, possibly reflecting the reconfiguration of plate movements following the final plate coupling in the Uralian. In the studied region, this is mainly expressed as an extension and marks the onset of the break-up of the Pangean supercontinent.

The Early Permian In Asselian times continued transgression established large platforms of mixed dolomite and clastics in the easternmost areas of the northern region (Figs. 2 and 12). In the central parts of the present Barents Sea the shallow marine carbonate environment persisted with evaporite deposition in basin areas, and marginal evaporite facies and carbonate buildups on previous emergent land (Gudlaugsson et al., 1998). Transgression and carbonate platform facies also dominated Spitsbergen and Bj~rmaya, though the BjornOya area experienced repeated inversions in the Asselian (Steel and Worsley, 1984). The Lower Permian carbonate platform facies also

20

H. Brekke et al.

Fig. 12. The palaeogeography of Asselian times plotted on the 150 Ma plate reconstruction. Some key references for compilation: Pegrum (1984); Hfikansson and Stemmerik (1984); Steel and Worsley (1984); Ziegler (1988); Stemmerik et al. (1993); Embry (1993); Olaussen et al. (1994); Glennie (1998); supplemented by in-house studies. See Fig. 5b for legend.

Fig. 11. Map view of Palaeozoic grabens in the Utsira High.

spread across northern Greenland (Hfikansson and Stemmerik, 1984). Further south in East Greenland the Lower Permian is a continuation of the alluvial and fluvial facies of the Carboniferous. The same is assumed for the areas of the present Norwegian shelf in the Norwegian Sea. The last stages of the Variscan Orogeny in Asselian times induced an E - W extension/transtension regime that opened up a graben system of mainly NW-SE trends by the reactivation of inherited lines

of weakness in the southern region south of the Midland Valley-Ling Depression rift line (Figs. 4 and 12) (Glennie and Underhill, 1998) However, prominent graben structures of NE-SW and N-S trends also developed. This E-W extension was accompanied with a volcanic flare-up of rhyolites, ignimbrites and basaltic flows (e.g. see Glennie, 1998). These Lower Permian volcanics and volcaniclastics, mixed with minor fluvial and lacustrine sediments, are mainly recorded onshore (Oslo area, northern Germany and Poland), but an increasing volume of well data show that there may be major areas of volcanic strata also offshore in the present North Sea, perhaps with a wider distribution in the eastern parts of the North Sea than in the southwestern (Glennie, 1998). In northern Germany it is possible to identify a Lower Permian, regional erosional unconformity that may have been caused by the uplift associated with the extension and magmatism, i.e. the Altmarkian Unconformity (Glennie, 1998). The major Saalian Unconformity that developed all across the southern region in middle to Late Permian times indicating a later phase of uplift, cuts through the Altmarkian Unconformity. This middle Permian uplift and erosion is also recorded as a major peneplanation in East Greenland (Surlyk et al., 1984). An apparently coincident regional unconformity is seen even as far as into the present Barents Sea (Johansen et al., 1993) (Fig. 2).

Sedimentary environments offshore N o r w a y - an overview

The Late Permian

Late Permian times brought about dramatic changes in sedimentary environments. It was the end of the carbonate deposition in the Barents Sea region and the onset of widespread fluvial deposition and later transgression in the Norwegian Sea region, and establishment of the two well known southern and northern Permian Basins in the North Sea region (Figs. 2 and 13). Late Kazanian/early Tatarian times saw the onset of a regional transgression and the closing of the ocean in the northeast by the Uralian Orogeny. This transgression was accompanied by (or responsible for) a transition from warm and arid to temperate and humid climate all across the northern region. That climatic change effectively ended the carbonate platform/evaporite environment of the northern ocean. The change to marine clastic environments was accompanied by a regional blooming of sponges that gave rise to spiculitic shales, siltstones and cherts (Figs. 2 and 13). In the central region there is evidence of renewed clastic sedimentation all across the Permian peneplain, possibly caused by the change to a more humid climate. Typically, the Upper Permian (upper Kazanian?) starts with a basal fluvial conglomerate with a remarkably widespread distribution (Figs. 2 and 13). Evidence of marine reworking towards the top of the

Fig. 13. The palaeogeography of late Kazanian-early Tatarian times plotted on the 150 Ma plate reconstruction. Some key references for compilation: Hfikansson and Stemmerik (1984)" Steel and Worsley (1984); Ziegler (1988); Surlyk (1990); Stemmerik et al. (1993); Embry (1993); Olaussen et al. (1994); Gudlaugsson et al. (1998); Glennie (1998); Taylor (1998); supplemented by in-house studies. See Fig. 5b for legend.

21

conglomerate is common and testifies to a southward transgression of the northern ocean (Stemmerik et al., 1993). In the south, the Permian Basins of the present North Sea were the sites of very rapid subsidence and the accumulation of the Upper Rotliegende 2 (Glennie, 1998), consisting of fluvial (wadi), aeolian, sabkha, and lacustrine facies. This reflects a desert climate that was in strong contrast to the newly established temperate humid climate in the northern region. The lacustrine muds and salts of the desert lakes in the Southern Permian Basin was flanked by sabkha facies which in turn passed into wadi deposits and aeolian dune sands. In its thickest parts these deposits presently amount to up to 2500 m which were deposited in the course of 4-8 Ma, giving an extreme rate of subsidence. The mechanism for such a rapid subsidence is not known. By late Tatarian times the continued transgression caused flooding of the Permian Basins of the present North Sea (Fig. 14). This flooding of the low-lying basins gave rise to the vast evaporite deposits of the Zechstein Group. The sea probably entered from the north via the early Viking Graben, Stord Basin, and Central Graben areas and/or the Pennines, and may have dramatically completed the flooding of the basins in the order of 6 years (Glennie, 1998). This

Fig. 14. The palaeogeography of late Tatarian times plotted on the 150 Ma plate reconstruction. Some key references for compilation: H~kansson and Stemmerik (1984); Steel and Worsley (1984); Ziegler (1988); Surlyk (1990); Stemmerik et al. (1993); Embry (1993); Olaussen et al. (1994); Gudlaugsson et al. (1998); Glennie (1998); Taylor (1998); supplemented by in-house studies. See Fig. 5b for legend.

22 also compares to the estimated rate of the catastrophic flooding of the Black Sea 7000 years ago through the Strait of Bosporus (Ryan et al., 1997). A catastrophic flooding event is indicated by the way the initial sapropelic black shales of the Kupferschiefer abruptly drapes delicate topographic features like aeolian sand dunes (Glennie, 1998). The first Zechstein cycle of black shale, marine carbonate, and anhydrites, has been correlated with upper parts of the Late Permian Foldvik Creek Group of similar facies in East Greenland (Stemmerik et al., 1993; Taylor, 1998) (Fig. 2). It is assumed that the Foldvik Creek Group was deposited during the southward transgression. Since the present Norwegian-Greenland Sea margins occupied the intermediate zone between the warm arid North Sea and the temperate humid Barents Sea, and with rather open marine conditions, the postulated shallow stable central block would be a good place for the establishment of a Tatarian shallow marine carbonate platform. Carbonate deposits, probably of this age, are recorded in the fringing areas: in wells on the northwest part of the TrCndelag Platform, in outcrop on And~ya (Dalland, 1981), and on East Greenland (Surlyk et al., 1984; Stemmerik et al., 1993). Such an environment would also include a good chance for source rock at that time. In the present Barents Sea region the marine spiculitic clastic facies, initiated in the Kazanian, persisted. The Triassic

Triassic strata are well known in all parts of the region, in some places amounting to 5-6 km thickness, e.g. the Stord Basin in the North Sea and the South Barents Basin in eastern Barents Sea (Johansen et al., !993; Fa~rseth, 1996) (Figs. 6 and 17). But the details of the Triassic tectonic development seem obscure and difficult to assemble into a complete picture (Fisher and Mudge, 1998; Dor6 et al., 1999; Errat et al., 1999). Except for the widespread Late PermianEarly Triassic crustal extension the Triassic apparently remained a period mainly of thermal relaxation (Fig. 1) (Surlyk, 1990; Johansen et al., 1993; Glennie and Underhill, 1998; Dor6 et al., 1999; Roberts et al., 1999). A characteristic feature of the Early Triassic seems to be that most of the extension and sediment volumes were accommodated by a limited number of major rift boundary faults trending north-northeast to north-northwest (Fig. 15) (Surlyk, 1990; Blystad et al., 1995; Fa~rseth, 1996). Major uplift and erosion of the southern Norwegian mainland seems to be linked to the subsequent thermal subsidence stage of the rift basin in Middle Triassic-Early Jurassic times (van der Beek, 1994; Rohrman et al., 1995; Riis, 1996).

H. Brekke et al.

In the present North Sea the Permian evaporite environments more or less continued into Early Triassic times, but with an increase in clastic input. Below the mid-Scythian Hardegsen Unconformity the Lower Triassic demonstrates a transition from marginal marine to alluvial and fluvial environments, by which the Zechstein Group was covered by prograding fluvial sandy deposits. Above the Hardegsen Unconformity, the Triassic environment of the southern North Sea diverges from that of the rest of the North Sea. Above the Hardegsen Unconformity, the present central and northern North Sea is dominated by continental alluvial and fluvial environments (Figs. 2 and 15). The southern North Sea, on the other hand, is characterised by several evaporite cycles through the rest of the Triassic, giving the R6t, Muschelkalk and Keuper evaporites (Figs. 2, 15 and 16). In the present Norwegian Sea and East Greenland, the Triassic is dominated by continental fluvial and alluvial environments, interrupted by short-lived marine incursions from the north. In the present Norwegian shelf area there is evidence of only one marine transgression, seen as a Ladinian-Carnian evaporite sequence (Jacobsen and van Veen, 1984). In East Greenland there are indications of an earlier, Anisian black shale of some source rock potential (Surlyk, 1990). This may then be the southern limit of the time equivalent widespread Barents Sea source rock of the Bottenheia Formation (Fig. 15). In the Barents Sea, the marine clastic environments of the Permian continued, but with a considerable increase in sandy influx which ended the dominance of the spiculitic facies. Large volumes of clastics came in from the east by the peak of the Uralian Orogeny in Novaya Zemlya (Fig. 1). 7-8 km of Permo-Triassic sediments accumulated in the progressively subsiding South Barents Basin at the foot of the orogen (Johansen et al., 1993) (Fig. 17). Clastics were also derived from the other flanks of the basin, all contributing to the filling in and shallowing of the northern ocean. Recent provenance studies show that in Early Triassic times the Uralian Mountains were the provenance area for the eastern Barents Sea, the Baltic Shield margin and the Caledonides for the Hammerfest Basin, while Svalbard had its provenance area in a palaeo-land area to the northwest (M~rk, 1999). In the central Barents Sea there is a mixture of Uralide and Caledonide provenance. In Middle Triassic times provenance areas became more localised. In the Upper Triassic of the Barents Sea there is a distinct increase in sandstone maturity on a regional scale, possibly due to extensive reworking and/or a change to more favourable climatic conditions for kaolinitisation (Bergan and Knarud, 1993; MCrk, 1999). In the central, deeper parts of the Boreal

Sedimentary environments offshore N o r w a y - - an overview

23

Fig. 15. The palaeogeography of late Anisian-Ladinian times plotted on the 150 Ma plate reconstruction. Some key references for compilation: Jacobsen and van Veen (1984); H~tkansson and Stemmerik (1984); Steel and Worsley (1984); Ziegler (1988); Surlyk (1990); Van Veen et al. (1993); N0ttvedt et al. (1993a,b); Leith et al. (1993); Stemmerik et al. (1993); Embry (1993); Fisher and Mudge (1998); supplemented by in-house studies. See Fig. 5b for legend.

Fig. 16. The palaeogeography of late Carnian times plotted on the 150 Ma plate reconstruction. Some key references for compilation: Jacobsen and van Veen (1984); Hfikansson and Stemmerik (1984); Steel and Worsley (1984); Ziegler (1988); Surlyk (1990); Van Veen et al. (1993); N0ttvedt et al. (1993a,b); Stemmerik et al. (1993); Embry (1993); supplemented by in-house studies. See Fig. 5b for legend.

Sea and the Sverdrup Basin, Anisian times saw the widespread deposition of black shales of very good source potential (e.g NOttvedt et al., 1993a,b; Leith et al., 1993) (Fig. 15). A number of transgressive-regressive cycles have been identified by MOrk (1994) in the Triassic stratigraphy in the Barents Sea. Only four of these are recognised as "simultaneous" (i.e. truly eustatic). These transgressions/regressions had great impact on the distribution of emergent land and facies in the shallowing Boreal Sea. This is illustrated by comparing the palaeogeography of the Barents Sea under the transgressive event in late Anisian-Ladinian times (Fig. 15) with that of the regressive period of late Carnian (Fig. 16). It is obvious that in regression periods like late Carnian times, there must have been a significant river system to transport clastics out to the distant shoreline to the west.

1999). However, a characteristic feature of Middle Jurassic times is the progradation of clastic wedges from regional and semi-regional areas of erosion in all parts of the region. A widely known example of this is the Mid-Cimmerian Unconformity and the associated building of the Middle Jurassic Brent Delta in the central and northern parts of the North Sea (Underhill and Partington, 1993). Underhill and Partington (op cit.) attributes this to the growth, erosion and deflation of a central thermal dome during the period from the Aalenian to the Oxfordian. It was noted that the doming was accompanied by a marked faunal provinciality between the northern Boreal and southern Tethyan Seas from Aalenian to late Bathonian/Callovian times, suggesting that the central North Sea constituted a significant barrier in that period (Callomon, 1979; Enay and Mangold, 1982; Dor6, 1992; Underhill, 1998)(Fig. 18). Dor6 et al. (1999) suggest that the North Sea dome is only one of a family of such uplifts extending across Northwest Europe. The latter authors base this on references to late-Early to Middle Jurassic unconformities west of Shetland and west of Ireland, and the coincident marine faunal separation between the northern and southern oceans. One may add to this "family" similar uplifted land areas of the same period between the sedimentary basins of East Greenland and the Norwegian continental margin as proposed by several au-

The Jurassic

The stratigraphy reflects a relative sea-level rise in the Early to Middle Jurassic that caused the whole study area to become dominated by shallow clastic shelf environments (Fig. 2). This may partly be attributed to a marine flooding of the passively subsiding old Permo-Triassic rift basins as Pangea started breaking up (Dor6 et al., 1999; Roberts et al.,

24

H. Brekke et al.

Fig. 17. Regional geoseismic east-west profile across the Barents Sea shelf. Note the pronounced role of the South Barents Basin as a sediment sink in Triassic times. See Fig. 6 for line location. Modified from Johansen et al. (1993).

of the main stages of the rifting process from Late Jurassic times onwards. This deflation and collapse may well be the mechanism behind the pronounced sag geometry of the deepest basins from this phase of rifting (e.g. the Vcring and Mere Basins). That would fit with the interpretation of Brekke (2000) that the flanks of the VCring and MOre Basins formed mainly by large-scale monoclinal downflexing of the crust rather than down-to-the basin faulting.

The Early and Middle Jurassic

Fig. 18. The palaeogeography of late Bajocian times plotted on the 150 Ma plate reconstruction. Some key references for compilation: Dalland (1981); Hgtkansson and Stemmerik (1984); Ziegler (1988); Surlyk (1990); Bergan and Knarud (1993); Ncttvedt et al. (1993a); Stemmerik et al. (1993); Embry (1993); Johannessen et al. (1995); Underhill (1998); Brekke et al. (1999); supplemented by in-house studies. See Fig. 5b for legend.

thors (e.g. Dalland, 1981; Larsen, 1987; Dor6, 1992; Brekke, 2000). From this review, treated in more local detail in the following, it seems that the axial areas for the subsequent Middle Jurassic to earliest Cretaceous rifting experienced precursory doming and uplifts all the way from the south of Ireland to the borders of the Barents Sea. The deflation and collapse of these domes and uplifts were then an important integral part

In the North Sea, large portions of the Lower Jurassic stratigraphy are represented by an erosional hiatus, probably due to the thermal updoming of the central parts of the area in late-Early Jurassic times (Whiteman et al., 1975; Leeder, 1983; Underhill and Partington, 1993). However, based on the erosional remnants from different parts of the North Sea, it seems that shallow marine environments had established all over the area in earliest Jurassic times following the transgression that had started in the Late Triassic (Fig. 2). Apparently, increased sediment input in late Aalenian to early Bajocian times forced the northward progradation of the Brent Delta (Rannoch, Etive, and Ness Formations) against the regional sea-level rise (Graue et al., 1987; HellandHansen et al., 1992; Johannessen et al., 1995). This probably reflects the increased erosion following the domal uplift of the central North Sea (Underhill and Partington, 1993, 1994), and the uplift and erosion of the adjacent Shetland area (Dor6 et al., 1999) and mainland Norway (van der Beek, 1994). The semicircular subcrop pattern beneath the Mid-Cimmerian Unconformity on the central North Sea dome shows

Sedimentary environments offshore N o r w a y - - an overview

that this was a dome of semi-regional scale with its central apex in the triple junction of the Viking Graben, Central Trough and Moray Firth Basins (Underhill, 1998). This observation, together with the reports of separated areas of erosional unconformities further west, implies that, within the widespread uplift on the regional scale suggested above (e.g. Dor6 et al., 1999) there were probably a set of several domes of semi-regional scales. The Early/Middle Jurassic dome-shaped uplift of the south Norwegian mainland (van der Beek, 1994) also fits into this pattern. The Early Jurassic coastal plain/delta plain deposits of the Norwegian Sea (the Are Formation), East Greenland (Kap Stewart Formation) and the Barents Sea (the Tubfien Formation) appear to be time equivalents. In all cases the sedimentation seems to involve the progradation of the coastlines by sediment influx onto the shelf. Evidence of influx of sand from the west onto the Halten and DOnna Terraces and the eastern margin of the MOre Basin during the Early to Middle Jurassic (Gjelberg et al., 1987; Jongepier et al., 1996), is taken to imply that the present deep Cretaceous MOre and V0ring Basins were areas of uplift, sub-aerial exposure and deep erosion in that period (Dor6, 1992; Brekke et al., 1999) (Fig. 18). Brekke (2000) argues that the areas of highest extension in the Middle Jurassic/earliest Cretaceous rifting phase (i.e. the MOre and V0ring Basins) were subject to the highest elevation and the deepest erosion in the Middle Jurassic. Hence, the areas of the MOre and V0ring Basins were the sites of thermal domes in that period and developed an erosional unconformity equivalent to the Mid-Cimmerian of the North Sea. This implies that Lower and Middle Jurassic deposits are missing in the deep MOre and V0ring Basins (Brekke et al., 1999; Brekke, 2000). A western hinterland is supported also by in-house studies in the NPD, which show a transition from proximal sands in the west to distal marine shales towards the northeast on the Halten Terrace and the Tr0ndelag Platform, including the Helgeland Basin, in Early to Middle Jurassic times (Fig. 18). The co-existing Early to Middle Jurassic basin of the Jameson Land area in East Greenland shows evidence of sediment influx from the east and north (Surlyk, 1990). Evidence from And0ya north of the Tr0ndelag Platform strongly indicates a Bajocian/Bathonian delta prograding towards the south being fed by clastic sediments from hinterlands to the west, north and east (Dalland, 1981). Together with the evidence of a western hinterland for the Halten Terrace and TrOndelag Platform, this implies a central landmass exposed for erosion between present Norway and East Greenland (Fig. 18). Such a landmass has been suggested by several authors (Dalland, 1981; Larsen, 1987; Dor6, 1992; Brekke et al., 1999;

25 Brekke, 2000). This hinterland configuration indicates that the seaway between the present Norwegian Sea and Barents Sea was to the east rather than along the overall rift axis between present Norway and East Greenland (Fig. 18). This also fits with the Middle Jurassic erosional hiatus in East Greenland (Fig. 2) (Surlyk, 1991). Recent apatite fission track data substantiate a phase of rapid uplift and erosion in the Middle Jurassic in East Greenland, the first phase of uplift since the Carboniferous of the area (Johnson and Gallagher, 1999). As argued above, the uplift of the central landmass between Norway and East Greenland seems to be part of a more regional uplift of the whole of the Norwegian-Greenland Sea and surrounding area, and linked to the onset of crustal extension and increased heat flow. In the Norwegian-Greenland Sea this extension mainly affected the axial area of postulated, long-lived stable, unrifted basement blocks. Being previously unrifted, and thereby not subjected to "strain hardening" at deep crustal levels during the Triassic thermal relaxation, the crust in this area would be the natural location for the subsequent Middle Jurassic to Early Cretaceous rifting episode. However, the crust of these stable basement blocks did not react uniformly to the extension and rifting. The foci of uplift (doming), extension and attenuation of the crust were along the axes of the MOre and V0ring Basins. The crust beneath the elevated flanks of these new domal areas, i.e. mainland Norway and the Halten Terrace/Tr0ndelag Platform to the east and the MOre and V0ring Marginal Highs to the west, was not attenuated. The thickness of the crystalline continental crust beneath the V0ring Marginal High west of the V0ring Escarpment is still in the order of 15 km, whereas it is reduced to 5 km beneath the V0ring Basin (Mjelde et al., 2001). Thus, the stable basements blocks west of the Faeroe-Shetland and V0ring Escarpments were unaffected by the Jurassic/Early Cretaceous extension and became the elevated flanking platforms to the west when the attenuated crust of the MOre and V0ring Basins started to subside in the Late Jurassic. During Cretaceous times, this western platform area constituted intermittent emerged land areas in the Norwegian Sea between the Norwegian and Greenland mainlands (e.g. Brekke et al., 1999). The continental crust beneath this platform area only became involved in extension at the end of the Cretaceous and was finally ruptured in the continental break-up in Early Eocene times. In the present Barents Sea sediments continued to pour into the basin from the east, keeping up the coastal plain/delta plain development. However, this coastal plain/delta plain environment was gradually transgressed from the west during Middle Jurassic

26

times. It seems probable that the Lower Jurassic Tubfien Formation and the Middle Jurassic Nordmela Formation (Dalland et al., 1988) are parts of the same time transgressive coastal plain/delta plain system, which had its maximum western extent in Toarcian times (Tubfien Formation) and which was finally transgressed in the east (South Barents Basin) in earliest Oxfordian times (Fig. 2). After the long period of erosional denudation of the hinterlands (e.g. Figs. 15 and 16) through Triassic and Early Jurassic times, the whole region was probably dominated by a low-lying peneplain (Riis, 1996). The local and semi-regional Middle Jurassic domes and uplifts therefore had great effects on the distribution of emerged land and sea, and the palaeogeography of Middle Jurassic times was probably complex (Fig. 18). The Late Jurassic

The initial phase of the major extensional tectonic period that caused the break-up of the Central Atlantic, started at the end of Middle Jurassic times in the North Sea and Norwegian-Greenland Sea (e.g. Blystad et al., 1995; F~erseth, 1996), and probably in early-Late Jurassic times in the central and western Barents Sea (Johansen et al., 1993). The major regional tectonic phase, however, started in late Oxfordian/early Kimmeridgian times and continued intermittently into Ryazanian/Valanginian times (e.g. Blystad et al., 1995; Underhill, 1998). In East Greenland the major phase is reported to be as late as middle Volgian to Valanginian times (Surlyk, 1990). During early- to middle-Late Jurassic times the North Sea dome deflated and the elevated areas in the central parts of the present MOre and VOting Basins subsided rapidly (Underhill, 1998; Brekke, 2000). These tectonic events caused a marked rejuvenation of the topography into a complicated system of tectonic highs and basins on a variety of scales. Amongst these was the emergent platform area separating the VCring and MOre Basins from the basins of East Greenland (Dor6, 1992; Brekke et al., 1999; Brekke, 2000). At the same time, from the early Bathonian to early Kimmeridgian, there was a major sea-level rise that flooded this topography. However, it did not succeed to drown all the new highs (Dor6, 1992; Underhill, 1998; Brekke, 2000) (Figs. 2 and 19). This period was entirely dominated by open marine claystone deposition (e.g. Heather, Melke and Fuglen Formations) (Fig. 2). The sea-level rise was followed by a regional sea-level fall in the early to mid-Volgian with a low-stand lasting till the mid-Ryazanian (e.g. Rawson and Riley, 1982; Surlyk, 1991; Dor6, 1992). In combination with the renewed and complicated

H. Brekke et al.

rift topography this caused temporary faunal provinciality (e.g. Dor6, 1991). This fluctuation in sea level under such tectonic circumstances seems to have been very favourable for the widespread accumulation of large volumes of black shales, of which large parts have very good source potential (Fig. 19). Source rock deposition seems to have been most pronounced during the relative sea-level fall and low-stand. Upper Jurassic sandy deposits include syn-rift clastic wedges and shallow marine sheet sands associated with deltas and coastal plains. Deep-water submarine fans of coarse clastics deposited on the hanging walls of major faults are reported from several locations, including the Brae trend of the Southern Viking Graben (Stow et al., 1982) and similar facies in East Greenland (Surlyk, 1990). Fan and bar deposits encapsulated in source rock shales are also typically found on the back of major tilted footwall fault blocks of the North Sea (e.g. Dahl and Solli, 1993; Underhill, 1994). The sands in these cases are believed to be derived from the erosion of the crest of the fault block itself. In the platform areas one may find high-energy shallow marine sheet sands and bar deposits like the Sognefjord Formation on the Horda Platform (Vollset and Dor6, 1984) and the Rogn Formation (Dalland et al., 1988) on the Tr0ndelag Platform, respectively (Fig. 2). These deposits are believed to have been sourced froria clastic shorelines of delta plains/coastal plains. Early Kimmeridgian marine sands are also reported from the Janusfjellet Formation of Svalbard (N0ttvedt et al., 1993a,b). The Cretaceous

The Early Cretaceous

The Ryazanian low-stand was followed by a renewed sea-level rise to an intermediate maximum in the Barremian (Fig. 20). The early Neocomian was still dominated by emergent structural highs and platform areas, and the Ryazanian/Berriasian erosional unconformity on top of the carbonaceous marine shales is observed across the entire study area (e.g. Vollset and Dor6, 1984; Dalland et al., 1988; Surlyk, 1990; Smelror et al., 1998), except in north Greenland (Hfikansson and Stemmerik, 1984). In Ryazanian through Hauterivian times deep basinal areas continued to develop by subsidence along the rift axis of the North Sea, in the Mere and V~ring Basins, Jameson Land, and the Harstad, Troms~ and SCrvestsnaget Basins, and probably their north Greenland conjugate parts. The present onshore outcrops in north Greenland show a laterally diverse development of shallow marine to fluvial sandy deposits (Hfikansson and Stemmerik, 1984), which may be the fringe de-

Sedimentary environments offshore N o r w a y - an overview

Fig. 19. The palaeogeography of late Oxfordian-early Kimmeridgian times plotted on the 70 Ma plate reconstruction. Some key references for compilation: Dalland (1981); H&kansson and Stemmerik (1984); Larsen (1984); Steel and Worsley (1984); Ziegler (1988); Surlyk (1990); Leith et al. (1993); Stemmerik et al. (1993); Embry (1993); Underhill (1998); Brekke et al. (1999); supplemented by in-house studies. See Fig. 5b for legend.

posits to deep, fault-bounded basins indicated in the offshore areas to the east by Larsen (1984). Also in the North Sea and in the Norwegian Sea the emergent highs and land areas were fringed by shallow marine sands as transgression progressed (e.g. Oakman and Partington, 1998; Brekke et al., 1999). The V~aring Basin area comprised several sub-basins at the time (Brekke et al., 1999; Brekke, 2000), and the same may have been the case for the Harstad, Troms~a and S~arvestsnaget Basins, and north Greenland. The central platform area of the Norwegian-Greenland Sea, separating the M~are and V~aring Basins from the Jameson Land Basin, seems to have existed throughout Cretaceous times (Brekke et al., 1999; Brekke, 2000). All these deep basin areas accumulated open marine mudstones and shales during the early Neocomian. The platform areas and structural highs were unconformably capped by a condensed sequence of limestone and marl, like the Lyr Formation in the Norwegian Sea (Dalland et al., 1988) and the Klipprisk Formation in the northern Barents Sea (Smelror et al., 1998; Gradstein et al., 1999). The shallow basins within the platform areas (e.g. the Helgeland, Jameson Land, Hammerfest and Nordkapp Basins) accumulated lime-rich open marine mudstones and shales. The increasing sea level that in this way led to widespread shale and marl deposition in platform

27

Fig. 20. The palaeogeography of Barremian times plotted on the 70 Ma plate reconstruction. Some key references for compilation: Dalland (1981); Hfikansson and Stemmerik (1984); Steel and Worsley (1984); Ziegler (1988); Surlyk (1990); Dor6 (1992); N~attvedt et al. (1993a); Johansen et al. (1993); Stemmerik et al. (1993); Oakman and Partington (1998); Larsen et al. (1999); Brekke et al. (1999); supplemented by in-house studies. See Fig. 5b for legend.

areas and in starved distal deeps, was halted by a sudden sea-level drop at the peak of the Barremian high-stand. This gave time to renewed delta progradations from the transgressed land areas (Fig. 20). These include the Wealdon paralics in southern England (e.g. Dor6, 1991), the Nordelva Member on AndCya (Dalland, 1981), and the major delta deposits of the Helvetiafjellet Formation on Svalbard (Steel and Worsley, 1984; Nemec et al., 1988). The Helvetiafjellet Formation may have been a response to a major uplift of the northwestern Barents Sea area associated with the break-up of the Amerasian Basin of the present Arctic Sea (Dor6, 1991; Ne~ttvedt et al., 1993a,b). This event is dated to Barremian times by Rowley and Lottes (1988) and was accompanied by magmatism in the platform areas around Kong Karls Land and Franz Josef Land (Steel and Worsley, 1984; in-house studies) (Fig. 20). A new pulse of regional transgression started in the Aptian and continued into the Late Cretaceous, slowly drowning emergent intrabasinal highs and surrounding land areas throughout the entire study area. In the North Sea area, however, this transgression formed the background to the rejuvenation of older landmasses caused by the Austrian tectonic phase. This rejuvenation caused a new pulse of progradation of shelfal greensands all around the fringes of the North Sea Basin in the Aptian. Progressive deepening

28

during Albian times subsequently re-established the starved deep shelfal and basinal marl facies environment. The Late Cretaceous

By the end of the Albian, the sea had flooded most of the lowlands surrounding the North Sea Basin, effectively cutting off the clastic input to the whole basin area (Oakman and Partington, 1998). This led to the establishment of the early Chalk Sea of Cenomanian to Santonian times as the basinal marl facies changed to pelagic chalky limestones and spread across the whole North Sea Basin (Fig. 2). The global sea level continued to rise until its maximum in the middle of Campanian times. This was the time span of the mature Chalk Sea that also transgressed the crystalline basement of southern Norway and southem Sweden (Riis, 1996) (Fig. 21). The increased area of submergence caused the mature Chalk Sea to produce carbonates of higher purity than the early Chalk Sea. The general deposits of the chalk seas were a combination of bioturbated homogenised chalk ooze and downslope redeposited, better sorted ooze. The shallower shelf area was dominated by deposits from benthic forms as bryozoa, echinoids and crinoids (Oakman and Partington, 1998). A number of hiati are identified within the chalk sequence in the North Sea, of which the intra-Campanian and the base Paleocene are the most prominent (Fritsen et al., 1999) (see Fig. 2). A Late Cretaceous polyphasal tectonic episode is documented in the Norwegian-Greenland Sea area (e.g. Surlyk, 1990; Blystad et al., 1995; Dor6 et al., 1999; Brekke, 2000) (Fig. 1). Brekke (2000) dated the onset of this tectonic episode to end Cenomanian/earliest Turonian times by seismic correlation with shallow water wells on the Norwegian Sea shelf. However, in-house studies on wells 6607/5-1, 6707/10-1 and 6706/11-1 indicate a latest Turonian age for the initial stages of this tectonism. These events are superimposed on the regional transgression and were probably linked to the incipient seafloor spreading in the Labrador Sea. This is also the timing of the onset of rifting between Greenland and the Rockall Plateau according to the models of Rowley and Lottes (1988) and Srivastava and Verhoef (1992). In East Greenland the crystalline basement was transgressed during the Albian (Stemmerik et al., 1993; Larsen et al., 1999), and it is likely that this also has been the case on the northern Norwegian mainland (Riis, 1996). The tectonism was expressed as faulting, accelerated basin subsidence and conjugate uplift, tilting and emergence of the bounding platform areas to the major basins, i.e. the

H. Brekke et al.

MOre, V0ring, Harstad and Troms0 Basins (Brekke and Riis, 1987; Brekke, 2000). Evidence of coincident flank uplift in the East Greenland basins is given by thin conglomerates and tidal deposits of Cenomanian/Turonian age abruptly overlying Albian shales (Stemmerik et al., 1993) (Fig. 2). This topographic rejuvenation gave rise to new basin-bounding platforms that, once again, showed up as axial emergent land areas from the Rockall Platform to And0ya (Fig. 21). The flank uplifts and platform areas were deeply eroded and gave a pronounced Turonian/Coniacian unconformity that was subsequently tilted and partly transgressed (e.g. Brekke and Riis, 1987). The deep Cretaceous basins of the Norwegian-Greenland Sea contain up to 13 km of sediments in their axial parts, of which the Cretaceous succession alone makes up 8-9 km (Skogseid et al., 2000; Brekke, 2000). The background sedimentation is deep marine mudstones, but in local depo-centres of rapid subsidence this may be overprinted by coarse-grained turbidites in stacks of considerable thickness. The deep-water wells 6707/10-1 and 6706/11-1 in the northern V0ring Basin proved a more than 1000-m-thick interval of Coniacian/Campanian sandy turbidites (Kittilsen et al., 1999). The provenance area for these sands may have been the emergent axial platform areas to the west and north. Alternatively, the clastics may have been transported from Greenland itself along regional channels along prominent NW-SE lineaments bisecting the axial highs (Fig. 21). This would fit with the Late Cretaceous shelf break margin model described by Whitham et al. (1999). In this model, fluvial point sources provided fine to coarse clastic material to the narrow shallow marine shelf. In periods of lowstand these clastics were transported more or less directly from the terrestrial source (e.g. a delta), with very little shallow marine reworking, into the deep basin to the east beyond the shelf break as fine- to coarse-grained sediment gravity flows, accumulating as basin floor fan deposits. The outcrops of the Upper Cretaceous of East Greenland are dominated by dark deep-water shales and the evidence of such sandy low-stand fans are scarce. However, the locations for the sediment transport routes for these fan deposits anticipated by Whitham et al. (1999) fit very well with the prominent NW-SE lineaments leading into the VOting Basin (Fig. 21). In the course of Late Cretaceous times, the Barents Sea shelf and Svalbard were finally de-coupled from the areas south of the de Geer Zone which is the broad zone of deformation along the present western continental margin of the Barents Sea, including the Wandel Sea Basin of north Greenland, the Tromsr and Vestbakken Basins, the Palaeo-Homsund Fault and the Senja Fracture Zone (Harland, 1969; Faleide

Sedimentary environments offshore Norway

an overview

Fig. 21. The palaeogeography of early Campanian times plotted on the 70 Ma plate reconstruction. Some key references for compilation: Hfikansson and Stemmerik (1984); Steel and Worsley (1984); Dalland et al. (1988); Ziegler (1988); Surlyk (1990); Stemmerik et al. (1993); Riis (1996); Oakman and Partington (11998); Brekke et al. (1999); Larsen et al. (1999); supplemented by in-house studies. See Fig. 5b for legend.

et al., 1993a,b) (Fig. 6). The whole of the Barents Sea shelf was uplifted while the deep Cretaceous basins of the Norwegian-Greenland Sea to the south continued to subside rapidly (e.g. Breivik et al., 1999). The regional Cretaceous transgression into the Barents Sea therefore only resulted in a shallow shelf leaving a condensed marine sedimentary sequence of calcareous sandstones, sandy and glauconitic mudstones and thin limestones of the Kviting Formation in the central parts of the Barents Sea (Fig. 21) (Dalland et al., 1988; N0ttvedt et al., 1993a,b). The degree of uplift increased northwestwards so that Svalbard and the whole of the northwestern Barents Sea platform areas were eroded during Late Cretaceous times. The timing of the regional uplift in relation to the onset of the tectonic phase in the Norwegian-Greenland Sea and its bearing on the interpretation of the Eurekan Orogeny is uncertain. Data from Svalbard indicate that the first phase of compression and folding of strata in north Greenland and Svalbard took place between Albian and Paleocene times (Steel and Worsley, 1984; Hanisch, 1984). Basin fill in pullapart basins of the Wandel Sea Mobile Belt seems to constrain the dating of the deformation to between middle Turonian and end Maastrichtian times. Further constraint may be inferred by the fact that the transgressive base of the Kviting Formation is a prominent hiatus of Cenomanian to Santonian age,

29

pointing to the initiation of tectonism and associated regional uplift at that time. This coincides with the onset of the Late Cretaceous tectonic episode and accelerated subsidence of the MOre and Vcring Basins, strongly pointing to a link with the rapidly subsiding basins along the Barents Sea margin and the Wandel Sea Basin through the de Geer Zone (as suggested by Brekke and Riis, 1987). Separate from the Turonian rifting and increased basin subsidence, Faleide et al. (1993a,b) argues for the initiation of the Eurekan Orogeny in the late Santonian, and refers to a related phase of uplift and faulting in And0ya in that respect (Dalland, 1981). This is also in agreement with the plate reconstruction models of Rowley and Lottes (1988) which predicts transcurrent movements between north Greenland and Ellesmere Island/Svalbard from anomaly 34 time onwards. In that case, the initiation of transpression on the de Geer Zone may have been coincident with the Campanian compressional phase of the Late Cretaceous tectonic episode of the MOre and V0ring Basins (Brekke and Riis, 1987; Bj0rnseth et al., 1997; Brekke, 2000).

The Tertiary The sedimentary environment of the Tertiary was a response to the Palaeogene transition from the continental rift setting to a drift and passive conti-

Fig. 22. The palaeogeography of Early Eocene times plotted on the 53 Ma plate reconstruction. Some key references for compilation: H~kansson and Stemmerik (1984); Steel and Worsley (1984); Larsen (1984); Ziegler (1988); Surlyk (1990); Livsic (1992); NOttvedt et al. (1993a); Knott et al. (1993); Stemmerik et al. (1993); Bowman (1998); Larsen et al. (1999); Brekke et al. (1999); supplemented by in-house studies. See Fig. 5b for legend.

30 nental margin setting, and the subsequent climatic change giving Neogene glaciations. The widespread Late Cretaceous polyphasal tectonic episode may be viewed as plate tectonic adjustments and rifting in the North Atlantic leading to the final continental break-up in the Norwegian-Greenland Sea area in the Late Paleocene/Early Eocene (e.g. Skogseid and Eldholm, 1989; Roberts et al., 1997; Brekke, 2000). This event was associated with a regional uplift of the whole of the Norwegian-Greenland Sea and its circumference. The cause of the regional uplift is uncertain; the debate involves the Iceland mantle plume (White, 1989; Skogseid, 1994) to intraplate stress (Cloetingh et al., 1990). The axial part of the Norwegian-Greenland Sea was highly uplifted due to increased heat flow along the future spreading axis just prior to break-up (Fig. 22). The general uplift drastically reduced the size of the basins and expanding the hinterland areas. The depositional area of the gross North Sea Basin has been estimated to have been reduced to 70% of that of the late Maastrichtian (Oakman and Partington, 1998). The Paleocene and Eocene

In the southern and central North Sea, chalk deposition continued through the Maastrichtian and Danian, typically filling in a slightly tectonically rejuvenated seafloor topography by the redeposition of pelagic ooze. But as the regional uplift, basin reduction and hinterland expansion initiated in the late Danian/early Thanetian, clastic input increased drastically and effectively ended the carbonate depositional environment (Fig. 2). The clastics that derived from the emergent Shetland Platform in Paleocene to Early Eocene times in the North Sea area may be divided into two sedimentary units (Bowman, 1998). The older unit consists of a sequence of aggradational submarine fan deposits. The younger, latest Thanetian to earliest Ypresian unit consists of a progradational sequence of muds and localised sands. In the late Thanetian the North Sea Basin was cut off from oceanic circulation giving a basin-wide anoxic phase. This was probably at the peak uplift of the Norwegian-Greenland Sea axis just prior to continental rupture. Subsequent to the final break-up, the whole area started to subside and the resulting relative sea-level rise ended the North Sea anoxic phase in the latest Thanetian, at the base of the progradational unit. The rest of the Eocene was characterised by deep-water sandy turbidite pulses as a response to fluctuations in relative sea-level and hinterland rejuvenation. The well-known sandy Frigg Fan is one of the major low-stand fan systems of this period (Bowman, 1998; Martinsen et al., 1999).

H. Brekke et al.

The regional uplift is recorded as a hiatus and erosional break of probable late Danian/early Thanetian age across the western bounding platforms and basin flanks of the MOre and Voring Basins, and also across highs and domes within the V0ring Basin (Brekke et al., 1999; Martinsen et al., 1999). The Danian/Thanetian was therefore a period of dramatic shallowing of the MOre and VCring Basins and the emergence of the surrounding areas. Isolated outcrops in East Greenland show Upper Cretaceous and Lower Paleocene offshore shales and submarine channel turbidites unconformably overlain by upper Danian and Thanetian fluvial conglomerates and sandstones, implying a coincident similar dramatic shallowing there (Larsen et al., 1999). In the Voring Basin, the basin flanks, highs and domes were eroded and sediments were deposited in shallow synclinal, perhaps circulation restricted, areas within the basin. In the deeper MOre Basin and in the northern North Sea, thick Paleocene/earliest Eocene sedimentary wedges prograded into the basin from the platforms on both flanks (Brekke et al., 1999; Martinsen et al., 1999). This symmetrical progradation of low-stand wedges downlapping the base Paleocene hiatus and thinning towards the basin axis, is seen throughout the northern North Sea and the MOre Basin (Martinsen et al., 1999). In the MOre Basin, there is a widespread sand-rich unit near the base of the Paleocene. The progradation from the western platform areas probably led to the eastward advancement of the western shoreline into the MOre Basin (Brekke et al., 1999). The final rupture of the continental crust within the axial platform area between Greenland and Norway was accompanied by the eruption of basaltic lavas (Fig. 22). Large volumes of tholeiitic flood basalts flowed across the whole of the eroded platform area and stopped at the newly established shoreline in the VOring and MOre Basins (Brekke et al., 1999). This shoreline is defined by the limit of the early flows, termed the "inner flows" by Talwani et al. (1983). By the subsequent subsidence and tectonic activity the shoreline retreated westwards to the present FaeroeShetland and VOting Escarpments building a line of lava deltas (Smythe et al., 1983; Planke et al., 1999). Due to recent uplift and erosion, the position of the Paleocene shoreline in East Greenland is not known. Subsequent to the latest Paleocene/earliest Eocene voluminous break-up magmatism the newly established spreading axis and surrounding lava platforms started to subside and eventually became submerged. The main part of the Eocene succession is a slope/basin floor system downlapping and thinning towards the east throughout the northern North Sea and the MOre Basin. This system is interpreted to be a low-stand wedge related to fall and subsequent rise in

31

S e d i m e n t a r y environments offshore N o r w a y - - an o v e r v i e w

relative sea level (Martinsen et al., 1999). The Eocene of the Norwegian Sea area is dominated by marine claystone (Dalland et al., 1988). After the Late Cretaceous uplift and erosion, the western Barents Sea shelf was transgressed in early Thanetian times leaving a Paleocene to Oligocene uniform sequence of outer sublittoral to deep-shelf claystone with minor siltstone, tuffaceous and carbonaceous horizons (Dalland et al., 1988; Faleide et al., 1993a). The Paleocene to Oligocene environments of Svalbard were much more complex because of the involvement in the Spitsbergen Orogeny (Fig. 2). The Late Cretaceous uplift and erosion is testified by Lower Paleocene strata overlying Albian and Aptian strata onshore Spitsbergen (Steel and Worsley, 1984). The Spitsbergen Orogeny was due to dextral movement on the de Geer Zone megashear in which Steel and Worsley (1984) record Paleocene transtension, Eocene transpression and Oligocene oblique separation and rifting. Further south, in the S~rvestsnaget Basin and Troms~ Basin, the timing of transtension and transpression was different (Faleide et al., 1993b). Although changing configuration considerably through time, the Paleocene and Eocene basins onshore Spitsbergen exhibited a complete lateral transition from alluvial fan, delta plain and fan delta into delta front and shallow to distant shallow shelf environments (Steel and Worsley, 1984). The Oligocene and Miocene The Oligocene and Miocene of the Norwegian continental margin reflect the sedimentation on a marine, subsiding passive margin overprinted by intermittent regional phases of tectonic movements and uplift. The two main phases of compression are associated with the formation of intrabasinal domes and arches in the V~ring Basin and around the Faeroe Islands, dated to latest Eocene/earliest Oligocene and Middle Miocene (Boldreel and Andersen, 1993; Blystad et al., 1995; Andersen and Boldreel, 1995; Dor6 and Lundin, 1996; Brekke, 2000). Vfignes et al. (1998) proposes that the compression was continuous from Eocene to Miocene without discrete phases. As documented by Gradstein and B~ickstr6m (1996), however, a Late Eocene/Early Oligocene phase of compression would coincide very well with a regional hiatus on the eastern basin margins across major parts of the North Sea and on the Halten Terrace. Martinsen et al. (1999) prefer an Early Oligocene date for this stratigraphic break. A prograding unit of coastal (deltaic?) deposits on the northeastern part of the Tr~ndelag Platform, dated to the Early Oligocene by Eidvin et al. (1998), constitutes a stratigraphic record of this event (Fig. 2). Seismic and biostratigraphical

evidence points to a regional uplift and erosion of the North Sea and Norwegian Sea region and surrounding mainlands just prior to the late-Middle Miocene phase of compression (Jordt et al., 1995; Gradstein and B~ickstr6m, 1996; Martinsen et al., 1999; Eidvin et al., 2000). This phase of uplift is reported to have lasted for about 7 million years (Anderton et al., 1979; Jordt et al., 1995). In the deep V~ring and M~re Basins the Miocene erosion was entirely submarine and sedimentation remained deep marine mud and siliceous oozes across the hiatus (Eidvin et al., 1998). In the V~ring Basin a considerable seafloor dome topography was filled in during the Late Miocene (Brekke et al., 1999). In the North Sea Basin, more proximal to the mainland, the uplift gave rise to widespread influx of sand (Utsira Formation) resting unconformably on the Lower and Middle Miocene shales (Rundberg et al., 1995; Eidvin et al., 2000). The Pfiocene and Pleistocene The periods of glaciation caused by the climatic deterioration in the Neogene had a significant impact on the sediment supply to the shelf in the Late Pliocene and Pleistocene. The onset of major glaciations at approximately 2.7 Ma led to deep mainland erosion and to the deposition of huge sediment volumes on the adjacent shelf (Riis and Fjeldskaar, 1992; Vfignes et al., 1992; Riis, 1996; Martinsen et al., 1999; Eidvin et al., 2000). The mainland and the Barents Sea area were uplifted tectonically in dome-shaped areas, and in general as an isostatic response to the erosion (Riis, 1996; Dehls et al., 2000). A regional hiatus is recorded in the lower part of the Upper Pliocene (2.7 Ma) along the Norwegian continental margin (Eidvin et al., 2000) (Fig. 2). The hiatus is mainly preserved as a surface of nondeposition downlapped by the sandy muds of the prograding Upper Pliocene sediment apron. In the V~ring Basin, the Upper Pliocene sediments rest on the Middle Miocene unconformity on the summits of large domes and arches and in large areas in the western parts of the basin and the adjacent marginal high (Brekke, 2000). On the mainland and on the shallow parts of the margin, there is an erosional unconformity below the glacigenic sediments. This unconformity was formed in the Pleistocene as a result of large ice sheets which extended to the shelf break. The largest sediment volumes are found in several kilometres thick fan systems adjacent to major submarine channel systems, like just northeast of Kvit~ya in the Arctic Ocean, west of Bj~rn~ya, and in the M~re Basin area at the mouth of the Norwegian Channel (Riis, 1996; Eidvin et al., 2000).

32

Summary and conclusions (1) In Carboniferous times the area, situated at low latitudes, developed a wide range of sedimentary environments through time and space. In that period, the area was situated between a northern and southern ocean. Through the regional drainage pattern sediments were transported into these oceans from a highland situated in the southern part of the Norwegian-Greenland Sea, probably governed by the tectonic framework. (2) Permian times saw both the final plate tectonic assemblage of Pangea and the subsequent onset of rifting of the supercontinent. The period was characterised by an early period of magmatism, followed by widespread erosion, and the subsequent development of shallow marine environments of low latitudes. (3) Triassic and Early Jurassic times were a period of peneplanation of hinterlands and clastic infilling and shallowing of basin areas. (4) Through renewed rifting and relative sea-level rise in the Middle to Late Jurassic, the seaway between the northern and southern oceans became permanent and shallow marine; clastic environments were established throughout the region, now at middle latitudes. Subsequent to sea-level oscillations during Late Jurassic tectonic activity, clastic-starved environments were established as large basin areas subsided and the sea level rose through the Cretaceous. (5) Through hinterland rejuvenation by early Tertiary rifting and subsequent continental separation and seafloor spreading, clastic marine environments were re-established, this time at high latitudes and cold waters. In the Neogene, major parts of the clastic input to the marine environments originated from glacial erosion. (6) An extension model has been proposed in which the crustal extension and rifting activity is grouped into three broad rifting episodes, the Carboniferous/Triassic, the Middle Jurassic/earliest Cretaceous, and the latest Cretaceous/Paleocene episodes. These episodes of extension were separated by two periods of thermal relaxation, from the end of the Early Triassic to the middle of the Early Jurassic, and from the earliest to latest Cretaceous. It is argued that the Carboniferous/Triassic episode resulted in a significantly larger share of the total accumulated extension than the Middle Jurassic/earliest Cretaceous episode. The latest Cretaceous/Paleocene rifting episode is considered to have contributed least to the total extension of the continental crust between Scandinavia and Greenland. (7) It seems that the Carboniferous/Triassic rifting was distributed across a broad area. For each of the two subsequent rifting episodes, the central axes of

H. Brekke et al.

rifting shifted towards the axial parts of the continental margins of the present Norwegian-Greenland Sea, thereby progressively narrowing the zone of actual rifting. (8) In the rift model proposed, the axial area of the continental margins of the present NorwegianGreenland Sea was underlain by long-lived, stable basement blocks. The contention is that these were established already in Early Carboniferous times as the Carboniferous to Triassic rifting activity took place symmetrically along both flanks, i.e. the inner parts of the present continental margins of mid-Norway and East Greenland, respectively, leaving the axial area non-rifted and non-attenuated. In such a setting, the progressive shift of the rift axes of the subsequent extension episodes towards the central parts of the Norwegian-Greenland Sea area is believed to reflect the preference for non-rifted crust in place of previously rifted and attenuated crust. The explanation may be that crust that has been through the cycle of rifting and attenuation, and subsequent thermal relaxation and subsidence becomes physically more resistant to further extension. By the time of the latest Cretaceous/Paleocene extension episode, the area of crust still unaffected by previous rifting was considerably reduced, so that the rifting activity was restricted to a narrow zone, along which the final continental break-up occurred. Such a narrow zone of crust available for extension would then not contribute much to the total extension accumulated through time, even with a high Beta-factor locally. The lack of a broad area of easily extendable crust, may also explain the apparently clean-cut, vertical rupture of the crust in the process of continental break-up, in which the extension was compensated by magmatism instead of attenuation of the crust., (9) Rejuvenation and subsequent denudation of hinterlands played an important role in the evolution of sedimentary environments and facies through time. In the Carboniferous and Permian, the uplifted hinterlands were of two categories by origin. The hinterlands of one category were caused by uplift by orogenic activity (Variscan and Uralian Orogenies). Uplifts of the other category were probably associated with the rifting between northern Europe and Greenland, including the flanking mainlands of Norway and Greenland and the elevation of the suggested central watershed area of Visean times. The Early to Middle Jurassic and the Paleocene were periods of prominent hinterland rejuvenation. Both were associated with increased heat flow in the initial stages of the major episodes of rifting and extension. In both cases, the whole area of the regional rift zones and their flanks experienced regional uplift causing widespread erosion and concomitant development of progradational

S e d i m e n t a r y environments offshore N o r w a y - - an o v e r v i e w

systems of clastic sediments. Typical of the Early and Middle Jurassic was the emergence of a number of semi-regional domes within the regionally elevated area from the North Sea and British Isles to the borders of the Barents Sea. This caused a complex configuration of emerged and submerged areas in that period.

Acknowledgements We want to thank our colleagues Fritjof Riis and Paul Grogan for their valuable input to the present study. We are also grateful to the referees, Roy H. Gabrielsen and Snorre Olaussen, for their very critical review of the first draft of the manuscript. We are indebted to the editors, for their patience.

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The alluvial cyclicity in Hornelen Basin (Devonian western Norway) revisited: a mu lti parameter

sedimentary analysis and stratigraphic implications Atle Folkestad and Ronald J. Steel

Accommodation space and sediment supply are the main factors controlling the spatial and stratigraphic pattern of the infill of sedimentary basins. The interaction of these factors, over periods of time, can be identified in the basin-fill succession by the changes of, for example, grain size, bed thicknesses and erosion-surface frequency, which are parameters easily measurable in outcrops as well as in cores and on image logs of subsurface successions. This study demonstrates how this approach can be used to recognise changes in the accommodation-space/sediment-supply ratio, based on an analysis of the alluvial succession in the Devonian Hornelen Basin of western Norway. Four of the basin-fill cyclothems have been logged (a total of 525 m of the basin-fill) with a systematic quantification of such parameters as grain size, bed thicknesses, erosion-surface density, occurrence of intraformational clasts and extraformational clasts, clast position within beds, and the degree of soft-sediment deformation. The analysis leads to a new understanding of the cyclicity and style of sedimentation in the Hornelen Basin. The deposits of the alluvial succession can be divided into three facies associations: (1) fluvial channels; (2) channel-mouth splays; and (3) distal floodbasin deposits. The dynamics of basin infilling can be expressed as an interplay of stratigraphic accommodation-space creation (A) and sediment supply (S), commonly expressed as the A/S ratio. In broad terms, an increasing A/S ratio implies increasing preservation potential of sediment infill, whereas a decreasing A/S ratio signifies a decrease of preservation potential and an increasing probability of sediment bypass or erosion. Because of the critical importance of erosion in the A/S ratio concept, cycles in the A/S ratio for the succession can be identified by using peaks in the frequency of occurrence of erosion surfaces (quantified as "erosion-surface frequency") to pick A/S ratio minima, and minima of erosion surfaces to pick A/S ratio maxima. One of the more interesting results from the study shows that peaks in grain size occur somewhat after A/S ratio minima, the offset being caused by continued high levels of sediment supply and flow competence despite a relative increase in A where the offset represents a time-lag. Bed thicknesses show low values close to the A/S ratio maxima and minima, and peak where A approaches S. The effect can be compared with the depositional pattern along the length of a clinoform with a low-angle trajectory. In a proximal position the clinothem, after a certain time period, is thin due to low A/S ratio conditions, in the distal part it is thin due to a high A/S ratio conditions, whereas the greatest thicknesses are recorded in between these two extremes. The soft-sediment deformation parameter follows the pattern of the bed thickness parameter and is thus interpreted as being linked to the bed thickness. The clast parameters (intra-, extra-formational clasts and clast position within the beds) follow the pattern of grain-size and erosion-surface frequency parameters where increasing clast occurrence reflects lower A/S ratio and decreasing clast occurrence indicates higher A/S ratio. The approach described here can be applied easily to subsurface successions. In cored intervals, parameters such as grain size, erosion surfaces and bed thicknesses can be extracted. The same approach has been used on a Formation Micro Image Log from a well in the North Sea where bioturbation, erosion-surface density, set density and angle of lamination were quantified and cross-analysed in terms of shallowing and deepening trends.

Introduction

The dynamics of a basinal stratigraphic system can be described in terms of the changing ratio of the rate of accommodation-space development and the rate of sediment supply (referred to below as the A/S ratio) as done by Shanley and McCabe (1994). Bars, hydraulic bedforms and other geomorphic elements of a depositional system are likely to be better preserved when the A/S ratio is high, but poorer preserved when this ratio is low. Erosion and sediment bypass may prevail in the latter case.

In an alluvial basin such as the Old Red Hornelen Basin in Norway, where the sediment accumulation rates are estimated to have been as high as 2 m/ka (Steel et al., 1977), it is appropriate to consider the accumulation/preservation potential of sediments in terms of A / S ratio changes. The conceptual A/S ratio defines the dynamic state of the sedimentary system, whether it evolves towards a higher degree of preservation (maximum A/S) or a greater degree of erosional destruction (minimum A/S) of the deposits (Fig. 1) (see also Shanley and McCabe, 1994). Consequently, it can be expected, as a working hypothesis,

Sedimentary Environments Offshore Norway - Palaeozoic to Recent edited by O.J. Martinsen and T. Dreyer. NPF Special Publication 10, pp. 39-50, Published by Elsevier Science B.V., Amsterdam. 9 Norwegian Petroleum Society (NPF), 2001. "

40

A. Folkestad and R.J. Steel

'"~

Bypass surface"-'~1

Accommodationspace creation -" decrease

Starvation

Bypass and | erosion I Quadrant L/1 Deposition i..i" /

~= -

Starvationsurface

Accommodation space creation increase

Erosion Starvation Quadrant Quadrant Sediment supply decrease

Fig. 1. Conceptual relationship ~etween the changes in the rates of accommodation-space creation (A) and sediment supply (S). The A/S ratio defines conditions of sediment erosion, deposition and bypass (non-deposition). The negative axis of the accommodation-space creation represents erosion, whereas the negative axis of sediment supply is included to account for the mass-balance of the system; when deposition occurs in one area, erosion must occur elsewhere. The equilibrium line 1:1 in the diagram represents equal rates of accommodation-space creation and sediment supply (A -- S). The area of bypass in the positive quadrant of the diagram (A > 0, S < 0), is meant to indicate the condition of non-deposition in a situation where the rate of sediment supply is disproportional higher than the rate of accommodation-space creation. Similarly, the area of starvation in the same quadrant indicates conditions where the rate of accommodation-space creation grossly exceeds the rate of sediment supply.

that some measurable sedimentary parameters reflecting variation in the degree of sediment preservation may indicate changes in the A / S ratio in a stratigraphic succession (see also Gardner, 1995). If this is the case, this crucial aspect of a depositional system's dynamics might then be deciphered from the sedimentary succession. The aim of the present study is to evaluate this hypothesis through an analytic approach in identifying time-trends in the A / S ratio and their turn-around points (i.e. changes from a decreasing to an increasing trend, or vice versa) in a thick, representative portion of the alluvial succession in the Hornelen Basin. The basin is probably ideal for testing such ideas because of its high but variable rates of sediment accumulation and its very thick stratigraphic succession. The quantitative method used for this purpose was originally developed by Folkestad (1995) and is presented here in a refined form.

Tectonic and stratigraphic setting The Hornelen Basin is a small (<2000 kin2), E - W oriented, oblong basin in southwestern coastal Norway. The basin-fill has a huge stratigraphic thickness (about 25 km), consisting of conglomerates, sandstones, mudstones and siltstones, and is probably of Middle Devonian age (Steel et al., 1977; Steel, 1988). Alluvial fans, and local fan deltas developed along the

margins (Steel et al., 1977; Larsen and Steel, 1978; Steel and Gloppen, 1980; Anderson, 1997) whereas a westward-directed, sandy fluvial system developed along the basin axis with a terminal floodbasin in the western part (Steel and Aasheim, 1978) (Fig. 2). The Hornelen Basin is bounded by a low-angle listric fault in the east and high-angle normal faults to the north and south. The basin-fill succession is gently folded, with several east-west trending axes, and generally tilted towards the east, except of the northern margin where the strata tend to be deformed and often occur in a vertical position (Steel and Aasheim, 1978). The surrounding and underlying basement rocks are Precambrian gneisses and Cambro-Silurian schists, metagreywackes, metabasalts, granodiorite and gabbro of the Caledonian orogen (Bryhni, 1964; Steel et al., 1977). The basin probably originated by gravitational collapse of the Caledonian orogen during the final stage of the orogeny (Hossack, 1984; Seguret et al., 1989). The Caledonian crust along the Fennoscandian border is thought to have been excessively thickened by the folding and thrusting, but returned to its normal (present-day) thickness b y a reactivation of some of the thrusts as listric, normal detachment faults (Norton, 1986; Steel, 1988; Seranne et al., 1989; Fossen and Rykkelid, 1992). The Hornelen sedimentary succession, which developed syntectonically on a west-gliding, hanging-wall basement block, slid discontinuously westwards while rotating slightly eastwards during the extensional collapse. A clear evidence of this development can be seen along the low-angle normal fault at the eastern edge of the basin, where the basin-infill strata show a systematic eastward onlap of the footwall through time. This syntectonic and migratory aspect of the basin was first recognised by Bryhni (1964) and documented in greater detail by Steel et al. (1977) and Steel and Gloppen (1980). Steel (1988) has demonstrated that the alluvial-fan wedges along the transfer faults of the northern and southern basin margins are geometrically skewed eastwards, opposite to the direction of the tectonic displacement, supporting the extensional model for the origin of the basin. By the mid-80s, an early notion that the basin was formed purely due to strike-slip tectonics (Steel et al., 1977; Steel and Gloppen, 1980) had been replaced by the gravitational collapse model (Hossack, 1984; Steel, 1988; Seguret et al., 1989; Seranne et al., 1989). However, more recent structural research (Osmundsen et al., 1998) has re-emphasised the importance of strike-slip faulting interacting with the regional extension. The thick infill of the Hornelen Basin is organised into more than 200 basin-wide cyclothems, 100200 m in thickness (Fig. 3). These cyclothems, orig-

The alluvial cyclici~ in Hornelen Basin (Devonian western Norway) revisited

41

Fig. 2. Sketch map of the Hornelen Basin, showing generalised palaeo-flow direction and facies distribution (from Steel, 1988). The location of the study area is indicated by the asterisk.

inally distinguished solely on the basis of facies and grain size, have been interpreted to reflect the changing subsidence rate of the basin floor during the gravitational collapse, with the maximum rates of accommodation-space development corresponding to the finegrained parts of the cyclothems (Steel, 1988). The cyclothems are generally asymmetrical as far as the upward change in grain size is concerned. This asymmetry is well displayed at the southern margin, where relatively thick coarsening-upward parts are capped by much thinner fining-upward parts. At the northern margin, cyclothems are less asymmetrical, with the fining-upward part constituting up to 35% of a cyclothem thickness (Gloppen and Steel, 1981; Steel, 1988). In the axial area of the basin, the cyclothems are often nearly symmetrical in terms of grain-size trend motif and facies stacking pattern, and also show abundant soft-sediment deformation (Steel and Aasheim, 1978). The alluvial succession

The sedimentary succession of the Hornelen Basin consists of conglomeratic alluvial fan and fan delta deposits, sandy braided-stream deposits, sandy to silty floodbasin deposits and silty to muddy lacustrine deposits. The fans at the northern margin are characteristically small and steep, debris-flow dominated systems, whereas those at the southern margin are stream-flow dominated, with greater radii and gentler slopes (Gloppen and Steel, 1981) (Fig. 2). The axial part of the basin-fill succession was deposited by an extensive, sand-prone fluvial system

debouching westwards and northwards into a floodbasin/lacustrine area, rendering it a kind of basinaxis terminal fan, or lacustrine braid-plain delta. The alluvium of the axial fluvial system shows a downstream lateral facies change, with trough cross-stratified coarse pebbly sandstones and conglomerates giving way to finer-grained sandstones characterised by planar cross-stratification and further passing into an assemblage of alternating fine-grained sandstones, siltstones and mudstones, dominated by ripple cross-lamination (Steel and Aasheim, 1978). The palaeo-current measurements in the axial alluvium are consistent with a westerly transport, but with secondary deviations towards both northwest and southwest (Fig. 2). The representative portion of the axial alluvial succession selected for the present study is in the western, relatively distal part of the basin (Fig. 2). The sedimentary succession logged is 525 m thick and comprises four sections, labelled 1-4 (Fig. 3). Four 2-D panels (Figs. 4 and 5) showing the sedimentary architecture of the alluvium have been constructed on the basis of section 2 (Fig. 3) to illustrate the character of facies associations, the stacking patterns of facies and the lateral persistence of beds and bed sets. The deposits have been grouped into three facies associations, interpreted to be floodbasin facies, channel-mouth splay facies, and fluvial channel facies. The three facies associations are described briefly below, and their broader basinal context and more detailed description are given by Steel and Aasheim (1978), and Folkestad (1995).

42

A. Folkestad and R.J. Steel

Fig. 3. An oblique eastward aerial view of the axial part of the Hornelen Basin, showing the spectacular basin-wide cyclothems 100-200 m thick, represented by the morphological steps. The measured Sections 1-4 are indicated by the white bars, with a total thickness of 525 m. Courtesy of Bremanger Kommune 9

The fluvial channel facies association consists of mainly medium-grained sandstones dominated by an alternation of planar tabular cross-stratification and tangential low-angle cross-stratification, with multiple erosion surfaces (Fig. 4C). The cross-stratified sandstones tend to be overlain by slightly finer-grained sandstones with medium-scale trough cross-stratification and/or planar-parallel stratification (commonly showing parting lineation), but this trend is less obvious where frequent erosion surfaces occur. Soft-sediment deformation is relatively rare in this facies association. The multiple erosional surfaces, the stratification types and the sheet-like geometry of the bed sets suggest relatively broad and shallow braided streams, possibly flashy and ephemeral, dominated by transverse and longitudinal bars (see also Cant and Walker, 1978; Miall, 1985). The planar-parallel stratification indicates channel shoaling, whereas the occasional occurrences of trough crossstratification scoured into the bar tops suggest some cross-cutting minor channels, formed at the falling stage of stream flood. The lack of vegetation in Devonian time is likely to have promoted the de-

velopment of such a fluvial system (Schumm, 1968; Macnaughton et al., 1997). The channel-mouth splay facies association is dominated by fine-grained, low-angle stratified sandstones with subordinate, but occasionally relatively thick, planar cross-stratified sets, accompanied by very fine-grained sandstones with planar-parallel stratification and abundant soft-sediment deformation features (Fig. 4B,C and Fig. 5). Sandstone units with trough cross-stratification and ripple cross-lamination, commonly convoluted, are minor components of this facies association. The ripple crosslamination occasionally indicates wave action. The depositional units are characterised by a coarseningupward transition from mudstone and siltstones or very fine-grained, wave ripple-laminated sandstone to fine-grained, low-angle cross-stratified sandstone, and eventually to planar cross-stratified, fine-grained sandstones (Fig. 4C and Fig. 5). The latter is typically a solitary cross-set, sigmoidal or top-truncated, tangential, often strongly deformed. This coarseningand thickening-upward motif is followed in turn by a fining-upward trend of a fine-grained, low-angle strat-

The alluvial cyclicity in Hornelen Basin (Devonian western Norway) revisited

43

Fig. 4. Field-measured stratigraphic logs and facies distribution panels. Panels A and B show the channel-mouth splay and floodbasin/lacustrine facies associations, and panel C show the fluvial channel-mouth splay facies associations. The three panels A-C are located in the lower half of outcrop Section 2 (see Figs. 3 and 6), with panel A showing the fine-grained part of an alluvial cyclothem.

ified or trough cross-stratified sandstone overlain by very fine-grained, ripple-laminated sandstone, planeparallel stratified sandstone and siltstone or mudstone (Fig. 5). The thicker coarsening-upward lower part of such units, with an outstanding set of planar cross-strata and soft-sediment deformation, overlain by the fining-upward motif of a waning flow, suggest jointly

the progradation and abandonment of a mouth-bar form, or a splay, in the area of terminal flow expansion (see also Elliott, 1974; Miall, 1985; Fielding, 1986; Pulham, 1989; Macnaughton et al., 1997). The wave-ripple cross-lamination in the alternating mudstone/siltstone deposits indicates a standing water environment, probably a floodbasin pond or shallow lake.

Fig. 5. Field-measured panel D, showing channel-mouth splay facies associations (yellow) and poorly developed floodbasin/lacustrine facies association (green). The outcrop shown is 50 m long and 11 m high, located in the upper half of Section 2 (see Figs. 3 and 6). Note the overturned soft-sediment deformation folds in the cross-stratified sandstone facies in the middle part of the panel. The panel illustrates all the sedimentary parameters used in the present study.

The alluvial cyclicity in Hornelen Basin (Devonian western Norway) revisited The floodbasin/lacustrine facies association consists of mudstones and siltstones thinly interbedded with very fine-grained, planar-parallel stratified sandstones (Fig. 4A and Fig. 5). The mudstone and siltstone layers show an alternation of uneven, subhorizontal parallel lamination and ripple cross-lamination, occasionally with convolutions, water-escape structures and desiccation cracks. The sandstones occur as isolated sheet-like layers, sharp-based or erosive, with plane-parallel stratification (showing parting lineation), ripple cross-lamination or low-angle parallel stratification, common normal grading, occasional desiccation cracks and local basal mudclasts. A common motif in these sandstone units is low-angle cross-stratification passing upwards into planar stratification, capped by ripple cross-lamination and siltstone/mudstone. The motif is often repeated in the form of composite, amalgamated units several metres thick (Fig. 4), traceable laterally for more than 50 m (Fig. 5). The individual sandstone sheets are thought to represent successive flood events in an area of shallow standing water. The desiccation cracks suggest ephemeral floodbasin ponds, occasionally turning into a shallow lake (e.g. see Elliott, 1974; Tunbridge, 1984). In the measured cyclothems, with their characteristic fining- and coarsening-upward trends (Steel and Aasheim, 1978), the facies associations show a stacking pattern where fluvial channels are succeeded by channel-mouth splays and floodbasin deposits in the fining-upwards part of a cyclothem. In the coarsening-upward part of a cyclothem, the facies associations are stacked as floodbasin deposits succeeded by channel-mouth splays and fluvial channels above (see Fig. 6).

Multiparametric analysis Stratigraphic changes in the A / S ratio are thought to be reflected in the variation of a number of sedimentological parameters, whose monitoring in the sedimentary basin-fill succession has been attempted in the present case. The selected parameters included mean grain size, intraformational and extraformational clasts occurrences, degree of soft-sediment deformation, bed thickness, erosion-surface frequency, and scour relief, all of which can be objectively measured and monitored through a sedimentary succession. An understanding of how the sedimentary parameters are related to the framework of the A / S ratio changes is crucial to an understanding of the depositional environment and associated allocyclic processes. The parameters measured are a continuous quantitative or semiquantitative variables of the sedimentary succession, with the values determined and

45

averaged for each 0.5 m and 2 m. Each parameter was then plotted as two graphs (Fig. 6), with the exception of bed thickness, which was only averaged over 2 m intervals. Mean grain-size parameter of sediment was determined visually for each 0.5 m interval and quantified as follows" coarse sandstone - 6; medium-grained sandstone - 5; fine-grained sandstone - 4; very fine-grained sandstone - 3; siltstone - 2; mudstone = 1; unexposed interval - 0. Intraformational and extraformational clasts were recorded in terms of the maximum particle size (MPS) and the dominant size for each 0.5 m interval was considered; separate graphs were made for intraformational clasts and for extraformational clasts. The basal-lag or dispersed clasts parameter indicates whether the intraformational clasts and extraformational clasts are concentrated as a lag at the lower bed boundary or scattered within the bed associated with a lag at the lower bed boundary. Intervals with no gravel clasts were given the value of 0, those containing a bottom lag were given a value of 1 and those with scattered clasts were given the value of 2. Soft-sediment deformation is a common type of sedimentary structure in the succession, and its intensity, or "degree", has been determined visually and averaged for each 0.5 m interval. Erosion-surface frequency was determined as the number of erosion surfaces per 0.5 m thickness interval. Bed thicknesses were measured and a mean thickness was calculated for each 2 m interval. Scour relief of the erosion surfaces was measured wherever significant, but the scours were mainly broad and shallow with little or no determinable relief. The sedimentary parameters were plotted, analysed individually for possible trends, and defined as systematic stratigraphic changes of either an increasing or a decreasing type. The scour relief parameter was not used due to the scarcity of (non-zero) data. The trends were determined by visual inspection for each parameter (Fig. 6).

The A/S ratio cycles Since the A / S ratio describes whether the sedimentary system moves towards a higher degree of preservation (maximum A/S) or a greater degree of erosional destruction (minimum A/S) of the deposits, the erosion-surface frequency parameter gives a good approximation of the changes of the A / S ratio. In times with a high degree of preservation of the deposits, little erosion takes place whereas during erosional destruction of the deposits, erosional processes dominates. Thus the large-scale trends of the erosion-surface frequency parameter is here interpreted to be equal to the large-scale changes of the

46

A. Folkestad and R.J. Steel

Fig. 6. The stratigraphic plots of the various sedimentary parameters measured in outcrop Sections 1-4 (Fig. 2) with the interpreted cycles of the A/S ratio. The sedimentary parameters were analysed individually for possible trends. Concentration of repeated high or low parameters values within a field were favoured over the occurrence of occasional extreme high or low values. The visually determined trends of each parameter is shown on the side of each graph. On the left hand side, the stacking pattern of the facies associations (FA) is shown, where: a -- fluvial channel; b = channel-mouth splay; c -- floodbasin. The lower part of Section 4 represents an unexposed interval.

A/S

ratio within the measured cyclothems (Fig. 6). The identification of the turn-around points of the A/S ratio minima and maxima allows a subdivision

of each cyclothem, within a 2-D framework, into two parts: a trend of an increasing A/S ratio, from a minimum to a maximum, and a trend of a decreasing

The alluvial cyclicity in Hornelen Basin (Devonian western Norway) revisited

Fig. 7. A generalised summary of the changes in sedimentary parameters and the A/S ratio in one complete cycle. Note the asymmetric trend of the grain-size parameter within the A/S ratio cycle.

A/S ratio, from the maximum to the next minimum (Figs. 6 and 7). Interestingly, the observed grain-size trend does not necessarily correspond strictly with the trend of the A/S ratio (Fig. 6). The coarsest sediment apparently tends to occur stratigraphically somewhat higher than the turn-around point of the A/S ratio minimum (Figs. 6 and 7). A possible explanation for the offset is that when the A/S ratio was reaching a minimum, coarse sediment was still being supplied to the upstream part of the depositional system, due to continued erosion and a time lag between the delivery of the sediment and its dispersal by the system. Consequently, coarse sediment continued to be deposited, extending the system's progradational signature, while the A/S ratio had already begun to increase. The geometric pinch-out, or architectural turn-around point, of the alluvial clastic wedge thus occurs at a higher stratigraphic level than the actual turn-around point of the A/S ratio in the succession. A similar observation has been reported by Wood et al. (1993) from a study of base-level fluctuations in sedimentation tank experiments. The latter authors used the base-level concept of Gary et al. (1972), equating base level with sea level, and observed that deposition of coarse sediment persisted into the early phase of base-level rise, rather than ceasing at the end of the base-level fall. The base-level minimum in that case can be said to correspond to a minimum A/S ratio, if the notion of this latter term is slightly simplified. A similar mismatch between the sediment grain-size maximum and the maximum regression level at a sequence boundary, is implicit in the sequence-stratigraphic models of Hunt and Tucker (1992) and Helland-Hansen and Gjelberg (1994). The stratigraphic distribution and mode of occur-

47

rence of gravel-size clasts, both intra- and extraformational, seem to follow the pattern of the A/S ratio changes without any significant or systematic offset. The increasing-upward abundance of these clasts is interpreted to reflect the decreasing A/S ratio. However, some other parameters, such as the intensity of soft-sedimentary deformation and bed thicknesses, show a notably different trend than the trend of the A/S ratio in a cyclothem. The latter two parameters tend to have low values at the turn-around points of the A/S ratio and higher values between these points (Figs. 6 and 7). The low values of these parameters at the A/S ratio minima support the notion of greater erosion efficiency and a decreased rate of accommodation-space creation at this turn-around point. At the turn-around point related to the A/S ratio maximum, the rate of the accommodation-space creation is high, but the sediment supply is low, resulting in deposition of thin beds, with bed-shear stresses (sand load) low enough to prevent substratum deformation. Between the turn-around points, the two parameters assume higher values because the A/S ratio approaches an equilibrium state, with the rate of sediment supply keeping pace with the rate of the accommodation-space creation (Fig. 1). An equilibrium state comparable to the one considered here has been discussed by Schlager (1993), Shanley and McCabe (1994), and Gardner (1995). In terms of the cyclic changes in the A/S ratio, an alluvial cyclothem in the Hornelen Basin can thus be divided into four parts, or stratigraphic intervals: a lower interval characterised by an increasing-upward trend of the A/S ratio (Fig. 7), with the A < S relation meaning system progradation, culminating in a transient state of equilibrium (A = S) and turning into the state of retrogradation (A > S) in the second interval; and a third interval recording the system's aggradation (A > S) towards another transient state of equilibrium (A = S), with the fourth interval representing renewed progradation (A < S), but within the context of decreasing rather than increasing A/S ratio compared to the first (lowermost) segment. These four stratigraphic intervals of an alluvial cyclothem can be considered in terms of the conventional sequence-stratigraphic models (e.g. Helland-Hansen and Gjelberg, 1994) within a 2-D framework.

Sequence-stratigraphic interpretation of the cyclic changes in the A/S ratio The application of the concept of sequence stratigraphy to continental sedimentary successions has become common in the recent years. However, it has mostly been applied to continental successions where

48 the sedimentation was controlled by changes in sea level (Shanley and McCabe, 1994; Burns et al., 1997; Willis, 1997; McCarthy et al., 1999). Sedimentation in the Hornelen Basin took place in an intra-cratonic basin (Steel et al., 1977) and involved no sea-level control. This renders the conventional concept and terminology of sequence stratigraphy non-applicable to this basin's alluvial succession (see also Olsen et al., 1995; Martinsen et al., 1999). It is worth noting that the four stratigraphic intervals of an alluvial cyclothem identified in the present study are comparable to the genetic sedimentary units (or systems tracts) used in sequencestratigraphic models (Helland-Hansen and Gjelberg, 1994). Since the changes in sediment-preservation potential are a key aspect of both the A/S ratio cycles and the conventional stratigraphic sequences, an interesting comparison can be made between the former and the sequence-stratigraphic models (Fig. 7). The subaerial unconformity and its correlative conformity bounding a stratigraphic sequence (sensu Van Wagoner et al., 1990) correspond to the phase of lowest sediment-preservation potential. With regard to the present study, this phase corresponds to a minimum of the A/S ratio. Furthermore, the maximum flooding surface (or condensed section) coincides with the phase of maximum sediment-preservation potential, which in comparison within the present study equals a maximum of the A/S ratio. The stratigraphic levels of minimum and maximum preservation potential represent the sedimentary system's maximum and minimum energy levels, respectively, recorded by the stratigraphic sequence or an alluvial basin-fill cyclothem. The equilibrium (A -- S) point of the increasing-upward (minimum to maximum) trend of the A/S ratio divides the lower part of the cyclothem into a progradational interval passing into a retrogradational interval. In a coastal setting, these two intervals could be referred to as the lowstand prograding wedge (sensu Hunt and Tucker, 1992) and the transgressive systems tract, respectively. Similarly, the equilibrium point (A -- S) of the decreasing-upward trend of the A/S ratio divides the upper part of the alluvial cyclothem into an aggradational interval followed by a progradational interval. As the A/S ratio begins to decrease and the sediment supply is barely sufficient to fill the available accommodation space, an aggradational style of deposition prevails in the system. But when the sediment supply eventually exceeds the rate of accommodation-space creation, the systems begins to prograde in a basinward-stepping pattern. In a coastal setting, this type of stacking pattern, following a maximum flooding phase, is referred to as a highstand systems tract in conventional sequence stratigraphy (Posamentier et al., 1988). A possible

A. Folkestadand R.J. Steel fifth segment, analogous to the forced regressive systems tract in a shoreline setting (Hunt and Tucker, 1992), is difficult to envisage in the sedimentary succession of a purely alluvial basin, except perhaps in incised valleys. Discussion and conclusions

The dynamic behaviour of an alluvial sedimentary system controlled by pulses of tectonic subsidence tends to show cyclic phases of progradation and retrogradation, with the corresponding decrease (minimum A/S ratio) and increase (maximum A/S ratio) of the sediment preservation potential. This oscillating behaviour results in systematic changes in the morphology and sedimentary characteristics of the depositional system, such as the upward coarsening and fining of the sediment grain size, due to the increasing and decreasing A/S ratio. A range of quantitative and semiquantitative sedimentary parameters appears to reflect changes in the A/S ratio controlling the depositional system, and the stratigraphic plots of these measurable characteristics can thus be used to recognise cyclic changes in the A/S ratio. On this basis, an alluvial succession can be analysed by a "sequencing" approach analogous to that used in conventional sequence stratigraphy. The full cycle of changes in the A/S ratio can be devised into four intervals, or phases, including progradation, retrogradation, aggradation and some "remnant" progradation. The cyclicity in the Hornelen Basin was previously defined on the basis of facies and grain-size changes alone. The present method documents a stratigraphic offset of the maximum grain-size points with respect to the turn-around points of minimum A/S ratio. This is interpreted in terms of a time lag between the A/S ratio minimum and the peak of the alluvial system progradation. The study shows that although the grain-size maximum and minimum in an alluvial succession may be a good proxy of the system's peak progradation and peak retrogradation, respectively, the grain-size parameter alone is insufficient to detect accurately the stratigraphic pattern of allocyclic changes in the A/S ratio. The resolution of the stratigraphic record appears to be improved when a wider range of sedimentary parameter are used. The bed thickness and soft-sediment deformation parameters show higher values between the A/S ratio minima and maxima, when the system passes through an equilibrium between the rates of accommodation-space creation and sediment supply. This effect can be compared to the longitudinal pattern of deposition along a clinoform with a low-angle trajectory. The clinothem tends to be thin in the source-proxi-

The alluvial cyclicity in Hornelen Basin (Devonian western Norway) revisited

mal part due to the low A/S ratio and in the distal part due to the high A/S ratio, while reaching the greatest thickness in the middle part, between the two extremes. The intensity of soft-sediment deformation follows the pattern of the bed thickness changes and is thought to be related to the latter parameter, probably by the specific sediment discharge controlling both the bed morphology and the bed shear stresses. The measured sections represent an 1-D data set that alone is insufficient to describe 3-D stratigraphic changes in the basin. However, in combination with the well-established palaeo-flow direction of the axial part of the basin (Steel and Aasheim, 1978) (Fig. 2), predictions of A/S ratio trends becomes feasible away from the measured 1-D profile. The A/S ratio maxima and minima represent correlatable zones and predictions about the A/S ratio trends in adjacent areas can be tested by measuring more sections. The approach described here can readily be applied to subsurface sedimentary successions, because parameters such as grain size, soft-sediment deformation, frequency of erosion surfaces and bed thicknesses can be determined from cores. A modified version of the technique has been applied to a Formation Micro Image Log from a North Sea well (Folkestad, 1999), where the degree of bioturbation, the frequency of erosion surfaces, the bed thickness and the angle of lamination were determined for half-metre intervals. The trends shown by these parameters were compared and used to recognise longer-term shallowing- and deepening-upwards cycles in the marine sedimentary succession.

Acknowledgements The authors wish to thank Jan Tveranger for discussions and improvements on the early versions of the manuscript. The NPF referees Wojtec Nemec and Michael Gardner are thanked for their very constructive criticism and suggestions for improvements. The authors would like to acknowledge the University of Bergen for their funding of the fieldwork. Liv Ims and Eden Potter have given technical assistance in the preparation of figures.

References Anderson, D.S., 1997. Sedimentary responses to base-level change in linked alluvial fan, lake and braidplain strata, Hornelen Basin, Western Norway. Unpublished Ph.D. Thesis, Colorado School of Mines, Golden, 269 pp. Bryhni, I., 1964. Sediment structures in the Hornelen series. Nor. Geol. Tidsskr., 5: 486-488. Burns, B.A., Heller, EL., Matzo, M. and Paola, C., 1997. Fluvial response in a sequence stratigraphic tYamework: example from the Montserrat fan delta, Spain. J. Sediment. Res., 67:311-321.

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Cant, D.J. and Walker, R.G., 1978. Fluvial processes and facies sequences in the sandy braided South Saskatchewan River, Canada. Sedimentology, 25: 625-648. Elliott, T., 1974. Interdistributary bay sequences and their genesis. Sedimentology, 21 : 611-622. Fielding, C.R., 1986. Fluvial channel and overbank deposits from the Westphalian of the Durham Coalfield, Northeastern England. Sedimentology, 33:119-140. Folkestad, A., 1995. Cyclicity in Hornelen Basin (Devonian), Norway. Unpublished Candidatus Scientiarum Thesis, University of Bergen, Norway, 96 pp. Folkestad, A., 1999. Compilation of sedimentological parameters from Formation Micro Image logs (FMI) to identify stratigraphic patterns. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway. Norwegian Petroleum Society (NPF) Extended Abstracts, 37. Fossen, H. and Rykkelid, E., 1992. Postcollisional extension of the Caledonide orogen in Scandinavia: structural expressions and tectonic significance. Geology, 20: 737-740. Gardner, M.H., 1995. Tectonic and eustatic controls on the architecture of mid-Cretaceous, stratigraphic sequences, Central Western Interior Foreland Basin of North America. In: S.L. Dorobek and G.M. Ross (Editors), Stratigraphic Evolution in Foreland Basins. Soc. Econ. Paleontol. Mineral. Spec. Publ., 52:243-281. Gary, M., McAffce, R. Jr. and Wolf, C.L., 1972. Glossary of Geology. American Geological Institute, Washington DC, 805 pp. Gloppen, T.G. and Steel, R.J., 1981. The deposits, internal structure and geometry in six alluvial fan-fan delta bodies (Devonian, Norway): a study in the significance of bedding sequences in conglomorates. In: F.G. Etheridge and R. Flores (Editors), Recent and Ancient Non-Marine Depositional Environments: Models for Exploration. Soc. Econ. Paleontol. Mineral. Spec. Publ., 3 1 : 4 9 56. Helland-Hansen, W. and Gjelberg, J., 1994. Conceptual basis and variability in sequence stratigraphy: a different perspective. Sediment. Geol., 92:31-52. Hossack, J.R., 1984. The geometry of listric normal faults in the Devonian basins of Sunnfjord, Western Norway. J. Geol. Soc., London, 141: 629-637. Hunt, D. and Tucker, M.E., 1992. Stranded parasequences and the forced regressive wedge systems tract: deposition during baselevel tall. Sediment. Geol., 81: 1-9. Larsen, V. and Steel, R.J., 1978. The sedimentary history of a debris flow-dominated alluvial fan: a study of textural inversion. Sedimentology, 25: 37-59. Macnaughton, R.B., Dalrymple, R.W. and Narbonne, G.M., 1997. Early Cambrian braid-delta deposits, MacKenzie Mountains, northwestern Canada. Sedimentology, 44: 587-609. Martinsen, O.J., Ryseth, A., Helland-Hansen, W., Flesche, H., Torkildsen, G. and Idil, S., 1999. Stratigraphic base level and fluvial architecture: Ericson Sandstone (Campanian), Rock Springs Uplift, southwest Wyoming, USA. Sedimentology, 46: 235-263. McCarthy, P.J., Faccini, U.E and Plint, G., 1999. Evolution of an ancient coastal plain: palaeosols, interfluves and alluvial architecture in a sequence stratigraphic framework, Cenomanian Dunvegan Formation, northeastern Columbia, Canada. Sedimentology, 46: 861-891. Miall, A.D., 1985. Architectural-element analysis: a new method of facies analysis applied to fluvial deposits. Earth-Sci. Rev., 22: 261-308. Norton, M.G., 1986. Late Caledonian extension in Western Norway: a response to extreme crustal thickening. Tectonics, 5: 195-204. Olsen, T., Steel, R.J., H0gseth, K., Skar, T. and ROe, S.-L., 1995. Sequential architecture in a fluvial succession: sequence stratigraphy in the Upper Cretaceous Mesaverde Group, Price Canyon, Utah. J. Sediment. Res., B65: 265-280. Osmundsen, ET., Andersen, T.B., Markussen, S. and Svendby, A.K., 1998. Tectonics and sedimentation in the hanging wall of a major extensional detachment: the Devonian Kvamshesten basin, Western Norway. Basin Res., 10:213-234.

A. Folkestad and R.J. Steel

50 Posamentier, H.W., Jervey, M.T. and Vail, ER., 1988. Eustatic controis on clastic deposition, I. Conceptual framework. In: C.K. Wilgus, B.S. Hastings, C.G.St.C. Kendall, H.W. Posamentier, C.A. Ross and J.C. Van Wagoner (Editors), Sea Level Change: An Integrated Approach. Soc. Econ. Paleontol. Mineral. Spec. Publ., 42: 109-124. Pulham, A.J., 1989. Controls on internal structures and architecture of sandstone bodies within Upper Carboniferous fluvial-dominated deltas, County Clare, Western Ireland. In: M.K.G. Whatley and K.T. Pickering (Editors), Deltas: Sites and Traps for Fossil Fuels. Geol. Soc. Spec. Publ., 41: 179-203. Schlager, W., 1993. Accommodation and supply a dual control on stratigraphic sequences. Sediment. Geol., 86:111-136. Schumm, S.A., 1968. River adjustment to altered hydrologic regime: the Murumbridge River and paleochannels, Australia. U.S. Geol. Surv. Prof. Pap., 598, 65 pp. Seguret, M., Seranne, M., Chauvet, A. and Brunel, M., 1989. Collapse basins: a new type of extensional sedimentary basin from the Devonian of Norway. Geology, 17: 127-130. Seranne, M., Chauvet, A., Seguret, M. and Brunel, M., 1989. Tectonics of the Devonian collapse-basins of Western Norway. Bull. Soc. Geol., Fr., 8: 489-499. Shanley, K.W. and McCabe, P.J., 1994. Perspectives on the sequence stratigraphy of continental strata. Am. Assoc. Pet. Geol. Bull., 78: 544-568. Steel, R.J., 1988. Coarsening-upward and skewed fan bodies: symptoms of strike-slip and transfer fault movement in sedimentary basins. In: W. Nemec and R.J. Steel (Editors), Fan Deltas: Sedimentology and Tectonic Setting. Blackie and Son, Glasgow, pp. 75-83.

A. FOLKEsTAD R.J. STEEL

Steel, R.J. and Aasheim, S., 1978. Alluvial sand deposition in a rapidly subsiding basin (Devonian, Norway). In: A.D. Miall (Editor), Fluvial Sedimentology. Can. Soc. Pet. Geol., Mem., 5: 385-413. Steel, R.J. and Gloppen, T.G., 1980. Late Caledonian (Devonian) basin formation, Western Norway: signs of strike-slip tectonics during infilling. In: H.G. Reading and EE Ballance (Editors), Sedimentation in Oblique-Slip Mobile Zones. Int. Assoc. Sedimentol. Spec. Publ., 4: 79-103. Steel, R.J., Ma~hle, S., RCe, S.-L., Spinnanger, A. and Nilsen, H.R., 1977. Coarsening-upwards cycles in the alluvium of Hornelen Basin (Devonian), Norway. Sedimentary response to tectonic events. Geol. Soc. Am. Bull., 88:1124-1134. Tunbridge, I.E, 1984. Facies models for a sandy ephemeral stream and clay playa complex; the Middle Devonian Trentishoe Formation of North Devon, U.K. Sedimentology, 31:697-715. Van Wagoner, J.C., Mitchum, R.M., Campion, K.M. and Rahmanian, D., 1990. Siliciclastic sequence stratigraphy in well logs, cores and outcrops: concepts for high-resolution correlation of time and facies. Methods in Exploration Series 7, American Association of Petroleum Geologists, Tulsa, OK, 55 pp. Willis, B.J., 1997. Architecture of fluvial-dominated valley-fill deposits in the Cretaceous Fall River Formation. Sedimentology, 44: 735-757. Wood, L.J., Ethridge, F.G. and Schumm, S.A., 1993. The effects of rate of base level fluctuation on coastal plain shelf and slope depositional systems: an experimental approach. In: H.W. Posamentier, C.E Summerhayes, B.U. Haq and G.E Allen (Editors), Sequence Stratigraphy and Facies Associations. Int. Assoc. Sedimentol. Spec. Publ., 18: 43-53.

Statoil Research and Technology, Department of Reservoir Characterisation, N-4035 Stavanger, Norway University of Wyoming, Department of Geology and Geophysics, Laramie, WY 82071, USA

51

Upper Permian Iowstand fans of the Bredehorn Member, Schuchert Dal Formation, East Greenland Mikkel Kreiner-Moller and Lars Stemmerik

Sedimentary facies analysis of the Upper Permian Bredehorn Member, Schuchert Dal Formation shows that it is composed of eight sand-rich facies that can be classified as Bouma sequences. The turbidites are grouped into six facies associations that form different parts of two lowstand fan systems in northern Jameson Land. The larger, Biskop Alf Gletscher-Oksedal fan extends for more than 30 km north-south. In the southern proximal part it is composed of channels and proximal lobe deposits with a composite thickness of 127 m. In the northern, more distal outcrops it consists of outer lobe deposits. Outcrops of the smaller, marginal Store Blydal fan are composed primarily of lobe sediments. Transport directions in both fans indicate north- and northeastward flow of the turbidites reflecting redistribution by axial turbidity currents of the transverse westerly infill. The northward tilting of the basin conforms with the data from the underlying Carboniferous fluvial systems but contrasts to the southward tilting of the Jurassic basin. Both fans are composed of three turbidite systems that can be correlated laterally. They are separated by shales. Each system represents deposition during sea-level lowstand, and the Bredehorn Member represents three events of lowstand fan deposition. The oldest part of the fan systems was deposited during latest Ravnefjeld Formation times and the sandstones are encased in organic-rich, source-prone shales.

Introduction The upper Palaeozoic-Mesozoic succession of the East Greenland rift basin is a classic analogue for the oil-bearing strata of the North Sea and Norwegian Shelf, and provides important data for new challenges in North Atlantic exploration. In central East Greenland, oil-prone source rocks are well documented from the Upper Permian Ravnefjeld Formation. Play models with the Ravnefjeld Formation as source rock and Upper Permian Wegener Halv~ Formation carbonates as reservoirs have been used for exploration both onshore Jameson Land and offshore mid-Norway. The source rocks of the Ravnefjeld Formation are limited to distal settings of transgressive and highstand system tracts and are most thickly developed in the central part of the Jameson Land basin and in structural lows, whereas the carbonate reservoir rocks were deposited along the basin margins or over structural highs. Renewed studies of the Upper Permian succession in northern Jameson Land, however, show that siliciclastic-dominated lowstand fans of the Bredehorn Member (Schuchert Dal Formation) are directly associated with potential source rocks in the proximal parts of the Late Permian basin forming an alternative stratigraphic play. The present study gives the first detailed description of the Bredehorn Member lowstand fans in northern Jameson Land with focus on their internal

composition, the sand distribution, and the sedimentation-controlling mechanisms. It is based on detailed sedimentological studies of eleven sections at four localities representing a transverse from a distal setting near the basin centre to the marginal parts of the basin (Figs. 1 and 2).

Geological setting In central East Greenland, the Late Permian transgression followed Carboniferous syn-rift sedimentation in narrow rotated half-grabens and a prolonged period of uplift and erosion in the Early Permian (Surlyk, 1990; Stemmerik et al., 1993a). Marine conditions were established as the Late Permian sea transgressed an elongate north-south-oriented depression approximately 400 km long and up to 80 km wide (Surlyk et al., 1986). The Late Permian depositional basin is separated in the west from the stable Greenland craton by the post-Devonian Main Fault (Fig. 1). The basin was landlocked towards the south, and towards the southeast it is separated by the Liverpool Land High from the rift basins that connected the Northern Permian Basin of Northwest Europe with the Permian basins of North Greenland and the Barents Sea (Stemmerik, 1995). Differential subsidence still occurred across major faults and influenced sedimentation during the Late Permian (Stemmerik et al., 1993a).

Sedimentary Environments Offshore Norway - Palaeozoic to Recent edited by O.J. Martinsen and T. Dreyer. NPF Special Publication 10, pp. 51-65, Published by Elsevier Science B.V., Amsterdam. 9 Norwegian Petroleum Society (NPF), 2001.

52

M. Kreiner-MOller and L. Stemmerik

of the Karstryggen Formation deposition was marked by a major fall in relative sea level and a prolonged period of subaerial exposure (Scholle et al., 1993). A second major Late Permian sea-level rise resulted in deposition of fully marine carbonates of the Wegener Halve Formation along basin margins and over structural highs and black shales of the Ravnefjeld Formation in the central parts of the basin (Fig. 3; Surlyk et al., 1986; Stemmerik et al., 1993a,b). Conodonts from these formations indicate a Kazanian age (Rasmussen et al., 1990) and most likely the Wegener Halve and Ravnefjeld Formations can be correlated with the Zechstein-1 a n d - 2 cycles of northwestern Europe (Stemmerik, 1995). The Wegener Halve and Ravnefjeld Formations correspond to three depositional sequences (Piasecki and Stemmerik, 1991; Stemmerik et al., 1992; Christiansen et al., 1993). Each sequence in the platform areas is characterised by a general shallowing-upwards facies trend and is capped by subaerial exposure surfaces (Scholle et al., 1991, 1993; Stemmerik, 1991; Stemmerik et al., 1993b). In the basin centre they are characterised by shifts between laminated organic-rich shales deposited during sea-level rise and highstand and bioturbated organic-lean shales deposited during sea-level lowstand (Piasecki and Stemmerik, 1991; Stemmerik et al., 1992). In more proximal basinal settings in northern Jameson Land, the bioturbated lowstand shales are seen to interfinger with more sand-rich sediments of the Bredehorn Member, thus indicating that the Bredehorn Member sandstones were deposited during sea-level lowstands (Stemmerik et al., 1997, 1998). The Bredehorn Member consists of up to three laterally extensive units, 2-50 m thick, of fine- to medium-grained sandstone separated by laminated and bioturbated shale (Fig. 3). The lower part of the sandy Bredehorn Member is regarded as time equivalent to the upper part of the Ravnefjeld Formation, whereas the upper part post-dates deposition of the Ravnefjeld Formation source rocks (Fig. 3) (Stemmerik et al., 1997). Bredehorn facies associations Fig. 1. Outline of the Late Permian sedimentary basin of central East Greenland showing Upper Permian outcrops and the proposed outline of the Bredehorn lowstand fans. Sediment input was from the west with an axial northwards flow more basinward. SBF = Store Blydal fan; BOF = Biskop Alf Gletscher-Oksedal fan.

The initial base level rise resulted in deposition of a fluvial conglomerate and red siltstone unit followed by a thick succession of shallow marine carbonates and evaporites of the Karstryggen Formation. Termination

The Bredehorn Member has been studied in eleven sections in Store Blydal, Blyryggen, Oksedal and Biskop Alf Gletscher in northern Jameson Land (Fig. 2). Ten facies have been recognised, including two shale facies, seven sandstone-dominated facies and one carbonate conglomerate facies (Table 1) (Kreiner-M011er, 1999). The sandstone-dominated facies and the carbonate conglomerates are gravity flow deposits and have been described using the classifications of Bouma (1962) and Lowe (1982). Description

Upper Permian lowstand fans of the Bredehorn Member, Schuchert Dal Formation, East Greenland

53

Fig. 2. Locality map of the studied area showing the position of the sections and the cross-sections in Figs. 9 and 12.

Fig. 3. Simplified cross-section of the western half of the Permian succession showing the stratigraphic relationships between the Bredehorn Member and the Ravnefjeld Formation. Modified from Stemmerik et al. (1998).

and interpretation of the facies are summarised in Table 1. The ten facies are grouped into six facies associations. Facies association A

This association is characterised by units 2-7 m thick, coarsening-upward from shale (facies 1), sand-

stone/mudstone beds (facies 3) to beds dominated by sandstone (facies 4). Occasionally the association is composed of sandstone/mudstone beds, where the sandstone/shale ratio increases upwards. The association is characterised by plane-parallel bedding and forms units that can be traced laterally for more than 3 kin, although individual beds can only be traced for shorter distances. Facies 3 is composed of beds up to

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M. Kreiner-MOller and L. Stemmerik

TABLE 1 Description and interpretation of facies from the Bredehorn fans and the interbedded highstand shales; sediments are deposited either by settling from suspension or by sediment gravity flows Facies

Description

Interpretation Settling from suspension in an anaerobic to anoxic environment. Fallout of suspension in an aerobic to dysaerobic environment. Deposited by low-density turbidity currents during lower flow regime conditions. Deposited by immature low-density turbidity currents with a little content of fine-grained sediments. Deposited by turbidity currents by fallout of suspension during successively lower flow regimes. Deposited by frictional freezing of a basal inertia-flow layer followed by gradual aggradation below a steady high-density turbidity current. Flow-sweeps driven thin inertia layers, which were deposited when the flow energy decreased; the turbidity currents only settle a minor part of the sediment load. Deposited by high-density turbidity currents; the currents were unsteady and fluctuated between erosion and deposition.

1.

Laminated shale

Clay-silt, black to dark grey, laminated.

2.

Silt, grey, bioturbated, occasionally laminated.

3.

Bioturbated siltstone Mudstone/sandstone

4.

Sandstone/mudstone

5.

Graded sandstone

6.

Massive sandstone

Fine-grained sandstone and silty shale, sandstone/shale ratio 1/5 to 1/1, sharp bases, Tcde divisions (Bouma, 1962). Fine-grained sandstone to silty shale, thickness 4-30 cm, sandstone/shale ratio 9/1 to 3/1, sharp occasionally erosional bases, Tbcde divisions (Bouma, 1962). Fine- to medium-grained sandstone and shale, thickness 20-100 cm, sandstone/shale ratio 4/1, normally graded, sharp erosional bases, flute marks occur, Tabcde (Bouma, 1962). Fine- to medium-grained, thickness 15-100 cm, massive, sharp erosional bases, flute marks occur, Ta (Bouma, 1962).

7.

Stratified sandstone

Fine- to medium-grained, thickness up to 4 cm, inversely graded $2 divisions (Lowe, 1982).

8.

Gravel and sandstone

Medium- to coarse-grained sandstone, lags and lenses of gravelstone, up to 3 m thick, amalgamation common, S 123 divisions (Lowe, 1982), erosional surfaces and scours occur, mud clasts common. Gravelstone to fine-grained sandstone, thickness 3-4 m, normally graded, sharp erosional bases, Tabcd divisions (Bouma, 1962), large "floating" clasts occur.

Graded gravelstone

10.

Carbonate conglomerate

Conglomerates with carbonate clasts and siliciclastic matrix, 10-100 cm thick, normally graded, sharp erosional bases, Tabcde divisions (Bouma, 1962). Carbonate clasts are only present in the Ta division, the upper divisions are siliciclastic.

30 cm thick, and has a sandstone/mudstone ratio in the range of 0.2 to 1. The individual beds have a sharp, occasionally erosional base and are normally graded with a lower sandy part showing parallel or climbing ripple lamination and an upper mudstone cap of laminated to bioturbated shale. Facies 4 is in many ways similar to facies 3 but has a higher sandstone/shale ratio. The lower sandy part often shows parallel lamination followed by climbing ripple lamination. The mudstone cap is occasionally missing. The sharp base, the normal grading, and the sedimentary structures suggest that facies 3 and 4 are deposited by turbidity currents (Bouma, 1962; Walker, 1965). Using the classification of Bouma (1962), facies 3 includes beds composed of the Tbde and Tcde divisions, whereas facies 4 beds can be classified as composed of the Tbc, Tbde, Tcde and Tbcde divisions. The lateral extent of the sandstone bodies, the finegrained nature, and the dominance of turbidites composed of upper Bouma divisions suggest that facies

Frictional freezing of inertia layer followed by fall-out of suspension during successively lower flow regimes; the clasts were trapped above the base because the entire inertia-flow phase of the high-density turbidity current froze. Deposited by frictional freezing of inertia layer followed by fall-out of suspension during successively lower flow regimes.

association A represents outer lobe deposition. The high sandstone/shale ratio and the limited lateral extent of the individual beds exclude facies association A as basin floor sheet-sand deposits. The associations are in many way similar to overbank deposits (e.g. Galloway and McGilvery, 1995) but, as discussed later, the overall basin configuration and the observed palaeoflow directions make it more likely that this association represents an outer lobe environment (see below). Facies association B

The association forms 7-18 m thick and more than 3 km wide lenticular units dominated by plane-parallel, medium- to fine-grained sandstone-mudstone beds, less than 30 cm thick (facies 3 and 4 in Table 1) (Fig. 4). Occasionally, up to 1 m thick carbonate conglomerates (facies 10 in Table 1) occur in the lowermost part of the association. Facies 3 and 4 resemble those described in facies associa-

Upper Permian lowstand fans of the Bredehorn Member, Schuchert Dal Formation, East Greenland

55

Fig. 4. Facies association B is a marginal middle-lobe facies association composed of fine-grained sandstones and silty shale (facies 3 and 4). The turbidites are characterised by upper Bouma Tbcde divisions. The association forms homogeneous units without evidence of coarsening- or thickening-upwards trends. Note overall planar bedding. Exposed section is approximately 3 m thick. Store Blydal.

tion A, although they are much more sand-rich with sandstone-mudstone ratios in the range of 9 to 3 (rarely as low as 0.2). Mutti (1992) has interpreted similar sandstone/mudstone beds to represent deposition from immature turbidity currents with a high sand content. The beds are internally less organised than classical Bouma sequences and are composed of the Tbcde, Tbde, Tcde, Tbc, Tb and Tc divisions. The associated carbonate conglomerates are composed of carbonate clasts in a siliciclastic matrix. The bases are sharp and erosional, and individual beds are normally graded with of a lower structureless conglomerate unit followed by a parallel and/or a climbing ripple-laminated sandstone unit and an upper mudstone cap. The sharp bases and the normal grading of the carbonate conglomerates indicate that facies 10 is de-

posited from turbidity currents. It can be classified as a Bouma sequence dominated by the conglomeratic Ta division and capped by siliciclastics of the upper Bouma divisions. Facies association B is interpreted to represent an outer lobe environment or the more proximal part of a relatively small lobe based on the dominance of beds formed of upper Bouma divisions, the fine-grained nature, and the bed thickness. The plane-parallel bedding, the non-channelisation, and the lenticular shape of the sandstone bodies support this interpretation.

Facies association C Facies association C is composed of up to 1 m thick, fine- to medium-grained sandstone beds (facies

56

M. Kreiner-Mr

and L. Stemmerik

Fig. 5. Facies association C consists mainly of 15 to 100 cm thick, massive or normal-graded sandstone turbidites (facies 5 and 6). The beds are organised into Bouma sequences with dominance of the Ta divisions. Note overall planar bedding. Exposed section is approximately 5 m thick. Oksedal.

5 and 6 in Table 1) (Fig. 5). The sandstones form 2 24-m-thick non-channelised bodies with great lateral extent. Individual beds are plane-parallel bedded with sharp erosional bases and common flute marks. The sandstones are massive or normally graded, occasionally with a thin cap showing parallel and/or climbing tipple lamination followed by mudstone. The sharp bases, the flute marks, and the sequences of sedimentary structures all suggest that facies association C is composed of turbidites. Individual beds can be classified as composed of the Ta, Tade, Tabcde, Tabde or Tacde divisions, respectively (sensu Bouma, 1962). The facies association is interpreted as lobe deposits based on the non-channelised and plane-parallel bedded nature of the sandstone bodies. Facies characteristics like the lack of scours, the uniform

grain size, the lack of pebbles and granules, and the dominance of Ta turbidites support this interpretation (e.g. Chen and Hiscott, 1999b). The association differs from association B by being more coarse-grained and dominated by turbidites composed of the Ta divisions. This indicates that association C either was deposited on a relatively bigger lobe than association B or represents a more proximal setting on the lobe. Facies association D The association is only present at B iskop Alf Gletscher where it reaches a thickness of 28 m. It consists of a lower coarse-grained unit of non-channelised pebblestones (facies 9 in Table 1) and an upper, sandy part of plane-parallel bedded massive

Upper Permian lowstand fans of the Bredehorn Member, Schuchert Dal Formation, East Greenland

57

Fig. 6. Facies association D shows an overall fining-upward trend from stacked gravelstones (facies 9) to medium-grained sandstones (facies 5 and 6) overlying black mudstones of the Ravnefjeld Formation (arrow). Exposed section is approximately 15 m thick. Biskop Alf Gletscher.

sandstones (facies 5 and 6 in Table 1) (Fig. 6). The pebblestone beds are up to 3 m thick and have an erosional base with scours. The beds are normally graded from pebblestone to fine-grained sandstone with a thick structureless lower unit composed of a massive lower part and a normal graded upper part. Up to 50-cm-long floating mudclasts occur in the lower massive part. The basal pebblestone is overlain by a sequence of parallel-, climbing-ripple-, and parallel-laminated sandstone that can be classified as the Tbcd divisions, suggesting that this facies represents deposition from a turbidity current. The lower pebblestone unit is accordingly classified as a R3 division following Lowe (1982). The upper part of the association is composed of graded or massive sandstone beds similar to those seen in facies association C. Facies association D is interpreted as representing channel mouth-inner lobe deposits. The lower coarse-

grained part o f the association has facies characteristics similar to channel deposits and the sandy upper part shows facies matching lobe deposits (e.g. Chen and Hiscott, 1999b). The fining-upward trend and the shift from channel to lobe deposits may represent an incision-backfill cycle deposited at the mouth of a channel (Chen and Hiscott, 1999a). Facies association E

Facies association E consists of 2-5-m-thick bodies, with a lateral extent of more than 3 km composed of stacked fine- to medium-grained sandstone beds or bands each up to 4 cm thick (facies 7 in Table 1). Small, 1-m-deep and a few metres wide channels occur locally and are filled with identical sandstone bands and lenses of ripple cross-bedded sandstone on the channel margins. Evidences of slumping are seen in the

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Fig. 7. Facies association E composed of spaced stratified sandstones, inversely graded and up to 4 cm thick (facies 7). Slumping is evident in the upper part. Hammer for scale. Store Blydal.

upper part of the association (Fig. 7). Individual sandstone bands are plane-bedded and massive or inversely graded with common mud clasts. They resemble stratification band deposits, and are accordingly interpreted as traction carpets believed to be deposited by pulsating turbidity currents (Hiscott, 1994). Facies association E is interpreted as a channel facies association, representing the feeding system of a small or marginal fan system. The dominance of traction carpet deposits is taken as evidence for a rather proximal setting since the currents transported most of the sediment load further out in the basin. The limited size of the channels and the overall finegrained nature of the association indicate deposition in a small marginal channel or in a small fan system.

Facies association F The association forms units, up to 26 m thick, of pebblestone and medium- to coarse-grained sandstone with erosional and sometimes scoured bases (facies 8 in Table 1) (Fig. 8). Individual beds are up to 3 m, but amalgamation is common. The bedding is irregular and channels up to 2 m deep and more than 30 m wide have been recorded. Facies 8 is characterised by a lower unit with traction structures, scours, and

small lenses or lags of pebblestone, followed by a unit composed of 3 to 15 cm inversely graded mediumto coarse-grained sandstone stratification bands, and finally a medium- to coarse-grained, massive or normally graded sandstone unit. Mud clasts are very common and concentrated on erosional surfaces. The three units are interpreted as part of turbidity current deposit units and can be classified as the S 1, $2 and $3 divisions according to Lowe (1982). However, most beds are not organised in ideal S 1-3 sequences and occasionally the channels are filled with massive or normally graded sandstone beds, which can be classified as composed of the Tabcde and Tade divisions (facies 5 and 6). The occurrence of channels indicates that facies association F represents the feeding system of a fan. Facies characteristics like strong amalgamation, basal scours, pebbles concentrated on the bed bases and the lack of upper Bouma divisions support this interpretation (Chen and Hiscott, 1999b).

Bredehorn turbidite systems Mutti and Normark (1987, 1991) established a hierarchical scheme to describe and analyse turbidite successions. Turbidite complexes refer to basin-fill

Upper Permian lowstand fans of the Bredehorn Member, Schuchert Dal Formation, East Greenland

59

Fig. 8. Facies association F composedof up to 4 m thick, coarse-grained sandstone locally with lags and lenses of gravelstone (facies 8) overlain by shales of the Ravnefjeld Formation(arrow). The exposed section is about 20 m high.

successions composed of several turbidite systems stacked one upon the other, where turbidite systems are bodies of genetically related gravity flow deposits commonly bounded by units of highstand mudstone (Mutti and Normark, 1987, 1991). The studied succession forms a turbidite complex composed of three laterally extensive sandstone units (Bredehorn Member), each regarded as a turbidite system in the terminology of Mutti and Normark (1987, 1991). The systems are named from the base and up: systems 1, 2, and 3 and the interbedded shale units form intersystems 1 and 2. The inter-system 1 shales belong to the upper laminated unit of the Ravnefjeld Formation and the inter-system 2 shales form the lower Oksedal Member of the Schuchert Dal Formation (Fig. 3) (Stemmerik et al., 1997, 1998).

Turbidite systems 1 and 2: description The thickness variations and facies distribution of the two lowermost turbidite systems are overall similar. Both systems are thickest, respectively 55 and 28 m, at B iskop Alf Gletscher in the southern part of the study area and pinch out north of Oksedal and Store Blydal (Fig. 9). The turbidite systems also become thinner south of B iskop Alf Gletscher and the

Bredehorn Member is missing in southern Schuchert Dal where the succession is composed entirely of shales (Surlyk et al., 1984; Piasecki and Stemmerik, 1991). In the Oksedal-Store Blydal area, the systems are thickest in eastern Oksedal downdip on the eastward tilted Oksedal fault block (Figs. 10 and 11). The systems become thinner westwards towards the crest of the fault block. The minimum thickness is at B lyryggen, and from there the systems increase in thickness westward in Store Blydal (Fig. 12). The coarsest deposits occur at the B iskop Alf Gletscher where turbidite systems 1 and 2 are dominated by sediments interpreted as belonging to the channel, inner-lobe, and lobe associations (C, D, and F). In Oksedal, the two turbidite systems consist of sediments belonging to the lobe association (C) and in Store Blydal they consist of sediments belonging to the lobe- and outer-lobe-basin floor facies associations (A, B) (Fig. 12). At Blyryggen, updip on the Oksedal fault block, the systems are composed of lobe- and outer-lobe-basin floor sediments (facies association A, B) (Fig. 12). The most significant difference in facies is related to the carbonate conglomerate facies of facies association B; it is restricted to Store Blydal and Blyryggen where it occurs both in turbidite system 1 and 2.

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Fig. 9. Sedimentological logs trough the Biskop Alf Gletscher-Oksedal fan from Biskop Alf Gletscher in the south to southern Traill 0 in the north. All three turbidite systems are present at Biskop Alf Gletscher and Oksedal but are missing at Traill 0, where the time-equivalent succession is composed of basinal shales. The cross-section spans approximately 60 km. Facies associations are indicated by letters (A, C, D and F) in the left column. For position of the section see Fig. 2. For legend see Fig. 12.

Fig. 10. E - W cross-section from Store Blydal to Oksedal showing the thickness variations of the Bredehorn Member. Deposition of the Store Blydal fan took place in the graben west of the Blyryggen High and the Biskop Alf Gletscher-Oksedal fan was confined to the eastern rotated half-graben.

Measurement of palaeoflow directions on flute marks and current ripples show transport directions toward the north-northeast in Oksedal and Store Blydal, and toward east at Blyryggen (Kreiner-MOller, 1999).

Turbidite system 3: description Turbidite system 3 displays the same overall thickness variations as the two underlying systems. It is the thickest of the three systems at Biskop Alf Gletscher where it attains a maximum thickness of 34 m. It is also thicker than the two underlying systems in

Upper Permian lowstand fans of the Bredehorn Member, Schuchert Dal Formation, East Greenland

61

Fig. 11. Line drawing of western Oksedal showing pre-Late Permian tilting and erosion of Carboniferous fluvial deposits. The basal Upper Permian Huledal Formation conglomerates are missing from the fault crest. The fault block topography also influenced sedimentation in the later parts of the Permian and subtle thickness variations indicate that tilting continued during the Permian. The fault block was also active during deposition of the overlying Triassic Wordie Creek Formation.

Oksedal and at Blyryggen but in Store Blydal it is the thinnest of the three turbidite systems (Fig. 12). The coarsest sediments occur at B iskop Alf Gletscher in the south, where the deposits are interpreted as belonging to the channel- and lobe associations (F and C) (Fig. 9). In Oksedal, turbidite system 3 also includes sediments belonging to the lobe- and channel associations with a dominance of lobe sediments. In Store Blydal, turbidite system 3 is composed entirely of outer channel deposits (facies association E), in contrast to the underlying systems where lobe deposits dominate (Fig. 12). At Blyryggen, sediments of the lobe association (B) dominate. Measured palaeoflow directions from turbidite system 3 show a pattern similar to that seen in the underlying systems with flow directions toward the north-northeast in Store Blydal and Oksedal and towards the east at Blyryggen (Kreiner-MOller, 1999).

Depositional evolution of the Bredehorn turbidite systems Thickness mapping of the Bredehorn Member turbidites and the observed distribution of facies associations in northern Jameson Land indicate that deposition took place in two fans separated by a structural high west of Blyryggen (Figs. 1 and 10). In Oksedal, deposition took place in an eastwardtilted half-graben as indicated by the gradual onlap of the Upper Permian sediments towards the crest of the fault block and the increasing thicknesses of individual units towards the east (Fig. 11). In the Store Blydal area, the turbidites thin towards the east suggesting the presence of a structural high west of

Blyryggen, and deposition of the Store Blydal fan in a separate graben (Figs. 9 and 10). The B iskop Alf Gletscher-Oksedal fan is regarded as the main fan extending for more than 30 km north-south (Fig. 1). The Bredehorn Member attains its greatest thickness, 127 m, at Biskop Alf Gletscher in the south (Fig. 12) where the facies associations representing the most proximal parts of the fan occur. The fan thins northwards to approximately 57 m in Oksedal and has disappeared before reaching southern Traill 0. The Store Blydal fan is regarded as a separate, much smaller marginal fan (Fig. 1). Compared to the B iskop Alf Gletscher-Oksedal fan, the turbidites are less organised, mainly composed of sandstone, and they are more immature and include coarse-grained carbonate conglomerates in the middle of turbidite system 1 and in turbidite system 2. Together with the eastwards thinning of the fan towards B lyryggen, these differences point towards deposition of the Store Blydal turbidites in a separate fan rather than in an isolated segment of the B iskop Alf GletscherOksedal fan. If the latter was the case, a better organisation and a lower sandstone/shale ratio should be expected and carbonate conglomerates should also be present in turbidite systems 1 and 2 elsewhere in the basin. On a regional scale, the turbidite systems form isolated bodies that basinwards interfinger with bioturbated shales. The interbedded inter-system shales form laterally extensive units that were deposited during sea-level rise and highstand at times when the fans were inactive (Piasecki and Stemmerik, 1991; Stemmerik et al., 1997, 1998).

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Fig. 12. Sedimentological logs trough the Store Blydal fan at Store Blydal and Blyryggen. The fan is clearly divided into three sandstone-dominated turbidite systems separated by transgressive and highstand shales. Note thinning towards the Blyryggen High (right). Cross-section spans approximately 9 km. Facies associations (A, B, E) are indicated in the left column. Position of the cross-section is shown in Fig. 2.

Turbidite systems I and 2 time

The two lowermost turbidite systems show similar spatial distributions of facies associations, thickness variations, and palaeoflow directions indicating minor changes of the fans during the first two lowstand events. Facies associations representing the more proximal parts of the B iskop Alf GletscherOksedal fan occur at B iskop Alf Gletscher where the two lowermost systems are composed of channel and lobe facies associations. In Oksedal the systems consist of sediments belonging to lobe facies association C. The palaeoflow directions show that in the B iskop Alf Gletscher-Oksedal fan sediment was transported northwards from B iskop Alf Gletscher towards Oksedal. In the Store Blydal fan, exposed at Store Blydal and Blyryggen the lower and middle turbidite systems consist of sediments belonging to the lobe facies association B and the outer lobe-basin plan association A. Flow directions towards the north and northeast dominate in the systems at Store Blydal indicating an overall configuration similar to that of the Biskop Alf Gletscher-Oksedal fan. The location at B lyryggen is

located east of the structural high that separated the two fans. The recorded transport direction towards the east at the B lyryggen and the dominant mudstone indicate that the turbidite systems at B lyryggen represent small overbank-lobes deposited by the biggest flows from the Store Blydal fan system. Turbidite system 3 time

The B iskop Alf Gletscher-Oksedal fan succession of turbidite system 3 is also thickest and coarsest in the south and disappears north of Oksedal indicating an overall similar configuration as turbidite systems 1 and 2. However, both at Biskop Alf Gletscher and in Oksedal, turbidite system 3 is composed of sediments representing a more proximal setting than seen in the underlying systems. At Biskop Alf Gletscher, the system is dominated by channel deposits (facies association F) and in Oksedal both channel and lobe deposits are present. The Store B lydal fan shows the same pattern of turbidite system 3 being composed of more proximal facies associations than the underlying systems. The fan is composed of sediments belonging to the outer channel facies association (E) in contrast

Upper Permian lowstand fans of the Bredehorn Member, Schuchert Dal Formation, East Greenland

to dominance of lobe deposits (facies associations A and B)earlier. At Blyryggen, the upper turbidite systems are dominated by lobe sediments like in the lower systems. The shift towards more proximal facies with time indicates progradation of both fan systems during turbidite system 3 time. However, flow directions are similar to those of the underlying systems both in the B iskop Alf Gletscher-Oksedal and the Store B lydal fan, indicating that the overall configuration of the fan systems remained unchanged. Discussion The sand distribution within the Late Permian basin of East Greenland reflects changes in relative sea level, sediment input and basin configuration as well as shifts in the climate (Stemmerik et al., 1998). Sandstones are not known from the lower, carbonate-dominated part of the Upper Permian succession. During deposition of the Kazanian Wegener Halv0 and Ravnefjeld formations, the basin was fringed by carbonate platforms during sea-level highstand and the central parts of the basin were sites of shale deposition (e.g. Stemmerik et al., 1992). During sea-level lowstands, the carbonate platforms became subaerially exposed and coarser-grained material was transported across the platforms in incised valleys (Scholle et al., 1993). However, coarse-grained deposits corresponding to sea-level lowstands are missing in most of Jameson Land, and only recently it became evident that the lower part of the Bredehorn Member at Biskop Alf Gletscher represents lowstand deposits of Wegener Halv0-Ravnef~jeld Formation age (Stemmerik et al., 1997). The present study shows that the inter-system shales, separating the lower and middle turbidite system, at all the studied localities are composed of black to dark-grey laminated shale (facies 1) of the Ravnefjeld Formation. Similarly, the shales between the middle and upper turbidite system have a very uniform composition: in this case grey, bioturbated siltstones (facies 2) of the Oksedal Member. The uniform lithology of the inter-system shales suggests that sedimentation took place simultaneously in the two fans; the lower turbidite system represents lowstand deposits of latest Wegener Halv~a-Ravnefjeld Formation age and the turbidite systems 2 and 3 are age-equivalent to the Oksedal Member. The facies and thickness distribution within the B iskop Alf Gletscher-Oksedal fan together with the overall basin configuration point towards a source for the fan near B iskop Alf Gletscher to the west of the post-Devonian main fault (Fig. 1). The fan is dominated by turbidites showing north- and north-

(33

eastwards transport directions, suggesting that the transversely infilled sand was distributed by axial, northwards-flowing turbidites. The turbidites primarily were focussed along structural depressions as evidenced by the eastward thickening of the turbidite systems in Oksedal. The outcrops of the Store Blydal fan mainly expose lobe facies suggesting that the more proximal parts of the fan were located to the south and southwest (Fig. 1). Transport directions from turbidites in Store Blydal again indicate deposition from axial, northward-flowing currents with some deviations in flow directions around the B lyryggen high where a small overbank-lobe was build by bigger flows. The situation with westerly source areas and axial northward flow existed throughout the Late Permian suggesting that the overall configuration of the basin did not change. Northwards tilting of the basin axis and northerly transport directions also characterised the Carboniferous and Lower Permian axial, fluvial systems (e.g. Surlyk et al., 1984), whereas the basin was dipping southwards during the Jurassic (Surlyk, 1991 ). The similarities between the two lower turbidite systems suggest minor changes in depositional and basinal conditions from lowstand 1 to lowstand 2. The more humid climate during lowstand (see Scholle et al., 1993) may have produced a higher sediment input caused by higher run-off from the hinterland. This increase in sediment input may have favoured formation of the fan systems. The upper turbidite system is dominated by more proximal deposits suggesting that the fans prograded further out into the basin at this stage. This shift in depositional patterns may be related to changes in basin configuration caused by the end of marginal carbonate-platform deposition. It is believed to be the result of a more permanent change towards a colder and more humid climate in the latest Permian and earliest Triassic, and the overlying Lower Triassic Wordie Creek Formation is also siliciclastic in composition. Deposition of the lower turbidite system took place during sea-level lowstand in latest Ravnefjeld Formation times and the sandstones are encased in laminated source-prone shales. The basinal equivalent to the Bredehorn sandstones are bioturbated, organiclean shales (cf. Piasecki and Stemmerik, 1991). The sandstones of turbidite system 2 represents the lowstand at earliest Schuchert Dal Formation time (Stemmerik et al., 1997), they directly overlie laminated organic-rich shales of the Ravnefjeld Formation and are overlain by organic-lean shales of the Oksedal Member. It means that the sandstones of turbidite systems 1 and 2 form interesting potential reservoirs directly associated with potential source rocks and sealed by shales.

M. Kreiner-MOller and L. Stemmerik

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The Bredehorn Member may form an interesting new analogue for Permian offshore mid-Norway since it is the first more extensive turbidite system of Permian age to be described from the rifted seaway between Greenland and Norway. It is an example on relatively small basin-margin-sourced turbidite systems formed during late stage of rifting (cf. Ravngts and Steel, 1998). The most important lesson to be learned from the Greenland outcrop example is that the underlying middle to Late Permian rift topography, although very subtle, has important control on sedimentation, and mapping of the fault blocks can be used to predict sand distribution. The initial rift topography was only modified slightly in the distal parts of the basin and acted as flow barriers during deposition of the Bredehorn lowstand fans. The lowstand fans are focussed in areas overlying the deeper axial parts of rotated fault blocks whereas the crestal areas are characterised by shales and only thin sands. The regional tilt of the Greenland Permian basin was towards the north, and by analogy, redistributed sand on the Norwegian shelf most likely has to be found north of their source area.

Conclusions (1) The studied succession in northern Jameson Land, East Greenland includes ten facies: two shale facies deposited by fallout of suspension, and eight sandstone-dominated turbidite facies deposited by low- and high-density turbidity currents. (2) Seven fan system facies associations are recognised in the studied interval. The facies associations are composed of channel- and channel-lobe transition zone to outer lobe associations. (3) Deposition took place in two separate fan systems, deposited in structural depressions separated by the Blyryggen High. The sediment input was from the west with a basin-axial N-S distribution. (4) The succession consists of three lowstand turbidite systems separated by transgressive and highstand shales. The facies distribution in the lower and middle turbidite system was generally the same suggesting a stationary position of the fans. The upper system consists of more proximal turbidites indicating that the fan prograded further basinward during this event. (5) The sand reached the proximal parts of the basin during sea-level lowstand when the marginal carbonate platforms were subaerially exposed. The higher sediment input was possibly linked to shortterm shifts in climate with a relatively more humid climate during sea-level lowstand and therefore expected higher sediment input (see Scholle et al., 1993).

(6) The fans are most thickly developed in the proximal parts of the basin in areas with poorly developed or missing carbonate platforms.

Acknowledgements This paper forms a contribution to the project "Resources of the Sedimentary Basins of North and East Greenland" supported by the Danish Research Council. MKM acknowledges supervision of Prof. F. Surlyk, Copenhagen. Published with approval of the Geological Survey of Denmark and Greenland.

References Bouma, A.H., 1962. Sedimentology of Some Flysch Deposits: a Graphic Approach to Facies Interpretation. Elsevier, Amsterdam, 168 pp. Chen, C. and Hiscott, R.N., 1999a. Statistical analysis of turbidite cycles in submarine fan successions: tests for short-term persistence. J. Sediment. Res., 69: 486-504. Chen, C. and Hiscott, R.N., 1999b. Statistical analysis of facies clustering in submarine-fan turbidite successions. J. Sediment. Res., 69:505-517. Christiansen, EG., Piasecki, S., Stemmerik, L. and Telnaes, N., 1993. Depositional environment and organic geochemistry of the Upper Permian Ravnefjeld Formation source rock in East Greenland. Am. Assoc. Pet. Geol. Bull., 77: 1519-1537. Galloway, W.E. and McGilvery, T.A., 1995. Facies of submarine canyon fill reservoir, Lower Wilcox Group (Paleocene), Central Texas coastal plain. In: R.D. Winn and J.M. Armentrout (Editors), Turbidites and Associated Deep-Water Facies. Soc. Econ. Paleontol. Mineral. Core Workshop, 20: 1-23. Hiscott, R.N., 1994. Traction-carpet stratification in turbidites fact or fiction? J. Sediment. Res., 64A: 204-208. Kreiner-Mr M., 1999. Sedimentologisk og hierarkisk elementanalyse af Schuchert Dal Formationen, Ovre Perm, OstgrCnland. Unpublished Cand. Scientific Thesis, University of Copenhagen, 180 pp. Lowe, D.R., 1982. Sediment gravity flows: II. Depositional models with special reference to the deposits of high-density turbidity currents. J. Sediment. Petrol., 52: 279-297. Mutti, E., 1992. Turbidite Sandstones. Agip, Instituto di Geologia Universith di Parma, 165 pp. Mutti, E. and Normark, W.R., 1987. Comparing examples of modern and ancient turbidite systems: Problems and concepts. In: J.K. Leggett and G.G. Zuffa (Editors), Marine Clastic Sedimentology. Graham and Trotman, London, pp. 1-38. Mutti, E. and Normark, W.R., 1991. An integrated approach to the study of turbidite systems. In: E Weimer and M.H. Link (Editors), Seismic Facies and Sedimentary Processes of Submarine Fans and Turbidite Systems. Springer, New York, pp. 75-106. Piasecki, S. and Stemmerik, L., 1991. Late Permian anoxia in central East Greenland. In: R.V. Tyson and T.H. Pearson (Editors), Modern and Ancient Continental Shelf Anoxia. Geol. Soc. London, Spec. Publ., 58: 275-290. Rasmussen, J.A., Piasecki, S., Stemmerik, L. and Stouge, S., 1990. Late Permian conodonts from central East Greenland. Neues Jahrb. Geol. Palaeontol. Abh., 178: 309-324. Ravn~s, R. and Steel, R.J., 1998. Architecture of marine rift-basin successions. Am. Assoc. Pet. Geol. Bull., 82:110-146. Scholle, EA., Stemmerik, L. and Ulmer, D.S., 1991. Diagenetic history and hydrocarbon potential of Upper Permian carbonate buildups, Wegener Halve area, Jameson Land Basin, East Greenland. Am. Assoc. Pet. Geol. Bull., 75: 701-725.

Upper Permian lowstand fans of the Bredehorn Member, Schuchert Dal Formation, East Greenland Scholle, P.A., Stemmerik, L., Ulmer-Scholle, D., Liegro, G.D. and Henk, EH., 1993. Paleokarst-influenced depositional and diagenetic patterns in Upper Permian carbonates and evaporites, Karstryggen area, central East Greenland. Sedimentology, 40: 895-918. Stemmerik, L., 1991. Reservoir evaluation of Upper Permian buildups in the Jameson Land basin, East Greenland. Rapp. Gr~nl. Geol. Unders., 149, 23 pp. Stemmerik, L., 1995. Permian history of the Norwegian-Greenland Sea Area. In: P.A. Scholle, T.M. Peryt and D.S. Ulmer-Scolle (Editors), The Permian of Northern Pangea, Vol. 2. Sedimentary Basins and Economic Resources. Springer, Berlin, pp. 98-118. Stemmerik, L., Surlyk, F., Scholle, EA. and Piasecki, S., 1992. Sequence stratigraphy of a carbonate-evaporite-siliciclastic basin, Upper Permian of East Greenland [abs.]. In: SEPM/IAS Research Conference on Carbonate Stratigraphic Sequences, Tremp, Spain, Abstract Volume and Conference Program, p. 97. Stemmerik, L., Christiansen, F.G., Piasecki, S., Jordt, B., Marcussen, C. and N~hr-Hansen, H., 1993a. Depositional history and petroleum geology of Carboniferous to Cretaceous sediments in the northern part of East Greenland. In: T.O. Vorren, E. Bergsager, O.A. Dahl-Stammes, E. Holter, B. Johansen, E. Lie and T.B. Lund (Editors), Arctic Geology and Petroleum Potential. Norwegian Petroleum Society (NPF), Special Publication, 2: 67-87. Stemmerik, L., Scholle, EA., Henk, F.H., Di Liegro, G. and Ulmer, D.S., 1993b. Sedimentology and diagenesis of the Upper Permian Wegener Halv~ Formation carbonates along the margins of the Jameson Land Basin, East Greenland. In: In: T.O. Vorren, E. Bergsager, O.A. Dahl-Stammes, E. Holter, B. Johansen, E. Lie and T.B. Lund (Editors), Arctic Geology and Petroleum Potential. Norwegian Petroleum Society (NPF), Special Publication, 2: 107119.

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Stemmerik, L., Clausen, O.R., Korstg~rd, J., Larsen, M., Piasecki, S., Seidler, L., Surlyk, F. and Therkelsen, J., 1997. Petroleum geological investigations in East Greenland: Project "Resources of the sedimentary basins of North and East Greenland". Geol. Greenl. Surv. Bull., 176: 29-38. Stemmerik, L., Dam, G., Noe-Nygaard, N., Piasecki, S. and Surlyk, F., 1998. Sequence stratigraphy of source and reservoir rocks in the Upper Permian and Jurassic of Jameson Land, East Greenland. Geol. Greenl. Surv. Bull., 180: 43-54. Surlyk, F., 1990. Timing, style and sedimentary evolution of late Palaeozoic-Mesozoic extensional basins of East Greenland. In: R.F.E Hardman and J. Brooks (Editors), Tectonic Events Responsible for Britain's Oil and Gas Reserves. Geol. Soc. London, Spec. Publ., 55: 107-125. Surlyk, F., 1991. Sequence stratigraphy of the Jurassic - - lowermost Cretaceous of East Greenland. Am. Assoc. Pet. Geol. Bull., 75: 1468-1488. Surlyk, F., Piasecki, S., Rolle, F., Stemmerik, L., Thomsen, E. and Wrang, E, 1984. The Permian Basin of East Greenland. In: A.M. Spencer, E. Holter, S.O. Johnsen, A. MCrk, E. Nysmther, E Songstad and A. Spinnanger (Editors), Petroleum Geology of the North European Margin. Graham and Trotman, London, pp. 303315. Surlyk, F., Hurst, J.M., Piasecki, S., Rolle, F., Scholle, P.A., Stemmerik, L. and Thomsen, E., 1986. The Permian of the western margin of the Greenland Sea - - a future exploration target. In: T.M. Halbouty (Editor), Future Petroleum Provinces of the World. Am. Assoc. Pet. Geol., Mere., 40: 629-659. Walker, R.G., 1965. The origin and significance of the internal sedimentary structures of turbidites. Proc. Yorkshire Geol. Soc., 35: 1-32.

Geological Institute, University of Copenhagen, ~ster Voldgade 10, DK-1350, Copenhagen K, Denmark Geological Survey of Denmark and Greenland (GEUS), Thoravej 8, DK-2400, Copenhagen NV, Denmark; E-mail: [email protected]

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Sedimentology and palaeogeography of the Statfjord Formation (Rhaetian-Sinemurian), North Sea AIf Ryseth

Facies analysis of the Statfjord Formation (Rhaetian-Sinemurian) in the Tampen Spur, Horda Platform and Utsira High areas of the North Sea Viking Graben, shows that alluvial sediments on the Utsira High are finer grained and contain a higher proportion of non-pedogenic, subaqueously deposited sediments than time-equivalent deposits located further north. This tendency is accompanied by regional variations in the lithofacies composition of the Statfjord Formation, indicating that the Early Jurassic drainage system was dominated by a southerly dipping continental palaeoslope that eventually terminated in shallow marine environments to the south. Furthermore, data from basins in Denmark and offshore Britain show that the Late Triassic-Early Jurassic period in northwest Europe was characterized by marine incursions from the south, with gradual onlap of marine strata to the north. This regional pattern supports the interpretation of a southerly dipping palaeoslope in the Viking Graben during the Early Jurassic. The proposed palaeogeographic model links the dominantly continental Statfjord Formation to age-equivalent coastal and marine deposits in the southern North Sea, and implies that the contemporary shoreline was located up to 350 km southwards of proximal fluvial environments in the north Viking Graben.

Introduction

Sandstones and mudrocks of the Statfjord Formation (Rhaetian-Sinemurian) are widely distributed throughout the North Sea Viking Graben and on bordering basin margins and intrabasinal highs such as the Tampen Spur, Horda Platform and Utsira High (Fig. 1; see Vollset and Dor6, 1984; R~e and Steel, 1985; Nystuen et al., 1989; Ryseth and Ramm, 1996). Well data from the Norwegian sector show that formation thicknesses range from about 30 m at basin margins to more than 500 m in the central parts of the basin, due to differential subsidence (Steel and Ryseth, 1990; Ryseth and Ramm, 1996). However, the formation is absent in the South Viking Graben and on the southern parts of the Utsira High due to erosional truncation. Vollset and Dor6 (1984) defined three formal members of the Statfjord Formation in its type area of the Statfjord Field: Raude, Eirikson and Nansen (see Fig. 2 for lihostratigraphic nomenclature). The base of the Statfjord Formation is defined by the level of an abrupt increase in the overall sandstone content relative to the underlying Triassic deposits (see Steel and Ryseth, 1990), whereas the lithostratigraphic top is represented by an abrupt transition into marine mudrocks of the Dunlin Group. R~e and Steel (1985) documented a vertical change from calcrete-bearing red-beds in the lower part of the formation (Raude

Member) into coal-bearing grey-beds associated with coarser sandstones in the middle to upper part (Eirikson Member). This transition can be ascribed to progressively more humid climatic conditions during Rhaetian-Sinemurian time, and also to possible hinterland rejuvenation due to early Cimmerian tectonics. R~e and Steel (1985) interpreted the Raude and Eirikson members in terms of fluvial deposition on alluvial fans and fan deltas within coastal basins, and related the Nansen Member to a phase of transgressive marine deposition and reworking prior to the deposition of the marine mudrocks of the Dunlin Group. Subsequent studies have altered the stratigraphic definitions and the environmental interpretations, although the overall vertical transition from continental to transgressive, shallow marine deposition is maintained. Nystuen et al. (1989) and Steel (1993) correlated the Raude Member (Fig. 1) with the uppermost part of the Lunde Formation. More recent studies have pointed toward an alluvial depositional setting (rather than fan delta systems) for the main part of the Statfjord Formation (Nystuen et al., 1989; Steel and Ryseth, 1990; MacDonald and Halland, 1993; Ryseth and Ramm, 1996). Ryseth and Ramm (1996) interpreted the Statfjord Formation in terms of three main facies assemblages, reflecting deposition within fluvial channels, interfluvial floodplains and shallow marine environments. A cored section illustrating the stacking of facies assem-

Sedimentary Environments Offshore Norway - Palaeozoic to Recent edited by O.J. Martinsen and T. Dreyer. NPF Special Publication 10, pp. 67-85, Published by Elsevier Science B.V., Amsterdam. 9 Norwegian Petroleum Society (NPF), 2001.

130

Fig. 1. (a) Map showing the position of wells and structural features. (b) Main structural elements of the Permo-Triassic northern North Sea Basin (from Fa~rseth, 1996). Note the prominent N-S orientation of the master faults and the easterly position of the main depocentre.

t...,

Sedimentology and palaeogeography of the Statfjord Formation (Rhaetian-Sinemurian), North Sea

Fig. 2. Late Triassic and Early Jurassic lithostratigraphic nomenclature for the Viking Graben area (after Vollset and Dot6, 1984).

blages is shown in Fig. 3. The bulk of the formation comprises intercalated units of fluvial channel and interfluvial deposits (Assemblages 1 and 2), which are capped by marginal marine deposits (Assemblage 3) and finally marine mudrocks (Dunlin Group). Together, Assemblages 1 and 2 make up the Eirikson Member of the Statfjord Formation, whereas Assemblage 3 corresponds to the Nansen Member. Steel and Ryseth (1990) developed a stratigraphic framework based on chronostratigraphic megasequences for the Triassic and Early Jurassic in the North Sea area, and presented a series of palaeogeographic models for the continental successions. These reconstructions focus on a northward directed axial palaeoslope with strong transverse sediment supply from both Fennoscandia (east) and Scotland (west), into a marine basin to the north. There is, however, a growing set of evidence from the Tampen Spur of sediment transport from source areas located to the west and northwest. Mearns et al. (1989) compared Samarium/Neodymium (Sm/Nd) isotope data from the Statfjord Formation and potential hinterlands, and found that the Statfjord Formation in part was sourced from an ancient source terrain, corresponding to the Lewisian basement of western Scotland, particularly the Rona Ridge to the west of the Shetland Islands. This indicated a sediment source located to the northwest of the Tampen Spur. Mearns et al. (1989) also suggested that the ancient source was activated due to early Cimmerian tectonic activity, related to thermal expansion and uplift along the North Atlantic Rift system. Dalland et al. (1995) inferred a possible northerly source for part of the Statfjord Formation in the Gull-

69

faks Field (Fig. 1) from analyses of Sm/Nd isotopes and dip-meter log data. Morton et al. (1996) considered zircon age spectra from the Statfjord Formation in the Brent field (Fig. 1) and suggested that part of the formation could be sourced from a possible exposed extension of the Caledonide foldbelt located to the north of the Tampen Spur. Dor6 and Gage (1987) inferred an uplifted "Rockall Land" high during the Triassic, and Knott et al. (1993) suggested that parts of the present Norwegian Sea were uplifted landmasses during the Late Triassic, thus forming a potential source to the north of the contemporary Viking Graben. Recent studies of the Norwegian Sea have also indicated that parts of the MOre and VOring Basins were uplifted areas until the Middle Jurassic (Brekke et al., 1999). It is the aim of this paper to examine the Early Jurassic (Rhaetian-Sinemurian) palaeogeographic setting in the Viking Graben by a comparative facies analysis of the Statfjord Formation on the Tampen Spur, Horda Platform and Utsira High (Fig. 1). Furthermore, the palaeogeographic setting is related to changes in depositional environments in adjacent basins.

Triassic and Early Jurassic structural setting Following the Caledonian orogeny and Devonian shearing, continental rifting commenced in the Viking Graben area during the Late Permian to Early Triassic (e.g. Badley et al., 1988; Yielding et al., 1992; Fa~rseth, 1996). Middle Triassic to Middle Jurassic successions are thus generally assigned to thermally driven, post-rift subsidence (Badley et al., 1988; Steel and Ryseth, 1990; Steel, 1993). Sedimentation throughout the Triassic and earliest Jurassic took place in an intracontinental setting, providing a succession of approximately 2-5 km of alluvial post-rift deposits, above rotated fault-blocks and clastic wedges of the Permo-Triassic syn-rift phase (e.g. Steel and Ryseth, 1990; Fa~rseth, 1996). The continuing alluvial post-rift sedimentation most likely produced a broad, saucer-shaped Late Triassic basin (see Yielding et al., 1992). Fa~rseth (1996) documented a predominant N-S orientation of the Permo-Triassic basin boundaries and intrabasinal faults, and noted that the TriassicEarly Jurassic basin formed a 130-150-kin-wide depression bounded to the east by the Oygarden Fault Complex, and to the west by major faults bounding the East Shetland Platform (Fig. 1). The main depocentre lay to the east, along the Horda Platform and the Asta Graben to the south (Fig. 1). Fa~rseth (1996) also demonstrated that the Triassic basin was segmented into three structural domains by two SWNE-oriented structural lineaments, representing the

A. Ryseth

70

Open marine Assemblage 3 Shallow marine Assemblage 1 Stacked fluvial channels Assemblage 2 Assemblage 1 Assemblage 2

Coal-bearing mudrocks and crevasse splay deposits

Assemblage 1

Assemblage 2

Pedogenic to lacustrine mudrocks and thin-bedded crevasse splay and channel sandstones

Assemblage 1

Stacked fluvial channels

Assemblaae 2 w

Assemblage 1 Stacked fluvial channels

Alluvial red beds

Sedimentology and palaeogeography of the Statfjord Formation (Rhaetian-Sinemurian), North Sea offshore extensions of the Nordfjord-Sogn detachment and the Hardangerfjord shear zone (Fig. 1). The compartmentalization of the Triassic northern North Sea Basin indicates that sub-basins with separate fluvial drainage systems may have existed. Whereas the structural difference between the three compartments is clear at the Early Triassic syn-rift level (see details in Faerseth, 1996), the several kmthick accumulations of alluvial post-rift deposits in all three compartments suggest that high sedimentation rates would remove the effects of potential structural barriers to produce a broad regional basin, as suggested by Yielding et al. (1992). Analysis of the effects of differential subsidence during deposition of the Statfjord Formation (Ryseth and Ramm, 1996) also shows that sedimentation kept pace with the contemporary subsidence. This is illustrated for the Utsira High in Fig. 4. Across this intrabasinal high, the Statfjord Formation is about 200 m thick, comprising laterally and vertically amalgamated fluvial sandstones alternating with overbank deposits. The increase in the formation's thickness to about 600 m in well 30/11-4 in the Viking Graben is not reflected in spatial changes in the depositional environment. Only a slight decrease in the stacking density of fluvial sandstones can be inferred. Very similar trends have been documented from the Tampen Spur and Horda Platform areas (Ryseth and Ramm, 1996). Thus, an alluvial depositional environment was maintained throughout the Hettangian and Sinemurian stages despite significant local variations in subsidence rate. The effect of structural compartmentalization and development of sub-basins on the fluvial drainage system during deposition of the Statfjord Formation is therefore considered to be minimal. The actual nature of the drainage system, consequently, has to be derived from sedimentary facies analysis.

Data base and methodology This study is based on sedimentological description and interpretation of about 1700 m of slabbed core in 23 wells from the Tampen Spur, Horda Platform and Utsira High (Fig. 1). These three areas, each with a reasonable spread of wells, are believed to be regionally representative for the Statfjord Formation. Core data are described and plotted on 1:50 and 1:500 scales, using the Wentworth grain size system and conventional classification of lithologies and sed-

71

imentary structures. The channel and floodplain assemblages are coded into ten primary lithofacies types (Fig. 5) in accordance with the facies scheme presented by Miall (1992), and related to depositional processes (see Table 1 for a summary of facies codes and processes). The discussion of palaeoslope will, in particular, be related to regional variations in the composition of the fluvial channel and floodplain assemblages. Two parameters, reflecting the "coarseness" of the channel fill deposits and the "wetness" of the associated floodplain deposits are defined. "Coarseness" reflects the proportion of Assemblage-1 deposits that is of medium size sand or coarser, whereas "wetness" reflects the proportion of Assemblage 2 that is deposited in a subaqueous floodplain environment (i.e. lakes and ponds) and has remained undisturbed by pedogenic processes. A division is made between the "upper" and "lower" parts of the formation. The majority of cores from the Horda Platform area are cut within the upper half of the formation, hence the sub-division of fully cored successions has been carried out to secure a stratigraphically adequate comparison between the three areas. The lower parts are generally characterized by more abundant red-beds and reduced sandstone content as compared to the upper parts.

Biostratigraphic resolution and dating Despite limitations due to poor preservation, a number of useful palynological and micropalaeontological observations have been made in the Late Triassic and Early Jurassic succession. The floral assemblage seen in the uppermost Lunde Formation is characterized by the taxa Granuloperculatipollis rudis, Riccisporites tuberculatus and Ovalipollis oval& which together indicate a Rhaetian age for this part of the succession (Eide, 1989). The lithological base of the Statfjord Formation occurs below the first downhole appearance of Riccisporites tuberculatus, hence a Rhaetian age can probably be assigned to its lowermost part. The Triassic/Jurassic transition in the three areas is apparently not associated with major stratigraphic breaks, as is also indicated by other regional studies (e.g. Lervik et al., 1989; Nystuen et al., 1989; Steel and Ryseth, 1990; Steel, 1993). The coexistence of Aratrisporites minimus and Trachysporites fuscus generally indicates an Hettangian age, whereas consistent Cerebropollenites spp.

Fig. 3. Stratigraphy, gamma-ray log response, lithology, facies assemblages and depositional environments in the Statfjord Formation, well 25/4-1 (Utsira High). The main part of the formation (Eirikson Member) comprises intercalated fluvial sandstones (facies Assemblage 1) and floodplain deposits (facies Assemblage 2), including lake and crevasse deposits, and palaeosols. Marginal marine deposits (Nansen Member) cap the formationat the transition to the overlyingmarine mudrocksof the Dunlin Group. For legend, see Fig. 5.

S

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Fig. 4. Correlation panel, Utsira High area, illustrating the architecture of the Statfjord Formation. The Statfjord Formation rests on generally fine-grained Triassic deposits 9 The alluvial part (Statfjord Formation) comprises multistorey/multilateral sandstone sheets and occasional, isolated fluvial sandstones and associated overbank deposits. The increase in the formation thickness to about 600 m in well 30/11-4 is not reflected in changes in the depositional environment.

o~

Sedimentology and palaeogeography of the Statfjord Formation (Rhaetian-Sinemurian), North Sea Core graph

Facies Code Gm Se Shc Sm Sx Shf Sr

FI

P

C

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II 20m

m II

15m

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'clay' ~i ~ ~ r~ c vcc;I Bioturbation

Current ripples :~

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Low angle cross-stratification Planar lamination .

.

.

.

Nodules (siderite, calcite, pyrite) Plant litter

.

~:~.. 9

Massive sandstones Lag conglomerates, clasts Coal

Fig. 5. Lithofacies composition of the two facies assemblages of the Eirikson Member. The fluvial channel sandstones (facies Assemblage 1) comprises lithofacies Gin, Se, Shc, Sin, St, and Shf, and minor proportions of lithofacies F1. The floodplain deposits (facies Assemblage 2) is characterized by lithofacies F1, P and C, with variable proportions of the sandy lithofacies, particularly Sr and Shf.

and Corollina spp. in the upper part of the Statfjord Formation indicates a S inemurian age (Raunsgaard Pedersen and Lund, 1980). Scattered occurrences of the marine dinocyst Liasidium variabile in the uppermost part of the Statfjord Formation (Nansen Member) on the Utsira High and Horda Platform imply a late Sinemurian age for the first marine influence in the Viking Graben. The regional distribution of this dinocyst suggests that the lithostratigraphic Statfjord/Dunlin boundary is diachronous and younging to the north (see Partington et al., 1993; Charnock et al., 2001). An ostracod assemblage recorded from the lowermost Dunlin Group, with occurrences of Ogmoconcha amalthei and Ogmoconchella danica, gives an early Pliensbachian age for the establishment of marine shelf conditions at all three locations (Ainsworth et al., 1989). Lithofacies

characterization

and interpretation

As shown in Fig. 3, the main part of the Statfjord Formation comprises an interstratification of composite fluvial sandstone bodies (Assemblage 1), and

73

finer-grained, interfluvial overbank deposits (Assemblage 2). The described core material shows that the sandstone bodies assigned to Assemblage 1 are composed of lithofacies Gm, Se, Shc, Sm, Sx, Shf and Sr (Table 1). The finer-grained intervals (Assemblage 2) are characterized by lithofacies F1, P and C, with variable proportions of the sandy lithofacies Se, Sin, Sx, Shf and Sr. Some notes are to be added to the interpretation of lithofacies which are of particular relevance to the environmental interpretation. (1) Lithofacies Gm and Shc contain quartzitic and other lithic fragments derived from a metamorphic/crystalline basement (external source), and is thought to reflect deposition in a rather proximal position within the Statfjord depositional system. In comparison, lithofacies Se contains deformed mudrock fragments, coal pieces and calcite/siderite nodules that may have been eroded from the floodplain environment (internal source). This lithofacies is seen in all three areas, and has no particular relevance to the palaeogeographic interpretation. (2) Lithofacies Sx embraces all types of crossstratified sandstones seen in the Statfjord Formation. Units can be further grouped into sets and cosets of troughs and planar-tabular cross-beds. Both types results from deposition from tractional currents in a river system, but the differentiation between planar sets and troughs is traditionally important to the interpretation of channel morphology (e.g. Miall, 1992). Hence, solitary planar sets and occasional cosets of planar sets are related to the formation of transverse bars, sand waves and sand flats, whereas troughs are related to persistent migration of dunes in relatively deep channels. (3) Massive sandstones (lithofacies Sm) are common throughout the fluvial sandstones. This lithofacies is not included in current facies schemes for fluvial sandstones (Miali, 1992), although such deposits have been reported from both modern and ancient systems, and related to rather variable processes of formation, including post-depositional deformation by liquefaction (Coleman, 1969; Jones and Rust, 1983), and primary depositional processes involving rapid deposition from sand-choked waters (e.g. Kirk, 1983; Johnson, 1984). Massive sandstones in the Statfjord Formation are invariably interstratified with well-structured deposits, indicating a primary origin, and are also associated with water-escape structures. Hence, lithofacies Sm is related to abrupt sediment dumping and vertical accretion within channels, associated with immediate water-escape. This reflects a rather flashy discharge regime. Common vertical shifts into coarse-grained horizontally stratified sandstones are related to upper stage flow conditions following phases of rapid sediment dumping.

A. Ryseth

74 TABLE 1 Summary of lithofacies types identified in the alluvial part of the Statfjord Formation Lithofacies Description

Association a

Interpretation

Gm

Shc, Sm

Channel bedload, lags, gravel bars.

Sm, Shc, Sx Sm, Sx

Channel bedload, lags, scour fills.

Se Shc Sm

Sx

Shf

Massive to crudely stratified, extraclast conglomerate. Sharply based, fining upwards into sandy lithofacies. Massive to crudely stratified intraclast conglomerate. Sharply based, with siderite nodules, mudclasts and coal pieces. Medium- to coarse-grained, pebbly sandstone. Crude to well-defined horizontal lamination, sets up to 250 cm. Fine- to coarse-grained, massive/disorganized sandstone. Units up to several m thick, with water-escape structures, convolute lamination. Fine- to coarse-grained and pebbly sandstone. Sets/cosets of trough and planar cross-stratification. Sets commonly 10-40 cm, maximum 200 cm. Very fine- to fine-grained sandstone. Planar lamination.

Shc, Sx, Shf

Channel bedload/saltation load, upper stage plane beds. Channel saltation/suspension load. Rapid deposition associated with water escape.

Shc, Sm, Shf

Lower flow regime dune/bar accretion by tractional currents.

Sr, Sm, F1

Upper flow regime plane beds, channel suspension load, waning flow. Lower flow regime current ripples. Interbedding with Shf indicates rapid shifts between upper and lower stage, waning flow. Low-energy suspension fall-out in standing water (lakes, ponds), weak agitation during floods. Palaeosols, reflecting oxidized (red) and reducive (grey) soil conditions.

Sr

Unidirectional ripple lamination. Very fine- to fine-grained sandstone.

Shf, F1

F1

Laminated, dark grey mudrocks, occasionally bioturbated, with sandy streaks and lenses. Pyrite and siderite nodules.

P

Destratified mudrocks and sandstones. Red/brown, mottled to grey with carbonaceous rootlets. Calcrete (red-beds) and siderite/pyrite (grey-beds). Coal and carbonaceous mudrocks, units up to 75 cm, rootlets below.

C

F1

In-situ peat and swamps, protected from clastic input.

a This column summarizes common vertical lithofacies transitions.

(4) Red and mottled red/grey palaeosols are characteristic of the lower half of the Statfjord Formation, and are also associated with nodular calcrete, whereas grey palaeosol types containing carbonaceous rootlets occur in the upper half, in association with coal beds (lithofacies C). The red coloration and presence of nodular carbonate (immature soil carbonate build-up) point toward pedogenesis in a well-drained (oxidizing) setting with a general moisture deficit (Blodgett, 1988; Ettensohn et al., 1988). Preservation of carbonaceous rootlets and peat (coal) in the upper half of the Statfjord Formation, in contrast, require permanently stagnant, water-saturated conditions (McCabe, 1984; Besly and Fielding, 1989). (5) Subaqueously deposited units of lithofacies F1 and associated sandstones are interstratified with palaeosols/coal beds in the Statfjord Formation. The proportion of pedogenic accumulations to non-pedogenic, subaqueously deposited sediment is of importance to the palaeogeographic reconstruction, as bodies of standing water are probably more likely to form on the low-gradient distal slopes relative to the steeper slopes of more proximal fluvial areas.

Variability of the continental deposits Representative core sections from the three study areas are shown in Figs. 6-8. Comparison of the core logs shows that fluvial sandstones from the Tampen Spur and Horda Platform are characterized by more irregular and differentiated vertical grain size distribution than those from the Utsira High. The former sections show abundant scour surfaces, abrupt vertical shifts between fine and coarse deposits and common upper flow regime deposits, as opposed to a more even vertical alternation between massive and cross-stratified sandstones on the Utsira High. Furthermore, trough cross-stratification is the dominant type on the Utsira High, whereas planar sets and occasional cosets are more common on the Tampen Spur and Horda Platform areas. The mean "coarseness" values of channel deposits in the upper half of the Statfjord Formation on the Tampen Spur and Horda Platform are rather high, 0.66 and 0.53, respectively. The equivalent deposits on the Utsira High are finer grained, with an average "coarseness" of only 0.14. A similar trend can also

Fig. 6. Cored section from the Tampen Spur area (well 34/10-30). The fluvial sandstones are rather generally coarse grained throughout the Statfjord Formation, with strong vertical alteration in sediment calibre. The observed cross-stratification comprises both troughs and planar sets. For legend, see Fig. 5.

Sedimentology and palaeogeography of the Statfjord Formation (Rhaetian-Sinemurian), North Sea Dunlin Group

~ _ ~

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Shallow marine (mouth bar?) ~'U"

, Transgression .,__-__.._

Fluvial channel

Oo-o-oO.-..o.o..-...

Swamp Crevasse Lake Stacked fluvial channels

Stacked fluvial channels Lake swamp Fluvial channel Flood plain/lake Stacked fluvial channels Palaeosol/lake Stacked fluvial channels

Lake Palaeosol Lake Fluvial channel Floodplain Stacked fluvial channels Lake Palaeosol Crevasse

Stacked fluvial channels

Crevasse Lake J GR (API)

Clay

150

Si vf

f

m

c vc cgl

75

A. Ryseth

76

Dunlin Group

Open marine Transgression

U Lr

Shallow marine

(Mouth bar?)

Lr

Transgression

Stacked fluvial channels 9 ! Q O~--lo

(majorstacked complex)

9

Lake

Swamp Crevasse Swamp Levee/lake Stacked

fluvial channels

t~

L

~

~

~ 0

~

~ GR (API)

Clay

Si vf

f

m

c

vc cgl

150

Fig. 7. Cored section from the Horda Platform area (well 30/6-15). As on the Tampen Spur, rather coarse-grained sandstones dominate 9Also, the fluvial sandstones are characterized by trough and planar cross-stratification. For legend, see Fig. 5.

be seen in the lower half of the formation, where "coarseness" values of 0.56 and 0.65 characterize the fluvial sandstones on the Tampen Spur and Horda Platform, respectively. On the Utsira High, the corresponding value is 0.24. This variability in sediment calibre is also reflected by the complete absence of the lithofacies Gm in the core material from the Utsira High, where all conglomerates are intraformational (lithofacies Se). Fig. 9 shows the relative distribution of lithofacies in the fluvial channel deposits in the upper half of

the Statfjord Formation. Fluvial sandstones on the Tampen Spur and on the Horda Platform have rather similar compositions. In both areas, the massive sandstone facies is most common, with the relative proportions of the other lithofacies types showing similar distributions. In comparison, fluvial sandstones from the Utsira High area are also dominated by lithofacies Sm, but contain a significantly lower proportion of the high-energy lithofacies Shc. Furthermore, the proportion of cross-stratified sandstones and the finergrained lithofacies Shf and Sr have increased.

Sedimentology and palaeogeography of the Statfjord Formation (Rhaetian-Sinemurian), North Sea

Dunlin

Group

itllll~silil

77

Open marine

Shallow marine Transgression

Swamp Crevasse

Stacked fluvial channels

Crevasse Swamp

Stacked fluvial channels

Lake "~

Palaeosol Crevasse splay Crevasse channel

~,

Palaeosol Lake/pond

Stacked fluvial channels

Floodplain

~

Clay 0

GR API

Si

vf

f

m

c

vc cgl

150

Fig. 8. Cored section from the northern Utsira High area (well 25/2-13). The fluvial sandstones are generally finer grained, and troughs are the dominant type of cross-stratification. See also Fig. 5 for comparison. For legend, see Fig. 5.

A. Ryseth

78

wells from the Utsira High gave an average channel deposit proportion of approximately 51%. Hence, the Statfjord Formation is least sandstone-prone in the Utsira High area.

Palaeoslope and channel morphologies

Fig. 9. Averaged volumetric facies composition of Assemblage 1 sandstone bodies from the upper part of the Statfjord Formation. Note the similarities between the Tampen Spur and Horda Platform, and the low proportion of lithofacies Shc and the absence of lithofacies Gm in the Utsira High area. T = total thickness of sediment included in the analysis.

Inspection of Assemblage-2 deposits shows that the floodplain environment of the Statfjord Formation was significantly more wet on the Horda Platform and Utsira High (wetness of 0.78 and 0.74, respectively) than on the Tampen Spur to the north (0.57) throughout the deposition of the upper part of the formation. This is accompanied by much more common occurrences of coal (approximately 8% of the measured Assemblage-2 intervals) on the Utsira High than on the Horda Platform and Tampen Spur (approximately 4% and 1%, respectively). In the lower part of the formation, a "wetness" value of 0.78 is recorded for the Utsira High, whereas a lower value characterizes the Tampen Spur (0.39). Ryseth and Ramm (1996) presented data from more than 60 wells from the Tampen Spur and Horda Platform, from which average fluvial channel deposit proportions of 57% (Tampen Spur) and 62% (Horda Platform) were calculated. A similar analysis of 11

From the data presented in Fig. 9, it seems that the fluvial systems on the Tampen Spur and Horda Platform were rather similar in terms of discharge regime and competence to transport and deposit coarse material. A different type of fluvial regime is likely to have persisted on the Utsira High, where the channel fill deposits are finer grained, and also characterized by a different lithofacies composition (Fig. 9). It is inherent to any evolved river system that its slope will generally decrease in the proximal to distal direction. Studies of modern rivers also show that average sediment particle sizes tend to decrease in the downstream direction due to upstream deposition of coarse-grained material. Hence, the calibre of sediment that is being transported toward the contemporary shoreline is likely to decrease downstream, with a subsequent increase in the proportion of fines which are being transported and deposited. Similar grain size/slope relations have also been documented from ancient fluvial systems in a variety of tectonic and climatic settings. For instance, Masson and Rust (1990) concluded that temporarily decreasing basin slopes during deposition of the coal-bearing Sydney Mines Formation (Pennsylvanian, Nova Scotia) resulted in a vertical decrease in the ratio of channel sandstones to floodplain mudrocks. Graham (1983) noted that downstream changes in Devonian Old Red Sandstone in the Munster Basin are characterized by a decrease in the maximum and mean grain size of the fluvial sandstones, an increase in the proportion of mudrocks and a decrease in the proportion of channel deposits. In a later study of this area, MacCarthy (1990) described a series of proximal to distal transects, and demonstrated that fluvial sandstone bodies became spatially thinner and finer grained in the downslope direction, accompanied by an increase in the abundance and thickness proportion of interfluvial deposits (both pedogenic and lacustrine) over distances of some tens of kilometres. From these interpretations, it appears that the Statfjord Formation on the Utsira High area, being characterized by finer-grained channel deposits, lower proportion of channel fill and generally more subaqueously deposited floodplain material than the equivalent deposits on the Tampen Spur and Horda Platform, was deposited on the lowest average slope gradient. Further, comparison of "coarseness" and "wetness" values suggests that the highest average slope gradient existed

Sedimentology and palaeogeography of the Statfjord Formation (Rhaetian-Sinemurian), North Sea

on the Tampen Spur to the north, and that the Horda Platform represents a medial area. Changes in river planform/channel morphology, such as sinuosity and braiding, are also at least partially controlled by the continuous downstream loss of gradient, so that braided fluvial patterns are likely to pass downslope into meandering patterns before the river system terminates in distributary channels within deltaic/coastal plains (Brierley and Hickin, 1991). Many summaries of alluvial deposition emphasize downslope transitions from gravel-dominated alluvial fans to braided streams, to meandering types and finally deltaic distributaries (Miall, 1992). It is generally problematic to reconstruct river planforms from vertical lithofacies sequences (Bridge, 1985). However, from the interpretation that relatively high slope gradients may have persisted in the Tampen Spur and Horda Platform areas, relative to the Utsira High, it is tempting to interpret the channel deposits in terms of braided and meandering systems due to the variations in the average slope. For instance, the existence of external gravel beds (lithofacies Gm) and the relatively high proportion of planar cross-stratified sets on the Tampen Spur and Horda Platform tend to characterize both modern and ancient sandy braided systems (e.g. Cant and Walker, 1978; Rust, 1978; Miall, 1992), and would clearly support a braided stream interpretation of the Statfjord Formation in these areas. In comparison, the dominance of trough crossstratification on the Utsira High seems more compatible with a meandering fluvial pattern. Trough cross-stratification formed by dune migration requires rather deep channels and persistent tractional currents to form, and this lithofacies is apparently most common within ancient meandering stream deposits (Stewart, 1981; Mossop and Flach, 1983; Johnson, 1984). The association of migrating dunes with massive sandstones (Sm; Table 1) can possibly be related to events of channel-plugging by rapid deposition due to the flashy discharge regime that apparently characterizes the entire fluvial system of the Statfjord Formation. These interpretations of average slope gradients and channel planforms suggest that the average basinal palaeoslope decreased in the north to south direction during the Early Jurassic. The likely implication of these interpretations is that the main drainage of the Early Jurassic fluvial system was directed to the south. The interpretation of fluvial transport to the south is further supported by published dip-meter log data from the Gullfaks Field (Dalland et al., 1995). Other studies from the Oseberg area (Horda Platform) and the Utsira High have yielded similar results (Norsk Hydro unpublished data).

79

Observations from the shallow marine deposits Environmental interpretation of the shallow marine assemblage capping the Statfjord Formation gives additional information about the palaeoslope associated with the marine transgression following the fluvial phase of deposition. Representative core sections of this assemblage on the Tampen Spur and Horda Platform (Fig. 10) show a vertical alternation between fine-grained, wave- and current-rippled sandstones and silty mudrocks, and sharp-based units of coarsergrained, massive, cross-stratified and current-rippled sandstones. These deposits are related to the deposition of mouth bar/distributary channels within wave-agitated embayments. In comparison, sections from the southern area (e.g. well 30/11-4, Fig. 10) are seemingly more intensively bioturbated, and contain beds of low-angle, hummocky/swaley cross-stratification that can possibly indicate a marine shoreface environment affected by stronger waves.

Late Triassic and Early Jurassic setting: implications from adjacent areas The Early Jurassic of northwest Europe is a period of general eustatic sea level rise (e.g. Hallam, 1988), which eventually caused the link up of northern (Boreal) and southern (Tethyan) seas, possibly through the Viking Graben (Dord and Gage, 1987; Ziegler, 1990). Hence, the direction of coastal onlap and timing of marine incursions in adjacent basins may give indications on the position and behaviour of contemporary marine environments. Late Triassic to Early Jurassic deposits in Denmark (onshore and offshore) comprise both marine and continental deposits with well-documented interfingering relationships (Vinding, Gassum and Fjerritslev Formations, Fig. 11; see also Bertelsen, 1978; Michelsen, 1989; Nielsen et al., 1989; Dybkjaer, 1991). Marine invasion from the south established a shallow brackish environment during Late Triassic times (Vinding Formation; Bertelsen, 1978). During the Early Jurassic Hettangian and Sinemurian stages, marine shelf conditions (Fjerritslev Formation) persisted in basinal areas, whereas contemporary shallow marine to fluvial conditions (Gassum Formation), sourced from the Fennoscandian hinterland to the north, prevailed along basin margins (Michelsen, 1989; Nielsen et al., 1989; Dybkjaer, 1991). The Gassum/Fjerritslev boundary has a Late Triassic (Rhaetian) age in the central part of the basin, but becomes younger towards the northern basin margin where shelfal conditions were established during the late Sinemurian due to coastal onlap (see Michelsen, 1989; Dybkjaer, 1991).

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Sedimentology and palaeogeography of the Statfjord Formation (Rhaetian-Sinemurian), North Sea

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Fig. 11. Late Triassic and Early Jurassic lithostratigraphy of the Danish Sub-basin/Norwegian-Danish Basin, Moray Firth Basin and Hebrides Basin, compared to the Viking Graben area.

In the inner Moray Firth Basin Early Jurassic (Hettangian-Pliensbachian) strata, resting unconformably upon Triassic deposits, have been described from both onshore and offshore areas (Batten et al., 1986; Stephen et al., 1993). Onshore, Hettangian-Sinemurian deposits comprise proximal, alluvial fan deposits (Dunrobin Pier Conglomerate, Dunrobin Castle Member; Fig. 11) that can be correlated with offshore packages of lacustrine mudrocks, coals and fluvial/estuarine channel deposits (Varicoloured and White Members; Fig. 11) of a more distal fluvial to marginal marine character. These deposits are buried below offshore mudrocks of late Sinemurian to Pliensbachian age (Lady's Walk Shale). Stephen et al. (1993) noted that the Hettangian-Sinemurian deposits form a conformable progression from lacustrine to marginally marine fluvial/deltaic and eventually offshore environments that can be related to relative sea level rise. Stephen et al. (1993) also used biostratigraphic data to define a late Sinemurian transgressive surface separating the fluvially influenced deposits from overlying marine mudrocks, equivalent to the Dunlin Group (Charnock et al., 2001).

To the west of the British Isles, Late Triassic and Early Jurassic (Rhaetian-Pliensbachian) sediments include fluvial deposits (New Red Sandstone) of Rhaetian-Hettangian age which are conformably overlain by Rhaetian-Hettangian offshore mudrocks (Penanth Group, Blue Lias Formation; Fig. 11) and Hettangian-Sinemurian nearshore limestones and sandstones (Broadford Beds Formation; Fig. 11; see Morton et al., 1987; Morton, 1989). Offshore mudrocks of the Pabba Shale (late SinemurianPliensbachian) rest unconformably upon these deposits in the Hebrides Basin. Morton et al. (1987) noted a strong diachroneity in the Late Triassic to Early Jurassic marine transgression. Rhaetian deposits are marine inthe southern part of the Hebrides Basin, whereas the first marine influence is dated as Hettangian further north, thus indicating marine incursion from the south. This was also the case during the Sinemurian, as the base of the Pabba Shale Formation is younging to the north. The combined evidence for a possible northerly source area for the Statfjord Formation, and the Late Triassic marine incursion to the south (Danish Sub-

Fig. 10. Cored sections of the shallow marine Nansen Member. The deposits on the Tampen Spur and Horda Platform are related to the development of fluvial-dominated mouth-bar and distributary channels within wave-agitated embayments. Further to the south (well 30/11-4), the l~resence of low-angle cross-stratification (hummocky and swaley types?) is related to deposition in a shoreface environment. For legend, see Fig. 5.

82 basin and Hebrides Basin) followed by subsequent migration of marine environments to the north during the Hettangian and Sinemurian stages serve to demonstrate that the contemporary fluvial environment of the Statfjord Formation may have drained into coastal and marine settings to the south.

Synthesis and discussion The sedimentological interpretation of the Statfjord Formation presented above accords with previous studies from the Tampen Spur area, which focus on braided fluvial systems as the principal depositional environment for the sandstones (Nystuen et al., 1989; Ryseth and Ramm, 1996). Also, the similarity seen in the lithofacies composition and grain size distribution in sandstone bodies from the Tampen Spur and Horda Platform areas suggest that fluvial systems here were rather similar in terms of river discharge, competence and morphology, although floodplain deposits from the Horda Platform area contain significantly higher proportions of non-pedogenic material. Fluvial sandstones from the Utsira High are finer grained, and the associated floodplain deposits contain higher proportions of non-pedogenic deposits. These observations suggest that the highest depositional gradients existed on the Tampen Spur to the north, and the lowest on the Utsira High to the south. The lithofacies composition of fluvial sandstones on the Utsira High is seemingly more compatible with meandering than braided systems. The palaeogeographic implication of the comparative facies study is that the main fluvial drainage system of the Statfjord Formation was directed to the south, with a strong transverse supply. Additional observations from the marine deposits capping the Statfjord Formation fit with a depositional model involving marine incursion from the south through the Viking Graben. Fully marine conditions were established in the Inner Moray Firth in the late S inemurian, and in the earliest Pliensbachian in the Viking Graben. This diachroneity is compatible with a gradual marine flooding from the south onto a southerly dipping palaeoslope existing throughout the Viking Graben and into the Norwegian-Danish Basin further south. By the palaeogeographic scheme presented in Fig. 12, Early Jurassic deposits of the northern and central North Sea can be linked to a basin-wide depositional system. However, the palaeogeographic model implies that Early and Middle Jurassic palaeoslopes in the Viking Graben were oppositely directed, as the evidence for northward progradation of the Middle Jurassic Brent delta system is indisputable (Graue et al., 1987; Helland-Hansen et al., 1992; Johannessen et al., 1995).

A. Ryseth

The postulated late-Early Jurassic reorganization of drainage patterns can be explained due to the rise of a thermal dome at the structural junction between the Moray Firth Basin, Viking Graben and Central Graben. This feature has clearly influenced the drainage pattern of the Middle Jurassic Brent delta, but its actual effect during the Early Jurassic is much more uncertain. Underhill and Partington (1993) found that the initial rise of the dome and the first phase of related shallowing occurred during the late Toarcian, or approximately 16 million years after the Pliensbachian transgression and termination of alluvial deposition in the Statfjord Formation. Probably, the dome did not exist during the Hettangian and Sinemurian stages. The southerly drainage direction may also apply for the older, Triassic alluvial system of the North Sea. Steel and Ryseth (1990) indicated that Triassic alluvial deposits of the Skagerrak Formation (Anisian-Norian) in the southern part of the Viking Graben/Horda Platform and NorwegianDanish Basin are replaced by fine-grained lacustrine deposits further to the south. Goldsmith et al. (1995) also indicated that continental Triassic deposits (Scythian-Norian) in the Central Graben correlates with marginal marine succession further to the south. These stratigraphic relationships are indicative of a southerly directed Triassic fluvial drainage system in the Viking Graben, prior to the deposition of the Statfjord Formation.

Conclusions Comparison of lithofacies compositions of the continental deposits of the Statfjord Formation suggests that the steepest depositional slope gradients possibly existed on the Tampen Spur to the north, and the lowest on the Utsira High to the south, with the Horda Platform representing an area of medial slope gradient. The palaeogeographic implication of these interpretations is that a southerly dipping palaeoslope existed in the Viking Graben during deposition of the Statfjord Formation, and that the continental environment of the Statfjord Formation was terminated by a marine inundation from the south during the late Sinemurian. By the earliest Pliensbachian, fully marine conditions were established in the Viking Graben. Data from the Danish Sub-basin, Norwegian-Danish Basin and the Moray Firth Basin, and from the Hebrides-West Shetland Basin to the west, confirms that Rhaetian-Sinemurian marine sediments show a successive younging to the north, thus supporting the interpretation of southerly dipping palaeoslopes at these stages. The palaeogeography may have been controlled by uplift to the north of the Viking Graben during the

Sedimentology and palaeogeography of the Statfjord Formation (Rhaetian-Sinemurian), North Sea

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83

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- Utsira High

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- Viking Graben

Fig. 12. Palaeogeographic map illustrating the proposed Early Jurassic (Hettangian-Sinemurian) sediment transport pattern. The Viking Graben may have acted as the alluvial conduit for clastic sediment shed off the Fennoscandian hinterland and source areas to the west and northwest of the Tampen Spur, linking up with shallow marine environments to the south (map based on Dord and Gage, 1987; Morton et al., 1987; Mearns et al., 1989; Ziegler, 1990).

Early Jurassic due to thermal expansion along the North Atlantic Rift system. The proposed palaeogeographic model implies that a reorganization of the drainage pattern in the Viking Graben occurred during the late-Early Jurassic, preceding the northward progradation of the Middle Jurassic Brent Group. This reorganization can be related to the rise of a thermal dome to the south of the Viking Graben during the late Toarcian.

The proposed palaeogeographic model links the Early Jurassic continental deposits of the Viking Graben to time-equivalent marginal marine sandstones (Gassum Formation) and offshore mudrocks (Fjerritslev Formation) in the Norwegian-Danish Basin to the south. It also indicates that the contemporary shoreline of the Early Jurassic fluvial drainage system was located about 300-350 km to the south of the Tampen Spur and Horda Platform areas. A further

84

implication of this is that the main area separating southerly and northerly drainage provinces lay not as far south as previously believed, but rather to the north, in the North Atlantic rift of the Norwegian Sea.

Acknowledgements The present paper is based on the author's Dr. Scient. dissertation at the Bergen University. Norsk Hydro ASA is thanked for financial support. I also wish to thank Ron Steel and Wojtek Nemec for their supervision and encouragement throughout the thesis work. Reviewers John Collinson and Ragnar Knarud provided many useful comments to the original manuscript. Jan Andsbjerg kindly supplied me with publications from the Geological Survey of Denmark and Greenland (GEUS).

References Ainsworth, N.R., O'Neill, M. and Rutherford, M.M., 1989. Jurassic and upper Triassic biostratigraphy of the North Celtic Sea and Fastnet Basins. In: D.J. Batten and M.C. Keen (Editors), Northwest European Micropalaeontology and Palynology. British Micropalaeontological Society Series, pp. 1-44. Badley, M.E., Price, J.D., Rambech Dahl, C. and Agdestein, T., 1988. The structural evolution of the northern Viking Graben, and its bearing upon extensional modes of basin formation. J. Geol. Soc., London, 145: 455-472. Batten, D.J., Trewin, N.H. and Tudhope, A.W., 1986. The TriassicJurassic junction at Golspie, inner Moray Firth Basin. Scott. J. Geol., 22: 85-98. Bertelsen, F., 1978. The Upper Triassic-Lower Jurassic Vinding and Gassum Formations of the Norwegian-Danish Basin. Geol. Surv. Denm. Ser., B3: 1-26. Besly, B.M. and Fielding, C.R., 1989. Palaeosols in Westphalian coal-bearing and red-bed sequences, central and northern England. Palaeogeogr., Palaeoclimatol., Palaeoecol., 70: 303-330. Blodgett, R.H., 1988. Calcareous paleosols in the Triassic Dolores Formation, southwestern Colorado. In: J. Reinhardt and W.R. Sigleo (Editors), Paleosols and Weathering Through Geologic Time: Principles and Applications. Geol. Soc. Am. Spec. Pap., 216: 103-121. Brekke, H., Dahlgren, S., Nyland, B. and Magnus, C., 1999. The prospectivity of the Vcring and MOre basins on the Norwegian continental margin. In: A.J. Fleet and S.A.R. Boldy (Editors), Petroleum Geology of Northwest Europe: Proceedings of the 5th Conference. Geological Society, London, pp. 117-131. Bridge, J.S., 1985. Paleochannel patterns inferred from alluvial deposits: a critical evaluation. J. Sediment. Petrol., 55: 579-589. Brierley, G.J. and Hickin, E.J., 1991. Channel planform as a non-controlling factor in fluvial sedimentology: the case of the Squamish River floodplain, British Columbia. Sediment. Geol., 75: 67-83. Cant, D.J. and Walker, R.G., 1978. Fluvial processes and facies sequences in the sandy braided South Saskatchewan River, Canada. Sedimentology, 25: 625-648. Charnock, M.A., Kristiansen, I.L., Ryseth, A. and Fenton, LEG., 2001. Sequence stratigraphy of the Lower Jurassic Dunlin Group, northern North Sea. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments of the Norwegian Continental Shelf: Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 145-174 (this volume).

A. Ryseth Coleman, J.M., 1969. Brahmaputra River: channel processes and sedimentation. Sediment. Geol., 3: 129-239. Dalland, A., Mearns, E.W. and McBride, J.J., 1995. The application of Samarium-Neodymium provenance ages to correlation of biostratigraphically barren data: a case study of the Statfjord Formation in the Gullfaks oilfield, Norwegian North Sea. In: R.E. Dunay and E.A. Hailwood (Editors), Non-biostratigraphical Methods of Dating and Correlation. Geol. Soc., London, Spec. Publ., 89:201-222. Dor6, A.G. and Gage, M.S., 1987. Crustal alignments and sedimentary domains in the evolution of the North Sea, Northeast Atlantic Margin and Barents Shelf. In: J. Brooks and K. Glennie (Editors), Petroleum Geology of North West Europe. Graham and Trotman, London, pp. 1131-1148. Dybkjam K., 1991. Palynological zonation and palynofacies investigation of the Fjerritslev Formation (Lower Jurassic-basal Middle Jurassic) in the Danish Sub-basin. Geol. Surv. Denm. Ser., A30: 1-150. Eide, F., 1989. Biostratigraphic correlation within the Triassic Lunde Formation in the Snorre area. In: J.D. Collinson (Editor), Correlation in Hydrocarbon Exploration. Norwegian Petroleum Society (NPF). Graham and Trotman, London, pp. 291-297. Ettensohn, F.R., Dever, G.R. Jr. and Grow, J.S., 1988. A palaeosol interpretation for profiles exhibiting subaerial exposure "crusts" from the Mississippian of the Appalachian Basin. In: J. Reinhardt and W.R. Sigleo (Editors), Paleosols and Weathering Through Geologic Time: Principles and Applications. Geol. Soc. Am., Spec. Pap., 216: 49-79. Fa~rseth, R.B., 1996. Interaction of Permo-Triassic and Jurassic extensional fault blocks during the development of the northern North Sea. J. Geol. Soc., London, 153:931-944. Goldsmith, RJ., Rich, B. and Standring, J., 1995. Triassic correlation and stratigraphy in the South Central Graben, UK North Sea. In: S.A.R. Boldy (Editor), Permian and Triassic Rifting in North West Europe. Geol. Soc., London, Spec. Publ., 91: 123-143. Graham, J.R., 1983. Analysis of the Upper Devonian Munster Basin, an example of a fluvial distributary system. In: J.D. Collinson and J. Lewin (Editors), Modern and Ancient Fluvial Systems. Int. Assoc. Sedimentol. Spec. Publ., 6: 473-483. Graue, E., Helland-Hansen, W., Johnsen, J.R., L~amo, L., Ncttvedt, A., Re~nning, K., Ryseth, A. and Steel, R., 1987. Advance and retreat of the Brent Delta system, Norwegian North Sea. In: J. Brooks and K. Glennie (Editors), Petroleum Geology of North West Europe. Geological Society, London, pp. 915-937. Hallam, A., 1988. A reevaluation of Jurassic eustasy in the light of new data and the revised Exxon curve. In: C.K. Wilgus, B.S. Hastings, C.G.St.C. Kendall, H.W. Posamentier, C.A. Ross and J.C. Van Wagoner (Editors), Sea-Level Changes: An Integrated Approach. Soc. Econ. Paleontol. Mineral., Spec. Publ., 42: 261273. Helland-Hansen, W., Ashton, M., L~amo, L. and Steel, R., 1992. Advance and retreat of the Brent delta: recent contributions to the depositional model. In: A.C. Morton, R.S. Haszeldine, M.R. Giles and S. Broom (Editors), Geology of the Brent Group. Geol. Soc., London, Spec. Publ., 61: 109-127. Johannessen, E.R, Mj~as, R., Renshaw, Dalland, A. and Jacobsen, T., 1995. Northern limit of the 'Brent delta' at the Tampen Spur - - a sequence stratigraphic approach for sandstone prediction. In: R.J. Steel, V. Felt, E.R Johannessen and C. Mathieu (Editors), Sequence Stratigraphy of the Northwest European Margin. Norw. Pet. Soc. Spec. Publ., 5: 213-256. Johnson, S.Y., 1984. Cyclic fluvial deposition in a rapidly subsiding basin, Northwest Washington. Sediment. Geol., 38: 361-391. Jones, B.G. and Rust, B.R., 1983. Massive sandstone facies in the Hawkesbury sandstone, a Triassic fluvial deposit near Sydney, Australia. J. Sediment. Petrol., 53: 1249-1259. Kirk, M., 1983. Bar development in a fluvial sandstone (Westphalian "A"), Scotland. Sedimentology, 30: 727-742. Knott, S.D., Burchell, M.T., Jolley, E.J. and Fraser, A.J., 1993. Mesozoic to Cenozoic plate reconstructions of the North Atlantic

Sedimentology and palaeogeography of the Statfjord Formation (Rhaetian-Sinemurian), North Sea and hydrocarbon plays of the Atlantic margin. In: J.R. Parker (Editor), Petroleum Geology of North West Europe: Proceedings of the 4th Conference. Geological Society, London, pp. 953-974. Lervik, K.S., Spencer, A.M. and Warrington, G., 1989. Outline of Triassic stratigraphy and structure in the central and northern North Sea. In: J.D. Collinson (Editor), Correlation in Hydrocarbon Exploration. Norwegian Petroleum Society (NPF). Graham and Trotman, London, pp. 173-189. MacCarthy, I.A.J., 1990. Alluvial sedimentation patterns in the Munster Basin, Ireland. Sedimentology, 37: 685-712. MacDonald, A. and Halland, E.K., 1993. Sedimentology and shale modelling of a sandstone-rich fluvial reservoir: upper Statfjord Formation, Statfjord Field, northern North Sea. Am. Assoc. Pet. Geol. Bull., 77: 1016-1040. Masson, A.G. and Rust, B.R., 1990. Alluvial plain sedimentation in the Pennsylvanian Sydney Mines Formation, eastern Sydney Basin, Nova Scotia. Bull. Can. Pet. Geol., 38: 89-105. McCabe, EJ., 1984. Depositional environments of coal and coalbearing strata. In: R.A. Rahmani and R.M. Flores (Editors), Sedimentology of Coal and Coal-bearing Sequences. Int. Assoc. Sedimentol. Spec. Publ., 7: 13-42. Mearns, E.W., Knarud, R., Ra~stad, N., Stanley, K.O. and Stockbridge, C.R, 1989. Samarium-Neodymium isotope stratigraphy of the Lunde and Statfjord Formations of Snorre Oil Field, northern North Sea. J. Geol. Soc., London, 146:217-228. Miall, A.D., 1992. Alluvial sediments. In: R.G. Walker and N.R James (Editors), Facies Models: Response to Sea Level Change. Geological Association of Canada, Toronto, pp. 119-142. Michelsen, O., 1989. Log-sequence analysis and environmental aspects of the Lower Jurassic Fjerritslev Formation in the Danish Sub-basin. Geol. Surv. Denm. Set., A25: 1-23. Morton, N., 1989. Jurassic sequence stratigraphy in the Hebrides Basin, NW Scotland. Mar. Pet. Geol., 6: 243-260. Morton, N., Smith, R.M., Golden, M. and James, A.V., 1987. Comparative stratigraphic study of Triassic-Jurassic sedimentation and basin evolution in the northern North Sea and northwest of the British Isles. In: J. Brooks and K. Glennie (Editors), Petroleum Geology of North West Europe. Geological Society, London, pp. 697-709. Morton, A.C., Claoue-Long, J. and Berge, C., 1996. SHRIMP constraints on sediment provenance and transport history in the Mesozoic Statfjord Formation, North Sea. J. Geol. Soc., London, 153: 915-929. Mossop, G.D. and Flach, RD., 1983. Deep channel sedimentation in the Lower Cretaceous McMurray Formation, Athabasca Oil Sands, Alberta. Sedimentology, 30: 493-509. Nielsen, L.H., Larsen, F. and Frandsen, N., 1989. Upper TriassicLower Jurassic tidal deposits of the Gassum Formation on SjaJland, Denmark. Geol. Surv. Denm. Set., A23: 1-30. Nystuen, J.E, Knarud, R., Jorde, K. and Stanley, K.O., 1989. Correlation of Triassic to lower Jurassic sequences, Snorre Field and adjacent areas, northern North Sea. In: J.D. Collinson (Editor),

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Correlation in Hydrocarbon Exploration. Norwegian Petroleum Society (NPF). Graham and Trotman, London, pp. 273-289. Partington, M.A., Copestake, E, Mitchener, B.C. and Underhill, J.R., 1993. Biostratigraphic calibration of genetic sequences in the Jurassic-lowermost Cretaceous (Hettangian-Ryazanian) of the North Sea and adjacent areas. In: J.R. Parker (Editor), Petroleum Geology of North West Europe: Proceedings of the 4th Conference. Geological Society, London, pp. 371-386. Raunsgaard Pedersen, K. and Lund, J.J., 1980. Palynology of the plant-bearing Rhaetian to Hettangian Kap Stewart Formation, Scoresby Sund, East Greenland. Rev. Palaeobot. Palynol., 31: 1-69.

R~e, S.-L. and Steel, R., 1985. Sedimentation, sea-level rise and tectonics at the Triassic-Jurassic boundary (Statfjord Formation), Tampen Spur, northern North Sea. J. Pet. Geol., 8: 163-186. Rust, B.R., 1978. Depositional models for braided alluvium. In: A.D. Miall (Editor), Fluvial Sedimentology. Can. Soc. Pet. Geol., Mem., 5: 605-625. Ryseth, A. and Ramm, M., 1996. Alluvial architecture and differential subsidence in the Statfjord Formation, North Sea: prediction of reservoir potential. Pet. Geosci., 2:271-287. Steel, R.J., 1993. Triassic-Jurassic megasequence stratigraphy in the Northern North Sea: rift to post-rift evolution. In: J.R. Parker (Editor), Petroleum Geology of North West Europe: Proceedings of the 4th Conference. Geological Society, London, pp. 299-315. Steel, R. and Ryseth, A., 1990. The Triassic-Early Jurassic succession in the northern North Sea: megasequence stratigraphy and intra-Triassic tectonics. In: R.F.E Hardman and J. Brooks (Editors), Tectonic Events Responsible for Britain's Oil and Gas Reserves. Geol. Soc., London, Spec. Publ., 55: 139-168. Stephen, EJ., Underhill, J.R., Partington, M.A. and Hedley, R.J., 1993. The genetic sequence stratigraphy of the Hettangian to Oxfordian succession, Inner Moray Firth. In: J.R. Parker (Editor), Petroleum Geology of North West Europe: Proceedings of the 4th Conference. Geological Society, London, pp. 485-505. Stewart, D.J., 1981. A meander-belt sandstone of the Lower Cretaceous of Southern England. Sedimentology, 28: 1-20. Underhill, J.R. and Partington, M.A., 1993. Jurassic thermal doming and deflation in the North Sea: Implication of the sequence stratigraphic evidence. In: J.R. Parker (Editor), Petroleum Geology of North West Europe: Proceedings of the 4th Conference. Geological Society, London, pp. 337-345. Vollset, J. and Dote, A.G., 1984. A revised Triassic and Jurassic lithostratigraphic nomenclature for the Norwegian North Sea. Norw. Pet. Direct. Bull., 3: 1-53. Yielding, G., Badley, M.E. and Roberts, G., 1992. The structural evolution of the Brent Province. In: A.C. Morton, R.S. Haszeldine, M.R. Giles and S. Brown (Editors), Geology of the Brent Group. Geol. Soc., London, Spec. Publ., 61: 27-43. Ziegler, EA., 1990. Geological Atlas of Western and Central Europe. 2nd ed., Shell Internationale Petroleum Maatschappij B.V., The Hague, 239 pp.

Norsk Hydro Research Centre, P.O. Box 7190, N-5020 Bergen, Norway Present address: Norsk Hydro Harstad, Storakern 11, Kanebogen, N-9401 Harstad, Norway

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Sedimentary facies in the fluvial-dominated Are Formation as seen in the Are 1 member in the Heidrun Field Knud Egil Svela

The ,~re 1 member in the Heidrun Field was deposited in a fluvial to deltaic setting where variations in accommodation space and sediment supply controlled the different sedimentary facies and reservoir properties. It shows an overall transgressive sequence from fluvial plain in the lower part to lower delta plain in the upper part. Fluvial channels and thick coal-bearing floodplain deposits dominate the lower part of the Are 1. Stacked (multi-storey) channels were deposited as a response to relative fall in sea level. These sands have large lateral continuity compared to the meandering (single-storey) channel sands that are more a function of autocyclic processes and are not correlatable between wells. The upper part of the Are 1 member is dominated by stacked bay/lake fill sequences and thin, but laterally continuous coals. These represent deposition in a lower delta plain setting where the dominant depositional agents were crevasse channels and splays (crevasse deltas) with subsequent wave reworking in places. The reservoir properties of the Are 1 member are strongly controlled by depositional processes and the sequence stratigraphic framework. Incised valley fill deposits (LST) have large lateral continuity and the internal reservoir properties are very good. The TST deposits are more dominated by thick, coal-bearing, floodplain deposits with no reservoir quality, thin crevasse splays with restricted lateral continuity and limited reservoir properties and single-storey channel sands with very good internal reservoir properties but limited lateral continuity. The individual parasequences in the stacked bay fill sequences (HST) have good lateral continuity but reservoir properties are limited, and the bases of individual bay fill sequences are vertical permeability barriers, at least on a local scale.

Introduction

The Heidrun Field is located offshore Mid-Norway (Fig. 1), and was discovered by Conoco in 1985. The field is now operated by Statoil and the current estimate of total STOIIP is 2733 mmbo, of which 465.5 mmbo is within the Are 1 member. Estimated total recoverable oil reserves are 1132 mmbo, of which 125.7 mmbo comes from the Are 1. This gives an estimated recovery factor of 27% from the Are 1 member. To date, there are no production data from the Are 1 member. The Are Formation overlies the Triassic Grey Beds and comprises a succession of sandstones, mudstones and coals of Rhaetian to early Pliensbachian age. In the Heidrun Field, the top Are Fro. (base Tilje Fro.) is defined by the first full marine flooding as seen from biostratigraphy and is clearly reflected by a gamma-ray peak on wireline logs. This definition is slightly different from the standard lithostratigraphic definition as described by Dalland et al. (1988), with the implication that the upper part of the Are Fm. in Heidrun is time-equivalent with the lower part of the Tilje Fm. elsewhere on the Halten Terrace. The Are Fro. has been informally subdivided into two members in the Heidrun Field; Are 1 (base) and

Are 2 (top) (Fig. 2). The top of coal-bearing strata approximately corresponds to the top of the Are 1 member. B iostratigraphic data show that the entire Are 1 member was deposited in a non-marine setting. The Are 1 member has a maximum observed vertical thickness of 486 m in well 6507/7-2, which is the only well in the Heidrun Field that has penetrated a complete Are Formation. Only two wells, 6507/7-6 and 6507/7-A38, have been cored in the ,~re 1 member in the Heidrun Field. Of these, well 6507/7-A-38 has the best coverage, and a summary of the core description is presented in Fig. 3. Detailed examples of the main facies associations are presented in Fig. 4. The main focus of this paper is a description and sedimentological interpretation of the Are 1 member as seen in the Heidrun Field. A sequence stratigraphic interpretation based on Exxon terminology (Van Wagoner et al., 1990) is also proposed. Due to the limited study area, it is difficult to conclude firmly on basinal effects on the sequence stratigraphic framework. Gjelberg et al. (1987) presented a regional interpretation of the Are Fro. (then the Hitra Fro.) and concluded that the Are Fro. had an overall transgressive nature. This study agrees with this interpretation of the overall vertical change in depositional environment, but a different interpretation of

Sedimentary Environments Offshore N o r w a y - Palaeozoic to Recent edited by O.J. Martinsen and T. Dreyer. NPF Special Publication 10, pp. 87-102, Published by Elsevier Science B.V., Amsterdam. 9 Norwegian Petroleum Society (NPF), 2001.

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Fig. 1. Heidrun Field location map. Hydrocarbon area, i.e. oil and gas distribution is composite for all reservoir formations and not specific for the Are Fm. (After Statoil).

the fluvial systems is presented. Finally, a short discussion of reservoir properties in the different facies associations is presented.

facies, can occur in more than one association (e.g. crevassing).

Stacked (multi-storey) fluvial channels Sedimentary facies associations In this study, individual lithofacies have been grouped together in facies associations that represent depositional environments. It is also attempted to define facies associations that represent reservoir (or seal) units that are of sufficient thickness that they can be recognized on wireline logs. This means that individual sedimentary processes, and thereby litho-

Description This facies association has an erosive lower boundary, and comprises thick units of dominantly cross-bedded, medium-grained sand (Figs. 4 and 5). The thicknesses of these units range from 7 to 18 m. Except for just above the erosive base, where beds with granules are seen in a few occasions (Fig. 5), there is little vertical variation in grain size. The

Sedimentary facies in the fluvial-dominated fire Formation as seen in the fire 1 member in the Heidrun Field

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Fig. 2. Wireline logs from well 6507/7-A-38 that show the boundary between the Tilje Fm. and Are Fm. as defined in the Heidrun Field.

lower parts of units often contain large coal clasts and smaller coal fragments are seen on foresets. The sands generally display sets of large-scale tabular cross-stratification, both tangential and angular. In places, small current ripples on toe-sets can be seen. Trough cross-bedding and current ripples are present in places, generally without any preferred vertical arrangement. Individual cross-sets are normally 20 to 50 cm thick, but are found up to 90 cm thick and occur in up to 4 m thick cosets. As far as it is possible to determine from a non-oriented core, palaeocurrent direction seems fairly constant. Set boundaries are normally erosional within cosets. The upper boundary to overlying rooted horizons and coal is commonly abrupt.

Interpretation The erosive lower boundaries, associated coal intraclasts and pebbles together with the internal vertical

arrangement of facies suggest that this facies association represents the in-channel or coarse member of fluvial channels. Tidal channels can also develop sequences which resemble those of fluvial channels (Oomkens, 1974; Barwis, 1978; Weimer et al., 1982). However, the lack of clay drapes and other tidal indicators make such an origin unlikely in the Are 1. The biostratigraphic data together with the overall setting with erosional surfaces into fluvial plain/delta plain associations and the abundance of coal intraclasts also suggest a fluvial origin. In this facies association, the dominance of largescale cross-bedding and the uniformity of grain size throughout the units make it likely that deposition took place by vertical accretion rather than lateral accretion (point bars) in meandering channels. The significant increase in grain size compared to the sands in the other facies shows a considerable increase in hydraulic regime. Together with the fact

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Fig. 3. Summary of core description and interpretation from well 6507/7-A-38.

that these channels can be correlated between wells, they have been interpreted to represent low-sinuosity, braided channels instead of ribbon-like, anastomosing channels (Collinson, 1986). The presence of coal intraclasts shows that part of the eroded topstratum included swamp deposits (peat). However, the fact that the channel sands often sit on top of coal layers also shows the resistive nature of peat to erosion. In exposures, like the Breathitt Group in Kentucky, channels are also often seen to erode down to coal layers without cutting through them. Because peat has such a strong resistive nature, it often controls the degree of incision such that erosion tends to take place laterally, a process that extends the lateral dimensions of these channel fill sequences. As it is difficult to observe a correlatable interfluvial sequence from the limited well data in the Heidrun Field, and because it has not been possible to see channel geometries on 3D seismic, it is difficult to use the term "incised valley fill". However, the indications of braided type deposition together with the correlatable nature of these sands in the Heidrun Field and the abrupt change from the underlying fine-grained floodplain deposits make it likely

that this facies association represents deposition following a relative fall in sea level. Such lowstand channel deposits generally have large lateral extent. Similar channels in the Breathitt Group in Kentucky, USA, are seen to have kilometres of lateral continuity (Aitken and Flint, 1995). This is also proposed to be the case for these channels in the Heidrun area, where some units are correlatable in all wells drilled to date (Fig. 6). Similar multi-storey fluvial channels have also been described in many studies of the Carboniferous in the UK and Germany. A summary of their character and recognition is presented by Hampson et al. (1999).

Meandering (single-storey) fluvial channels Description This facies association has an erosive lower boundary and comprises cross-bedded and current-rippled sand (Figs. 4 and 7). The thickness of these units ranges from 6.2 to 9.0 m. They display a clear fining-upward trend from medium- to fine-grained sand at the base, grading upwards into very fine sand and silt in the upper part (Fig. 4). This facies association has a gradual boundary with the overlying sediments,

Sedimentary facies in the fluvial-dominated Are Formation as seen in the Are 1 member in the Heidrun Field

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Fig. 4. Detailed core descriptions of the main facies associations in the Are 1 member from well 6507/7-A-38. The depths for these intervals within the Are Fm. succession are shown on Fig. 11.

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Fig. 5. Core photo from stacked (multi-storey) channels in 6507/7-A-38 showing: (A) erosional channel base; (B) granule size channel lag; (C) tabular cross-bedding; (D) fine-grained sand; (E) coal clasts. See Figs. 4 and 11 for location of photo.

which normally comprises rooted floodplain deposits. Internally, the sands display mostly trough cross-bedding. The upper parts of the sand bodies show current ripples and lenticular bedding. Coal intraclasts are abundant at the bases of units.

Interpretation The erosive lower boundary fining-upward trend and the internal vertical arrangement of facies suggest that this facies association represents deposition by lateral accretion in high-sinuosity, meandering fluvial channels. In contrast to the stacked (multi-storey) fluvial channels, this facies association shows a clear vertical decrease in grain size and change in sedimentary structures upwards to current ripples. This indicates decreasing current strength, something that is seen in meandering channels where lateral accretion develops point bars (Collinson, 1986). The average grain size is also finer than in the stacked (multistorey) fluvial channels. Diagnostic features, such as

lateral accretion surfaces, are however impossible to identify in core. The lateral extent of these channels is much more limited than the stacked (multi-storey) fluvial channels. From wireline logs, these single-storey channels are seldom seen to be correlatable between wells. Similar channels in the Breathitt Group in Kentucky, USA, are seen to have a wide range in lateral continuity, some of them down to 10s of metres (Fig. 8). Heterolithic channel fills with very limited lateral extent, as frequently seen in the Breathitt Group, are however not recognized in any of the cores from the Are 1 member. Gjelberg et al. (1987) interpreted the fluvial channels in the Are Fm. to be anastomosing. The channels seen in cores in the Heidrun Field contradicts this interpretation. Anastomosing channels are characterized by extremely stable channel positions (low lateral mobility) and vertical accretion in areas of very low downstream slopes (Collinson, 1986).

Sedimentary facies in the fluvial-dominated Are Formation as seen in the Are 1 member in the Heidrun Field

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Fig. 6. Stratigraphic correlation of the Are Fm. in the Heidrun Field. Note that the lower part of Are 1A has a high density of meandering channels. This probably reflects lower rate of accommodation space generation, but still within the TST. Porosity plug data show cored intervals. See Fig. 1 for location of wells.

Minor (crevasse) channels Description This facies association has a fining-upward nature with an erosive base and grades upward from fine to very fine sand into silt (Fig. 4). These units have a thickness range from 2.0 to 4.0 m. Sands displays mostly current ripples, often climbing, and smallscale cross-bedding is sometimes seen toward the base of units. Bed boundaries are sometimes sharp, occasionally erosive and minor soft-sediment deformation is seen in places. Mud drapes between sand sets are also frequently seen. The tops of the units are both gradational and sharp.

Interpretation As for the previous facies associations, this facies has been deposited by erosive, channelized processes, although the transition to crevasse splays may be gradual and difficult to determine. The thickness of the units, fining-upward trend and presence of largescale cross-bedding make it reasonable to interpret these features as crevasse channels. On the other hand, the often thin, lateral margin of meandering (single-storey) fluvial channels could be misinterpreted as crevasse channels. Crevasse channels emanate from a relatively prominent break in the levee of the main channel during flood

events. Due to levee development and vertical aggradation of the main channel, the channel often becomes elevated above the surrounding areas (Elliott, 1974). This produces a gradient difference and thereby a tendency for crevasse channels to develop. These channels probably fed crevasse splays and crevasse deltas (Fielding, 1984, 1986). Flow through such a crevasse channel is not always continuous. After the flood, the channel may be abandoned due to lowering of the water level in the main channel (Elliott, 1974). Crevasse channels are often only active during succeeding flood events, thus leading to several reactivation surfaces and fine-sediment drapes. These crevasse channels are often of very limited lateral extent and become less confined and more like crevasse splays downcurrent. The Breathitt Group shows numerous examples where metre-thick crevasse channels have a width of only a few metres. Crevasse channels in the Heidrun Field are found both within thick floodplain deposits and sometimes in the upper part of the bay fill sequences.

Crevasse splay complexes Description This facies association comprises sequences with stacked beds of very fine- to fine-grained sandstone. Individual beds are from a few cm to 60 cm thick.

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Fig. 7. Core photo from fining-upward meandering (single-storey) channel in 6507/7-A-38 showing: (A) large-scale trough cross-bedding; (B) current ripple lamination. See Figs. 4 and 11 for location of photo.

Fig. 8. Photo showing how meandering (single-storey) fluvial channels may have very limited lateral continuity. Note lateral accretion surfaces. Example from the Breathitt Group, Kentucky, USA.

Complexes of stacked beds, showing mainly current ripple lamination, reach thicknesses of up to 9.7 m (Fig. 3). Beds of fine-grained sand showing

cross-bedding are occasionally seen. Climbing ripples and evidence for decreasing current strength (horizontally laminated sandstone that grades into

Sedimentary facies in the fluvial-dominated f~re Formation as seen in the f~re 1 member in the Heidrun Field

current ripples) are frequently seen. Individual beds have sharp contacts, often defined by thin clay and coal fragment layers. The top of these units is often rooted and the facies association is always found within units of rooted floodplain deposits.

Interpretation The sharp-bedded nature of beds together with the dominance of current ripples, frequent climbing ripples and indications of decreasing current strength, indicate that this facies was deposited by crevasse splays. The close association with fine-grained floodplain deposits and the thin individual beds support this. Individual sandstone beds are probably the result of a single crevasse splay episode. These are the product of discrete incursions of sediment-laden waters onto the floodplain after a breach in the levee of the main distributary channel (Elliott, 1974). In the proximal part, close to the main channel, this facies often grades laterally into crevasse channels. Crevasse splays from modern fluvial and delta plains are known to be erosive and channelized in their proximal part, becoming less confined and less erosive as they splay out downcurrent (Fielding, 1984). The stacked nature of beds in this facies association suggests proximity to a nearby distributary, although the non-erosive nature of many of the individual crevasse splay beds indicates some distance from the distributary. Many of the sequences show several crevasse splay sandstones interbedded with thin mud and silts layers deposited by overbank flooding, thus showing that current activity was not continuous, but occurred as several distinct events. This might indicate that the crevasse splays are the product of major flood events, while minor, maybe more "normal" flood events resulted in overbank flooding and deposition of mud and silt over large parts of the floodplain. Aitken and Flint (1996) described a stacked crevasse splay complex in an interfluve setting that clearly correlates with incised valley fill sandstones in the Breathitt Group. This crevasse splay complex is interpreted to have been deposited by overspill of the incised valley margins during major flood events. Such an interpretation is not proposed for the Are Fm. as there are no indicators that the crevasse complexes correlate to incised valley fill sandstones. Thin, single-bed crevasse splays are also found in the floodplain facies association, but here they are thin and often strongly rooted and have no reservoir potential. Crevasse splays and thin crevasse channels also occur in the upper part of bay/lake fill sequences (see below).

95

Floodplain (overbank fines) Description This facies association comprises up to 10 m thick, intensely rooted units of claystone, siltstone, carbonaceous shale and coal (Fig. 9). Individual coals are up to several metres thick and are underlain by rooted horizons. The fine-grained, rooted horizons (palaeosols) are dark grey and often have a high organic content. Only minor horizontally laminated dark grey claystones are preserved in places. Thin, very fine sand and silt beds with sharp boundaries and some sedimentary structures, mostly current ripples, are occasionally seen.

Interpretation The fine-grained nature and strong abundance of strongly rooted sections (palaeosols) and in-situ coals show that this facies association represents deposition mainly by overbank flooding. The most likely environment was an intensely vegetated, swampy inter-channel area with shallow lakes and occasional crevasse splays on a fluvial/delta plain (Coleman and Prior, 1982; Fielding, 1984; Elliott, 1986). The hydromorphic nature (after Besley and Fielding's 1989 classification) of the palaeosols and the presence of numerous and thick coal horizons, shows that deposition took place under reducing conditions, most likely in an area of low topographic relief with poor drainage and a constantly high and rising, reducing water table (Duchaufour, 1982; Retallack, 1983; Besley and Fielding, 1989). Some of the palaeosols with high organic content can be classified as humic gley's after Duchaufour's (1982) classification. The total lack of desiccation cracks and red (oxidized) sediments indicates that the site was submerged, without significant lowering of the water table. Mature soil profiles that show oxidation and concretions suggesting well-drained conditions are often regarded good indicators for interfluves, i.e. areas between fluvial channels during a lowering of base level. However, better-drained palaeosols may have been overprinted by a later rise in ground water level during transgression, thereby giving such soils a hydromorphic imprint (Aitken and Flint, 1996). This can make it difficult to identify interfluves in cores, at least from macroscopic descriptions. More detailed geochemical analysis can potentially detect two-stage palaeosol developments, and thereby interfluvial sequence boundaries like those described by Gardner et al. (1988) in the Breathitt Group (Aitken and Flint, 1996). The thick intervals of organic-rich, rooted sediments imply vertical accretion strata characterized by continuous vegetation. Sediments brought into the vegetated areas of the fluvial/delta plain by over-

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Fig. 9. Core photo from floodplain deposits (overbank fines) in 6507/7-A-38 showing: (A) organic-rich (carbonaceous) shale with rootlets and thin silt layers; (B) thin crevasse splay sands; (C) roots; (D) in-situ coal. See Figs. 4 and 11 for location of photo.

bank flooding were not sufficiently thick to terminate vegetation growth. More than 1 m of sediment from catastrophic flood events will normally terminate vegetation (Retallack, 1983). Normal floods do not deposit more than a few centimetres of sediment by overbank flooding (Collinson, 1986), and these sediments will easily be incorporated into the pre-existing soil (Retallack, 1983). The transition into coals clearly shows that clastic input ceased for longer periods, thus allowing highquality peat to be deposited (McCabe, 1984). This was probably due to removal of the clastic sources (McCabe, 1984). Significant volumes of peat can only be preserved to form coal when the overall increase in accommodation space approximately equals the accumulation rate of peat (Bohacs and Suter, 1997). Intense vegetation could also decelerate the flow during floods, causing the clastic sediments to be deposited at the margins of the swamp, allowing high-quality peat to be deposited in central parts (Staub and Cohen, 1979; Collinson, 1986).

The limited thickness of some of the coal layers makes an origin in low-lying backswamps, relatively close to distributaries, likely (Gersib and McCabe, 1981). This is supported by the fact that several "sheet sandstones", which are thought to be crevasse splays, interrupt the section, thus strongly suggesting the presence of a nearby distributary. The sharp, but non-erosive bases of these crevasse sandstones indicate deposition on distal parts of splay lobes. The presence of thin organic-rich, but still welllaminated mudstones shows that deposition also took place in a shallow anoxic lake environment (Coleman and Prior, 1982; Fielding, 1984). The very finegrained nature of these lake deposits was probably due to intense vegetation surrounding the shallow lake, which protected the lake from coarser clastic input. The lakes were slowly filled by fine-grained sediments, and became the site for re-establishment of vegetation. It is difficult to distinguish upper and lower delta plain from facies alone, but this facies association

Sedimentary facies in the fluvial-dominated Are Formation as seen in the ,3.re 1 member in the Heidrun Field

probably represents deposition in inter-channel areas of the upper to middle delta plain and fluvial plain. In modern deltaic environments the boundary between lower and upper delta plain is defined as the limit of marine influence (Coleman and Prior, 1982). The total lack of marine fauna could thereby indicate upper delta plain and fluvial plain. However, this could not be used as a criterion if the water in the receiving basin was not fully marine. The limited thickness of lake deposits also suggests an upper delta plain environment. Relative thick lake and bay fill sequences are more common on the lower delta plain (Elliott, 1986), while lakes of the upper delta plain are commonly extremely shallow (Coleman and Prior, 1982).

Bay~lake fill Description This facies association shows overall coarsening (shallowing) upward sequences with a lower muddy unit and an upper sandy unit (Figs. 4 and 10). It often

97

has rootlets and coal at the top. Thicknesses range from 1.1 to 5.4 m. Horizontally laminated silty claystone in the lower parts often grades upward into lenticular-laminated sandy siltstone and then to very fineand fine-grained sandstone. Sandstone layers in the uppermost part sometimes have erosional and sharp bed boundaries. Internally, sand beds display current ripples that are sometimes climbing and show indication of decreasing current strength. Coal fragments are abundant in this facies association. Towards the upper part of Are 1, both claystones and sandstones sometimes comprise small sub-vertical and horizontal burrows. Increasing wave reworking and sometimes hummocky cross-bedding is also seen towards the upper part of Are 1 (Fig. 4). Different sequences display great variations in abundance and thickness of individual lithofacies. The number and thickness of sandstone beds in each sequence also vary greatly. Some sequences show several thin beds interbedded with mudstones, while others show only a single, thick sandstone layer. This facies association is generally

Fig. 10. Core photo from a full shoaling (coarsening) upward bay/lake fill sequence in 6507/7-A-38 showing: (A) finely laminated organic-rich (carbonaceous) shale; (B) crevasse splay sands; (C) roots; (D) in-situ coal. See Figs. 4 and 11 for location of photo.

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found in stacked aggradational successions of up to 40 m thick in the upper part of the Are 1 section.

Interpretation This facies association is thought to represent sediments deposited in an inter-distributary bay or lake environment on the lower delta plain. Inter-distributary areas form the major land areas on the lower delta plain (Coleman and Prior, 1982). Sediments are derived mainly from nearby fluvial channels by various flood-generated processes. Some sediments may have been transported into open bays by wave action or littoral drift from the distributary mouth area (Elliott, 1974, 1986). The laminated dark and organic-rich mudstones that are sometimes found at the base of sequences suggest very slow rates of sedimentation. These are interpreted as anoxic bay/lake floor deposits, and represent deposition in an early phase of bay/lake development through drowning of peat swamps due to subsidence or rise in sea level. More silty and laminated grey mudstones, which are often found in lower parts of the sequences, are deposited from overbank flooding in areas beyond the influence of significant traction current activity. No breaching of the channel margin or levee, i.e. crevassing, was necessarily involved in this process (Coleman et al., 1964). The sands generally found in the upper parts of the sequences are interpreted to be the products of crevassing caused by breaching of the levee during flood events. Both crevasse splays and sometimes thin crevasse channels are seen. These sands have often led to complete filling of the bay, resulting in the establishment of intense vegetation as seen from the rooted horizons and overlying coals. The relatively thin palaeosols and coals (compared with those found in the floodplain facies association) indicate little or no sediment aggradation after complete infill of the bay. The overlying laminated black mudstones, seen in some sequences, show that peat production was terminated by drowning of the swamp due to subsidence or rise in sea level. Sequences with more wave ripples and hummocky cross-bedding represent more open bays. This is supported by the fact that these sequences show some bioturbation and a more brackish water fauna. This continues into the overlying Are 2 where tidal influence is also seen. Kj~erefjord (1999) describes these bay fill sequences in more detail.

Sequence stratigraphic development The interpretation of system tracts in this study follows the same principles as used by Aitken and

Flint (1995) in the Breathitt Group, Kentucky, USA, and is based on Exxon terminology (Van Wagoner et al., 1990). Lowstand systems tract (LST) deposits are identified by the presence of erosional surfaces and stacked channel fills. Thick intervals of floodplain deposits (overbank fines) and thick correlatable coals with single-storey channels have been interpreted as transgressive systems tracts (TST). Thick units of stacked bay/lake fill sequences of mainly aggradational nature overlying TST are believed to be parasequences and are interpreted to be highstand systems tracts (HST). It is difficult to define the maximum flooding surface (MFS) in these alluvial deposits, but it could be defined at a change in parasequence stacking pattern from retrogradational to aggradational (Van Wagoner et al., 1990). No attempt has been made to classify sequences according to "third order", "fourth order", etc., but it seems clear that "smaller" sea-level fluctuations are superimposed on "larger" changes in relative sea level. The overall Are Fm. succession grading into the Tilje Fm. is clearly transgressive (large scale), as seen from the vertical change in sedimentary facies and fauna. This interpretation was also presented by Gjelberg et al. (1987). Several relative falls in sea level (reduction in accommodation space) are superimposed on this larger-scale cycle, resulting in the deposition of low-sinuosity (braided) channels. To establish a more precise picture of what kind of sequences this represents, more detailed biostratigraphic control (time control) and more regional data are needed. Although it can be difficult to identify sequence boundaries in predominantly alluvial strata, several sequence boundaries have been interpreted in the Are 1 member. The criteria used in this study follow those summarized by Hampson et al. (1999). Flooding surfaces and maximum flooding surfaces are also difficult to interpret in alluvial deposits. However, the transgressive flooding surfaces probably coincide with the thick, laterally continuous coal beds capping the incised valley fills (Figs. 6 and 11) (Aitken and Flint, 1995; Flint et al., 1995; Hampson et al., 1999). The main criterion used to identify fluvial sequence boundaries is the occurrence of laterally consistent, abrupt vertical change from fine-grained floodplain deposits into channel sands of a low-sinuosity "braided" nature. Such lateral correlatable change in fluvial style and thereby hydraulic character indicates a relative fall in sea level. The stacked, multi-storey channels in the * r e Fm. are not seen to be incised into marine deposits. As such, facies tract dislocations are not observed. Although the stacked multi-storey channel units represent a relative fall in sea level, and sit on top

Sedimentary facies in the fluvial-dominated Are Formation as seen in the ,3,re 1 member in the Heidrun Field

99

Fig. 11. Summary log of well 6507/7-A-38. Note the different vertical change in core plug permeability between the meandering channels and the stacked channels (incised valley fills). The trends are highlighted by blue lines. Also note the aggradational nature and slight upward increase in sandiness in the stacked bay fill sequences. This stacking pattern is the reason for interpreting this section as HST. The wide scatter in permeabilities in this zone clearly reflects the heterolithic nature of this facies association.

of sequence boundaries, they cannot automatically be classified as incised valley fills if lateral correlatable interfluves are not identified. Numerous multi-storey fluvial sandstone bodies have been described in the Upper Carboniferous in the UK and Germany (Hampson et al., 1999). Many of them have been proven to have lateral correlation with an interfluve surface, and therefore represent incised valley fills with widths of 5-25 km (Hampson et al., 1999). Others have much more sheet-like geometry and exceed 35-70 km in

width. In some cases these represent laterally amalgamated incised valley fills. Others may have a non-valley origin, but still occur above sequence boundaries and may represent fluvial braidplains formed as a result of a relative sea-level fall (Hampson et al., 1999). The limited number of well penetrations, and even more limited core coverage, make it difficult to identify correlatable interfluves in the Heidrun Field area. However, within the Are 1A reservoir zone in well

1 O0

6507/7-A-38 (Figs. 6 and 11), a 32 m thick multistorey channel sandstone has all the sedimentological characteristics of an incised valley fill (Fig. 4). A thick coal that probably represents the associated initial flooding surface overlies it. This sandstone unit is not seen to be present in any of the few other wells that have penetrated the Are 1A (Fig. 6). As such, this sandstone unit most likely represents an incised valley fill, and the associated interfluve should be found in the uncored sections in the other wells (Fig. 6). A strong seismic reflector, named Coal Marker 1, is mapped all over the Heidrun Field. This reflector is closely associated with the interval where this initial flooding surface is identified (Fig. 6). A seismic reflector like this probably does not represent reflection from one single coal bed. It is more likely that it corresponds to a unit with several thick coal layers representing the early part of the TST where the rate of peat production balanced the rate of accommodation space generation (Bohacs and Suter, 1997). The overlying TST unit, dominated by fine-grained floodplain deposits with less coal, probably reflects that the rate of accommodation space generation exceeded the rate of peat production. Assuming that this strong, correlatable seismic event represents a time line, it gives good aid to the stratigraphic correlation in these alluvial deposits with otherwise poor biostratigraphic control. The basal sandstone in ,~re 1B reservoir zone is seen in all wells drilled to date in the Heidrun Field (Fig. 6). It also represents a relative fall in sea level, but no potential interfluve has been identified within the limits of the field. However, since this unit is not found in exploration well 6507/8-2 located approximately 5 km to the east (Fig. 1), an incised valley fill interpretation is proposed rather than a multi-storey, multilateral braided system similar to those described by Hampson et al. (1999). On the other hand, the meandering (single-storey) distributary channels, identified both in cores and from wireline logs, are clearly not correlatable between wells (Fig. 6) and deposition is most likely due to autocyclic processes during TST rather than changes in relative sea level. The whole of the .~re 1A zone and in the lower part of the Are 1B zone (subzones .~IB 1-A1B5) are dominated by fluvial channels and thick floodplain deposits with no marine indicators (Figs. 6 and 11). This is interpreted to represent a fluvial plain to upper delta plain setting where relative changes in sea level have produced several sequences containing incised valley fills deposited during LST and thick fine-grained floodplain deposits with thick coals and single-storey channels during TST. The upper part of the .~re 1B zone (subzones ,~IB6-,~IB 11) is dominated by stacked bay/lake fill

K.E. Svela

sequences and thin but laterally continuous coals. This represents a lower delta plain setting where fluctuations in sea level caused the development of several parasequences during a HST. It is difficult to identify a maximum flooding surface in these deposits and thereby to define the exact boundary between the TST and HST of the .3,re 1B unit. However, the boundary has been placed at the top of reservoir zone ,~e 1 B7 (Fig. 11) where there is a change in parasequence stacking pattern from slightly retrogradational to aggradational (Van Wagoner et al., 1990). Reservoir characteristics

Reservoir properties in the Are Fm. are controlled at two levels. The first is the lateral connectivity of the reservoir units and the second is the internal character of the sandstones. In the Heidrun Field, depth of burial is limited and diagenesis has not significantly altered reservoir properties (Olsen et al., 1999). Internal reservoir properties are therefore governed primarily by original permeability, which in turn is a function of pore throat radius. In sedimentological terms, this is basically a function of grain size and sorting. As a consequence, variations in porosity and permeability are mainly a function of depositional processes. In addition, reservoir performance is also influenced by the fact that (1) the sandstones are unconsolidated, (2) the Heidrun Field is normally pressured, (3) the ,~e oil is viscous with an API gravity of 22 ~ and (4) the field is strongly segmented by faults. The multi-storey fluvial channels (incised valley fills) have superb reservoir quality with few internal barriers and little internal variation in permeability (Figs. 11 and 12). These LST sand deposits also have large lateral continuity and where present in the hydrocarbon column, they should be easy to target with wells. Large lateral continuity will also make it easier to provide pressure support and sweep by water injection. The meandering (single-storey) distributary channels deposited during TST also have very good internal reservoir quality, although an upward reduction in permeability reflects the fining-upward trend in grain size (Fig. 11). The lateral extent of these fluvial sand bodies is more restricted (Figs. 6 and 8) and it will be difficult to target specific channels. This is also reflected in Fig. 12, where the individual reservoir zones representing TST deposits have high permeabilities, but the relatively low net/gross shows the limited lateral distribution. Limited lateral extent also limits effective pressure support and sweep. Crevasse splays and crevasse channels have good horizontal permeabilities, but show significantly lower

Sedimentary facies in the fluvial-dominated f~re Formation as seen in the f~re 1 member in the Heidrun Field

101

Fig. 12. Summary of average reservoir properties for all Are 1 wells in the Heidrun Field and their relationships to system tracts. Note that permeability values (plotted on linear scale) are from net sand, and as such largely represent grain size and sorting. Also note that because the number of cores are limited, there is a difference in number of samples between the wireline-log-generated horizontal permeabilities and the vertical permeabilities from core plugs.

vertical permeability due to mud drapings. Individual sands within such complexes are likely to be discontinuous and the bodies themselves will sometimes have limited lateral extent. Bay/lake fill sequences also have good reservoir properties, although not as good as in the channel deposits. Within the Heidrun Field they are thought to have large lateral extent and stacked sequences make up thick units. This is shown in Fig. 12 where deposits in the upper part of Are 1B are seen to have good horizontal permeabilities and very high net/gross. One should bear in mind that the thin mudstones at the base of individual sequences are barriers to vertical flow. Unfortunately, these mudstone intervals are not easily detected on wireline logs. Fig. 12 clearly shows that vertical permeability from core plugs is much lower than horizontal permeability. Finally, one of the main challenges in producing the Are Fro. in the Heidrun Field is not how these reser-

voir facies will perform individually. The question is how they will perform when completed together. Due to the unconsolidated nature of the sands, production wells have to be gravel packed. The contrast in flow properties between different sedimentary facies will then limit economic recovery, especially when gravelpacked wells limit the possibility to selectively isolate zones after water or gas breakthrough.

Acknowledgements I want to express my thanks to Conoco for allowing me to write this paper and to Statoil and Fortum Petroleum for giving permission to publish it. The majority of these data was gathered when I was seconded to Statoil's Heidrun Petek group in Stj0rdal during 1994 and 1995. During my time with Statoil I had many fruitful discussions with Jostein Kja~refjord, Lars-Magnus F~ilt and Arne Dalland. I would

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also like to thank Ed Clifton with whom I worked closely on the sedimentology of the Heidrun Field.

References Aitken, J.E and Flint, S.S., 1995. The application of high-resolution sequence stratigraphy to fluvial systems: a case study from the Upper Carboniferous Breathitt Group, eastern Kentucky, USA. Sedimentology, 42: 3-30. Aitken, J.F. and Flint, S.S., 1996. Variable expressions of interfluvial sequence boundaries in the Breathitt Group (Pennsylvanian), eastern Kentucky, USA. In: J.A. Howell and J.F. Aitken (Editors), High Resolution Sequence Stratigraphy: Innovations and Applications. Geol. Soc. Spec. Publ., 104: 193-206. Barwis, J.H., 1978. Sedimentology of some South Carolina tidalcreek point bars, and a comparison with their fluvial counterparts. In: A.D. Miall (Editor), Fluvial Sedimentology. Can. Soc. Pet. Geol., Mem., 5: 129-160. Besley, B.M. and Fielding, C.R., 1989. Palaeosols in Westphalian coal-bearing and red-bed sequences, central and northern England. Palaeogeogr., Palaeoclimatol., Palaeoecol., 70(4): 303-330. Bohacs, K. and Suter, J., 1997. Sequence stratigraphic distribution of coaly rocks: fundamental controls and paralic examples. Am. Assoc. Pet. Geol. Bull., 81(10): 1612-1639. Coleman, J.M. and Prior, D.B., 1982. Deltaic environments. In: EA. Scholle and D. Spearing (Editors), Sandstone Depositional Environments. Am. Assoc. Pet. Geol., Mem., 31: 139-178. Coleman, J.M., Gagliano, S.M. and Webb, J.E., 1964. Minor sedimentary structures in a prograding distributary. Mar. Geol., 1: 240-258. Collinson, J.D., 1986. Alluvial sediments. In: H.G. Reading (Editor), Sedimentary Environment and Facies. 2nd ed., Blackwell, Oxford, pp. 20-62. Dalland, A., Worsley, D. and Ofstad, K., 1988. A lithostratigraphic scheme for the Mesozoic and Cenozoic succession offshore midand northern Norway. NPD Bulletin 4. Duchaufour, E, 1982. Pedology. Allen and Unwin, London, 448 pp. Elliott, T., 1974. Abandonment facies of high-constructive lobate deltas, with an example from the Yoredale series. Proc. Geol. Assoc., 85: 359-365. Elliott, T., 1986. Deltas. In: H.G. Reading (Editor), Sedimentary Environments and Facies. 2nd ed., Blackwell, Oxford, pp. 113154. Fielding, C.R., 1984. Upper delta plain lacustrine and fluvio-lacustrine facies from the Westphalian of the Durham coalfield, NE England. Sedimentology, 31: 547-567. Fielding, C.R., 1986. Fluvial channels and overbank deposits from the Westphalian of the Durham coalfield, NE England. Sedimentology, 33:119-140. Flint, S., Aitken, J.F. and Hampson, G., 1995. Application of sequence stratigraphy to coal-bearing coastal plain successions: implications for the UK Coal Measures. In: M.K.G. Whateley and D.A. Spears (Editors), European Coal Geology. Geol. Soc. Spec. Publ., 82: 1-16. Gardner, T.W., Williams, E.G. and Holbrook, EW., 1988. Pedogenesis of some Pennsylvanian underclay: ground water, topo-

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Norske Conoco A/S, Tangen 7, N-4070 Randaberg, Norway

graphic and tectonic controls. In: J. Reinhardt and W.R. Sigleo (Editors), Palaeosols and Weathering through Geologic Time: Principles and Applications. Geol. Soc. Am., Spec. Pap., 216: 81-102. Gersib, G.A. and McCabe, EJ., 1981. Continental coal-bearing sediments of the Port Hood Formation (Carboniferous), Cape Linzee, Nova Scotia, Canada. In: F.G. Ethridge and R.M. Flores (Editors), Present and Ancient Nonmarine Depositional Environments: Models for Exploration. Spec. Publ. Soc. Econ. Paleontol. Mineral., 31: 95-108. Gjelberg, J., Dreyer, T., H0ie, A., Tjelland, T. and Lilleng, T., 1987. Late Triassic to Mid-Jurassic sandbody development on the Barents and Mid-Norwegian shelf. In: J. Brooks and K. Glennie (Editors), Petroleum Geology of North West Europe. Graham and Trotman, London, pp. 1105-1129. Hampson, G.J., Davies, S.J., Elliott, T., Flint, S.S. and Stollhofen, H., 1999. Incised valley fill sandstone bodies in Upper Carboniferous fluvio-deltaic strata: recognition and reservoir characterization of Southern North Sea analogues. In: A.J. Fleet and S.A.R. Boldy (Editors), Petroleum Geology of Northwest Europe: Proceedings of the 5th Conference. Geological Society, London, pp. 771-788. Kja~refjord, J.M., 1999. Bayfill successions in the Lower Jurassic ,~re Formation, offshore Norway: sedimentology and heterogeneity based on subsurface data from the Heidrun Field and analog data from the Upper Cretaceous Neslen Formation, eastern Book Cliffs, Utah. In: T.F. Hentz (Editor), Advanced Reservoir Characterization for the 21st Century. Gulf Coast Section, Society of Economic Paleontologists and Mineralogists Foundation 19th Annual Research Conference, Tulsa, OK, pp. 149-158. McCabe, EJ., 1984. Depositional environments of coal and coalbearing strata. In: R.A. Rahmani and R.M. Flores (Editors), Sedimentology of Coal and Coal-Bearing Sequences. Spec. Publ. Int. Assoc. Sedimentol., 7: 13-42. Olsen, T., Rosvoll, K.J., Kj~erefjord, J.M., Arnesen, D.M., Sandsdalen, C., JCrgenvfig, S.H., Langlais, V. and Svela, K.E., 1999. Integrated reservoir characterization and uncertainty analysis, Heidrun Field, Norway. In: A.J. Fleet and S.A.R. Boldy (Editors), Petroleum Geology of Northwest Europe: Proceedings of the 5th Conference. Geological Society, London, pp. 1209-1220. Oomkens, E., 1974. Lithofacies relationships in the late Quaternary Niger Delta Complex. Sedimentology, 21: 195-222. Retallack, G.J., 1983. A palaeopedological approach to the interpretation of terrestrial sedimentary rocks: the Mid-Tertiary fossil soils of Badlands National Park, South Dakota. Bull. Geol. Soc. Am., 94: 823-840. Staub, J.R. and Cohen, A.D., 1979. The Snuggy Swamp of south Carolina: a back-barrier estuarine coal-forming environment. J. Sediment. Petrol., 49: 133-144. Van Wagoner, J.C., Mitchum, R.M., Campion, K.M. and Rahmanian, V.D., 1990. Siliciclastic Sequence Stratigraphy in Well Logs, Cores and Outcrops. Methods in Exploration Series 7, American Association of Petroleum Geologists, Tulsa, OK. Weimer, R.J., Howard, J.D. and Linsay, D.R., 1982. Tidal flats and associated tidal channels. In: EA. Scholle and D. Spearing (Editors), Sandstone Depositional Environments. Am. Assoc. Pet. Geol., Mem., 31: 191-245.

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Sedimentology of the heterolithic and tide-dominated Tilje Formation

(Early Jurassic, Halten Terrace, offshore mid-Norway) Allard W. Martinius, Inge Kaas, Arve Naess, Geir Helgesen, Jostein M. Kjaerefjord and Deborah A. Leith

The Early Jurassic Tilje Formation on the Halten Terrace, offshore mid-Norway, was deposited in a relatively narrow but long seaway connecting the Boreal Ocean in the north and the Tethys Ocean in the south. Sediments of the Tilje Formation in the SmOrbukk and Heidrun Fields have been classified into ten facies associations. Many of the lithofacies are mud-rich and typified by strong grain-size contrasts. In addition, all but two of the facies associations are tidally influenced or dominated. As a result, 80% of the total rock volume consists of very heterolithic sediments characterised by rapidly alternating grain-size changes between mudstone or siltstone and fineto medium-grained sandstone. In the Tilje Formation, the recognition and interpretation of heterolithic facies is crucial to understanding the depositional conditions and stratigraphic architecture. The classification scheme and the associated facies breakdown in cored wells served to define two successive conceptual depositional models that are placed in a sequence stratigraphic framework. The lower part of the Tilje Formation (T1 and most of T2) are envisaged to have formed in response to base-level fall, creating a series of low-relief valleys, and subsequent base-level rise resulting in the formation of a tide- and wave-dominated estuarine system. The upper part of the Tilje Formation (top of T2 to T6) is interpreted to have formed as a tide- and fluvial-dominated delta-like system. These two contrasting depositional styles resulted in different three-dimensional facies architectures, relative facies proportions, and facies stacking patterns, which have implications for reservoir model-building methods.

Introduction Hydrocarbon exploration in the Haltenbanken area, offshore mid-Norway, began in the early 1980s, with the discovery of the Midgard Field in 1981. Since then several major discoveries have been made and the area today is regarded as a fairly mature hydrocarbon province. Production is mainly from siliciclastic sequences deposited in shallow marine environments. These comprise either relatively homogeneous sands, or heterogeneous packages formed by an intercalation of mudstone, siltstone, and sandstone. The Tilje Formation, on which this paper will focus, forms a reservoir interval in several hydrocarbon-producing fields on the Halten Terrace, a smaller section of the Haltenbanken area (Fig. 1). Production from the Tilje Formation is significantly affected by the strongly heterolithic nature of many of the depositional facies and the complicated reservoir architecture. Various bedding styles are found at several scales and approximately 80% of the total rock volume is formed by heterolithic facies. Over the years, the complex sedimentology of the Tilje Formation has resulted in various published palaeoenvironmental interpretations (Karlsson, 1984; Gjelberg et al., 1987;

Pedersen et al., 1989; Ekern, 1990; Dreyer, 1992, 1993; Taylor and Gawthorpe, 1993; Van de Weerd, 1996), commonly tailored to accommodate field-specific observations, although the overall strongly tidal nature has been recognised by most workers. This paper will discuss the Tilje Formation as it is specifically developed in the Heidrun and SmOrbukk Fields (Fig. 1). However, the presented facies classification scheme and interpretations are based on selected observations from most of the hydrocarbon fields on the Halten Terrace where the Tilje Formation is present, and hence represent an overview of subregional character. The aim of the paper is, firstly, to present a conceptual depositional model for the Tilje Formation that is based on a generally applicable facies classification scheme. In that sense, it differs from previous studies. The second aim is to analyse in more detail depositional conditions for the different stratigraphic intervals of the Tilje Formation.

Data base and methodology This study is based on core observations and consistent facies breakdowns in 25 wells from 9 hy-

Sedimentary Environments Offshore Norway - Palaeozoic to Recent edited by O.J. Martinsen and T. Dreyer. NPF Special Publication 10, pp. 103-144, Published by Elsevier Science B.V., Amsterdam. 9 Norwegian Petroleum Society (NPF), 2001.

1 04

A.W. Martinius et al.

Fig. 1. Location of hydrocarbon fields producing from the Tilje Formation on the Halten Terrace.

drocarbon fields on the Halten Terrace (Nome, Heidrun, SmCrbukk, SmCrbukk SCr, Midgard, Trestakk, Tyrihans, Lavrans, and Njord; Fig. 1) with emphasis on the Heidmn and Sm0rbukk Fields. Lithofacies characterisation and interpretation were supported by detailed trace-fossil and ichnofabric analysis. In addition to data obtained from numerous published reports, conventional log data, in-house biostratigraphic data, samarium-neodymium (147Sm/144Nd and 143Nd/144Nd) isotopic data, and dip-meter data were used for correlation purposes, to support facies interpretations and/or identify provenance areas. Biostratigraphic data were obtained from several separate studies by different companies and, as a result, the findings are not always directly comparable. Two other complicating factors are: (1) the recurring dilemma of

whether the same events in different wells are time equivalent, or whether they are simply facies dependent and thus time transgressive; and (2) the fact that preservation of palynological material in deep wells (>4000 m in the SmCrbukk Field) is so poor that not only is the number of observations extremely low but the results may be uncertain. Nevertheless, Sm/Nd isotopic markers fit well with correlations based on biostratigraphic events and sedimentological criteria.

The Tilje Formation Geographical occurrence The Early Jurassic Tilje Formation (Dalland et al., 1988) is geographically confined to a NNE-

Sedimentology of the heterolithic and tide-dominated Tilje Formation

105

Fig. 2. Paleogeographic setting of the early Pliensbachian seaway during times of deposition of the Tilje Formation (after Dor6, 1991). Inset shows global configuration of continents during the Pliensbachian (190 Ma; after Smith et al., 1994). The boxed area indicates the area covered by Fig. 3.

SSW-oriented belt that is known to be at least 470 km long and 150 km wide judging by data from hydrocarbon exploration and nearshore shallow drill holes (Figs. 2 and 3). It is inferred that in an eastward direction the Tilje Formation passes into time-equivalent upper delta plain and alluvial fan environments typified by kaolinitic claystones alternating with both mud- and grain-supported conglomerates (Bugge et al., 1984). The formation has an average thickness of 120 m in the Heidrun Field and 150 m in the SmCrbukk Field, but reaches maximum thicknesses of more than 300 m in the central western part of the Halten Terrace (Lavrans Field). It has been traced north as far as 68o02 ' (where it gradually pinches out

on the R0st High; Fig. 3) and to the south as far as 63006 ' (where the facies pass into deeper marine sediments). It is also inferred that the Tilje Formation may extend to the east as far as approximately 09~ ' (eastern part of the Tr0ndelag Platform; Bugge et al., 1984; Fig. 3). The facies similarities between the comparable stratigraphic intervals in the various Halten Terrace oilfields suggest that deposition occurred while the source terrains were relative stable or underwent equal isostatic change. The inferred paleolatitude of the Halten Terrace was between 49 ~ and 53 ~ north (Dor6, 1991; Smith et al., 1994) during deposition of the Tilje Formation. In general, mid-paleolatitudinal temperature and pre-

106

A.W. M a r t i n i u s et al.

Fig. 3. Proposed paleogeographic sketch of the seaway in which the Tilje Formation sediments were deposited.

cipitation data indicate that the Jurassic climate was warm with rather equable global conditions (Hallam, 1985, 1994). Strong seasonal differences in temperature and rainfall may have existed (Hallam, 1994).

Structural setting and provenance The Early Jurassic seaway was approximately 1500 km long and 250 km wide, and during deposition of the Tilje Formation connected the Boreal Ocean in the north to the Tethys Ocean in the south (Fig. 2). Both oceans were bordered by shallow shelves with numerous emergent terrains (Dor6, 1991). Deposition of sediments occurred in association with (a) minor tectonic pulse(s), which resulted

in the development of N-S-oriented growth faults (inhouse data; Ehrenberg et al., 1992). The smaller Tilje seaway was separated from the main Early Jurassic seaway by the Helland-Hansen Arch-Bode High (Fig. 3). Connections with the Early Jurassic seaway existed both in the north, south of the ROst High and the uplifted Ribban Basin structure, as well as in the south (Fig. 3). The structural history of the Halten Terrace is typified by a long period of rifting and subsidence, that commenced in the Triassic, or earlier, and that continued up to early Eocene time. Bukovics et al. (1984) suggested that the late Palaeozoic to mid-Mesozoic subsidence of the Norwegian-Greenland Sea Rift was mainly governed by mechanical stretching of the

Sedimentology of the heterolithic and tide-dominated Tilje Formation

crust. Crustal extension accelerated during the Early Triassic with deposition of mostly continental strata on the TrOndelag Platform. Paleogeographical studies suggest that the Tilje Formation was deposited in the Halten-TrOndelag Basin, which encompassed the Halten Terrace and the TrOndelag Platform (Fig. 3). This coast-parallel riftgenerated basin, which became the locus of the Early Jurassic seaway (Gjelberg et al., 1987; Dor6, 1991; Fig. 3), was associated with the Kristiansund-BodO Fault Complex, a NNE-SSW-oriented fault system that included the FrOya High, the Bremstein Fault Complex, and the Nordland Ridge (Fig. 3). In a plate-tectonic perspective, the basin is part of the continental passive margin of the Northern Atlantic Rift Domain (Dor6, 1991). The Halten Terrace (Fig. 3) is located in the centre of the fault complex, between the TrOndelag Platform to the east and the Sklinna Ridge to the west (Bukovics et al., 1984; Schmidt, 1992; Fig. 3). In-house structural data suggest early activity of a domal structure to the west of the basin that became uplifted as part of the Early Jurassic proto-rift phase forming the Helland-Hansen Arch. The fossil rift underlying the basin formed during the Triassic as part of the proto-Atlantic Ocean. The Early Jurassic basin in which the Tilje Formation was deposited can be characterised as a late-stage pre-rift basin (cf. NOttvedt et al., 1995) in which subsidence was caused mainly by thermal sagging and sediment compaction (Fig. 4), and in which syn-sedimentary faults locally played an important role. Application of this model is further supported by: (a) absence of evidence for basin-bounding fault margins; (b) absence of volcanic

107

rocks; (c) relative uniformity of facies; and (d) large size of depositional systems. Provenance age data derived from Sm/Nd isotope analysis are considered totally independent of sample grain size because minerals from rocks in a specific source area and with a certain age weather down to a range of grain sizes (from mud to coarse sand) which maintain their similar provenance age (Dalland et al., 1995). This characteristic allows Sm/Nd isotope data to be used without restriction. Generally, in-house dip-meter and Sm/Nd isotope data suggest three possible source areas for the Tilje Formation. The main source area is considered to be the mainland east of the Tr~ndelag Platform (Fennoscandian Shield; Dalland et al., 1995). Additionally, sediment was shed from the Ribban Basin structure in the north (Fig. 3), which was dominated by erosion during the entire Jurassic. Furthermore, an emergent area in the west, the Helland-Hansen-Bod~ High (Fig. 3) is also inferred to have supplied sediment to the Halten Terrace area.

Stratigraphy The late Pliensbachian to early Toarcian Tilje Formation is one of four formations comprising the Early Jurassic Bfit Group (Fig. 5A). The formation is underlain by the Hettangian-Sinemurian Are Formation, dominated by lower delta plain and bay deposits (including coal seams), and overlain by the Toarcian Ror Formation, composed of marine mudstones. In the northernmost part of the Halten Terrace, the Rot Formation is replaced by the Tofte Formation, which is interpreted as a fan-delta deposit (in-house data). The

Fig. 4. Cross-sectional sketch of the depositional basin at Tilje time showing the position of the Middle Triassic fossil rift and Late Triassic salt layers.

108

A. W. Martinius et al.

A

GENERAL LITHOSTRATIGRAPHY

MID-NORWAY

] HALTENBANKEN INORDLANDRIDGEI CHRONOSTRATIGR. GP,

TR/ENABANKEN

FORMATIONS p

Callovian

LM

~

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" "

i

J

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aa

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<~

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(.gZ '

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~

,,r n,--j

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Pliensbach.

Sinemurian

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I-

9 ". " . "

,

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_

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~ ILE

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m

~_. . o< rn . . . .

~

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.

. _

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.

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."

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~

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r

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"GREY BEDS"

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Low

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Short term~ j

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'~

"RED BEDS"

High Longte)

Haq " et al 198

Dating uncertain

~t~

Hallam 1988

( A

Surlyk 1990

.---- JamesonLandEmbayment

/

Sinemurian

/

/

/

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i

Fig. 5. (A) Generalised time- and lithostratigraphic column of the Jurassic deposits on the Halten Terrace (modified from Dalland et al., 1988). Gp. = Group. (B) Sea-level curve for the Early Jurassic (after Surlyk, 1990). The Jameson Land Embayment is located on the East Greenland side of the Early Jurassic seaway (Fig. 2).

Tilje Formation corresponds to the upper part of the Aldra Formation, a name that was used to describe the same succession prior to publication of the revised lithostratigraphic scheme for the Mesozoic and Cenozoic succession of offshore mid-Norway and northern Norway (Dalland et al., 1988). An overall sea-level rise is inferred for the late Pliensbachian to early Toarcian (Haq et al., 1987; Hallam, 1988; Surlyk, 1990; Fig. 5B). The type section for the Tilje Formation is in well 6507/11-1 in the Midgard Field (Asgard license; Dalland et al., 1988; Fig. 1). Over the Halten Terrace area, the thickness and depositional facies of the Tilje Formation differ considerably from the type section due to local variations in subsidence rate and depositional patterns. The stratigraphic subdivision of the Tilje Formation differs slightly from field to field. In the Heidrun

Field, the Tilje Formation is divided into four reservoir units or members (Tilje 1 to 4) based on local lithostratigraphic correlation supported by biostratigraphic data (mainly palynofacies; in-house data; Pedersen et al., 1989; Whitley, 1992; Hemmens et al., 1994). Ehrenberg et al. (1992) divided the Tilje Formation of the SmCrbukk Field into three units based on the correlatability of offshore shales and bioturbated siltstones. In the operator's (Statoil) current reservoir zonation, the formation has been divided into six reservoir units, Tilje 1 to 6, some of which have been further subdivided (see below). Deposition of sediments of the Tilje Formation occurred simultaneously with deposition of the Neill Klinter Group sediments on the western side of the Early Jurassic seaway (Fig. 2). The Neill Klinter Group (Dam and Surlyk, 1998) is composed of four

Sedimentology of the heterolithic and tide-dominated Tilje Formation

formations that were deposited in a wide, shallow, wave-, storm- and tide-influenced paralic embayment. The lower part of the Neill Klinter Group has many features in common with the Tilje Formation. Facies association classification scheme

Descriptions and definitions The sedimentary facies of the Tilje Formation are grouped in ten facies associations, of which one contains only one lithofacies. The classification scheme is inferred to be generally valid for the Tilje Formation on the Halten Terrace. However, local variations will cause departure from the generalised descriptions. The descriptive characteristics of the facies associations are summarised in Table 1. A sedimentological summary log of the formation in the Heidrun Field illustrates the range of lithologies (Fig. 6). Special emphasis has been devoted to systematically characterise the various types of heterolithic facies encountered in the Tilje Formation. It is believed that recognition of commonly subtle but genetically significant variations in the intralayer characteristics of heterolithic facies is one of the key factors for their subdivision, for understanding the depositional processes, to establish a depositional model, and to understand the general stratigraphic architecture. The systematic description of heterolithic facies includes grain-size variation and grain sorting, structures in the coarse-grained layers, trace-fossil types and ichnofabrics, variability of fine-grained and coarse-grained layer thicknesses and descriptive statistical data, frequency of occurrence, average thickness, stratigraphic position, and (inferred) lateral extent. These characteristics are listed in Table 2 and illustrated in Fig. 7. A few terms used in the facies discussions need to be clarified in this introductory part. A barform is defined as a submerged or (partially) emerged aquatic feature, formed by mud, sand, and/or gravel at the interface of water currents and sediment bed, that predominantly produces "lateral accretion" deposits and tends to scale with flow width rather than flow depth (slightly modified after Dalrymple and Rhodes, 1995). Barforms typically have flow-parallel spacings, which are several times the channel width, and a flow-transverse dimension which is a large fraction of the channel width. Orientations can be flow-transverse,-oblique, and-longitudinal, sometimes combined in a single, composite bar. Barforms can be upward coarsening or fining and can be classified according to three main types (cf. Dalrymple and Rhodes, 1995). (1) Repetitive barforms: exhibiting a quasi-regular, repetitive spacing in the direction parallel to the flow. (2) Elongate tidal bars: features

109

that are developed at locations where there is strong, rectilinear, tidal flow. They are typically oriented at a small angle (<20 ~) to the predominant tidal flow, and are commonly asymmetric. Elongate tidal bars may be considered as levee-like features that accumulate in the zones of no net transport between channels with opposing directions of residual transport. (3) Delta-like bodies: isolated features that have a lobate morphology, and are situated at the end of a channel, or where a channel widens appreciably (point of flow expansion). A dune is defined as an aquatic flow-transverse feature, formed at the interface of water currents and sediment bed, with a crestline oriented perpendicular (+15 ~) to the resultant transport vector, that creates "foreward accretion" deposits, occupies a predictable stability field in bedform phase diagrams, and scales with flow depth (after Allen, 1982; Dalrymple and Rhodes, 1995; see also Anastas et al., 1997). Simple dunes are formed by single-dune bedforms whereas compound dunes are formed by the superimposition of smaller dunes on a large dune.

Interpretations Facies association 1 (storm-influenced prodelta facies) Facies association 1 is found at the base of the formation in Tilje 1.1 in the Heidrun Field and the timeequivalent stratigraphic interval (the uppermost part of the Are Formation) in the SmOrbukk Field. The most important occurrence, however, is in the Tilje 2 stratigraphic interval in both the Heidrun and SmCrbukk Fields where it composes approximately 80% of the rock volume. The main difference between the two occurrences is that lithofacies 1.2 and 1.3 are almost absent in the Tilje 2 stratigraphic interval. Intense bioturbation and the fine-grained texture of the muddy, very fine-grained sandstone component of lithofacies 1.1 (Table 1; Fig. 8) indicates prevailing low-energy Conditions during deposition. The trace-fossil suite of lithofacies 1.1 is indicative of an equilibrium colonisation style and fair-weather conditions. It is dominated by a variety of deposit feeders such as annelids and crustaceans, among which abundant bivalves exist. In general, the association of marine palynomorphs (for example Nannoceratopsis and Acritarcs), the co-occurrence of faintly laminated coarser-grained sandstone layers, and the lack of wave-generated structures suggest a distal shoreface, low-energy environment of deposition. Lithofacies 1.1 is interbedded with deposits of lithofacies 1.2 and 1.3. Regular occurrences of lithofacies 1.2 are interpreted to be proximal storm beds (Fig. 7), whereas relatively thick, hummocky cross-

o TABLE 1 Main characteristics of the lithofacies and facies associations Facies association

Lithofacies

Dominant texture and sedimentary structures

Dominant ichnofabrics

Diagnosis

1. Storm-influenced prodelta facies

1.1

Cycles of up to 1.2 m thick formed by muddy, very fine-grained sandstone or silty mudstone at the base to well-sorted fine-grained sandstone at the top; cycle boundaries gradual. Homogenised lenticular and wavy bedding discernible; strongly bioturbated. Up to 1 m thick calcified layers occur.

Teichichnus, resting traces, Rhizocorallium, Siphonichnus, Planolites, Palaeophycus heberti, Taenidium, Chondrites, Skolithos, burrow mottling,

Association formed by hummocky cross-stratified sandstones and thoroughly bioturbated, upward-coarsening muddy sandstone successions.

1.2

Well-sorted, normal-graded (medium- to fine-grained) thin sandstone beds with a sharp, parallel-laminated base and a ripple-laminated top.

As lithofacies 1.3.

1.3

Moderately to well-sorted, homogeneous and fine-grained sandstone forming on average 1.7 m thick beds. Lamination is parallel or undulous, including convex-up sets, often with laterally variable set thicknesses, often hummocky cross-stratified. Beds are sharp-based, occasionally erosional with cm-thick granule lag. Rich in muscovite and biotite.

Palaeophycus, escape traces, Planolites, and horizons with escape traces, Planolites, Berguaeria, Skolithos,

2.1

Moderately to poorly sorted, muddy, fine- to medium-grained matrix-supported sandstone, occasionally in coarsening-up successions of up to 70 cm. Less bioturbated sections preserve tabular-stratified sets or lenticular bedding. Floating, rather angular bioclastic calcite grains are occasionally observed and limited to certain intervals.

Escape traces, Planolites, Berguaeria, Skolithos, excavation traces, Siphonichnus, alternations between strongly and poorly bioturbated intervals.

Strongly bioturbated siltstone or fine- to coarse-grained sandstone with floating calcite grains and mostly homogenised bedding.

2.2

Moderately sorted, muddy siltstone occasionally including medium-grained, matrix-supported sandstone. No sedimentary structures preserved; often showing reddish colour. Floating, angular bioclastic calcite grains occur abundantly.

Thalassinoides, Palaeophycus, Skolithos, burrow mottling (hiatal surfaces) and (occasionally abundant) bivalve burrows, Taenidium, Planolites, Chondrites, Skolithos, burrow mottling, faecal pellets.

Total thickness: average = 4.6 m; Std. -- 4.5. Volume fraction: 11%.

3

Wavy-bedded facies composed of well-sorted, very fine-grained sand- or siltstone layers irregularly alternating with often sharp-based mud- or siltstone layers which commonly are lenticularly bedded at fine scale (Table 2; Figs. 7, 9 and 10). Wave ripples, often with chevron lamination, dominate but current ripples with tabular stratification and sigmoidal cross-stratification occur. Mud layers cover single sets; cracks are common, siderite layers are infrequent, and microslumps are rare. Medium-grained sandstone layers with small-scale trough cross-stratification, or rarely several cm-thick swaley cross-stratified layers may occur.

Type A: Palaeophycus, Teichichnus, Phycosiphon, Rhizocorallium, Chondrites, Siphonichnus, Skolithos, Planolites, P. heberti. Type B: Spreiten traces, resting traces, Planolites, Skolithos, Siphonichnus. Zones

Heterolithic facies formed by wavy- to lenticular-bedded (cf. Reineck and Wunderlich, 1968), regularly alternating silt- or sandstone and mudstone layers. Total thickness: average -- 3.3 m; Std. -- 2.7. Volume fraction: 7%.

2. Transgressive facies

3. Laminated delta-front facies

faecal pellets. Bivalve burrows common.

excavation traces.

with more and less intense bioturbation occur.

Thickness 1.1: average -- 4.8 m; Std.-- 5.0. Thickness 1.3: average = 1.7 m; Std. -- 1.4. Volume fraction: 12%.

~..~~ ~..,~

t~

TABLE 1 (continued) t...,

Facies association

Lithofacies

Dominant texture and sedimentary structures

Dominant ichnofabrics

Diagnosis

4. Delta-front lobe facies

4.1

Mica-rich, fine- to medium-grained sandstone with parallel lamination and double mud drapes. Intercalated with thin mud and lenticular-bedded layers (Fig. 11). Ripple-laminated and low-angle cross-stratified cosets between mudstone layers. Sorting within sandy layers good, among layers poor. Siderite layers common. Occasionally, cm-thick sandstone layers with floating pebbles occur.

Palaeophycus, Skolithos, Planolites, escape traces;

Upward-coarsening, parallel-laminated or low-angle cross-stratified heterolithic sandstone with mud or silt layers and/or drapes.

Heterolithic unit typified by abundant mudstone layers (Table 2; Figs. 7 and 12) and fairly well sorted sandstone. An upward change from ripples to small trough cross-stratified sets, associated with a coarsening-up from fine- to medium-grained sandstone occurs. Dominated by wave ripples with minor current ripples.

Teichichnus, resting traces, Rhizocorallium, Siphonichnus, Planolites, Palaeophycus heberti, Palaeophycus sp., faecal pellets, Phycosiphon, Skolithos, bivalve burrows.

Upward-coarsening succession of fine-grained, wavy-bedded, 11 to 12 mm thick sandstone with disturbed (double?), 7 to 8 mm thick mud layers in which uncommonly intercalations of 18 to 19 mm thick coarse-grained sandstone layers occur (Table 2). Overlain by medium- to coarse-grained, normally graded and moderately sorted, angular to subangular sandstone organised in small- and large-scale cross-stratified sets of up to 2 m thick showing progressively steeper inclined laminae and clay drapes (Fig. 13); rip-up clasts and internal erosion surfaces present.

Palaeophycus, Skolithos, Planolites, escape traces.

5.2

Clean, fine- to medium-grained and moderately sorted sandstone with erosive base, often rip-up mud clasts, vague trough cross-stratification at base, and common and wavy flaser-bedded, wave ripple-dominated sandstone in upper parts. Mud drapes are discontinuous at core-width scale and < 1 mm thick.

Type 1" Rosselia dominated. Type 2: Gyrochorte dominated.

5.3

Finely laminated and moderately bioturbated dark lenticular-bedded mudstone, and muddy, wavy-bedded heterolithic sandstone. Bioturbation causes feather-like appearance of mud layers.

As lithofacies 5.2.

6.1

Poorly sorted succession of coarse- to fine-grained sandstone, often with sideritised layers. Very coarse-grained sandstone lenses (occasionally with small granules), small-scale cross-stratified sets, and rip-up clasts occur at base. Passes upward into ripple-laminated, dominantly flaser- but occasionally wavy flaser-bedded sets at the top.

Diplocraterion, Planolites, escape traces, resting traces, Skolithos, faecal pellets, Berguaeria, Skolithos,

4.2

5. Outer estuarine and delta-front bar facies

6. Channelised delta-front lobe facies

5.1

moderate to intensely bioturbated zones.

t...,

e~

Total thickness: average -- 4.1 m; Std. - 4.0. Volume fraction: 15%.

t...,

excavation traces" moderate bioturbation intensity.

Upward coarsening then fining sandstone with gradational, wavy-bedded base and progressively more steeply inclined laminae in cross-stratified sets. Topped by wavy and ftaser-bedded unit.

Volume fraction: 8%.

Fining-upward unit with erosive, cross-stratified sandstone at the base, passing upward into heterolithic ripple-laminated, wavy flaser-bedded sandstone. Swaley sandstone beds not uncommon.

t...,

....k

po

TABLE 1 (continued) Facies association

7. Inshore estuarine facies

8. Heterolithic tidal-channel facies

Lithofacies

Dominant texture and sedimentary structures

Dominant ichnofabrics

6.2

Swaley or hummocky cross-stratification, formed by mm-thick, fine-grained and well-sorted micaceous laminae. Beds have sharp bases and gradual tops.

Absent

6.3

Composed of poorly and variably sorted, medium-grained and small-scale cross-stratified sandstone beds with a chaotic texture in which double mud drapes are not uncommon, and heterolithic beds formed by ripple-laminated fine-grained sandstone and mudstone layers, often in doublets (Table 2; Figs. 14 and 15). Heterolithic beds are dominated by wave-ripple structures but current ripples are also present. Bioturbation causes feather-like appearance of mud layers.

Planolites, resting traces, Skolithos, excavation traces, Palaeophycus, Teichichnus, Phycosiphon, Rhizocorallium, Chondrites, Siphonichnus; low to

7.1

Relatively well-sorted, fine-grained, simple flaser-bedded sandstone. Small-scale dunes at base and ripples at the top, set size decreasing upward from 23 to 2 cm (on average 3.5 cm; Fig. 16). Flasers vary considerably in length, occasionally double mud drapes occur, and rarely rooted horizons are found.

Escape traces, Gyrochorte, Planolites, Berguaeria.

Wavy flaser- or flaser-bedded sandstone, or small-scale cross-bedded sandstone with flasers. Topped by distinctly wavy-bedded unit.

7.2

Wavy bedded heterolithic facies formed by fine- to coarse-grained sandstone layers regularly alternating with mudstone layers with sharp boundaries (Table 2; Figs. 7, 15 and 17). Mud layers mostly cover single current-ripple sets. Sandstone layers are parallel laminated, or sigmoidal cross-stratified. Erosion surfaces, typically lined with small pebbles and organic debris, are common.

Bioturbation rare (Planolites, Palaeophycus)

Thickness 7.1" average = 3.2 m; Std. = 2.3. Thickness 7.2: average = 3.0 m; Std. = 1.6. Volume fraction: 18%.

8.1

Poorly sorted, coarse-grained, tabular and trough cross-stratified sandstone upward-fining from coarse- or medium-grained to fine-grained sandstone with common reactivation surfaces (Fig. 18). Often, up to 30 cm thick pebble lags occur at the base.

Uncommon horizons with Diplocraterion

Generally fining-upward succession formed of trough and tabular cross-stratified sandstone and inclined mud layers in the upper heterolithic part.

8.2

Strongly heterolithic facies formed by alternations of mudstone and fine- to medium-grained sandstone layers (Table 2). No gradual changes in inclination angle of mudstone layers. Dm-thick siltstone intervals at top.

Escape traces, Gyrochorte, Planolites, Berguaeria, Diplocraterion (large burrows), Planolites, resting traces, Skolithos, faecal pellets.

Total thickness: average = 5.1 m; Std. = 5.7. Volume fraction: 6%.

Diagnosis

Total thickness: average -- 4.5 m; Std. -- 4.3. Volume fraction: 16%.

moderate bioturbation intensity.

TABLE 1 (continued) Facies association 9. Tidal-channel facies

t,,..

Lithofacies

Dominant texture and sedimentary structures

Dominant ichnofabrics

Diagnosis

9.1

Variably sorted, fining-upward unit from coarse- to medium- or fine-grained sandstone (Fig. 19). Erosive base, often with granule or pebble lag. Siderite rip-up clasts and granules occur at base. Large, low-angle trough cross-stratified or parallel-laminated sets, up to 35 cm thick common. Occasionally, thin organic-rich layers and oxidised pyrite occur.

Almost no trace fossils.

Multiple units of fining-upward, trough cross-stratified, very coarse- to medium-grained sandstone with scoured base. Topped by heterolithic, wavy-bedded unit.

Medium- to coarse-grained, poorly sorted, angular to subangular sandstone organised in large-scale cross-stratified sets showing progressively steeper inclined laminae. Organic-rich mud streaks and layers mostly in sections with lowest laminae inclination.

Skolithos, Diplocraterion.

Heterolithic, wavy-bedded, ripple-laminated fine-grained sandor silt- and mudstone (Table 2; Fig. 7). Organic fragments occur. Commonly, the frequency of mud- and/or siltstone layers diminishes towards the top, and grain size of individual layers and lenses may vary.

Diplocraterion (large burrows common at the top of upward-fining units), Planolites, escape traces, resting traces, Skolithos, Berguaeria, faecal pellets, Siphonichnus.

10.1

Upward-fining sandstone, average thickness 60 cm (but up to 3 m), from fine-grained at the base (with some coarser-grained laminae) to very fine-grained at the top. Rather indistinct cross-stratified structures, each fining-up and decreasing in size upward. Erodes occasionally into thin, in-situ coal layer and is sharply covered by coal layer.

Sparse and indistinct.

10.2

Silt and (very) fine-grained sandstone. Dark, organic-rich mud layers or cm-thick carbonaceous layers. Mostly homogenised; parallel lamination in coaly layers.

Rooted sandstone layers with unidentified trace fossils.

10.3

Silty mudstone with swirly texture, frequently with thin, ripplelaminated sets resulting in lenticular bedding. Root traces are common and relatively thick siderite layers may occur in the top.

Monotypic, unidentified trace fossils.

9.2

9.3

10. Lower delta-plain facies

~,~~

~..~~

t..., ~,~~

~,~~

Total thickness: average = 4.4 m; Std. = 4.9. Volume fraction: 5%.

r Homogeneous, fine-grained sometimes rooted sandstone with coal layers, topped by black, silty and finely laminated mudstones.

Total thickness: 10.2; average = 2.7 m; Std. = 1.7. Volume fraction: 2%.

Volume fraction is expressed in percentage of the normalised total amount of metres of core studied. Std. -- standard deviation.

too

~,,~~

Fig. 6. Sedimentological summary log of the Tilje Formation in the Heidrun Field.

Sedimentology of the heterolithic and tide-dominated Tilje Formation

115

Fig. 7. Comparison of four heterolithic facies of the Tilje Formation (examples taken from the Heidrun Field; see also Table 2). 1 -- lithofacies 3 (laminated delta-front deposits); 2 = lithofacies 4.2 (delta-front lobe facies); 3 = lithofacies 7.2 (inclined, accretionary tidal point-bar deposits); 4 = lithofacies 9.3 (accretionary tidal-channel margin deposits). See text for further discussion. All (a) taken with normal light, all (b) taken with ultraviolet light; core sections are 1 m long.

stratified and commonly amalgamated beds of lithofacies 1.3 (Table 1) are interpreted to represent deposition during episodic storm events (cf. Dott and Bourgeois, 1982; Leckie and Krystinik, 1989). The ichnofabric of lithofacies 1.3, being relatively poor in species (Table 1), and the palynomorph fauna are typical for offshore sand layers and indicate a lower shoreface setting. In conclusion, deposition probably occurred in a prodelta setting between storm wave base and fair-weather wave base, reflecting periodic influxes of terrestrially derived coarse silt and very fine-grained sand. The lowermost part of the Tilje Formation was storm dominated. Sand is interpreted to have been derived by erosion of nearby lower delta-plain and delta-front deposits through seaward-returning high-energy wind-driven flows.

Facies association 2 (transgressive facies) Facies association 2 is developed in three zones in the Heidrun Field that are several kms wide and extend over the full N-S length of the field. The first occurrence (lithofacies 2.1) is located at the base of the formation, the second at the base of Tilje 2.2

(lithofacies 2.1), and the third composes the entire Tilje 4 reservoir zone (dominantly lithofacies 2.2). In the SmOrbukk Field, the association is found at the base of the formation, and at the base of Tilje 2 and 6 (all occurrences of lithofacies 2.1). The matrixsupported, muddy and homogenised character of the sandstones of facies association 2 (Table 1), together with its generally thin (0.5 to 2.5 m, with the exception of Tilje 4 in the Heidrun Field) but widespread (on the scale of kms) occurrence at specific stratigraphic intervals, imply reworking processes. This association typically is found at the transition from shallow, marginal marine sediments (for example facies 7.2, discussed hereafter) to open marine environments typified by normal salinities and relatively deep water (for example association 1 and/or 4). The poor sorting and local abundance of angular calcite grains reflect mixing of mud, sand, and shell debris. These features suggest that the association formed during a relatively rapid transgression in a shoreline and nearshore shelf environment. The transgressive events may be linked to tectonic causes, possibly associated with syn-sedimentary fault movements in the late-stage pre-rift basin.

11 6

A. W. Martinius et al.

TABLE 2

Characteristic features of the various types of dominantly heterolithic facies encountered in the Tilje Formation Lithofacies

3

4.2

5.1

Grain-size variation between layers

mudstone and coarse siltstone to very fine-grained sandstone

mudstone and fine- to medium-grained sandstone

mudstone and fine-grained sandstone

Grain sorting (coarse fraction)

good

fairly good

poor

Flow structures (coarse layers)

wave (dominant) and current ripples, single sets between mudstone layers, synaeresis cracks common

wave and current ripples, first dominate and change upward into small dunes

wave (dominant) and current ripples

Colonisation style

type A: mixed shallow to mid-tier deposit-feeding and suspension-feeding community (equilibrium colonisation), typical of T1.2 (H) type B: short-lived deposit- and suspension-feeding community (opportunistic colonisation), typical of

shallow to mid-tier deposit-feeding community (equilibrium colonisation)

mixed shallow to mid-tier deposit- and suspension-feeding community

n.a.

sandstone on average

T3.1 (H)

Layer statistics

see Fig. 10

11 ram, mudstone 7 mm

Frequency of occurrence Thickness Main stratigraphic position(s) Lateral extent

11% 3.1 m/2.6 H: T1.2, T3.1, T4 S: top Are Fm.; base T3.1 kms (up to 10 km?)

2% 3.7 m/2.3 H: top T1.2; base T2.2; T2.5 S" T1.2, top T2 and T3.1 100s of m

1% n.a. H: T3.2; T3.3 S" TI.1; T3.2; T4 to T6 10s-100s of m

For additional facies descriptions and trace-fossil names, see Table 1. Frequency of occurrence is volume fraction of the normalised total amount of metres of core studied. Descriptive layer statistics includes average (in mm), standard deviation (Std), and net-to-gross (N/G); the thickness row gives average (in metres) and Std. The main stratigraphic position refers to Fig. 20 and a discussion follows in the section on the sequence-stratigraphic framework, n.a. = not available; H = Heidrun Field; S = SmCrbukk Field; T = Tilje.

Lithofacies 3 (laminated delta-front facies) The stratigraphic occurrences of lithofacies 3 are listed in Table 2. The most striking characteristic of this lithofacies is its well-developed and persistent wavy bedding (cf. Reineck and Wunderlich, 1968) with approximately equal thicknesses of mudstone and siltstone or very fine-grained sandstone layers (Tables 1 and 2; Figs. 9 and 10). Both wave- and current ripples are present, indicating relatively lowenergy conditions during deposition. Regularly occurring, up to 5 cm thick, fining-upward (medium- to fine-grained) sandstone beds with sharp, parallel-laminated base and ripple-laminated top are interpreted as the products of minor storm events. The trace-fossil suite is typified by two colonisation styles. The first (Type A; Tables 1 and 2) is characteristic of T1.2 in the Heidrun Field and the uppermost part of the ,~re Formation in the Sm0rbukk Field is indicative of subtidal, species-rich, equilibrium communities. The second (Type B; Tables 1 and 2) is characteristic of T3.1 in both fields and indicative of a subtidal relatively species-poor, opportunistic colonisation style reflecting hostile conditions possibly imposed by high deposition rates and/or variable salinity. In conjunction with in-house palynofacies data, these differences in ichnofabric development may be a function of paleosalinity whereby the second colonisation

style could be associated with greater terrestrial input. Couplets are a common constituent of lithofacies 3 (see also Dreyer, 1992). The development of a series of couplets (one couplet is defined as a pair of layers consisting of one sandstone or siltstone and one mudstone layer) is generally interpreted as being the result of deposition in tidal regimes (Reineck, 1967; Reineck and Wunderlich, 1968; Terwindt and Breusers, 1972; Boersma and Terwindt, 1981; Baker et al., 1995; Gastaldo et al., 1995). Published examples indicate that couplet successions often do not show direct evidence of daily tidal cyclicity, but tend to be dominated by effects of longer-term tidal cycles (e.g. neap-spring or equinoxal cyclicity). Comparing couplet thickness variation of lithofacies 3 of the Tilje Formation with an example from the Fly River delta (Baker et al., 1995; Fig. 10), analogies are apparent although average layer thickness is twice as high in the Tilje Formation as in the Fly River delta. A common feature of lithofacies 3 are small-scale cracks developed in the mud layers that are filled with sand from the overlying layer. In most described cases, such cracks have been interpreted as either desiccation cracks if upward curling of mud polygons could be proven (cf. Van Straaten, 1959), or as subaqueous shrinkage ("synaeresis") cracks generated by

Sedimentology of the heterolithic and tide-dominated Tilje Formation

117

6.3

7.2

8.2

9.3

mudstone and upper fine- to lower medium-grained sandstone

mudstone and fine- to coarse-grained sandstone

mudstone and fine- to medium-grained sandstone

mudstone and silt- to fine-grained sandstone

variable to fairly good

good

poor

good

current ripples and small dunes with double mud drapes dominate; wave ripples present

current ripples and small-scale dunes; mudstone layers mostly cover single sets

small- and large-scale dunes

current ripples, and smalland large-scale dunes

short-lived deposit- and suspensionfeeding community present (opportunistic colonisation) + shallow to mid-tier deposit-feeding community (equilibrium colonisation) established

short-lived deposit-feeding (opportunistic) community

short-lived suspension- (at top of fining-upward units: deposit-) feeding community (opportunistic colonisation) + mixed shallow to mid-tier depositand suspension-feeding community -+ not necessarily coeval colonisations

as 8.2

see Fig. 15

see Fig. 15

n.a.

n.a.

6% 4.4 m/3.1 H: typical for T3.4; T3.2; T3.3 S: T3.2 10s-100s of m

3% 3.0 m/1.6 H: top T2.1 S: T 1.1 10s of m

3% 2.4 m/2.6 H: T3.2; T3.3 S: T3.2 to T6 10s of m

5% 1.4 m/1.6 H: T3.2; T3.3 S: T3.2

salinity changes if no decisive evidence for an intertidal flat depositional environment could be found (for example Plummer and Gostin, 1981). Astin and Rogers (1991) suggested an origin partly as gypsum crystals and partly by limited subaerial desiccation, while Pratt (1998) argued that the cracks could be formed as the result of syn-sedimentary earthquakes. The interpretation as synaeresis cracks is favoured because: (1) lithofacies 3 has a large lateral extent (at least kilometres, possibly tens of kilometres); (2) the two main occurrences of lithofacies 3 (T1.2 in the Heidrun Field and the uppermost part of the ,a,re Formation in the Sm~arbukk Field, and T3.1) are up to 5 m thick; (3) in most cases, facies association 3 overlies facies association 2, and/or is interbedded with facies association 1 and 4; (4) no evaporitic minerals or residues are found in facies association 3; and (5) the characteristics of the two colonisation styles. These arguments suggest a shallow subtidal setting and relatively open marine depositional conditions. A relation of cracks with syn-sedimentary fault movements could not be demonstrated conclusively. In conclusion, lithofacies 3 is interpreted as a quiet-water, mud-rich delta-front facies typified by alternating periods of relatively rapid and slow deposition from suspension and occasionally disrupted by deposition of somewhat coarser sediment by traction currents, or, when cross-bedded, caused by strong river floods. The delta-front facies zone may have protruded into the distal parts of distributaries which

were not serving as the main sediment fairways. Sedimentation patterns are inferred to have been governed by tidal processes, possibly associated with seasonally driven periodic influxes of freshwater and the generation of hyperpycnal plumes. Wave sheltering to promote couplet formation was probably provided by the narrow and elongate shape of the seaway.

Facies association 4 (delta-front lobe facies) Facies association 4 (Table 1) is commonly found in association with facies associations 1 and 3; its stratigraphic distribution is listed in Table 2. Lithofacies 4.1 (Fig. 11) is typified by the interbedding of sandstone and relatively thin and irregularly spaced mudstone layers, whereas lithofacies 4.2 is typified by abundant, relatively thick and more regularly spaced mudstone layers with a variable thickness (Fig. 12). Unlike lithofacies 3 and facies association 7 (discussed below), no regular interbedding of thin mudand sandstone layers is encountered. The bioturbation intensity of the association decreases upward. The trace-fossil suite is indicative of a mixed suspension and deposit-feeding community, but the assemblage in lithofacies 4.2 is more varied than that of 4.1. It is envisaged that lithofacies 4.2 is the distal equivalent of lithofacies 4.1. Both lithofacies 4.1 and 4.2 commonly exhibit a coarsening-upward grainsize profile expressed by an upward decrease in the number of mud layers, an increasing proportion of sand, and a slight increase of grain size (Figs. 11

118

A. W. M a r t i n i u s et al.

Fig. 8. Core photograph of facies association 1. Example taken from the Heidrun Field.

and 12). This trend is evident on gamma-ray log profiles. Although wave-generated ripple structures are dominant, structures indicative of tidal currents (single and double mud drapes, mud layers) are present. Tidal indicators are found in association with significantly coarser-grained, thin sandstone lenses (in some places with small pebbles; Table 1). The thick mud layers may be formed as a result of high suspended-sediment concentrations (for example as fluid mud layers) during one slack-water period, coupled with rapid consolidation after deposition to prevent significant erosion, or from amalgamation of mud layers formed during several slack-water periods.

Deposition of facies association 4 is envisaged to have occurred at low-relief delta-front barforms with a lobate geometry in front of the mouth of a distributary channel and associated with the outcoming flow. Facies association 3, in contrast, is envisaged to have been formed more distal and/or outside the direct influence of an active river mouth. Based on the more prominent coarsening-up grain-size trend, a somewhat shallower and more proximal setting is envisaged for lithofacies 4.1 as compared to lithofacies 4.2. The coarse-grained layers are inferred to have been formed as the result of deposition in the vicinity of a fiver mouth where deposition of fine-grained

Sedimentology of the heterolithic and tide-dominated Tilje Formation

119

Fig. 9. Lithofacies 3: well-log responses and core description (example taken from the Heidrun Field). CALl = calliper (bore-hole width in inches); KLHEST = horizontal estimated permeability (mD); GR = gamma ray (API); NPHI = neutron/porosity (fraction); RHOB = bulk density (g/cm3); DT = sonic (ms/ft); MD/RKB = measured depth/rotary kelly-bushing.

sediment was punctuated by discharge of coarser sediment during periods of increased run-off.

Facies association 5 (estuarine and delta-front bar facies) The stratigraphic distribution of facies association 5 is listed in Table 2. The most distinctive feature of this heterolithic facies association is its regular stacking of tabular, overall upward-coarsening, sets, each showing horizontally layered mudstones at the base and grading to cross-stratified sandstones via a continuum of progressively sandier and more steeply inclined heterolithic beds (Table 1; lithofacies 5.1; Fig. 13). Two different bedding-pattern styles occur in lithofacies 5.1. In the first case, individual set height generally does not exceed a few decimetres and sets stack to form a succession of a variable number of sets. These are found mainly in the Tilje 3.2 and 3.3 stratigraphic intervals in the Heidrun Field and the Tilje 3.2 and 4 stratigraphic intervals in the Sm0rbukk Field. In the second case, individual sets may reach a thickness of up to 2 m and may stack to form a succession of up to 15 superimposing sets to form barforms of several metres thick in some places. In addition, double mud drapes are locally preserved. These occurrences are mainly found in the Tilje 1.1 stratigraphic interval in the Sm0rbukk Field. Thick mud layers in bottomsets and thinner mud lay-

ers draping individual foresets are interpreted as an indication of tidal influence. The low degree of bioturbation and the low trace-fossil diversity, with both (opportunistic) suspension and deposit feeders, may be related to a combination of a high sedimentation rate and a salinity-stressed setting. Frequently, lithofacies 5.1 is underlain by fine- to medium-grained sandstone beds with erosive base. These beds show vague trough cross-stratification at the base and waveripple lamination in the upper parts (lithofacies 5.2). The upward-coarsening trend and the large-scale tabular cross-stratified sets of lithofacies 5.1 (Table 1; Fig. 13) are thought to have been generated by the migration of dunes with muddy toesets, which together formed subtidal sand bars. Occasionally, dune migration was preceded by scouring and subsequent filling of the scours. The two styles of dune superposition are likely to reflect geometrical and size differences between the two barform types. Bars composed of relatively small dunes are interpreted to have been formed in distributary channels, whereas bars composed of relatively large dunes and flaser-bedded (cf. Reineck and Wunderlich, 1968) intervals are inferred to have been formed at the mouth of an estuary. The barforms appear to have been fully submerged. In plan view they may have had an elongated or ribbon shape, with their long axis approximately parallel to the tidal flow (cf. Kenyon, 1970), or a disc shape without a preferred flow-oriented axis. Superimposed

120

A. W. Martinius et al.

Fig. 10. Descriptive statistics of a series of successive couplets of lithofacies 3 (example from the Heidrun Field) compared with the descriptive statistics of delta-front heterolithics of the Fly River delta (core 111; after Baker et al., 1995). One couplet is defined as a pair of layers consisting of one sand/silt- and one mudstone layer.

dunes migrated during periods of maximum tidal flow, which was replaced by erosion, ripple migration and the deposition of mud during other parts of the tidal cycle. In many cases, facies association 5 is topped by a fining-upward wavy flaser-bedded (cf. Reineck and Wunderlich, 1968; Terwindt, 1981) sandstone unit (lithofacies 5.3). This unit is interpreted to have formed as the result of (lateral) migration of tidedominated mid-channel or bank-attached barforms with the preservation of flaser bedding on (muddy) sand flats (Reineck, 1967; Terwindt, 1981) in rel-

atively shallow water. Intercalated coarse-grained, fining-upward sets may reflect periods of higher discharge (floods) from the feeder system. Occurrences of beds resembling facies association 5 in the Njord Field have previously been interpreted as large migrating bedforms, in some places forming part of subtidal shoals developed in inshore inlets, although no tidal-channel facies were recognised (Van de Weerd, 1996). Similar barforms have been described by Maguregui and Tyler (1991) from the Eocene Frag area of the Lagunillas Field (Lake Maracaibo, Venezuela) as distal tidal sand ridges.

Sedimentology of the heterolithic and tide-dominated Tilje Formation

121

Fig. 11. Lithofacies 4.1: well-log responses and core description (example taken from the Heidrun Field). See Fig. 9 for abbreviations.

Facies association 6 (channelised delta-front lobe facies) Facies association 6 is composed of three lithofacies, displays an overall upward-fining grain-size trend, and is specifically characteristic of T3.4 in the Heidrun Field. Additional stratigraphic locations are listed in Table 2. Lithofacies 6.1 at the base exhibits a well-developed upward-fining grain-size trend and a distinct vertical change in bedform size from small-scale cross-stratification, often underlain by a lag composed of granules and siderite rip-up clasts and interpreted as dunes, through flaser-bedded intervals, in some places with double mud drapes, to clean, ripple-laminated sand at the top (Table 1). Common flaser and wavy-flaser bedding, double mud drapes, and clay layers are interpreted as tidal indicators but lithofacies 6.1 is dominated by wave-generated bedform structures (Table 1). The impoverished trace-fossil suite of lithofacies 6.1 indicates an opportunistic colonisation style and relatively stressed conditions, possibly caused by frequent salinity changes resulting from fluvial discharge fluctuations, and/or high sedimentation rates. Frequently occurring hiatal colonisation indicators (Diplocraterion; Table 1) suggest non-coeval colonisation styles. In some cases, lithofacies 6.1 includes, or is overlain by, relatively thin (up to 2 or 3 dm thick) swaley or hummocky cross-stratified sandstone beds (lithofacies 6.2) interpreted as the result of minor storm events (compare with lithofacies 1.3).

The overlying lithofacies 6.3 (Tables 1 and 2) is dominated by heterolithic deposits but at the base it also contains medium-grained, small-scale crossstratified sandstone often with double mud drapes. Sedimentary structures are interpreted to have been formed by migrating small dunes and current tipples that were tide influenced. Lithofacies 6.3 resembles lithofacies 3 but the volume fraction of sandstone is significantly higher (Table 2; Figs. 14 and 15). Wave ripples are found to regularly erode into the previously deposited mud layer. The sedimentary structures resemble the regular internal stratification of successions found in the lower reaches of the Oosterschelde and interpreted as shoals (Boersma and Terwindt, 1981; Van den Berg, 1981). The ichnofabric structures of lithofacies 6.3 suggest the presence of both a short-lived deposit and suspension-feeding community with an opportunistic colonisation style as well as a shallow to mid-tier deposit-feeding community indicating an equilibrium colonisation style (Tables 1 and 2). Facies association 6 is envisaged to have been formed associated with channel thalwegs that were located at the seaward end of delta distributaries. These were occasionally affected by storm-related wave reworking. Initially, shallow channels were eroded during outflow periods and subsequently filled with progressively finer sediment and migrating dunes forming complex barforms. Subsequently, finer-grained and more heterolithic deposits formed on top as the

IX3 IX3

r..,~

Fig. 12. Core expression of a proximal prodelta lobe (lithofacies 4.2) in the Heidrun Field; dashed lines indicate base and top. Storm-dominated shelf facies of facies association 1 occur at the base and top of the succession shown. Note the gradual coarsening up and the, in some places, regular alternation of mud- and sandstone layers (e.g. 3464.7-3464.1 m).

Sedimentology of the heterolithic and tide-dominated Tilje Formation

123

Fig. 13. Lithofacies 5.1: well-log responses and core description (example taken from the SmOrbukk Field). See Fig. 9 for abbreviations.

Fig. 14. Lithofacies 6.3: well-log responses and core description (example taken from the Heidrun Field). See Fig. 9 for abbreviations.

result of the development of channel margin and inter-channel deposits due to laterally or obliquely migrating channel thalwegs. Alternatively, distributary channels may have been abandoned and disconnected from fluvial input thereby becoming tide-dominated, and gradually became filled in.

Facies association 7 (inshore estuarine facies) The stratigraphic position where facies association 7 most commonly occurs is in Tilje 2.1 in the Heidrun Field and Tilje 1.1 in the SmCrbukk Field. Furthermore, it is found in Tilje 3 in the SmOrbukk Field.

A. W. M a r t i n i u s et al.

124

Lithofacies 6.3 8070E

E 60O9

O9 tD c-

50-

o 40-

.i c-

.~ 30Q_

= 200

0

I00-

10

0

20

30

Couplet number

Lithofacies 7.2 35~" 30 E o9 2 5 -

v

O9 ID

c 200

~15.i-., ID

~10O

O

50 0

10

20

30

40

Couplet number

n sum avg geomean var std N/G

= = = = = = =

Lithofacies 6 . 3 Mudstone Sandstone 30.0 30.0 90.0 mm 431.0mm 3.0 mm 14.4 mm 2.5 mm 9.9 mm 2.9 172.4 1.7 13.1 0.83

Lithofacies 7 . 2 Mudstone Sandstone 44.0 45.0 189.0 mm 290.0 mm 4.2 mm 6.3 mm 3.1 mm 4.2 mm 9.7 29.7 3.1 5.4 0.61 J

Fig. 15. Descriptive statistics of a series of successive couplets of lithofacies 6.3 and 7.2 (examples from the Heidrun Field). One couplet is defined as a pair of layers consisting of one sand/silt- and one mudstone layer.

Particularly in Tilje 2.1 in the Heidrun Field, lithofacies 7.1 (Table 1) is laterally extensive and uniform, limited in its occurrence to an approximately 3 km wide facies belt (deduced from a geomorphologic reconstruction based on 29 Heidrun wells). Lithofacies 7.1 occurs on top of a sharp erosive surface which forms the stratigraphic separation between Tilje 1.2 and 2.1 (see below). Each flaser of lithofacies 7.1 (Fig. 16) is thought to have formed as the result of partial erosion of a mud layer by an ebb or flood tidal current, whereby the mud layer was deposited during the foregoing slack-water period. Mud was preserved only in the deepest parts of a ripple trough (Terwindt, 1971; Terwindt and Breusers, 1972, 1982). Vertical alternations in the number and thickness of

flasers can be attributed to neap-spring tidal cycles, and the length and width of these flasers depend on the size and spacing of the ripples (Terwindt, 1971; Terwindt and Breusers, 1982). The sedimentary structures of lithofacies 7.1 are interpreted to be indicative of an environment in which both waveand tidal-current processes were operating. In modern environments, flaser bedding is frequently interpreted to reflect deposition in the deeper, subtidal parts of ebb- or flood-dominated channels (up to 30 m deep in tidal channels), similar to deposits found in estuaries like the Oosterschelde (Reineck, 1967; Terwindt, 1971; Boersma and Terwindt, 1981; Van den Berg, 1982; Terwindt and Breusers, 1982) or Willapa Bay (Clifton, 1983; Clifton and Gingras, 1997). As the result of (lateral) migration of a tidal channel, one of the channel banks will preserve flaser bedding (Reineck, 1967; Terwindt, 1981; Van den Berg, 1982). Lithofacies 7.2 (Table 2; Figs. 7, 15 and 17) is typified by wavy-bedded heterolithic facies formed by on average medium-grained (but ranging from upper fine- to lower coarse-) sandstone layers regularly alternating with mudstone layers with sharp boundaries. This lithofacies resembles lithofacies 3 (Fig. 7) but three characteristics are significantly different. Firstly, lithofacies 7.2 almost exclusively occurs in the top of Tilje 2.1 in the Heidrun Field, where its is always found on top of lithofacies 7.1. A short (5 to 30 cm) but gradual interval typifies the transition. Secondly, the grain size of the sandstone layers is considerably coarser (sandstone of lithofacies 3 is very fine grained). In addition, the net-to-gross is higher due to, on average, thicker layers of the coarse grain-size fraction of lithofacies 7.2 (Figs. 10 and 15). Thirdly, bioturbation and recognisable trace fossils are rare, suggesting a stressed environment (lithofacies 3 as occurring in Tilje 1.2 is characterised by a subtidal, species-rich, equilibrium community). Lithofacies 7.2 is envisaged to have been deposited as inclined, laterally accreted point bars within meandering tidal channels in a shallow, possibly intertidal environment. Erosion surfaces occurring in successions of lithofacies 7.2 with their associated lag deposits may have resulted from small run-off channels draining the intertidal flats that existed at the top of the accretionary channel bank. The characteristics of lithofacies 7.2 are similar to heterolithic deposits described from, for example, Willapa Bay (Clifton, 1983; Clifton and Gingras, 1997), the Lower Jurassic of Bornholm (Sellwood, 1972, 1975), the Upper Cretaceous of Alberta (Rahmani, 1988), and the Waddenzee (Reineck, 1967). Based on a comparison with these analogues, the angle of inclination of the wavy-bedded point-bar deposits of lithofacies 7.2 is inferred to have been less than 5~.

Sedimentology of the heterolithic and tide-dominated Tilje Formation

125

Fig. 16. Lithofacies 7.1: well-log responses and core description (example taken from the Heidrun Field). See Fig. 9 for abbreviations.

Fig. 17. Lithofacies 7.2: well-log responses and core description (example taken from the Heidrun Field). See Fig. 9 for abbreviations.

Facies association 7 is interpreted as representing the regressive fill of the inshore part of a relatively large but shallow estuary. The association exhibits, in ideal cases, an upward transition from compound, aggradational or stable subtidal bars and accretionary channel banks (classified as elongate tidal bars cf. Dalrymple and Rhodes, 1995), to heterolithic tidal

point-bar deposits (classified as repetitive barforms cf. Dalrymple and Rhodes, 1995) formed in meandering channels during the final stages of infilling. The intimate relationship between these two genetically related sub-environments of deposition has been documented by several authors (for example Sellwood, 1975; Terwindt, 1981; Clifton and Gingras,

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1997). The rather uniform grain size and good sorting is interpreted to reflect dominant sourcing from the seaward side of the system.

Facies association 8 (heterofithic tidal-channel facies) The stratigraphic positions of facies association 8, which is composed of two lithofacies, are listed in Table 2. Lithofacies 8.1 at the base is underlain by an erosive basal surface above which upward-fining, dm-scale sandstone units are present. The sandstone contains occasional mudstone clasts, quartz pebbles and organic debris, and oppositely migrating sets and foreset mud drapes are common (Table 1; Fig. 18). These units are interpreted to have been formed by migrating simple and compound dunes. In some places, inclined units with alternating mud- and sandstone layers directly overlie a basal pebble lag although more commonly this heterolithic lithofacies is restricted to the upper parts of the association (lithofacies 8.2). The inclined mudstone layers are on average up to 3 cm thick and separated by up to 20 cm sandstone layers, and interpreted as (lateral or oblique) accretion surfaces. The uppermost part of the facies association (lithofacies 8.2) is typified by heterolithic, ripple-laminated aggradational channelmargin deposits formed by medium-grained, up to 2 cm thick sandstone layers interbedded with up to 1 cm thick silt- or mudstone layers (Table 1). The presence of mudstone clasts, pebbles and organic debris indicate that erosion was prevalent

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during periods of relatively high flow competence. Combined with the abundance of good tidal indicators as, for example bi-directional foresets and mud drapes, and fining-upward tendencies at set- and bed scales, it is suggested that these deposits most likely represent the products of heterolithic, possibly sinuous, tidal channels.

Facies association 9 (tidal-channel facies) Deposits of facies association 9, dominantly found in Tilje 3 in both the Heidrun and the SmCrbukk Field (Tables 1 and 2), are characterised by two superimposed upward-fining successions: one smallscale variant (1 to 2.5 m) comprising occurrences of lithofacies 9.1 (Fig. 19), and a large-scale variant (4 to 5 m) at the scale of the association. Just above the erosive base of the larger fining-upward units, the sandstones are clean, poorly sorted, and coarseto very coarse-grained. In the central parts, they are slightly better sorted and medium- to coarse-grained, while in the upper parts they are fairly well sorted and fine- to medium-grained. A great variety of sedimentary structures are present, their sizes generally decreasing upwards. Large-scale trough and tabular cross-stratification dominates the lower sections, whereas ripple lamination and small-scale cross-stratification are more common in the upper sections. Ripples and cross-stratified sets commonly show opposite migration directions, and they are often separated by mud drapes. In general, bioturbation is sparse, and it is commonly represented by large, iso-

Fig. 18. Facies association 8: well-log responses and core description (example taken from the Heidrun Field). See Fig. 9 for abbreviations.

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Fig. 19. Lithofacies 9.1: well-log responses and core description (example taken from the Heidrun Field). See Fig. 9 for abbreviations.

lated, Diplocraterion burrows indicating a non-coeval opportunistic colonisation style. Frequently, towards the top the succession passes via a series of bedforms interpreted as simple dunes (lithofacies 9.2) into lithofacies 9.3, becoming increasingly more heterolithic (Table 2). The ichnofabrics indicate the presence of a short-lived suspension-feeding community with an opportunistic colonisation style. The large-scale fining-upward successions were probably formed by migration of low-sinuosity or straight tidal channels, mudstone clasts being introduced by erosion of channel banks. The smallscale fining-upward units may result from waning flows within the channels in combination with lateral channel migration causing preservation of channelbank deposits. Tidal influence is most evident in the heterolithic deposits that contain numerous mud drapes and abundant indicators of bimodal currents such as herringbone structures. Double mud drapes suggest deposition as simple dunes in subtidal environments experiencing tidal cyclicity. The tidal channels experienced rapid fluctuations in current energy and water level resulting in the occurrence of periodic slack-water episodes leading to deposition of mud. The low degree of bioturbation and limited burrow diversity, indicating stressed environments, is interpreted to reflect a low salinity level, resulting from mixing of marine and fresh water, in a rather proximal, perhaps fluvially dominated setting. Some occurrences of facies association 9 are almost devoid of tidal indicators, and are often accom-

panied by the occurrence of thin gravel layers on top of the channelised base. These examples are interpreted as dominated by fluvial processes and represent the most proximal facies encountered in the Tilje Formation.

Facies association 10 (lower delta-plain facies) Well-developed examples of this association are found in Tilje 3.1 and 3.2 of the SmOrbukk Field and in the majority of cases in the stratigraphic equivalents of Tilje 3 to 6 in the Lavrans and Njord Fields. The association is composed of an up to 3 m thick fine- to very fine-grained and fining-upward sandstone unit (lithofacies 10.1; Table 1), overlain by either a single coal layer, a thin succession of interbedded thin sandy coal and sand layers (lithofacies 10.2), or a rooted silty mudstone (lithofacies 10.3). The basal sandstone layer is found to erode into an underlying coaly layer in some occurrences. In the Lavrans Field, rooted horizons are often preserved on top of flaser-bedded sandstones, interpreted as part of a tidal-channel abandonment fill possibly in medial and distal settings of distributary channels. The siltand mudstone beds are found immediately below coal layers, or below or above the fining-upward sandstone units, commonly with a gradational basal contact. Assuming only primary deposition and no reworking by Ca-rich fluids, the relatively thick siderite-cemented upper parts are interpreted to indicate brackish or freshwater conditions depleted of Ca 2+, rich in Fe z+,

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and related to bacterial decomposition of organic matter (Curtis and Coleman, 1986). The position of the sandstone layer between two coal layers, the near absence of cross-bedding, the erosive base and fining-upward grain-size trend, which indicate waning and non-stationary depositional conditions, suggest deposition as crevassesplay deposits. In conjunction with the other observations, an origin on the lower delta plain in swampy environments is envisaged.

Key stratal surfaces Recognition of key stratal surfaces serves to present the conceptual depositional model of the Tilje Formation in a sequence stratigraphic framework. Identification of key stratal surfaces is based on the interpretations of the above-described facies associations, their stratigraphic distribution and stacking patterns, biostratigraphic correlations, and 147Sm/144Nd and 143Nd/144Nd isotopic data. Two maximum flooding surfaces (MFS), ten flooding surfaces (FS), and two sequence boundaries (SB) have been recognised in the Tilje Formation in the Heidrun-Sm0rbukk area. They represent a second-order sequence-stratigraphic framework within an overall transgressive Early Jurassic succession (first-order framework; Haq et al., 1987; Hallam, 1988; Surlyk, 1990)and are described hereafter. The stratigraphic position of key surfaces is summarised in Fig. 20.

Second-order sequence boundaries In the Heidrun Field, SB2, underlying Sequence 2 (Fig. 20), is expressed by an incision surface, which erodes into mud-rich delta-front facies. In a 2D crosssection through part of the Heidrun Field, the incisive nature of SB2 is evident (base Tilje 2.1; Figs. 21 and 22). It is also seen to be relatively shallow (depth up to 14 m) as inferred from the almost completely preserved fill. Although cored at a few places only, it is observed that at those locations where no fluvially influenced tidal-channel deposits are preserved, the incision surface is sharp and covered with very coarse and poorly sorted sediment that contains sideritised mudstone clasts (Fig. 22). The surface itself was found to be penetrated by large Diplocraterion burrows in the most distal location cored, where wavy-flaser-bedded (cf. Terwindt, 1981) sandstones are preserved above the surface. These sandstones are typified by relatively abundant freshwater pollen taxa. In-house Sm/Nd isotopic data indicate a marked change in provenance age from 1500-1800 Ma below to 1200-1400 Ma above the incision surface. In the Sm0rbukk Field, SB2 at the base of Tilje

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1.1 is characterised by a marked transition from laminated delta-front facies (lithofacies 3) overlain by shallow subtidal, dominantly wavy-flaser-bedded sands of facies association 7 and intersected by small tidal channels. In places, an up to l0 cm thick pebble lag is found at the base of this succession but no paleovalleys are recognised. SB2 has an apparent dip towards the south with an approximate depth of erosion of approximately 9 m in a southwestern direction over a distance of 16 km, and approximately 4 m in a southeastern direction over a distance of 5 km. In addition, high 143Nd/144Nd ratios (in-house data) clearly indicate the presence of young igneous rocks in the source area, which became progressively more important during deposition of Tilje 1.1 (A. Dalland, pers. commun.). In both the Heidrun and Sm0rbukk Field, SB3, at the base of the Tilje 3.2 (Fig. 20), marks the incision of heterolithic tide-dominated channels (facies association 9) into intra-field wide laminated delta-front deposits of Tilje 3.1 (lithofacies 3). This pronounced lithological change, interpreted to reflect a basinward shift of facies, is supported by a preceding change in provenance ages from younger (1200-1400 Ma) below to older (1500-1800 Ma) ages above SB3 (the mid-Tilje Formation marker; Figs. 20 and 23). This change is found over the entire Heidrun-Sm0rbukk area. Additionally, above the mid-Tilje Formation marker terrestrially derived humic debris dominate, and biostratigraphic data suggest proximity to a fluvial source and freshwater environments in the catchment area. Above SB3, bioturbation becomes weak to nonexistent, and the ichnofabrics found indicate rapid deposition above, and slow deposition below SB3. At the transition from T3.2 to T4 in the SmCrbukk Field, an important grain-size change is observed that is associated with a change from heterolithic tidal-channel deposits at the top of T3.2 to very coarse-grained and poorly sorted channel sands at the base of T4 (Fig. 20). This lithological change is interpreted as a basinward shift in facies. However, not enough conclusive evidence is available to allow it to be interpreted as a (candidate) sequence boundary.

Second- and third-order (maximum) flooding surfaces A well-developed maximum flooding surface (MFS 1) is developed at the base of the Tilje Formation (Fig. 20) and forms part of Sequence 1. MFS 1 is well-cemented, thoroughly bioturbated, has a reddish colour, and is overlain by facies association 2 in the Heidrun Field. A second maximum flooding surface (MFS2; Fig. 20), part of Sequence 2, is picked within a succession formed by facies association 1 at a strati-

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Fig. 20. Sequence stratigraphic framework of the Tilje Formation in the Heidrun and Smorbukk Fields interpreted based on the combination of in-house biostratigraphic data, samarium/neodymium isotopic data, and sedimentary facies analysis. Absolute stage ages after Gradstein et al. (1995). See text for discussion.

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Fig. 21. Stratigraphic cross-section through a part of the Heidrun Field covering the Tilje 1 and Tilje 2 members. See Table 1 for lithofacies description and text for interpretations.

graphic position where the lithological characteristics indicate the most distal (as judged from grain size and degree of bioturbation) depositional position. A third maximum flooding surface is developed within the Ror Formation overlying the Tilje Formation and is part of Sequence 3. In Sequence 1, FS1/1 (Fig. 20) indicates a transition from delta-front facies (e.g. association 6) to overlying laminated heterolithic delta-front facies (association 3). Sequence 2 includes four flooding surfaces. FS2/1 (Fig. 20) is associated with a thin but laterally extensive, heterolithic and muddy sandstone layer, which separates two outer estuarine aggradational successions in the SmCrbukk Field (Tilje 1.1, containing facies associations 5 and 7). I n - h o u s e 143Nd/144Nd data support the lateral correlatability by indicating a consistent provenance signal that is younger than that of the underlying and overlying rocks. In the Heidrun area, FS2/1 is typified by the transition from facies associations 9 to 7. Between the Heidrun and Sm~arbukk Fields, the succession between SB2 and FS2/1 shows a lateral facies transition from fluvialdominated inner estuarine channel deposits (facies association 9) to heterolithic estuary-mouth deposits (lithofacies 6 and 7.1). Subsequently, the succession between FS2/1 and FS2/2 shows a lateral facies relationship from heterolithic inshore estuary deposits (lithofacies 7.2) to heterolithic estuary-mouth deposits (lithofacies 6 and 7.1). FS2/2 (Figs. 20 and 24) is the most prominent of three closely spaced (several decimetres) flooding surfaces. It is underlain by accretionary tidal-channel deposits (lithofacies 7.2), and is overlain by a well-developed, and laterally extensive, occurrence of transgressive facies associ-

ation 2 with calcite-cemented layers. Based on the isotopic composition of strontium dissolved in the formation waters, St~alum et al. (1993) were able to show the laterally extensive nature of the combined facies associations 2 and 3 above FS2/2 in the Sm~arbukk S~ar Field, in which marine pollen taxa (e.g. Nannoceratopsis) are generally common. In-house biostratigraphic data, in combination with the above-described spatial distribution of facies associations, indicate that this interval can be correlated to the equivalent stratigraphic intervals in the Sm~rbukk and Heidrun Fields. FS2/3 and 2/4, finally, are similar in nature to FS 1/ 1. Sequence 3 includes five flooding surfaces. FS3/1 occurs on top of Tilje 3 in both the Heidrun and Sm~arbukk Fields (Fig. 20). In the Heidrun Field, FS3/1 is overlain by transgressive deposits of facies association 2 and laminated delta-front deposits of facies association 3 containing abundant marine indicators (e.g. Nannoceratopsis; in-house data), whereas the deposits below FS3/1 are dominated by terrestrial fauna elements (e.g. Botryococcus; in-house data). In the SmCrbukk Field, however, delta-front deposits (associations 5 and 6) overlie distributary channel facies (association 9). In the SmCrbukk Field, FS3/2 and FS 3/3 (Fig. 20) are similar in nature and are typified by a heterolithic interval on top of a coarse-grained but upward-fining succession that is dominantly formed by facies associations 7 and 8. Inhouse biostratigraphic data indicate that FS3/2 and FS3/3 are not preserved in the Heidrun Field. FS3/4, however, is picked at the base of a thick succession (up to 5 m) formed by transgressive facies association 2. Between the Heidrun and SmCrbukk Fields, the succession between FS3/3 and FS3/4 shows a lateral

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1 {31

Fig. 22. Core expression of SB2 and the basal transgressive interval of the estuarine fill in two wells of the Heidrun Field across the Tilje 1.2-Tilje 2.1 reservoir intervals. Facies B is developed between SB2 and FS2/1 (Fig. 20) that is often a calcite-cemented layer. Note the distinct ichnofabric change which is associated with the flooding event. Well XX shows two superposed incision surfaces; see text for discussion.

facies relationship from delta-front to lower deltaplain distributary channels. FS3/5 marks the transition to the Ror Formation. In the westernmost part

of the Halten Terrace, in-house biostratigraphic data indicate a period of non-deposition between FS3/4 and FS3/5. An abrupt change occurs across FS3/4-5

~,.~~

Fig. 23. Correlation panel between 4 wells in the Sm~rbukk and 1 well in the Heidrun Field based on 147Sm/144Nd and 143Nd/144Nd isotopic data. The mid-Tilje Formation marker divides Tilje 1 and 2 from Tilje 3 to 6. Younger source rocks are found below, older source rocks are found above the marker.

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] 33

Fig. 24. Core expression of the upper surface of the estuarine fill (FS2/2) in three wells of the Heidrun Field across the Tilje 2.1-Tilje 2.2 reservoir intervals.

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from delta-front and distributary channel facies to deep marine mudstone. In-house biostratigraphic data support the time equivalence of FS3/5.

Conceptual depositional models Introduction Facies models for tide-dominated deltas are at present relatively immature (Bhattacharya and Walker, 1992; Dalrymple, 1992), whereas the estuary model is relatively mature (Dalrymple et al., 1992; Zaitlin et al., 1994). A river delta is defined as a progradational sediment body at the mouth of a river, formed of sediment supplied by the fiver, and containing fluvially influenced deposits (Dalrymple, 1999). Riverdelta deposits display an overall progradati0nal stratigraphic signature, with delta-plain and fluvial deposits overlying mouth-bar and prodelta sediments. A fiver delta is composed of three facies zones: a lower delta plain intersected by relatively wide and shallow, funnel-shaped distributaries, a delta front, and a prodelta area. Boundaries between facies zones are envisaged to not necessarily be straight and/or parallel to the average coastal trend. Following Dalrymple et al. (1992), an estuary is defined as the seaward part of a drownedvalley system which receives sediment from both fluvial and marine sources and which contains facies influenced by tide, wave and fluvial processes. The estu-

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ary is considered to extend from the landward limit of tidal facies at its head to the seaward limit of coastal facies at its mouth. Estuaries can only form during a transgression and may lie within an incised valley. Two different conceptual depositional models are proposed for the studied succession: an estuary model for Sequence 2, and a fiver-delta model for Sequence 3. Both cases are discussed below.

Estuary model Sequence 2 is characterised by a mud-rich, waveand tide-dominated estuary model (Fig. 25) formed during a rapid base-level fall and subsequent rise. The estuary model is assumed to be applicable to Tilje 1.2 to 2.3 in the Heidrun Field and Tilje 1.1 and 1.2 in the Sm~arbukk Field (Fig. 20).

Characteristics In the Heidrun Field, a fully preserved estuary fill in an axial position (for example, in well F of Fig. 21; Fig. 26) exhibits a characteristic shallowing-up vertical facies succession. In addition, the morphology of the estuary is inferred to be shallow and wide. The succession is typified by the following characteristics. (1) Thinly developed occurrences of lithofacies 9.1 at the base, interpreted to have been formed after initial incision of a fluvially dominated channel segment into delta-front deposits by aggradation.

Fig. 25. Sketch of the conceptual depositional model for a mud-rich, mixed wave/tide-dominated estuary as proposed for Tilje 2.1 and 2.2 in the Heidrun Field and Tilje 1 in the SmCrbukk Field. Numbers in boxes refer to approximate positions of facies associations (see Table 1 for lithofacies description and text for interpretation).

Q

e~

~,,,~

~,,,~

~,,~~

Fig. 26. Core expression of the estuarine fill in the Heidrun Field. Dashed lines indicate base and top.

r {31

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(2) Flaser-bedded sandstone of lithofacies 7.1 in the middle, interpreted to be derived from landwarddirected sediment transport (see discussion hereafter). This interval is interpreted to have been formed after initial flooding and during the subsequent drowning of the valley, by mid-channel accretionary bars and subaqueous channel-thalweg banks in the funnelshaped outer reaches of the estuary. These deposits are similar to deposits found in the middle and outer reaches of the modern Oosterschelde (Boersma and Terwindt, 1981; Van den Berg, 1982), and form approximately 60% of the estuary fill. (3) Highly regular, wavy-bedded heterolithics at the top, deposited on the accretionary banks of tidal point bars (lithofacies 7.2; Table 1) such as those in Willapa Bay (Clifton, 1983; Clifton and Gingras, 1997), or as the heterogeneous fill of "straight", tide-influenced channels. In the case of the Tilje Formation, relative sealevel fall and rise is assumed to be second-order to a general (first-order) period of eustatic sea-level rise, characteristic for the late Pliensbachian to early Toarcian (Haq et al., 1987; Hallam, 1988; Surlyk, 1990; Fig. 5B and Fig. 20). During rapid base-level fall, low-relief depressions were formed by erosion into the preceding wave- and storm-dominated shoreline deposits of Sequence 1 (SB2; Fig. 20). These were drowned during the subsequent transgression to form several adjacent, waveand tide-dominated estuaries. The shallowing-up vertical facies succession described above is interpreted to reflect superimposition of channel-planform patterns of successive estuary segments, the "straightmeandering-straight" morphology of Woodroffe et al. (1989) and Dalrymple et al. (1992). In some of the wells in the Heidrun Field, the base of the low-relief valley shows two superposed incision surfaces (for example, well XX of Fig. 22). Superposition of incision surfaces may be the result of the formation of terraces in the valley that were cut during punctuated base-level fall (cf. B lum, 1993). In that case, the valley-fill basal surface is more likely formed by a composite- instead of a single-incision event and, therefore, .it has been suggested that no unique surface exists to which sequence boundaries can be correlated (cf. B lum, 1993). However, in the case of well XX the transgressive facies above incision surface 1 (facies B in Fig. 22) is overlain by lithofacies 7.1 that is truncated by incision surface 2. Based on these observations, it might be speculated that the valley was filled by a multiple base-level fall and rise and, therefore, can be classified as a compound fill (cf. Zaitlin et al., 1994). Other well locations, however, show a single-incision surface (for example, well D in Figs. 21 and 22).

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Model definition The conceptual model of valley formation and filling defined here for the lower part of the Tilje Formation in the Heidrun Field assumes that valleys were formed as the result of at least two relative sealevel fall and rise cycles. The sea-level fall period of the first of these relative sea-level cycles is assumed to have been significantly longer than the associated sea-level rise period. A valley was formed as a result of the combined effect. During relative sea-level lowstand, most sediment bypassed the Heidrun area and was deposited out on the submerged parts of the shelf. Minor proportions of sediment were preserved in the deepest parts of the valley (Fig. 27A). The second (or last) relative sea-level rise period is assumed to have been significantly longer than the preceding one. The early stages of this relative sea-level rise caused flooding of the valley which resulted in infilling of the valley and estuary development (Fig. 27B). At locations where no fluvial facies are preserved, the transgressive surface and the initial flooding surface (cf. Zaitlin et al., 1994; FS2/1 in Fig. 20) coincide with the incision surface underlying the transgressive facies (SB2; Fig. 20). At locations where fluvial facies are preserved (e.g. well F of Fig. 21), the transgressive surface is either located within, or directly overlying the fluvial facies (cf. Zaitlin et al., 1994). In the latter case, the transgressive surface and the initial flooding surface coincide. The surface overlying the transgressive facies and underlying the outer estuarine flaser-bedded sands (lithofacies 7.1) is defined as the tidal ravinement surface (cf. Zaitlin et al., 1994; Dalrymple, 1997). The transgressive phase of the second-order relative sea-level rise discussed here is inferred to have been extended relative to the sea-level fall stage because it occurred during the overall (first-order) eustatic sea-level rise of the Early Jurassic (Haq et al., 1987; Hallam, 1988; Surlyk, 1990). As a result, the paleotopographic expression of the incision surface that formed during base-level fall is inferred to have been relatively moderate. Assuming constant speed of relative sea-level rise, within the low-relief valley the speed with which relative sea-level rise occurred is interpreted to have diminished through time. The reason for this is that the width of a valley is not constant from base to top but instead widens upward. In addition, the depth of a valley along its longitudinal profile is not constant either because it is deeper in a distal position than in a proximal position. If equal volumes of water per time unit are added to the valley continuously, it is assumed that the width-to-depth ratio of the estuary will gradually increase. This process will be associated with a local decrease of accommodation space and decrease of the

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Fig. 27. Model for estuary evolution during transgression based on the Tilje Formation. After initial incision relatively distal parts of the valleys are filled with fluvial channel deposits (incision stage) and seaward-derived marine sand during ongoing relative sea-level rise (incompletely filled, stage A). During the last stages of relative sea-level rise and filling of accommodation space, a local relative shallowing occurred, associated with the progressively shallower water depths in the estuary and decreasing tidal prism. This favoured progradation of heterolithic meandering tidal channels of somewhat more proximal areas of the estuary (mostly filled, stage B). Finally, ongoing relative sea-level rise overtops the estuary and the system is drowned (flooding stage). The inset shows the spatial location of the seven wells used for the construction of the stratigraphic cross-section of Fig. 21. In total, 29 wells of the Heidrun Field were used to reconstruct paleogeography at the time of incision, and define the model, fa = facies association.

tidal prism. It is proposed that this process favoured progradation of heterolithic meandering tidal channels of more proximal areas of the estuary (Fig. 27B) during filling of the estuary. Consequently, deposits of both straight and meandering tidal-channels cover the tidal bars (lithofacies 7.1). Finally, when transgression continued and the valley was overtopped, the succession was covered by a wave-reworked surface, the wave ravinement surface of Zaitlin et al. (1994), which is overlain by transgressive shoreface deposits (lithofacies 2.2, Table 1; Figs. 21, 24 and 27D). In the Heidrun Field, three narrowly spaced flooding surface can be identified (FS2/2; Fig. 20), two within and one above the transgressive shoreface deposits. A maximum flooding surface (MFS2; Fig. 20) is identified within the overlying succession (facies association 1) at a stratigraphic position where the lithological characteristics indicate a most distal (as judged from grain size and degree of bioturbation) depositiona! position.

The model proposed here (Fig. 28C) shows similarities with the flood-capped valley-fill model proposed for the Fall River Formation (Willis, 1997; Fig. 28B), although there the proportion of sandy and muddy fluvial facies is much higher and the proportion of estuarine facies is relatively small. The proposed model is significantly different from the flood-based valley-fill model of Dalrymple et al., 1992 (Fig. 28A). The fluvial channel deposits at the base of the fill are not encountered at all locations along the cross-sectional profile but only at the deepest point of incision (e.g. well F in Fig. 21) and even there it is only a lag of almost indeterminate origin. Elsewhere, the downward shift of facies is suggested by transgressive and distal estuarine facies (lithofacies 7.1) erosionally overlying layered delta-front heterolithics (lithofacies 3). The lithological composition and depositional interpretation of sediments forming the top of Sequence 1 and Sequence 2 are analogues to the facies charac-

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Fig. 28. Stratigraphic cross-sectional sketch of three models for valley-fill evolution. (A) Flooding-based model of an estuary resulting from rapid relative sea-level rise and subsequent gradual filling as relative sea-level rise slows down (Dalrymple et al., 1992, drawn after Willis, 1997). (B) Flooding-capped model of an estuary which is the result of filling during relatively slow sea-level rise as sediments fill the new accommodation, and subsequent flooding (after Willis, 1997). (C) Flooding-based and -capped model of an estuary resulting from a relatively slow transgression and estuary filling which keeps pace with increasing accommodation space. Flooding occurs after relatively sudden increase of tidal prism and decrease of tidal amplitude. See text for further discussion.

teristics and depositional setting of the lower part of the Elis Bjerg Member that spans the upper part of the Ra~vekloft Formation and lower part of the Gule Horn Formation (Neill Klinter Group; Dam and Surlyk, 1998) on Jameson Land, eastern Greenland (Fig. 2). It is proposed that SB2 of the Tilje Formation may be correlatable with SB2 of the Elis Bjerg and Ra~vekloft Members (Dam and Surlyk, 1998, their figs. 3 and 46).

River-delta model

It is suggested that Sequence 3 is characterised by a tide- and fluvial-dominated fiver delta (Fig. 29) which formed while sediment supply exceeded the combined effect of eustatic sea-level rise, sediment compaction, and the sum of tectonic and isostatic subsidence. The river-delta model is characteristic for Tilje 3 in the Heidrun Field and for Tilje 3 to 6 in the Sm~rbukk Field

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Fig. 29. Sketch of the conceptual depositional model for a prograding mud-rich, mixed tide/fluvial-dominated river delta as proposed for Tilje 3 in the Heidrun Field and Tilje 3 to 6 in the SmOrbukk Field. Numbers in boxes refer to approximate positions of facies associations (see Fig. 1 for lithofacies description and text for interpretation).

(Fig. 20). However, although one model is applied, the delta is differently developed in the Heidrun and Sm0rbukk Fields (see the discussion section). At the largest scale, the succession evolved in four stages, of which the lower two are developed in both the Heidrun Field (Tilje 3.2 and 3.3) and Sm0rbukk Field (Tilje 3.2 and T4). The third stage is exclusively developed in the SmCrbukk Field (T5) while a period of non-deposition typifies the Heidrun Field. Stage 4 is present in both fields (T6 in the SmOrbukk Field and T3.4 in the Heidrun Field). Stage 1 serves to discus the inferred depositional setting in more detail. Stage 1 can be divided in four units interpreted to reflect different river-delta facies zones. Consequently, their vertical stacking patterns are interpreted to reflect subsequent phases of development of the depositional system. The basal unit 1 is composed of several occurrences of facies association 8 (Fig. 18). Unit 2 is formed by a series of occurrences of facies association 9 (Fig. 19) but is better developed in the Heidrun Field than in the SmOrbukk Field. Facies associations 8 and 9 are inferred to have been formed simultaneously but in different reaches of the distributaries, association 8 somewhat more distal than association 9. The overlying unit 3 is thin (several metres) and composed of an upward-fining channel-fill succession similar to unit 1. In the Heidrun area, the top of unit 3 is formed by a dark, lenticular-bedded

(cf. Reineck and Wunderlich, 1968) mudstone bed that is found over the entire field area and interpreted as being formed in a semi-enclosed embayment. Unit 4 is dominantly composed of migrating compound dunes (facies association 5) interpreted to represent delta-like bars (cf. Dalrymple and Rhodes, 1995). Units 1, 2 and 3 are interpreted to reflect repetitive periodic high fluvial discharge causing cutting and filling of distributaries. After discharge ceased, tidal processes took over as the main control on sediment deposition, migration and development of the distributaries. It is envisaged that the seaward end of the river-delta plain was typified by wave- and tide-dominated distributary channels (facies associations 6 and 7). In the middle and upper reaches of the distributaries, sediment influenced by the combined effect of tide- and fluvial-dominated currents accumulated (facies associations 8 and 9), while the (near) emergent inter- and supratidal interdistributary areas were formed by intertidal flats, channel levees, and swamps (facies association 10). Indications for vegetated upper intertidal flat and/or salt-marsh deposits (facies association 10) are not found in the Heidrun Field, uncommonly in the SmOrbukk Field but more commonly in the Lavrans Field (Fig. 1) where sediment preservation is more complete. Heterolithic barforms (facies association 5) were associated with numerous channel thalwegs, forming the proximal delta-front

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extension of the distributary channels. These channels were filled with heterolithic and abandonment facies forming "muddy troughs" between the mouth bars. Closer to the mouth of a distributary channel, extension and bifurcation of distributaries took place, and bars formed by aggradation of migrating compound dunes (cf. Anastas et al., 1997; facies association 5), similar to the situation in many modem tide-dominated deltas (Coleman et al., 1970; Kenyon, 1970; Yang, 1989; Dalrymple, 1992). The Tilje delta is inferred to have had a somewhat convex shoreline profile (Fig. 29) which was modified by shore-oblique wave approach and shore-normal tidal currents. As mentioned earlier, the upper part of the Elis Bjerg Member (Neill Klinter Group; Dam and Surlyk, 1998) on Jameson Land, eastern Greenland (Fig. 2), may represent time-equivalent deposits to the Tilje Formation. It is proposed that SB3 of the Tilje Formation may be correlatable with SB3 recognised in the Ells Bjerg and Ra~veklCft Members (Dam and Surlyk, 1998, their figs. 3 and 46). Surfaces correlatable to SB4 and SB5 of Dam and Surlyk (1998) are not readily recognised in the Tilje Formation. However, based on the well-established time control in the Tilje Formation (in-house biostratigraphic data) and the Neill Klinter Group, it is hypothesised that SB5, which underlies the Astartekloft Member of the Neill Klinter Group, is possibly correlatable with the base of T4 in the Sm0rbukk Field. Discussion

The estuary stage The differences in the nature of SB2 and the rock succession above it between the Heidrun and Sm0rbukk Fields are interpreted to reflect the fact that two different estuarine systems were active in the two areas. Moreover, it is envisaged that Tilje 1.1 in the SmOrbukk Field was located in a relatively more distal estuary-mouth setting as compared to the time-equivalent Tilje 2.1 in the Heidrun Field. Arguments supporting this view include the fact that in-house palaeocurrent data indicated that in the Heidrun area the system was most likely fed from the (north)east, whereas in the SmCrbukk area the system was most likely fed from the (north)west. Furthermore, the spatial well configuration in the Heidrun Field clearly shows a proximal-distal facies trend for the Tilje 2.1 interval from estuarine deposits in valleys (facies association 7) in the east to fine-grained and well-bioturbated deposits (lithofacies 1.1) in the west. Secondly, in the SmOrbukk Field, lithofacies 7.1 has proportionally more wavy-flaser bedding than lithofacies 7.1 in the Heidrun Field where simple

A. W. Martinius et al.

flaser bedding (cf. Terwindt, 1981) is prominent. This observation is interpreted to reflect the effects of funnelling of tidal flows as a result of the landwarddecreasing width of the estuary. It is assumed that, as a result, tidal energy increases and the potential for mud preservation decreases landward. Simple flasers form just above the current-velocity threshold value (ca. 0.45 m/s), wavy flasers just below. This may indicate that slightly higher current velocities were dominant in the Heidrun area as compared to the SmOrbukk area, and that the Heidrun Field was located more proximal than the Sm0rbukk Field at this time of deposition. Additionally, relatively more interbar heterolithic fines are found in the Sm0rbukk area, as well as relatively small tidal channels, with their associated channel-margin deposits, that dissected the tidal bars. Thirdly, if the second-order relative sea-level fall is assumed to have created an equal degree of accommodation space decrease across the HeidrunS me~rbukk area, then it may have caused incision and valley formation in the Heidrun area because the gradient of the river course was larger than the equilibrium longitudinal profile of the fiver. As a result, significant volumes of sediment were bypassed during the relative sea-level fall. This sediment was deposited at mouth bars and subtidal flats in front of the distal end of the valley. An equal magnitude of relative sea-level fall in the SmCrbukk area, however, may have resulted in erosion without valley formation, either because the gradient of that river course was smaller than its equilibrium longitudinal profile, and/or because the SmCrbukk area was located in a comparatively more distal position where erosion, but no incision, was recorded. In the latter case, T I.1 in the SmOrbukk Field classifies as a forced-regressive wedge (cf. Posamentier et al., 1992) with inferred valley deposits updip and two parasequences downdip in the Sm~arbukk Field area. Arguments to support the forced-regression interpretation include: (1) deposition occurred on a ramp margin; (2) a downward shift in facies is observed; (3) SB2 at the base is a sharp erosive surface, with in places is an up to 10 cm thick pebble lag, overlain by tide- and waveinfluenced sandstones; (4) the top of Tilje 1.1 in the SmCrbukk Field is a transgressive surface. However, no conclusive data are available in support of any of the two options. If the second-order relative sea-level fall was related to movement along one or more (growth) faults, then the magnitude of the change in accommodation space may have been different between the Heidrun and SmOrbukk areas. Relatively rapid downthrown movement of the hanging wall along a growth fault in the Heidrun area, for example, may have caused a

Sedimentology of the heterolithic and tide-dominated Tilje Formation

relative rapid sea-level fall and associated basinward shift of facies associated with fluvial incision. Alternatively, hinterland uplift may have increased the river course gradient of the depositional system in the Heidrun area causing disequilibrium and incision. Unfortunately, no data are available to conclusively settle this issue.

The river-delta stage The Tilje 3 succession in the Heidrun Field has some characteristics that differ from the time-equivalent Tilje 3 and 4 in the Sm~rbukk Field (Fig. 20). It is found that in the Heidrun Field (1) the grain-size variation is larger, (2) the maximum grain size encountered is considerably larger, (3) the relative proportion of tidal-channel facies (lithofacies 9) is significantly higher, (4) inshore estuarine facies (lithofacies 7.1) occur less frequently, (5) channelised delta-front lobe facies (facies association 6) are dominant in Tilje 3.3, whereas in the time-equivalent Tilje 4 unit in the SmOrbukk Field inshore estuarine facies (lithofacies 7.1) and tidal-channel facies (lithofacies 9) dominate. Tilje 3 (lower part of Sequence 3) in the Heidrun Field shows a more distinct progradational and retrogradational trend (at a scale of tens of metres) than the SmOrbukk Field. In the Heidrun Field, the stratigraphic level at which facies association 9 is most common, coarsest (medium pebble size), and poorly sorted is interpreted as reflecting the time of maximum delta progradation, and a river-delta setting is proposed. Progradation and retrogradation likely occurred during the first-order period of eustatic sealevel rise, characteristic of the late Pliensbachian to early Toarcian (Fig. 5B and Fig. 20). In order to produce a progradational succession, locally reflecting relative sea-level fall, the rate of sediment supply has to surpass the rate of eustatic sea-level rise, a situation similar to processes in the Atchafalaya Bay (Wells, 1987). Progradation ceased most likely as a result of the combined effect of ongoing eustatic sea-level rise, the development of too many small distributaries, and diminishing sediment supply. This process resulted in retrogradation (Tilje 3.3 in the Heidrun Field), followed by sediment cessation and the development of muddy embayment depositional conditions at the top of Tilje 3 (Fig. 6). The Tilje 3 succession in the SmOrbukk Field is aggradational in nature and no point of maximum progradation can be recognised. A short retrogradational phase characterises the upper part of the Tilje 3 succession. The Tilje 4 stratigraphic unit in the SmOrbukk Field represents an initial progradation (with a pebble lag at the base) followed by an overall aggradational

141

phase. The aggradational phase is formed by a series of 1 to 5 m thick upward-fining units which are generally formed by inshore estuarine facies (lithofacies 7.1), commonly flaser- to wavy-flaser-bedded, overlying tidal-channel facies (facies association 9). Thin pebbly layers (up to 50 cm thick) infrequently occur throughout the succession. Tilje 4 in the SmOrbukk Field can be correlated to Tilje 3.3 in the Heidrun Field (Fig. 20) which is mainly composed of channelised delta-front lobe facies (facies association 6). In addition, the Tilje 3.3 in the Heidrun Field is half as thick as Tilje 4 in the SmOrbukk Field. Based on these observations, it is assumed that the Heidrun Field was located in a more distal position than the SmOrbukk Field at this time of deposition. In order to argue in favour of or against the existence of one or two delta depositional systems active in the Heidrun-SmOrbukk area at the time of Tilje 3, and reliably define a depositional model, knowledge and understanding of the feeder system(s) is essential (Orton and Reading, 1993). Unfortunately, very little is known about the Tilje feeder systems. Most classification schemes of deltas are unimodel grainsize-driven schemes which are mainly based on the coarse grain-size fraction because it is assumed to be most important in determining the steady-state pattern of deposition (Orton and Reading, 1993). However, the Tilje system is a heterolithic system typified by a bimodal grain-size distribution, which may be a reason that the Tilje delta system does not fit comfortably in one of the existing classification schemes. In addition, modem depositional coast lines, for example the southern coast of Papua New Guinea, show lateral changes in coastal morphology with various delta and delta-like depositional systems adjacent to each other. The morphological appearance of some of these delta-like systems resembles more a series of adjacent estuaries than a delta. Local differences in receiving basin characteristics for example, variations in subsidence rate due to (syn-sedimentary) fault movements or hinterland and catchment area characteristics may have governed sediment dispersal patterns, grain-size characteristics, and the volume of sediment supplied and preserved. Unfortunately, there is a lack of modern data that can be used to analyse fossil analogue depositional systems as many of these modern systems are not (yet) studied in detail. Conclusions

(1) Ten facies associations are recognised in the Tilje Formation on the Halten Terrace in the Heidrun and SmOrbukk Fields. (2) The Tilje Formation encompasses the highstand systems tract of Sequence 1 at the base (T1 in the

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Heidrun Field and the uppermost Are in the Sm0rbukk Field), the complete Sequence 2 (T1 in the Sm0rbukk Field, T2 and T3.1; Sequence 2), and the lowstand and transgressive systems tract of Sequence 3 (T3.2 to the top of the formation in both fields). Three maximum flooding surfaces are identified, the first at the base of the formation, the second in the middle of T2, and the third in the overlying Ror Formation. Two sequence boundaries are recognised, the first at the base of Tilje 2 in the Heidrun Field and the base of Tilje 1 in the Sm0rbukk Field, the second at the base of T3.2. (3) The highstand systems tract of Sequence 1 is wave-dominated and influenced by storm deposition. Two different conceptual depositional models are proposed for Sequences 2 and 3 of the Tilje Formation in the Sm0rbukk and Heidrun area to accommodate significant stratigraphic differences. Sequence 2 is characterised by several laterally adjacent low-relief valleys formed during base-level lowstand and subsequently transformed into waveand tide-dominated estuaries during transgression. Sequence 3 is typified by a mud-rich, tide- and fluvial-dominated delta-like system composed of three facies zones: a lower delta plain intersected by relatively small distributaries, a delta front, and a prodelta area. In the Heidrun area, the system is progradational/retrogradational, whereas in the SmCrbukk area it is aggradational/retrogradational. The two depositional systems developed in response to changes in the nature of the generation of accommodation space, the volume of sediment supplied, grain size, and the interplay between wave, tidal and fluvial energy. (4) Significant differences exist in three-dimensional facies architecture between Sequence 2 and Sequence 3. In Sequence 2, relatively thick and clean sandstones are developed in wide estuaries underlain by an incision surface. These are capped by sealing offshore muddy siltstones. In Sequence 3, a tens of metres thick succession of a thin (m scale) heterolithic fining-upward channel-fill developed with a complex internal 3D architecture and which prograded over layered prodelta deposits. These characteristics have implications for the way reservoir models need to be built because such models depend on dimensional properties, and geometries of genetic elements, relative facies proportions, and facies stacking patterns.

Acknowledgements The original core descriptions used as a basis for this study have emanated from five companies (Amoco, Conoco, Neste, Saga, and Statoil) over a

A.W. Martinius et al.

period of some 12 years. Opinions expressed herein, however, are the authors responsibility. We would like to acknowledge A. Dalland (deceased), L.-M. F~ilt, M. Kj011eberg, H. LCseth, T. Olsen, S. Olsson, D.K. Renshaw, A. R0muld, M. Amark (all Statoil), T. Sa~ther (formerly at Saga Petroleum), H.E. Clifton, K.E. Svela (Conoco, Houston), G. Dam, E Surlyk (Geological Survey of Greenland) and J.C. Harms for data analysis and/or discussions. Ichnofabrics and additional palynofacies data were analysed by Drs. S. Gowland, N. Hogg, and A.M. Taylor (Ichron Ltd.), whom we also thank for fruitful discussions. Sm/Nd isotopes were analysed by Dr. E. Mearns (IAS Ltd.). The manuscript benefited from discussions with Profs. R.W. Dalrymple (King's University, Ontario), S.S. Flint (Univ. of Liverpool) and H.D. Johnson (Imperial College, London), and Drs. D. McIlroy (Univ. of Liverpool), A.E Oost (Univ. of Utrecht), S. Yoshida (Imperial College, London) and B.J. Willis (State Univ. of New York). L. Reistad skillfully computer-drafted the figures. Valuable and constructive reviews by R.W. Dalrymple, T. Dreyer, and W. Helland-Hansen significantly improved the clarity and structure of the manuscript. The Heidrun and Asgard (Sm0rbukk Field) Licences are acknowledged for permission to publish the data.

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A.W. MARTINIUS I. KAAS A. NA~SS G. HELGESEN J.M. KJtEREFJORD D.A. LEITH

tures in inshore mesotidal deposits of the North Sea. In: S.D. Nio, R.T.E. Schtittenhelm and Tj.C.E. van Weering (Editors), Holocene Marine Sedimentation in the North Sea Basin. Spec. Publ. Int. Assoc. Sedimentol., 5: 4-26. Terwindt, J.H.J. and Breusers, H.N.C., 1972. Experiments on the origin of flaser, lenticular and sand-clay alternating bedding. Sedimentology, 19: 85-98. Terwindt, J.H.J. and Breusers, H.N.C., 1982. Flume experiments on the origin of flaser bedding: discussion. Sedimentology, 29: 903907. Van de Weerd, A.A., 1996. Reservoir geology of the shallowmarine Early Jurassic Tilje Formation of the Njord Field, offshore Mid-Norway. In: H.D. Johnson, J.E Wonham, R. Gupta, M.E. Donselaar, A.A. van de Weerd, J. Mutterlose, A. Stadler and A.H. Ruffell (Editors), Geological Characterisation of Shallow Marine Sands for Reservoir Modelling and High Resolution Stratigraphic Analysis. Final Report EC-Joule Prog., 1, 64 pp. Van den Berg, J.H., 1981. Rhythmic seasonal layering in a mesotidal channel fill sequence, Oosterschelde mouth, the Netherlands. In: S.D. Nio, R.T.E. Schtittenhelm and Tj.C.E. van Weering (Editors), Holocene Marine Sedimentation in the North Sea Basin. Spec. Publ. Int. Assoc. Sedimentol., 5: 147-159. Van den Berg, J.H., 1982. Migration of large-scale bedforms and preservation of cross-bedded sets in highly accretional parts of tidal channels in the Oosterschelde, S.W. Netherlands. Geol. Mijnbouw, 61: 253-263. Van Straaten, L.M.J.U., 1959. Minor structures of some Recent littoral and neritic sediments. Geol. Mijnbouw, 21: 197-216. Wells, J.T., 1987. Effects of sea-level rise on deltaic sedimentation in south-central Louisiana. In: D. Nummendal, O.H. Pilkey and J.D. Howard (Editors), Sea Level Fluctuations and Coastal Evolution. Spec. Publ. Soc. Econ. Paleontol. Mineral., 41: 157-166. Whitley, EK., 1992. The geology of Heidrun: a giant oil and gas field on the mid-Norwegian shelf. In: M.T. Halbouty (Editor), Giant Oil Fields of the Decade 1978-1988. Am. Assoc. Pet. Geol., Mem., 54: 383-406. Willis, B.J., 1997. Architecture of fluvial-dominated valley-fill deposits in the Cretaceous Fall River Formation. Sedimentology, 44: 735-757. Woodroffe, C.D., Chappell, J., Thom, B.G. and Wallensky, E., 1989. Depositional model of a macrotidal estuary and floodplain, South Alligator River, Northern Australia. Sedimentology, 36: 737-756. Yang, C.-S., 1989. Active, moribund, and buried tidal sand ridges in the East China Sea and southern Yellow Sea. Mar. Geol., 88: 97-116. Zaitlin, B.A., Dalrymple, R.W. and Boyd, R., 1994. The stratigraphic organization of incised-valley systems associated with relative sea-level change. In: R.W. Dalrymple, R. Boyd and B.A. Zaitlin (Editors), Incised-Valley Systems: Origin and Sedimentary Sequences. Spec. Publ. Soc. Econ. Paleontol. Mineral., 51: 45-60.

Statoil Research Centre, Arkitekt Ebbellsveg 10, N-7005 Trondheim, Norway Present address: c/o Statoil Venezuela- Sincor Project, N-4035 Stavanger, Norway Statoil Bergen, Sandslihaugen 30, N-5020 Bergen, Norway Statoil Research Centre, Arkitekt Ebbellsveg 10, N-7005 Trondheim, Norway Present address: Statoil StjOrdal, Strandveien 4, N-7501 StjOrdal, Norway Statoil Stavanger, Forusbeen 50, Stavanger, Norway Statoil Bergen, SandsIihaugen 30, N-5020 Bergen, Norway Statoil StjOrdal, Strandveien 4, N-7501 StjOrdal, Norway

145

Sequence stratigraphy of the Lower Jurassic Dunlin Group, northern North Sea M.A. Charnock, I.L. Kristiansen, A. Ryseth and J.P.G. Fenton

Sequence stratigraphic analyses of the Lower Jurassic Dunlin Group of the North Viking Graben provides a consistent basinwide framework for a better understanding of the sedimentary facies distributions. Nine sequences are defined on the basis of temporal changes in depositional environments, extensive biostratigraphic information and log correlation. The bounding surfaces of these sequences represent maximum transgressive surfaces (rots) and mark the points between a series of regionally identifiable regressive-transgressive episodes in the Early Jurassic (latest Sinemurian to Toarcian) interval of the North Viking Graben. The biostratigraphic information has enabled the sequences to be calibrated to standard schemes established elsewhere within the North Sea Basin and the ammonite-dated Lower Jurassic type section on the Yorkshire coast. The sequences have thicknesses of between 10 m and 100 m and a duration of approximately 1-4 Ma. They consist of facies representing deposition in a variety of shelf, shoreline and estuarine systems. The significance of estuarine facies within this interval is highlighted, most notably within the J14 and J15 sequences. The sediments generally have a source direction from the Horda Platform and Lomre Terrace areas on the eastern basin margin. The sandstone distribution within these sequences is primarily controlled by variations in sediment supply, accommodation potential and tectonic subsidence and represents a history of repeated progradation and retrogradation.

Introduction

An integrated sequence stratigraphic study, utilising well-log correlations, detailed core descriptions and extensive biostratigraphic information provides a chronostratigraphic framework to analyse and predict the distribution of potential reservoir sandstones within the Lower Jurassic Dunlin Group of the Norwegian sector in the northern North Sea. The biostratigraphic database utilises palynological and micropalaeontological information from 44 wells in Quadrants 30, 31, 34 and 35 of the North Viking Graben (Fig. 1). Very little detailed sedimentological and stratigraphical data have been published on the Dunlin Group (Fig. 2). The boundary between the Dunlin Group and the subjacent Statfjord Formation (Fig. 2) records a change in depositional environment, from alluvial to marginal marine in the uppermost Statfjord Formation, to rather deep marine (shelf) conditions in the lowermost Dunlin Group. Recent accounts of Early Jurassic stratigraphy (Parkinson and Hines, 1995; Marjanac and Steel, 1997) indicate that the Statfjord/Dunlin transition straddles the Sinemurian/Pliensbachian stage boundary on structural highs and platforms, whereas it falls within the latest Sinemurian in the deeper axial parts of the basin and

also in the southern parts of the Viking Graben (this study). The Dunlin/Brent transition (Fig. 2) separates Toarcian marine strata (Drake Formation) from overlying Aalenian-early Bajocian deltaic deposits. As this records a basinward shift of facies that can be related to the mid-Cimmerian thermal event (see below), it is usually regarded as an unconformity with a possible correlative conformity in the deeper parts of the basin (e.g. Steel, 1993; Underhill and Partington, 1993). The Dunlin Group, as originally defined by Dor6 et al. (1984) and subsequently redefined by Marjanac (1995), comprises two main lithologies: (1) mudrocks and associated heterolithic sandstone/shale alternations, and (2) cleaner sandstones. These two lithosomes were grouped into three fine-grained formations (Amundsen, Burton and Drake Formations, Fig. 2), and two sandier units (Johansen and Cook Formations), with the Johansen Formation being restricted to the eastern basin margin. Additional sandstone units, informally designated "Drake sandstones" are also commonly identified on the eastern margin. A marine shelf environment is usually inferred for the fine-grained formations, whereas the sandy formations were related to more marginal marine and tidally influenced settings (Dor6 et al., 1984; Livbjerg and Mjc~s, 1989). More recent studies have

Sedimentary Environments Offshore Norway - Palaeozoic to Recent edited by O.J. Martinsen and T. Dreyer. NPF Special Publication 10, pp. 145-174, Published by Elsevier Science B.V., Amsterdam. 9 Norwegian Petroleum Society (NPF), 2001.

146

M.A. Charnock et al.

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Fig. 1. Studyarea of the NorthVikingGrabenshowinglocations of wells and correlation lines illustratedin Figs. 9-14. maintained the general shelf/marginal marine environmental setting for the Dunlin Group (Dreyer and Wiig, 1995; Parkinson and Hines, 1995; Marjanac and Steel, 1997). However, new interpretations suggest that deposition of the main sandstones can be related to forced regressive events (Parkinson and Hines, 1995) and cutting and filling of estuarine valleys (Marjanac and Steel, 1997). The regional compilation of Parkinson and Hines (1995) suggests that a late Pliensbachian regressive event corresponding in part to the Cook Formation (Fig. 2) can be recognised throughout western Europe, thus indicating a regional relative sea-level fall at this stratigraphic position. Petrographic data on the Dunlin sandstones are largely absent from the public domain. However, Ehrenberg (1993) documented the preservation of anomalously high porosities in some late Pliensbachian sandstones in the Veslefrikk Field, due to the presence of grain-coating chlorite. Development of this mineral phase may be favoured by deposition in marine environments with a significant supply of fresh water, and can lead to the preservation of

reservoir properties at burial depths below 4 km. Hence, a predictive model to explain the occurrence of chlorite in terms of depositional environment and stratigraphic position within a sequence stratigraphic framework may be highly valuable for ongoing exploration in the area. This study of the Early Jurassic succession of the northern North Sea outlines the sedimentology of the main reservoir sandstones. Furthermore, sequence stratigraphic models using refined biostratigraphic data to correlate these units are used to develop a sequence stratigraphic model for the Dunlin Group. Special emphasis is on the spatial and temporal distribution of sandstones within the Dunlin Group, and on factors controlling the distribution of potential reservoir sandstones to the north, into relatively unexplored areas of the Viking Graben (Fig. 1). Geological setting The Brent province of hydrocarbon discoveries (Fig. 1) is geographically coincident with the northern

Sequence stratigraphy of the Lower Jurassic Dunlin Group, northern North Sea

Ma

AGE

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] 47

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early

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(part)

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.......................................................... :::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::: ..........................................................

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Fig. 2. S u m m a r y of the lithostratigraphy and sequence stratigraphy of the L o w e r Jurassic D u n l i n Group.

Viking Graben and its flanking platforms (Yielding et al., 1992). Siliciclastic successions of Triassic, Early Jurassic and Middle Jurassic age form the principal reservoir rocks, with deposits belonging to the Middle Jurassic Brent Group being by far the most economic. However, significant hydrocarbon volumes are also stored in sandstones of the Lower Jurassic (Pliensbachian-Toarcian) Dunlin Group as defined by Dor6 et al. (1984) (Fig. 2). Hydrocarbon trap development is related to late Mesozoic extensional rotation of fault blocks, and subsequent seal formation by mainly Cretaceous and Tertiary mudrocks. The main hydrocarbon source (Draupne Formation) is of Oxfordian-Volgian age, and accumulated in stagnant, riftrelated sub-basins during Late Jurassic/Early Cretaceous rifting. Hydrocarbons are mainly derived from the deeper, axial parts of the Late Jurassic rift system, and were emplaced throughout Late Cretaceous and Tertiary times. Crustal extension and rifting in the northern North Sea commenced in Late Permian/Early Triassic times and were followed by post-rift thermal subsidence throughout Triassic to Middle Jurassic times (Badley

et al., 1988; Steel and Ryseth, 1990). A second phase of crustal extension associated with widespread fault block rotation occurred in Late Jurassic/Early Cretaceous times, and was followed by thermal postrift subsidence during Cretaceous and Tertiary times. Recently, it is documented that the Permo-Triassic and Late Jurassic-Early Cretaceous rift phases involved approximately similar magnitudes of crustal extension (~-factors of about 1.4-1.5; Fa~rseth, 1996). However, the two phases are spatially offset and associated with different fault patterns. The PermoTriassic phase led to the development of a broad depression bounded by N-S-trending faults, with a subsidence maximum located within the present Horda Platform. In contrast, the Late Jurassic rift axis is laterally offset to the west, and associated with mainly NW-SE-trending faults (Fa~rseth, 1996). The Early and Middle Jurassic successions were deposited during the post-rift phase of the first PermoTriassic rift phase (Badley et al., 1988; Yielding et al., 1992), hence their thickness distributions are potentially controlled by the N-S-trending fault pattern of this tectonic phase.

148 Following the widespread alluvial deposition in the Viking Graben throughout Triassic and earliest Jurassic (Hettangian-Sinemurian) times (Steel and Ryseth, 1990; Ryseth and Ramm, 1996; Ryseth, 2000), a marine connection between the northern Boreal and southern Tethyan seas was established in the area in Early Jurassic (Pliensbachian-Toarcian) times (e.g. Dor6 and Gage, 1987; Ziegler, 1990). This marine seaway was probably bounded to the west by the Shetland Platform, and to the east by the Fennoscandian hinterland. Both these areas are probable sources to the sediments of the Dunlin Group (Fig. 2), which may have been transported and distributed into the seaway rather transverse to the N-S-trending basin axis (Ziegler, 1990). Domal uplift associated with volcanism occurred in the triple junction of the main Jurassic North Sea rifts (Central Graben, Moray Firth, Viking Graben) in late Toarcian-Aalenian times (Ziegler, 1990). Rise of the dome centre may have continued into Bathonian times, accompanied by concurrent deflation of dome margins from late Bajocian times (Underhill and Partington, 1993). The rising dome altered the basin physiography, by blocking the marine Early Jurassic (Pliensbachian-Toarcian) connection between the northern (Boreal) and southern (Tethyan) seas in the Aalenian (Ziegler, 1990), and by exposing Lower Jurassic and older deposits to subaerial erosion, with development of a widespread unconformity (mid-Cimmerian event; Underhill and Partington, 1993). Outside the exposed dome, the uplift is associated with major progradation of the contemporary shorelines, whereas the subsequent deflation coincides partly with major transgressions (e.g. Graue et al., 1987; Underhill and Partington, 1993). The pattern of Lower Jurassic deposition is markedly different from the Middle Jurassic setting. The thermal anomaly associated with the mid-Cimmerian doming is further associated with hinterland rejuvenation and uplift of the margins of the Viking Graben during the Aalenian (Mitchener et al., 1992; Rattey and Hayward, 1993) with initiation beginning within the latest Toarcian (this study). This phase of uplift is related to deposition of the initial coarsegrained deposits of the Brent Group (Broom and Oseberg Formations, Fig. 2; see Graue et al., 1987; Cannon et al., 1992; Steel, 1993). By AalenianBajocian times, the widespread uplift to the south of the area had closed the marine connection between the Boreal and Tethys oceans, and the North Viking Graben became the site of deposition of the deltaic Brent Group (Fig. 2). The Brent delta prograded northwards through the Viking Graben during Bajocian times, supplied from both westerly (Shetland Platform) and easterly (Fennoscandia) source areas

M.A. Charnock et al.

(Eynon, 1981; Graue et al., 1987; Helland-Hansen et al., 1992; Mitchener et al., 1992; Johannessen et al., 1995; Fjellanger et al., 1996).

Sequence stratigraphy: definitions The sequence stratigraphic approach of the present study is a refinement of the regional template established by Partington et al. (1993) but is otherwise based on the theoretical framework of HellandHansen and Martinsen (1996). The main bounding surfaces are defined at successive maximum transgressive surfaces (mts) such that each sequence is characterised by a regressive-transgressive component. They are therefore equivalent to maximum flooding surfaces of Galloway (1989). These surfaces represent marine condensed horizons and have been identified on wireline logs by high gamma ray and neutron values, low density and low sonic velocities and are considered to be the most reliable correlative events in shallow/marginal marine environments. An attempt has been made to identify, wherever possible, transgressive surfaces (TS), regressive surfaces of marine erosion (RSME) and Vailian-type sequence boundaries (SB). The maximum transgressive surfaces identified on wireline logs and sedimentologically in core have been calibrated with a biostratigraphic zonation (Fig. 3). The geochronological ages of the maximum transgressive surfaces utilise the time scale of Hardenbol et al. (1998). The maximum transgressive surfaces have been defined, wherever possible, using the biostratigraphic criteria established by Partington et al. (1993). Consequently, this study has provided an independent test on the stratigraphic calibration by Partington et al. (op. cit.) and is in effect the first detailed practical application of their framework to be published for the North Viking Graben area. Six of the maximum transgressive surfaces recognised are in common with those identified by Partington et al. (op cit.) namely (from oldest to youngest) J12 mts, J 14 mts, J16A mts, J 16B mts, J 18 mts and J22 mts (see Fig. 4). Three additional sequences, J 13, J 15 and J20, have also been defined in this study. The sequence stratigraphic framework established here has been compared to other published schemes (Fig. 4) and is discussed below. Other studies of the Dunlin Group have also attempted to provide sequence stratigraphic frameworks (see Fig. 4). Steel (1993) recognised three regressive-transgressive megasequences (Johansen, Cook and Drake megasequences) within the Dunlin Group bounded by maximum flooding surfaces. Parkinson and Hines (1995) also identified three regressive-transgressive cycles in the Dunlin Group

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laeviuscula 175.81 p j ' discites 176.5 4

I

(part)17(+.5 (part)

VIAX. TRANS3RESSIVE .SURFACES

VIICROPALEONTOLOGY

Tr

reticulatus 1C 2o9.6 Tr2

L.lundbladii, O.ovalis, A '" hncr. R.tuberculatus "'~q R.tuberculatus 9 9

B 1

2 3

"~k-1Tfitetes pinguis

).hagenowi,l,. 4.hopliticus

/

.~incr. O. pseudoalatus, / G. zwolinskai 'irk-]incr. P.microcorpus ' ~ O.pseudoalatus 9

VITr

9 - common

1

9 - abundant

9! V.ignacii, E.vigens 9

~.

%

- ammonite calibrated ~

R.penarthensis R.wiched, D.major

Fig. 3. Early Jurassic biostratigraphic zonation, northern North Sea.

r

not ammonite calibrated

LO

(.,1-1 c) MY

BIO-

(Hardenbol et al. 1998)

BAJOCIAN

(pare L

AMMONITES

EARLY (part) !

177.33i

laeviuscula

pj

discites

4

176.5

BOREAL T-R CYCLES

MEGA SEQUENCES

GULLFAKS FIELD SEQUENCES

SEQUENCES

(Partington et al., 1993)

(Jacquin & de Graciansky, 1998)

(Steel, 1993)

(Dreyer & Wiig, 1995)

(Marjanac & Steel, . . . . . 1997)

'

70R

opalinum

ej 3D

.

.

.

J24 (part)

J24 (part) .

T7 (part)

Brent Group

7 J22

,

.

J22

64R -->

.

180.1 !

i

----> .

Brent 7 (part)

M J!

177.33, bradfordensis178 1~ sonae 179.29;

EARLY

180.1

ZONES

concavum

179.29 '

i

ZONES

murchiMIDDLE sonae murchi-"

AALENIAN

180 _

PALYNOLOGY MICROPAL.

175.81

LATE

,

GENETIC SEQUENCES

E V E N T S

i

,

SEQUENCES (this study)

KEY

BIOSTRATIGRAPHY

AGE

Drake 4

i

aalensis 180.88

R6

pseudoradiosa

.==

181.67

LATE ,

levesque~82 47 .., pJ thouarsense

184.06, I

TOARCIAN

MJ 6

4 Drake 6

62R --->

variabilis

J18 185.25

186.84

MJ 5

60R

PJ 2 3B . 1 188.04,

9

55R - ~ \

falciferum

EARLY

tenuicostatum 189.6! 190

pj

189.63 !

!

spinatum

-'

!

3A

190.38 I margari" " LATE tatus 191,1~ 191.51.tatus91.,~11stokesi ,

'margari-

davoei

PLIENSBACHIAN EARLY i 195 -

I LATE (part) I

;

1

Cook

]

Sst. C

3

/

het.

J16

36 ~ . 40R

' PJ 193.77 2C

32 - - > MJ 3

jamesoni

[pj

.

.

.

J13

J14

.

O ok

~axe.1.... ...... I Sst. A . Cook 1 her.

'

2A Cook 1

.

_

2B

Cook 5

:

J12

Drake 2

3A 2C

A A . . . . . . . . . 714-

I

196.831

J16

'

3B

T6

~------

A __~ ' ' B , 49R, 45R I MJ ~

192.64

raricostatum

5OR

Cook 3

'

195.3..

195.3

SINEMURIAN (part)

ibex

J18

186.84

.

Burton Fm. ..... ~-""]

">J

J12 J6

I

T5 (part)

.... Z .....

Johansen 4

R5

.... 30R

h~t.

Drake 3

Drake

MIDDLE i bifrons

/

J20

63 --->

3C

184.06

185 -

~ooF- 1-- ~-st 5--

i

Statfjord 3

!

.....

Statfjord Fm.

2B c~

Fig. 4. Comparison of sequence stratigraphic schemes for the Lower Jurassic Dunlin Group. t~

Sequence stratigraphy of the Lower Jurassic Dunlin Group, northern North Sea

(denoted J12/J14, J16 and J18 using the framework established by Partington et al., 1993; Fig. 4), with the Johansen, Cook and intra-Drake Formation sandstones representing regressive maxima. Dreyer and Wiig (1995) applied the sequence stratigraphic concepts established by Van Wagoner et al. (1987) to the reservoir zonation of the Cook Formation in the Gullfaks Field. They noted the incised nature of the major sandstone units with the transition from Cook-2 to the main reservoir Cook-3 being interpreted as a sequence bounding unconformity (within J16B of this study; see Fig. 4) being associated with a major change in drainage pattern. Finally, Marjanac and Steel (1997) pointed at the incised nature of late Pliensbachian-late Toarcian sandstones in the North Sea, and argued for the existence of four higher-order sequences (denoted Cook-1 to Cook-4; see Fig. 4).

Sedimentary facies and environments A number of 21 cored sections have been studied for sedimentology. Fig. 5 shows that the cores cover almost the entire stratigraphic interval of the Dunlin Group and provides an important foundation for the facies and sequence stratigraphic interpretation. In addition, there is a wide geographic spread of cores from sequences J14 to J16, comprising the Cook Formation, providing a good basis to determine spatial variations in sedimentary facies type at this stratigraphic level. Horda Platform and Sogn area

Figs. 6 and 7 illustrate a number of cored sections from the eastern part of the basin, including the Horda Platform, the Lomre Terrace (well 35/10-1) and the Sogn area. Well 35/9-2 (Fig. 6) in the Sogn area has full core coverage of the Dunlin Group and serves to outline the main temporal changes in facies and depositional environments. Sequences J13 and J14 are characterized by a vertical alternation of coarse-grained to pebbly sandstones and sections dominated by mudrocks and fine-grained sandstones. The main sandstone bodies, corresponding to the Johansen and Cook Formations (Fig. 6) are sharply based, composite fining-upwards units dominated by cross-stratified troughs and current ripple lamination. In this proximal setting, these bodies are related to fluvial and/or estuarine deposition. The intercalated finer-grained intervals comprise a variety of biogenic traces (Chondrites, Helminthopsis) and wave-generated structures including hummocky/swaley types of cross-stratification, and are related to deposition in the offshore and lower shoreface of a wave-dominated shoreline. The alternation of fluvial/estuarine

151

and shelf/lower shoreface deposits in this part of the succession testify to rather significant variations in relative sea level as a major factor controlling the distribution of sediment in the Pliensbachian. The late Pliensbachian-Toarcian part of the succession (sequences J15-J18; Fig. 6) can be related to successively deeper marine depositional environments, reflecting an overall transgressive setting at this stage. The Brent Group rests with marked unconformity on the Dunlin Group in this well (Fig. 6), probably reflecting Aalenian uplift of the basin margin. Additional cores from the Horda Platform and Lomre Terrace are presented in Fig. 7. The basal surface of the Dunlin Group is cored in well 30/3-A-5 (Fig. 7). Notably, the lithological contact with the underlying Statfjord Formation is sharp and marked by a sideritic crust, separating fluvial and shallow marine sandstones from intensively bioturbated very fine-grained deposits of the lowermost Dunlin Group (assigned to the J13 sequence). The lowermost part of the Dunlin Group (below J13 mts in well 30/3-A-5) comprises pervasively burrowed, fossiliferous and sideritic very fine-grained sandstones, that may record reworking and condensation during the initial marine transgression of the Early Jurassic basin. Above J13 mts, the sediment becomes significantly finer grained, with lenticular wave ripples and a relatively diverse burrowing assemblage. These deposits are related to deposition in a low-energy marine shelf environment, showing that relatively deep marine conditions were established early in the Pliensbachian. The cored section in well 31/2-3 (Fig. 7) is also from the J13 sequence, but stratigraphically younger than the section in well 30/3-A5. The core comprises medium- to coarse-grained sandstones that can be correlated to the Johansen Formation (Dor6 et al., 1984). Internally, the sandstone can be divided into a series of sharply based fining-upwards units (storeys) characterized by a basal quartzite pebble lag succeeded by cross-stratified, laminated and massive sandstones, sparsely bioturbated by Skolithos and Ophiomorpha. The burrowing traces and sedimentary structures indicate deposition in a marginal marine environment (estuarine, see discussion below) characterized by strong traction currents. The superposition of this sandstone facies above shelf deposits demonstrates significant shoreline regression within the J 13 sequence. The J14 sequence is fully cored in wells 30/3-4, 30/6-16 and 35/10-1. A sharply based, coarsening-upward sandstone interval is present above J14 mts, terminated by the J15 mts in well 35/10-1 (Fig. 7). Wells 30/3-4 and 30/6-16 (Fig. 7) com-

MY

AMMONITES

PALYNO. ZONES

+

SEQUENCES

KEY BIOEVENTS

31OSTRATIGRAPHY

AGE [Hardenbol et al. 1998)

+

+

~

+

+

~+

+

,

EARLY

MICROPAL. ZONES

(part)

(part)

176.5

LATE

-

177.33

AALENIAN -

180 -

MIDDLE

laeviuscula175.81 PJ discites 4

,~O

O'~

CO

03

03

0'3

Or;

CO

O r)

concavum

177.33 bradford-

murchisonae

ensis 178.11 murchi-

180.1

03

("3

03 i

VIJ 7

3D

sonae

179.29

EARLY

pj

03 .......

I

J24

~OR

176.5

+

+

SEQUENCES

i

BAJOCIAN

I ~

J22

179.29

opalinum 180.1

34R

aalensis180.88

pseudoradiosa

-

181.67

-

LATE

dispansum 182.47

.houarsense 84.06

pj

VIJ 6

3C

,~3

>

J20

J20

184.10,.6

32R ---> 185 -

TOARCIAN

variabilis 185.25

VIlDDLE bifrons

B 186.84

pj 3B

falciferum

89.6

tenuicostatum PJ 189.6 3 A spinatum 19o.38 .ATEE ari-margarimargari- imarg ~91.51 tatus ~ I stokesi davoei

PLIENSBACHIAN EARLY

ibex

INEURIAN

--B

~oR~--~

|

l

I------

,,,

I . . !i. .

"

,,- -

-

-

---B

~,9R, ---> MJ 4

B A

~5R ~sF~ 44R

)

.... ~-

.....

.....

A _,I/L . J14

.

-i

.

i

.

-

-_

.

.,--~L~_--~.~_~

= .

n 11-i

m

-

-- -- =_---

.-

- ----

_I --

-

,

--

~----'-

I-H1 1 -

-

l i

,-"

i

J13

B

i

"aricostatum 196,83

--

9

|=

I

MJ 3 30R

LATE (part)

|

5

PJ 193.77 2 C

195.3

195.3

)

J18

l

----

-

A ":Z_~

. . . . . J14

192.64

jamesoni 195 -

60R

j1

188.04

EARLY

190-

MJ

- -- J12--

-- "

l

i

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

J13

J12

PJ 2B

IR'I~V

oxynotu m 197.61

Wells illustrated on correlation figures + Cores illustrated - - Maximum transgressive surfaces (MTS) Vailian type sequence boundaries (SB)

Fig. 5. Stratigraphic distribution of studied cored intervals within the Lower Jurassic Dunlin Group.

153

Sequence stratigraphy of the Lower Jurassic Dunlin Group, northern North Sea

35/9-2 KEY SYMBOLS

~;~ F l u v i a l c h a n n e l fill ~" ~~ Q T

e-

.o

Current ripples Wave ripples Lenticular bedding Trough cross-strat. Low angle cross-strat. Tabular cross-strat. Belemnites Shells Nodules Bioturbation

Delta plain

(I= El __

_(

- ~__.~. o ~ _ _ ~ . ~

!

! i

i .....

I

~

Wave-dominated shoreface

HCS

!

e-

-

El

Hummocky cross-strat. Anconichnus Asterosoma Chonddtes Diplocraterion Ophiomorpha Palaeophycus Pelecypodichnus Planolites Polyclodichnus Rosselia Skolithos Teichichnus Terebellina Zoophycos

Foreshore/ upper shoreface ....................

2650 -- ~ _

ABBREVIATIONS

HCS An Ast Ch Dip Oph Pal Pei Piano Poly Ros Sk Teich Ter Zoo

HCS

RS

Wave-dominated shoreface

L_

70R

SB m.

i~

roarc,

c o

63

~--'~

Offshore zone

i

transition

60R 55

..... I~ .

'~

M,

J16B MTS

- ~'z-f-r-f d

Wave-dominated shoreface

45

.cs RS/TS

t,O "a

Estuarine 2700

c

e'~

E ,-

"-~-

g

~

9

~ n

~

---'-v"r"'~

mouth

bar

J 15 MTS

-.-:rT.~.

44R

1

,,,~ .~ .

Estuarine/ f l u v i a l v a l l e y fill

o

4--- 40R

- ~ o o

~

m !

SB

|

Wave-dominated shoreface 4--- B44

e.o

RS/TS ~ 3 4

L O [,t,. Estuarine

e-

g

b a s i n fill ( ? )

SB ,~

E

~ ~ ' ~ Lower shoreface

2750 ~"

"-a m Q., ~,~

,

~=

SB

e--

o "~ "" o.,,,_,,

~g

TS

Estuarine/ f l u v i a l v a l l e y fill ( ? )

o

tat.

G,~

-

o

e-

r

~ - ~ / -

T T T--gQ

,..

.-~ i._ e.m

-

I Oh

'

_xe__~" -

45---o,,

dh o o o

s

Lower shoreface/ offshore transition zone TS

o

Proximal braid plain/ alluvial fan

Fig. 6. Cored section of well 3 5 / 9 - 2 on the M~l~ay Terrace.

prise very fine- to fine-grained sandstones with wave ripples and lenticular lamination, associated with a diverse burrowing assemblage (Chondrites, Palaeophycus, Planolites, Ophiomorpha, Rosselia) as well as shell debris and scattered belemnite tails. The

same section in well 35/10-1 is significantly coarser

grained, comprising bioturbated and cross-stratified fine- to medium-grained sandstones. The coarseningupward grain size distribution of the J14 sandstones is related to shoreline progradation (Fig. 11). Furthermore, the fine-grained nature of the sandstones, the burrowing style and small-scale sedimentary struc-

154

M.A. Charnock et al.

Sequence stratigraphy of the Lower Jurassic Dunlin Group, northern North Sea

155

Fig. 7. Selective cored sections: facies and sequence interpretation, Horda Platform and Lomre Terrace.

tures in wells 30/3-4, 30/6-16 are related to deposition within the offshore transition to lower shoreface. The coarser grain sizes and cross-stratification seen in well 35/10-1 may reflect a shallower part of the shoreline, and deposition in the middle to upper shoreface is inferred. Notably, a regressive surface of marine erosion (RSME) is inferred at the base of the J 14 sandstones to account for the sharp basal contacts with underlying deeper marine facies (Fig. 7), indicating forced shoreline regression at this stratigraphic level.

Mts J 15 re-established deep, shelfal deposits above the J14 shoreline. However, fine- to medium-grained sandstones of the J15 sequence (wells 30/3-4, 30/6-16, 35/10-1; Fig. 7) rest with a sharp, pebble-strewn contact on fine-grained, bioturbated shelf deposits immediately above J 1 5 rots. These sandstones are characterized by common din-size crossstratification, ripple lamination and planar lamination, with numerous cm-size mudrock partings and drapes throughout. Burrowing is moderate and of low diversity (mainly Planolites and Skolithos). The

156 cross-stratification is generally indicative of sustained traction currents, whereas the mudrock laminae and partings reflect common periods of slack-water deposition. Thus, an estuarine environment is inferred for the J15 sandstones, with the basal surface representing a sequence boundary cut during the preceding relative sea-level lowstand (see Dalrymple et al., 1994). In well 30/6-16 on the western margin of the Horda Platform (Fig. 7), the J15 sandstones are capped by generally dark muddy deposits of sequences J16A, J16B, Jl8 and J20 (basal part). The entire succession is organized into a series of stacked coarsening-upwards units of laminated dark mudrocks in the lower part and bioturbated, occasionally tipple-laminated siltstones and very fine- to fine-grained sandstones in the upper parts. Burrowing is weak to moderate and dominated by Anconichnus. The dark colour and preserved lamination of the mudrocks probably reflect deposition in a slightly oxygen-depleted marine shelf environment, with the coarsening-upwards units reflecting cycles of (distal) shoreline progradation. In contrast to the succession in well 30/6-16, the J16A sequence in well 35/10-1 (Lomre Terrace) contains significant sandstones of inferred shoreface and estuarine origin (Fig. 7). Above J16A rots, wavetippled very fine-grained sandstones burrowed by Helmintoidea are related to deposition in the lower shoreface. These deposits are abruptly succeeded by fine-grained, laminated, cross-stratified and massive sandstones with numerous mudrock partings and laminae, thought to record deposition in an estuarine environment. The emplacement of estuarine sandstones above the initial J16A shoreface indicates a significant base level fall and possible valley incision at this level. The upper part of the J16A sequence shows a regressive to transgressive development. Above the estuarine interval in well 35/10-1, fine-grained sandstones with hummocky cross-stratification and relatively diverse burrowing record deposition in a wave-influenced lower shoreface, whereas coarsergrained sandstones are thought to represent the upper shoreface. Below J16B mts in well 35/10-1, finergrained sandstones with hummocky cross-stratification related to the lower shoreface reflect deepening prior to the flooding of J 16B mts. Similar to well 30/6-16, the J16B and J18 sequences in well 35/10-1 comprise coarsening-upward sequences of mudrocks and sandstones. However, the sections are more bioturbated and contain hummocky cross-stratification as well as belemnite fragments. These deposits are related to deposition in the lower shoreface (J16B) and inner shelf (J18). The facies of sequences J16A to J20 in well 30/6-16 (Fig. 7)

M.A. Charnock et al.

show that a significant deepening occurred in the late Pliensbachian, and that a rather deep marine environment was maintained throughout the early to middle Toarcian. The onset of the J16A sequence is also marked by a significant deepening on the Lomre Terrace (well 35/10-1), but shallower marine conditions seemingly persisted here, particularly during deposition of the J 16A sequence. This may be due to a more proximal setting in this part of the basin. Whereas the J20 mts marks a significant deepening and establishment of anoxic shelf conditions, significant shoreline progradation occurred in the subsequent phase of the J20 sequence. The cored section of sequence J20 in well 30/6-22 contains coarsening-upwards siltstones and sandstones with wave ripples, small-scale hummocky cross-stratification and a diverse trace fossil assemblage (see Fig. 7). Furthermore, the J20 sequence in 30/6-22 is capped by coarse-grained sandstones with cm-size quartzite pebbles. The sandy section is related to deposition in the offshore transition zone to the upper shoreface, reflecting a major seawards shift of the shoreline in the middle to late Toarcian. Notably, the J20 sandstones seen in well 30/6-22 represent the basal parts of the Oseberg Formation as defined by Graue et al. (1987). The cored section above J22 mts in well 30/6-22 shows the main part of the Oseberg Formation and the base of the Rannoch Formation. The Oseberg Formation (above J22 rots) comprises four coarsening-upwards facies sequences of medium- to coarsegrained, pebbly sandstones and associated mudrocks and siltstones. Cross-stratification and crude parallel lamination characterize the otherwise rather massive sandstones, which also are weakly bioturbated (Skolithos, Ophiomorpha). The finer-grained intervals contain wave ripples and are more intensively burrowed. A blackish mudrock interval with abundant Helmithoidea traces is present at the base of the uppermost coarsening-upward sequence (Fig. 7). This mudrock interval is dated as late Toarcian-earliest Aalenian. The primary features of the Oseberg Formation are related to repeated phases of shoreline progradation punctuated by transgressions, demonstrating significant sand emplacement into the basin during late Toarcian times. The blackish mudrock at the base of the last facies sequence is related to deposition in a rather deep shelf environment, and is a candidate for another maximum transgressive surface.

TampenSpur Fig. 8 shows a number of representative core sections from the Tampen Spur. The lower part of the

Sequence stratigraphy of the Lower Jurassic Dunlin Group, northern North Sea

] 57

Fig. 8. Selective cored sections: facies and sequence interpretation, western margin of the North Viking Graben.

J13 sequence, immediately above the Statfjord Formation, is cored in well 34/8-1. As on the Horda Platform, intensively bioturbated siltstones are related to an offshore environment. The overlying fine- to

medium-grained sandstones rest abruptly on the underlying lithology, and are also well structured with common current ripples and dm-size trough crossstratification. Both the sediment type and structures

O'! Oo

Fig. 9. Sequence stratigraphic interpretation and facies correlation of line 1, Oseberg area.

t~

e~

~,~~

~,.~~ e~

~,~~

Fig. 10. Chronostratigraphic interpretation of line 1, Oseberg area.

0-1

M.A. Charnock et al.

160

indicate deposition from strong, sustained traction currents, and a distributary channel/estuarine environment is inferred. Consequently, the base of the sandstone unit is taken as a potential sequence boundary, reflecting a significant basinward shift of the shoreline. The upper part of sequence J13 is cored in well 34/4-5, comprising intensively bioturbated siltstones with scattered wave ripples and belemnite tails. These deposits are related to an offshore environment, reflecting subsequent re-establishment of deep marine conditions upon the distributary/estuarine deposits. Sequence J14 (Fig. 8, well 34/4-5) comprises two facies associations organized into a coarseningupwards sequence. The lower association commences with laminated silty mudrocks near J14 mts, passing vertically into fine-grained, wave-rippled and bioturbated (Chondrites) sandstones and silty interbeds deposited in the lower shoreface. Potentially, a regressive surface of marine erosion (RSME) is indicated at the base of the J14 shoreface, as a similar event can be picked on the Horda Platform (e.g. wells 30/3-4 and 30/6-16; see Figs. 6 and 9). The upper facies association 014; well 34/4-5) comprises sharply based fine- to coarse-grained, cross-stratified and massive sandstones with scattered ripples and sparse bioturbation (Planolites, Skolithos). Furthermore, numerous muddy drapes and partings occur throughout the sandstone body. An estuarine depositional environment is inferred for this facies association with the sharp lower boundary representing a sequence boundary cut during a preceding lowstand. Sequence J15 is partly cored in well 34/2-4 (Fig. 8). The medium-grained sandstones below J16A rots are characterized by dm-size cross-stratification and scattered trace fossils of a low-diversity assemblage (Skolithos, Planolites). As for the J14 sandstones, an estuarine environment is suggested. Above the J16A mts in wells 34/2-4 and 34/10-9 (Fig. 8), intensively bioturbated fine-grained sandstones and wave-rippled/hummocky cross-stratified sandstone/mudrock interbeds are related to deposition in offshore and lower shoreface environments. As on the Horda Platform, the J16A mts seems to mark a relatively significant deepening of the basin in the late Pliensbachian. Sequence J16B (Fig. 8: well 34/10-9) comprises intensively bioturbated siltstones and fine-grained sandstones of assumed offshore and lower shoreface environments. Potentially, the J16B sequence boundary in well 34/8-1 is represented by a coarse-grained sandstone bed immediately below the J 18 mts. Sequence J18 (34/10-9, Fig. 8) is represented by fine-grained sandstones of assumed offshore and lower shoreface origin and is succeeded by a sharply based medium-grained sandstone unit. The sand-

stones are cross-stratified with mudrock partings, and are related to an estuarine environment (see also Dreyer and Wiig, 1995), with the basal surface representing a sequence boundary (SB). This estuarine facies is not present in well 34/8-1 (upper core, Fig. 8), where the sequence comprises fine-grained sandstones, intensively bioturbated by high-diversity trace fossil assemblages (including Palaeophycus and Zoophycos). These deposits reflect deposition in a marine environment, and are placed in the lower shoreface. The upper part of the sandstone body in well 34/8-1 is regarded as a transgressive sandstone as a possible ravinement surface is indicated within the sandy unit due to the presence of Glossifungites trace fossils. Sequences J20 and J22 are not covered by core data in the Tampen area. However, these sequences are dominated by mudrocks on the Tampen Spur, and a deep offshore environment is inferred. Importantly, shallow marine sandstones of J20 and J22 are seemingly restricted to the eastern part of the basin.

Biostratigraphy The maximum transgressive surfaces identified on wireline logs and sedimentologically in core have been calibrated biostratigraphically using a combination of palynology and micropalaeontology. The biostratigraphic zonation (Fig. 3) has been directly calibrated to a standard ammonite biostratigraphy following an outcrop study on the Yorkshire coast. The geochronological ages of the maximum transgressive surfaces are based on the time scale of Hardenbol et al. (1998), which in some instances differs slightly from that of Partington et al. (1993). The following terminology is used to describe the micropalaeontological and palynological events: LO: LCO: LAO: FAO: FCO: FO:

last occurrence (i.e. species "top" or extinction) last common occurrence last abundant occurrence first abundant occurrence first common occurrence first occurrence (i.e. species "base" or evolutionary inception)

The maximum transgressive surfaces have been defined, wherever possible, using the biostratigraphic criteria established by Partington et al. (1993). It has been possible to directly relate most of the bioevents seen in the North Viking Graben to the ammonitecalibrated Lower Jurassic sequence exposed along the Yorkshire coast (to the level of ammonite subzones) suggesting a provincial similarity that is useful for large-scale correlation. Consequently, this study has

Sequence stratigraphy of the Lower Jurassic Dunlin Group, northern North Sea provided an independent test on the stratigraphic calibration by Partington et al. (op. cir.) and is in effect the first detailed practical application of their framework to be published for the North Viking Graben area.

Sequence stratigraphy: calibration and interpretation The sequence stratigraphic interpretation of the Dunlin Group in the North Viking Graben is illustrated by means of three correlation lines depicting key wells relating most of the principal fields in the area (Fig. 1). Wells in the Oseberg area are shown in correlation line 1 (Figs. 9 and 10), correlation line 2 trends approximately N W - S E from the Tampen Spur to the Lomre Terrace and Horda Platform on the eastern margin of the Viking Graben (Figs. 11 and 12), and correlation line 3 trends approximately in an N-S direction along the western margin from the Tampen Spur southwards to the Gullfaks area (Figs. 13 and 14). For each of these lines there are two figures that show (1) a facies and sequence stratigraphic interpretation, and (2) a chronostratigraphic interpretation (Wheeler diagram).

J12 sequence

Age" late Sinemurian-earliest Pliensbachian. J12 mts calibration: latest Sinemurian, Raricostatum Zone, 196.05 Ma. Primary bioevent: below LO Liasidium variabile (event 30R). Example: 30/6-18, 3141 m. Comment: on a supra-regional scale the J12 mts probably equates to the Boreal third-order sequence Si4 mfs of Jacquin and De Graciansky (1998). Lithostratigraphic associations" Statfjord Formation (upper part) and "lowermost" Amundsen Formation. Regional correlation: upper part of the Statfjord megasequence 3 (Steel, 1993). The base of the Dunlin Group is interpreted as being slightly diachronous (latest Sinemurian-earliest Pliensbachian). It is not possible (within the biostratigraphic framework) to equate the J12 mrs to the inundation of the Statfjord Formation in this area as Partington et al. (1993) indicated for the South Viking Graben. If their biostratigraphic calibration is correct then the flooding of the Statfjord Formation is diachronous and the onset of Dunlin deposition is later in the North Viking Graben than further south and suggests that this boundary is even more diachronous than indicated in this study.

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Interpretation: Regressive part: the regressive part of the sequence was not studied in detail since it is represented by sediments of the "upper" Statfjord Formation. Transgressive part: the J12 sequence is only partly represented by sediments of the Dunlin Group. Basal sediments of the Dunlin Group constitute only the later stage transgressive component of the sequence. In some wells (e.g. 30/3-A-5, Fig. 7), it is represented by a marine shale within the lowermost few metres of the Amundsen Formation directly overlying the Statfjord Formation and/or transgressive marine sandstones of the Nansen Member. A series of minor sandstones with fining-upward cycles and representing transgressive backstepping facies is locally developed, such as in well 30/6-22. Significantly, these represent the first marine transgression within the Dunlin Group and the inundation of the Statfjord Formation. This transgressive event is dated as being earliest Pliensbachian since late Sinemurian markers are present within or below the Nansen Member (e.g. 30/6-18) and therefore post-dates the J12 maximum flooding surface of Partington et al. (1993). A thin marine shale unit separates the major progradational units of the Statfjord and Johansen megasequences (Steel, 1993). This represents a phase of marine transgression that according to Steel (op cir.) penetrated far towards the Norwegian hinterland and to at least block 31/3. In this study it is clearly identifiable as far east as well 31/2-3 on the Horda Platform (Fig. 11).

J13 sequence (new; this study)

Age: early Pliensbachian. J13 mts calibration: earliest Pliensbachian, Jamesoni Zone, 194.89 Ma. Primary bioevent: above LO Liasidium variabile (event 30R). Secondary bioevents: below LCO Ogmoconcha spp., notable LCO O. danica, LCO O. amalthei and LO Ogmoconcha sp. B Apostolescu (event 32); below LCO Cerebropollenites cf. thiergartii (event 34). Examples: 30/6-18, 3120.5 m; 30/3-2, 3227 m (see Fig. 9). 31/2-19S (Fig. 11). Comment: the base of the new sequence J13 is defined above the LO Liasidium variabile (event 30R). Thus this does not correspond to the J 12 mfs of Partington et al. (1993) which is defined below this event and calibrated to the late Sinemurian Raricostatum zone. The LO of L. variabile has in the present study been found at a level within the Nansen Member within the Statfjord Formation (e.g. 30/6-18). This seems to be a general feature in the studied area and therefore it is not possible to equate the J12 rots to the inundation of the Statfjord Formation as

Fig. 11. Sequence stratigraphic interpretation and facies correlation of line 2, Troll/Fram area.

e5

~..~~

~..~~ e5

t...,

Fig. 12. Chronostratigraphic interpretation of line 2, Troll/Fram area.

r

(3)

Fig. 13. Sequence stratigraphic interpretation and facies correlation of line 3, Gullfaks area.

~..~~

Q t...,

e5

r~ ~..~~ e~

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Fig. 14. Chronostratigraphic interpretation of line 3, Gullfaks area.

~n

M.A. Charnock et al.

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indicated by Partington et al. (op. cit.) for the South Viking Graben area. This suggests that the termination of the Statfjord Formation results from a series of transgressive events rather than one single event as indicated by Partington et al. (op cit.; Fig. 12). On a supra-regional scale the J13 mts correlates with the Boreal second-order transgressive-regressive cycle 5 mts and third-order sequence Si5 mfs of Jacquin and De Graciansky (1998). This transgressive event was also recognised by Van Buchem and Knox (1998) and Hesselbo and Jenkyns (1998) in Yorkshire. The former authors considered it to be a major, i.e. 2nd order, early Pliensbachian transgressive event of worldwide significance. Lithostratigraphie associations: Amundsen and Johansen Formations. Regional correlation" Johansen megasequence 4 of Steel (1993). Interpretation. Regressive part: Steel (1993) correctly noted that the lower boundary of his megasequence 4 (equivalent to sequence J13 of this study; see Fig. 4) is well defined by a marine shale unit (assigned to the Amundsen Formation) that separates the Statfjord Formation from the overlying Johansen Formation over many areas of the Horda Platform. On the Horda Platform and the Lomre Terrace, e.g. wells 31/2-19S, 35/10-1 and 35/11-2 (Fig. 11), the regressive phase is characterised by a series of coarsening-upward units. These form a wedge of marine shoreface-dominated sandstones that represents the initial progradation of the sequence. Marjanac and Steel (1997) attributed these Johansen Formation sandstones to a relative fall of sea level and deposition in a large delta confined within a broad incised valley. Within the sequence are mudstone units that may represent minor transgressive surfaces but these are not readily correlatable within current biostratigraphic resolution. The regressive phase, therefore, is probably characterised by a series of progradational cycles that reflect minor base level changes or variations in sediment supply within the shore zone. Elsewhere, and west of the Horda Platform, the regressive part of the sequence is mud-dominated and deposited in a marine shelf setting. However, on the Tampen Spur area (e.g. 34/2-4) the sequence is represented by a thick, stacked unit of aggradational sands that may have been sourced locally. Transgressive part: within the J13 sequence, sandstones of both estuarine and marine origin have been interpreted, e.g. 31/2-3 and 35/10-1 (Fig. 7). The superposition of estuarine sandstone facies above shelf deposits indicates significant shoreline migration with the J13 sequence. The interpretation of estuarine facies is in contrast to others (e.g. Steel, 1993) and is open to debate since core data are lim-

ited. Steel (op. cit.) noted the possible presence of alluvial sediments and more brackish-water conditions in the Johansen Formation in the most easterly wells on the Horda Platform but he generally considered the succession as representing nearshore and inner shelf deposits following Dor6 et al. (1984). An interval representing a possible estuarine/fluvial valley fill is also interpreted in the cored section of well 35/9-2 (Fig. 6) on the Mhl0y Terrace that may indicate a more easterly source of sediments into the Fram area. In general, the later transgressive component of the J13 sequence is represented by a fining-upward retrogradational cycle of backstepping facies (e.g. well 31/2-19S) which is overlain by a thin but widespread shaly level associated with the next marine transgression and the J14 mts (cf. Marjanac and Steel, 1997). In some wells (e.g. 35/10-1) the transition from aggradational sandstones to transgressive mudstones is relatively abrupt and represents abandonment. A major regionally extensive transgressive event tied to the J 14 mts and approximately coinciding with the early/late Pliensbachian boundary terminates this sequence and Johansen sandstone deposition. J14 sequence

Age: late Pliensbachian. J14 mts calibration: "earliest" late Pliensbachian, basal Margaritatus Zone, 191.5 Ma. Primary bioevents: above LO Gramminacythere ubiquita and LO Ogmoconchella mouhersensis (event 36).

Secondary bioevents: above FO Nannoceratiopsis senex/gracilis (event 40); above LCO Cerebropollenites cf. thiergartii (event 34); below FO Leuhndea spinosa (event B55R); above LCO Ogmoconcha spp. notable LCO O. danica, LCO O. amalthei and LO Ogmoconcha sp. B Apostolescu, 1961 (event 32). Examples: 30/6-9 and 30/3-2 (both with only moderate biostratigraphic constraints). Comment: this maximum transgressive surface corresponds to the J14 mfs of Partington et al. (1993) and the surface lies at a level closely associated with the first downhole appearance of early Pliensbachian marker species and above the inception (FO) of a suite of dinocyst species associated with "late" early Pliensbachian and younger sediments, notably, FO N. gracilis/senex (event 40R). The J14 mfs may coincide with the Boreal third-order sequence P14 mts of Jacquin and De Graciansky (1998) although the stratigraphic calibration is not certain. Lithostratigraphic associations: Burton and Cook (lower part) Formations. Regional correlation: lower part of the Cook mega-

Sequence stratigraphy of the Lower Jurassic Dunlin Group, northern North Sea sequence 5 of Steel (1993) based on direct comparison with well 31/2-5. Interpretation. Regressive part: the base of the Cook Formation in many instances (e.g. well 30/6-11, Fig. 9; 35/10-1, Fig. 7 and 34/4-5, Fig. 8) is represented by a sharply based coarsening upward sandstone. This surface is interpreted as a regressive surface of marine erosion (RSME) (cf. Marjanac and Steel, 1997) or marine downshift surface sensu Dreyer and Wiig (1995) that marks the initiation of relative sea-level fall (falling stage systems tract of their terminology). In proximal parts of the Gullfaks Field, Dreyer and Wiig (1995) noted that this section may be absent due to erosion from the subsequent maximum regressive phase. This is considered likely although it has not been possible to verify this suggestion due to the lack of appropriate well coverage in this study. The sequence is laterally extensive on the Horda Platform and Lomre Terrace and is typically represented by a coarsening upward sandy unit (e.g. 31/2-19S and 35/10-1, Fig. 7) interpreted as representing deposition in a marine middle to upper shoreface setting. Transgressive part: a widespread estuarine/fluvial valley fill unit (e.g. 34/2-4, Fig. 11; 35/4-1, Fig. 11; and 35/9-2, Fig. 6) is interpreted in the northern part of the study area. However, in many instances (e.g. wells 30/3-2 and 31/4-3) the transgressive part of this sequence is either condensed or absent as a result of erosion from the overlying sequence (see discussion of this event in the next sequence).

J15 sequence (new; this study)

Age: late Pliensbachian. J15 mts calibration" late Pliensbachian, intra-Margaritatus Zone, 191 Ma. Primary bioevent: below LCO Botryococcus spp. (event 44R). Secondary biovents: below LO Dentalina terquemi (event 42); above FO Luehndea spinosa (event B55R). Example: 30/6-7, 3022 m (see Fig. 9). Comment: it is uncertain how the J15 mts relates to the supra-regional Boreal third-order sequences of Jacquin and De Graciansky (1998) since they recognise a series of sequences around this stratigraphic level which are poorly defined in terms of palynology or micropalaeontology. The most likely candidate is the P15 mfs (or possibly P16 mfs). Lithostratigraphic associations: Cook Formation (part). Regional correlation" upper part of the Cook megasequence 5 of Steel (1993) based on direct comparison with wells 30/6-9 and 31/2-5 (see his fig. 8); upper

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part of the Cook-1 of Marjanac and Steel (1997) based on comparison with well 30/3-2. Interpretation. Regressive part: in wells 34/2-4 and 35/8-1 (Fig. 11) the regressive part is characterised by a coarsening-upward profile but elsewhere this part of the sequence is poorly represented or possibly truncated (beyond biostratigraphic resolution). Transgressive part: the sequence is dominated by a laterally extensive unit of "blocky" sandstones (e.g. wells 30/6-7 and 35/4-1) that frequently have sharp bases (e.g. well 30/6-7). Evidence from core (e.g. wells 30/3-4, 30/6-16 and 35/10-1, Figs. 6 and 7) suggests that these sandstones are of estuarine origin and the basal surface represents a sequence boundary cut during the preceding relative sea-level lowstand. Marjanac and Steel (1997) came to a similar conclusion based on the analysis of well 30/3-2 (see their fig. 7). These sand units appear to be correlatable over a large area of the North Viking Graben and were deposited onto a surface representing a marked intra-late Pliensbachian lowering of relative sea level. Steel (1993), however, interpreted the sandstones to represent a wave/storm-dominated shelf system and suggested that this unit is characteristic of a lowstand or forced regression. In either case they represent a period in which sandstone deposition extended much further beyond the Horda Platform. A time-equivalent unit is also present on the Tampen Spur (e.g. 34/2-4) which herein is also interpreted as of estuarine origin. but it is not clear whether this has been generated from an easterly source (as suggested by Steel, op. cit.) or westerly source. The late stage transgressive part of this sequence is poorly developed and is represented by either an abrupt termination of sand (e.g. well 30/3-4) or a thin, fining-upward unit (e.g. well 30/3-2, Fig. 9). Steel (op. cit.) also observed the abrupt termination of sandstone deposition and the thinness of the transgressive component. A major, regionally extensive transgressive event culminating in the J16A mts of intra-late Pliensbachian age terminated Cook sandstone deposition on the Horda Platform and most noticeably within the Oseberg area (Fig. 9). In this area significant sandstone deposition was not resumed until middle Toarcian times.

J16 sequence

Age: late Pliensbachian-early Toarcian. Comment: this study follows Partington et al. (1993) who distinguished two sequences within J16 and these are discussed separately below.

J16A sequence

Age: late Pliensbachian

M.A. Charnock et al.

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J16A mts calibration" late Pliensbachian, intra-Margaritatus Zone, 190.5 Ma. Primary bioevent: below LO Ogmoconcha/Ogmoconchella spp. (event 49R). Secondary biovents: below LCO Luehndea spinosa (event 50R); below LO Kraeuselisporites reissingeri (event 45R); above LCO Botryococcus spp. (event 44R); above LO Dentalina terquemi (event 42). Examples: 30/6-7, 2956 m (see Fig. 9), 35/9-2, 2694.5 m (core) (Fig. 6 and 11). Comment: Partington et al. (1993) discussed in detail the problems in distinguishing the J 16A and J 16B mts and stressed the importance of integrating both palynological and micropalaeontological data sets. In this study we suggest that this surface can be recognised in the North Viking Graben based on its relationship below the LO Ogmoconcha/Ogmoconchella spp. (event 49R) and the LCO Luehndea spinosa (50R). Its relationship below the FAO sphaeromorphs as suggested by Partington et al. (op. cit.) is difficult to determine in offshore settings without the use of cores or sidewall cores because of frequent cavings. As a guide we have also identified the surface using its relationship above the inception (FO) of the dinocysts Nannoceratopsis gracilis/senex and FO L. spinosa (B55R). Other events suggested by Partington et al. (op. cit.) namely FO Scriniocassis weberi and FO Maturodinium inornatum are less well documented. Likewise the intra-late Pliensbachian events LO Dentalina matutina and LO Haplophragmoides lincolnensis are known to be important elsewhere in the North Sea but are rarely recorded in this area of the North Viking Graben. The calibration of the J16A mfs to the supra-regional Boreal scheme of Jacquin and De Graciansky (1998) is unclear since they recognise many third-order sequences over this stratigraphic interval which are poorly defined in terms of palynology or micropalaeontology. Based on the available data the most likely candidates are the P16 mfs or P17 mfs although the latter, more pronounced event, is interpreted as early Spinatum Zone. Lithostratigraphic associations: Cook Formation (part). Regional correlation: this is equivalent to the lowest part of the Drake megasequence 6 of Steel (1993) based on direct comparison with wells 30/6-9 and 31/2-5 (see his fig. 8). Note in particular that Steel's Cook megasequence 5 only corresponds to the lower part of the Cook Formation and the J14/J15 sequences of this study. This exemplifies the problem with the broad usage of the term "Cook". This sequence also equates to the upper part of the Cook-2 of Marjanac and Steel (1997) based on comparison with well 30/3-2.

Interpretation. Regressive part: the J16A mts is a regionally significant transgressive event and marks a significant deepening of the basin during the late Pliensbachian which on the Tampen Spur and Horda Platform areas was maintained throughout the early and middle Toarcian. In the Oseberg area significant sandstone deposition terminated at this event, e.g. 30/6-22 (Figs. 7 and 9) which is earlier than in the Gullfaks area where Cook sandstones are commonly found until the J18 sequence. On the Horda Platform, the regressive part of the J16A sequence is represented by the first of a series of stacked coarsening-upward units of offshore, marine sediments reflecting cycles of (distal) shoreline progradation (e.g. 30/6-16, Fig. 7; 30/6-22, Fig. 9). In the Veslefrikk area (e.g. well 30/3-4, Fig. 9), laterally discontinuous sandstones are developed at the top of a much thicker, underlying coarsening-upward progradational unit. Lower shoreface and offshore, fine-grained sandstones and mudstones also dominate the sequence in the Tampen (e.g. 34/2-4, Fig. 11) and Gullfaks areas (e.g. 34/10-9, Fig. 8). Transgressive part: a sandstone-dominated unit in the core of well 35/10-1 (Fig. 7) on the Lomre Terrace is interpreted as representing deposition in an estuarine environment. The emplacement of estuarine sandstones above a shoreface sequence indicates a significant base level fall and possibly valley incision prior to the subsequent transgressive phase (cf. Marjanac and Steel, 1997). It is possible that this area provides the easterly source for sediments in the Gullfaks Field inferred by Dreyer and Wiig (1995). On the Horda Platform the transgressive component is typically represented by a thin fining-upward unit overlying a much thicker coarsening-upward unit (e.g. 30/6-22, Fig. 9). J16B sequence

Age: earliest Toarcian. J16B mts calibration: earliest Toarcian, Tenuicostaturn Zone, 189 Ma. Primary bioevent: above LO Ogmoconcha/Ogmoconchella spp. (event 49R).

Secondary bioevents: below LCO Luehndea spinosa (event 50R); above LO Kraeuselisporites reissingeri (event 45R). Example: 31/4-3, 2511 m (see Figs. 9 and 11); 30/6-9, 2784 m (Fig. 9). Comment: this corresponds to the J 16B MFS of Partington et al. (1993). They noted that palynological data alone are insufficient to distinguish this surface and suggest placing emphasis on the ostracods and

Sequence stratigraphy of the Lower Jurassic Dunlin Group, northern North Sea foraminifera. They defined the surface as being below the LCO Luehndea spinosa (50R) but above LO Ogmoconcha/Ogmoconchella spp. (event 49R) defining the Toarcian/Pliensbachian boundary. In the Yorkshire outcrop data the ostracods become extinct in the youngest ammonite subzone of the Spinatum Zone. We differ from Partington et al. (op. cit.) in not attaching weight to the benthic foraminifer LO Marginulina prima which we consider to range younger than the ostracods into the early Toarcian (Tenuicostatum Zone) and is only rarely recorded in this area. This is supported by onshore ranges (see Copestake and Johnson, 1989). In addition to the events cited above, this study also has used the LO Kraeuselisporites reissingeri (event 45R) to define Pliensbachian sediments and facilitate the positioning of this surface. We consider it necessary to use a (total) suite of bioevents to calibrate the surface rather than one single event. On a supra-regional scale the J 16B mts equates, with a high degree of certainty, to the Boreal third-order P18 mfs of Jacquin and De Graciansky (1998). Lithostratigraphic associations: Cook Formation, lower part of the Drake Formation. Regional correlation: lower part of the Drake megasequence 6 of Steel (1993) based on direct comparison with wells 30/6-9 and 31/2-5 (see his fig. 8 and the discussion under the J 16A sequence). It is also equivalent to the upper part of the Cook 2 sequence of Dreyer and Wiig (1995) (see Fig. 4). Note that the Cook-3 unit of Marjanac and Steel (1997) is considered to equate to the J18 sequence and not J16B sequence as indicated on their fig. 2 based on calibration with well 30/3-2. Interpretation. Regressive part: sandstones associated with the J16B sequence are rarely developed on the Horda Platform and Lomre Terrace and the J 16B mts reflects a significant deepening event within the early Toarcian. Its effect is most marked on the Lomre Terrace (e.g. 35/10-1; Figs. 7 and 11) and MfilOy Terrace (e.g. 35/9-2, Fig. 6) where more marine conditions were established. In the Oseberg area the sequence is locally condensed and mud-dominated as a result of either sediment by-pass or starvation (e.g. 30/6-16, Fig. 7). Dreyer and Wiig (1995) indicated that the boundary between Cook reservoir units 2 and 3 in the Gullfaks area corresponds to a Vailian-type sequence boundary and is associated with a major change in drainage pattern. This corresponds to a level within the J 16B sequence of this study based on calibration with well 34/10-B-4 (equates with well B in fig. 12 of Dreyer and Wiig, op. cit.). Transgressive part: the transgressive part of the J 16B sequence is generally represented by a thin, condensed, fining-upward sequence (e.g. 30/3-4, Fig. 9) that culminated in the regionally identifiable J 18 rots.

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In the Gullfaks area, the lack of a well-defined transgressive component may be explained by erosion from the overlying J 18 regressive phase (e.g. 34/11-1) where the interval is represented by a series of stacked coarsening-upward sandstones. J18 sequence

Age: early-middle Toarcian. J18 mts calibration: early Toarcian, Falciferum Zone, 187.75 Ma.

Primary bioevent: below LCO sphaeromorph acritarchs (event 60R).

Secondary bioevents: below LO Luehndea spinosa (event 55R); above LCO Luehndea spinosa (event 50R); above LO Kraeuselisporites reissingeri (event 45R). Examples: 30/6-9, 2766.5 m (see Fig. 9), 35/9-2 2682.5 m (core) (Figs. 6 and 11). Comment: details of the age calibration are discussed by Partington et al. (1993) who argue that the record of LO Luehndea spinosa above the surface suggests a calibration to at least the early Toarcian Tenuicostaturn Ammonite Zone based on the onshore borehole data of Riding (1987). Partington et al. (1993) indicated that the surface could be calibrated to the Exaratum Ammonite Subzone by extending the range of L. spinosa for unspecified reasons. Our study of the Yorkshire section provides positive support for this interpretation since the extinction of this species is, in fact, within the Exaratum Subzone. This correlates the transgressive surface to the global anoxic event of Jenkyns (1988). On a supra-regional scale that the J 18 mts coincides (on an ammonite zonal level) to the peak transgression of the Ligurian cycle of Jacquin and De Graciansky (1998, chart 6) and cycle 6 of De Graciansky et al. (1998, not fig. 2), the first of two transgressive-regressive second-order cycles that characterise the Jurassic of Europe. However, on a subzonal ammonite scale, i.e. Exaratum Subzone, the J18 mts appears to correlate to the Boreal third-order sequence Toal mfs rather than the more significant Toa2 mfs of Jacquin and De Graciansky (1998). Lithostratigraphic associations: Cook and Drake Formations. Regional correlation: lower part of the Drake megasequence 6 of Steel (1993) based on direct comparison with well 31/4-3 (see his fig. 9); Cook-3 of Marjanac and Steel (1997) based on comparison with well 30/3-2. The sequence is widespread in the northern North Sea, although sandstones are less common north of the Gullfaks area. Sandstones in the Oseberg, Troll and Fram areas (e.g. 35 / 10-1) have been lithostratigraphically assigned as units of the Drake Formation.

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Interpretation. Regressive part: the J18 mts is connected with the regional "mid"-early Toarcian Falciferum Zone anoxic event established over large areas of northwest Europe. On the Horda Platform (e.g. 30/6-22, Fig. 9) and Lomre Terrace (eg. 31/2-19S, Fig. 11) the regressive part of this sequence consists of a series of coarsening-upward sandstones. In the Gullfaks area (e.g. 34/11-1, Fig. 13) a lowstand wedge of deltaic to marginal marine deposits (Cook-3B/C) has been interpreted by Dreyer and Wiig (1995). Transgressive part: a series of backstepping transgressive sandstones are present on the Horda Platform and Lomre Terrace (e.g. 31/2-19S, Fig. 11). In the Tampen area a transgressive sandstone has also been identified in core in well 34/8-1 (Fig. 8). In the Gullfaks area sandstone deposition was terminated abruptly (e.g. 34/10-9) and the late stage transgressive part is poorly represented or condensed above an inferred estuarine sandstone unit (e.g. 34/10-9, Fig. 8). J20 sequence (new; this study)

Age: middle-late Toarcian. J20 mts calibration: middle Toarcian, Variabilis Zone, 184.86 Ma.

Primary bioevent: below LAO Chasmatosporites spp. (event 63R).

Secondary biovents: below LO Camptocythere toarciana (event 62R); above LCO sphaeromorph acritarchs (event 60R); above LO Luehndea spinosa (event 55R).

Example: 30/3-4, 2984 m (see Fig. 9). Comment: the ostracod Camptocythere toarciana provides a means of age calibration since this species which occurs above the rots (e.g. 34/8-1, 3007 m) ranges no younger than the middle Toarcian Variabilis Zone onshore (Lord, 1978 and Yorkshire outcrop data). On a supra-regional scale the J20 mts coincides with the Toa4 mfs of Jacquin and De Graciansky (1998). Lithostratigraphic associations: Drake Formation. This sequence includes the informally named "Drake sand unit" and the basal part of the Oseberg Formation as defined by Graue et al. (1987). Regional correlation" part of the Drake megasequence 6 of Steel (1993) based on direct comparison with well 31/4-3 (see his fig. 9); Cook-4 of Marjanac and Steel (1997) based on comparison with well 30/3-2; J18 cycle (part) of Parkinson and Hines (1995) based on direct comparison with well 34/8-1. There is potential to subdivide this unit into two separate sequences although this is currently beyond stratigraphic resolution. In some instances (e.g. 30/6-11, 3590 m) there is a potentially correlatable

M.A. Charnock et al. log signature that may represent a maximum transgressive surface but this requires further study. Interpretation. Regressive part: the lower part of this sequence is dominated by fine-grained sediments and consists of one or sometimes two (e.g. 30/3-2, Fig. 9) coarsening-upward units. In the Veslefrikk and Brage areas the regressive part of this sequence is represented by a progradational wedge of marine, lower and upper shoreface sandstones (e.g. 30/3-2 and 30/6-22, Fig. 7). The base of these sandstones may be represented by a marked change in log character interpreted as possibly representing an erosional hiatus and a basinward shift in deposition that Steel (1993) related to the tectonic uplift of the Horda Platform. This interpretation is considered likely since these sandstones are locally restricted to this area. In the Tampen area the sequence is dominated by mudrocks that were deposited in a relatively deep marine environment. Transgressive part: this part of the sequence maybe represented by a fining-upward mudstone-dominated interval (e.g. 30/6-7 and 31/4-3, Fig. 9) although in many wells on the Horda Platform (e.g. 30/6-11) it is truncated by later progradational phases of the Oseberg Formation.

J22 sequence

Age: late Toarcian-Aalenian. J22 mts calibration: late Toarcian, Levesquei Zone, 182 Ma.

Primary bioevent: below LCO Parvocysta/Phallocysta spp. (event 64R).

Secondary biovents: below LCO Haplophragmoides spp. (event 65R); above FCO Parvocysta/Phallocysta spp. (event B64R); above LAO Chasmatosporites spp. (event 63). Example: 31/4-3, 2385 m (Figs. 9 and 11). Comment: this transgressive event is defined using the same criteria as Partington et al. (1993). The age calibration of this surface is primarily based on its relationship within the total range of Parvocysta nasuta which ranges from the "earliest" Aalenian Opalinum Zone to the "latest" middle Toarcian Variabilis Ammonite Zone. In the context of the supra-regional Boreal third-order sequences of Jacquin and De Graciansky (1998) there is apparently no equivalent surface. Lithostratigraphic associations: Drake Formation and part of the Oseberg Formation as defined by Graue et al. (1987). Regional correlation: upper part of the Drake megasequence 6 of Steel (1993) based on direct comparison with well 31/4-3 (see his fig. 9).

Sequence stratigraphy of the Lower Jurassic Dunlin Group, northern North Sea

Interpretation.

Regressive part: in some wells (e.g. 31/4-3, Fig. 9) a coarsening-upward unit of sandstones is present but more often there is an abrupt change in facies that corresponds to the boundary between the Dunlin and Brent Groups and in many cases, the base of the Oseberg Formation which is interpreted as representing a marked basinward shift in deposition (e.g. 30/6-7; compare with Rider, 1996, fig. 15.3). Note, however, that the initial deposition of sandstones conventionally assigned to the Oseberg Formation started within the previous J20 sequence in the Veslefrikk and Brage areas and that this represents a later phase of progradation. In the cored section of well 30/6-22 (Fig. 7) the section comprises four coarsening-upward, sandstone-dominated, lower to upper shoreface sequences that represent a complicated history of repeated shoreline progradation and subsequent transgression in an overall regressive regime. Transgressive part: the transgressive part of the sequence is outside the scope of the present study being represented by deposition within the Brent Group.

Reservoir potential In the North Viking Graben area the Cook Formation has proven hydrocarbons and is a secondary reservoir in the Statfjord, Gullfaks (Dreyer and Wiig, 1995), Oseberg and Veslefrikk Fields. The Drake Formation is also known to be gas-prone. The discussion of depositional environments and stratigraphic sequences have significant bearings on the reservoir prediction in the Dunlin Group. The sedimentological study suggests that the main reservoir sandstones were deposited within estuarine environments. Particularly, the vertical and lateral alternation between these sandy systems and deposits of shelfal and lower shoreface environments (Figs. 9, 11 and 13) indicates that sand emplacement throughout the Pliensbachian was controlled by major falls in relative sea level. These major sea-level falls brought about major basinwards translations of the contemporary shoreline, with pronounced valley incision. This suggests that forced regression is the main mechanism by which reservoir quality sand is distributed into the basinal areas. One of the problems of reservoir potential within the Dunlin Group is maintaining high porosity and permeability values at deep burial depth. Ehrenberg (1993) documented the preservation of anomalously high porosities in some late Pliensbachian sandstones in the Veslefrikk Field, due to the presence of graincoating chlorite. The presence of grain-coating chlorite is an important factor in preserving good reservoir

171

quality at great burial depths although its occurrence alone does not guarantee these factors. The development of this mineral phase may be favoured by deposition in marine environments with a significant supply of fresh water, and can lead to the preservation of reservoir properties at burial depths below 4 km. Consequently, the estuarine setting is ideal for the formation of this mineral phase. A predictive model to explain the occurrence of chlorite in terms of the depositional environment and its position within a sequence stratigraphic framework may be highly valuable for ongoing exploration in the area. Chlorite coatings have been widely observed in both the northern and southern parts of the study area in sands restricted to sequences J 14 and J 15. These sequences are interpreted as representing an intra-late Pliensbachian period that was strongly affected by relative sea-level lowstands and the widespread development of estuarine systems. The forced regressive mechanism for sandstone emplacement in the Dunlin Group may also have led to deposition of reservoir lithologies outside the area of economic deposits in the Brent Group. For instance, the estuarine facies identified in well 34/2-4 (J15, Fig. 8) occurs well outside the northern pinchout of the Brent Group (see Graue et al., 1987; Johannessen et al., 1995). Similar observations have also been made in the Sogn area, where coarsegrained estuarine sandstones of sequences J14 and J15 (well 35/4-1, Fig. 11) occur below non-reservoir lithologies age-equivalent to the Brent Group.

Conclusions The main purpose of this study has been to define a sequence stratigraphic framework of the Lower Jurassic Dunlin Group within Norwegian Quadrants 30, 31, 34 and 35 of the North Viking Graben. It is intended that this regional study provides a framework in which more focused field and block studies can be viewed in a wider context and that the sequence stratigraphic methods employed provide a better understanding of sediment distribution patterns than conventional interpretations based on the lithostratigraphic scheme. Nine sequences are defined on the basis of their depositional systems and bounding surfaces utilising well-log correlations, detailed core descriptions and extensive biostratigraphic information. The boundaries represent maximum transgressive surfaces (mts) and mark the points between a series of regionally identifiable regressive-transgressive episodes in the Early Jurassic (latest Sinemurian to Toarcian) section of the North Viking Graben. These surfaces are calibrated to standard schemes established elsewhere

1 72

within the North Sea Basin and the ammonite-dated Lower Jurassic section exposed along the Yorkshire coast. More specifically, the final conclusions of the present study are: (1) The sequences have thicknesses of between 10 m and 100 m and a duration of approximately 1-4 million years. They consist of facies representing deposition in a variety of shelf, shoreline and estuarine systems. The significance of estuarine facies within this interval has been generally underestimated. (2) The J13 rots represents the terminating transgression of the Statfjord Formation. Sandstones of the Johansen Formation and deposited during the J13 sequence are locally developed on the Horda Platform. (3) The J13 sequence was terminated by a regionally identifiable transgressive event (J14 mts) which is related to the deposition of the Burton Formation. (4) The Cook Formation was deposited in five separate sequences (J14 to J18) which are characterised by thick, geographically widespread, estuarine and marine shoreface sandstones deposited during periods of lowstand incision, progradation and transgression. (5) The J16A mts (intra-late Pliensbachian) was a significant transgressive event that terminated Cook sandstone deposition in the Oseberg-Veslefrikk area although sandstone deposition persisted longer, into sequence J 18 (early Toarcian) in the Gullfaks field. (6) Two early Toarcian transgressions, namely J16B mts and J18 mts profoundly affected the later depositional styles since later sequences are dominated by marine sediments and more locally restricted sandstones deposited in positions centred on the Lomre Terrace and the Horda Platform on the eastern basin margin. (7) The relationship between the Dunlin Group and overlying Brent Group is more complex than previously envisaged. The J20 sequence includes sandstones that conventionally have been assigned to the Oseberg Formation. It can be demonstrated that, in the Veslefrikk and Brage areas, sandstone deposition began within the late Toarcian and that these sandstones represent the first of a series of progradational episodes on the Horda Platform and which continued across the Early-Middle Jurassic boundary.

Acknowledgements The authors would like to thank Norsk Hydro and Robertson Research International Ltd for permission to publish this study. We are also indebted to the partners of PL194 (Elf, Saga and Statoil) for permission to publish the results from this licence area. We would particularly like to thank Roger Davey (Robertsons) and the two reviewers Lars-Magnus Ffilt (Statoil) and Ron Steel (University of Wyoming) for

M.A. Charnock et al.

constructive criticisms and suggestions on improving the manuscript. This study was carried out within the research project on Deep Structures under the guidance of Harald Flesche and Mogens Ramm. They are gratefully acknowledged for help and encouragement. Within the context of the Deep Structures project we would especially like to thank Ruth Elin Midtbr for discussions on the reservoir potential of the sequences. Thanks are also directed to Gry Arnesen and Tom Thorstensen in the Norsk Hydro drafting department for support in the preparation of this paper.

References Badley, M.E., Egeberg, T. and Nipen, O., 1984. Development of rift basins illustrated by the structural evolution of the Oseberg feature, Block 30/6, offshore Norway. J. Geol. Soc., London, 141: 639-649. Cannon, S.J.C., Giles, M.R., Whitaker, M.E, Please, EM. and Martin, S., 1992. A regional reassessment of the Brent Group, U.K. Sector, North Sea. In: A.C. Morton, R.S. Haszeldine, M.R. Giles and S. Brown (Editors), Geology of the Brent Group. Geol. Soc., London, Spec. Publ., 61: 81-107. Copestake, E and Johnson, B., 1989. The Hettangian-Toarcian. In: D.G. Jenkyns and J.W. Murray (Editors), A Stratigraphical Atlas of Fossil Foraminifera. Ellis Horwood, Chichester, pp. 129-188. Dalrymple, R.W., Zaitlin, B.A. and Boyd, R., 1994. Estuarine facies models: conceptual basis and stratigraphic implications. J. Sediment. Petrol., 62:1130-1146. De Graciansky, E-C., Jacquin, T. and Hesselbo, S.P., 1998. The Ligurian cycle: an overview of Lower Jurassic 2nd-order transgressive/regressive facies cycles in western Europe. In: R-C. De Graciansky, J. Hardenbol, T. Jacquin and ER. Vail (Editors), Mesozoic and Cenozoic Sequence Stratigraphy of European Basins. Soc. Econ. Paleontol. Mineral. Spec. Publ., 60: 467-479. Dor6, A.G. and Gage, M.S., 1987. Crustal alignments and sedimentary domains in the evolution of the North Sea, North-east Atlantic Margin and Barents Shelf. In: J. Brooks and K. Glennie (Editors), Petroleum Geology of North West Europe. Graham and Trotman, London, pp. 1131-1148. Dor6, A.G., Hamar, G.E, Lilleng, T., Shaw, N.D., Skarpnes, O. and Vollset, J., 1984. Revised Jurassic lithostratigraphy of the Norwegian North Sea, northern area. In: J. Vollset and A.G. Dor6 (Editors), A Revised Triassic and Jurassic Lithostratigraphic Nomenclature for the Norwegian North Sea. Norwegian Petroleum Society (NPF) Special Publication, 3. Elsevier, Amsterdam, pp. 2-53. Dreyer, T. and Wiig, M., 1995. Reservoir architecture of the Cook Formation on the Gullfaks field based on sequence stratigraphic concepts. In: R.J. Steel, V. Felt, E.E Johannessen and C. Mathieu (Editors), Sequence Stratigraphy of the Northwest European Margin. Norwegian Petroleum Society (NPF), Special Publication 5. Elsevier, Amsterdam, pp. 109-142. Ehrenberg, S.N., 1993. Preservation of anomalously high porosity in deeply buried sandstones by grain-coating chlorite: examples from the Norwegian continental shelf. Am. Assoc. Pet. Geol., Bull., 77: 1260-1286. Eynon, G., 1981. Basin development and sedimentation in the Middle Jurassic of the northern North Sea. In: L.V. Illing and G.D. Hobson (Editors), Petroleum Geology of the Continental Shelf of North-West Europe. Heyden, London, pp. 196-204. Fa~rseth, R.B., 1996. Interaction of Permo-Triassic and Jurassic extensional fault-blocks during the development of the northern North Sea. J. Geol. Soc., London, 153: 931-944. Fjellanger, E., Olsen, T.R. and Rubino, J.L., 1996. Sequence stratigraphy and palaeogeography of the Middle Jurassic Brent and Vestland deltaic systems, Northern North Sea. Nor. Geol. Tidsskr., 76: 75-106.

Sequence stratigraphy of the Lower Jurassic Dunlin Group, northern North Sea Fleet, A.J., Clayton, C.J., Jenkyns, H.C. and Parkinson, D.N., 1987. Liassic source rock deposition in western Europe. In: J. Brooks and K. Glennie (Editors), Petroleum Geology of North-West Europe. Graham and Trotman, London, pp. 59-70. Galloway, W.E., 1989. Genetic stratigraphic sequences in basin analysis 1: Architecture and genesis of flooding-surface bounded depositional units. Am. Assoc. Pet. Geol., Bull., 73: 125-142. Graue, E., Helland-Hansen, W., Johnsen, J.R., L0mo, L., N0ttvedt, A., Ronning, K., Ryseth, A. and Steel, R., 1987. Advance and retreat of the Brent Delta system, Norwegian North Sea. In: K. Brooks and K. Glennie (Editors), Petroleum Geology of North-West Europe. Graham and Trotman, London, pp. 915-937. Haq, B.U., Hardenbol, J. and Vail, ER., 1987. Chronology of fluctuating sea levels since the Triassic. Science, 235:1156-1166. Hardenbol, J., Thierry, J., Farley, M.B., Jacquin, T., De Graciansky, E-C. and Vail, RR., 1998. Mesozoic and Cenozoic sequence chronostratigraphic framework of European Basins. In: E-C. De Graciansky, J. Hardenbol, T. Jacquin and ER. Vail (Editors), Mesozoic and Cenozoic Sequence Stratigraphy of European Basins. Soc. Econ. Paleontol. Mineral. Spec. Publ., 60: 3-13. Helland-Hansen, W. and Martinsen, O.J., 1996. Shoreline trajectories and sequences: description of variable depositional-dip scenarios. J. Sediment. Res., 66:1670-1688. Helland-Hansen, W., Ashton, M., LOmo, L. and Steel, R., 1992. Advance and retreat of the Brent delta: recent contributions to the depositional model. In: A.C. Morton, R.S. Haszeldine, M.R. Giles and S. Brown (Editors), Geology of the Brent Group. Geol. Soc., London, Spec. Publ., 61: 109-127. Hesselbo, S. and Jenkyns, H.C., 1998. British Lower Jurassic sequence stratigraphy. In: E-C. De Graciansky, J. Hardenbol, T. Jacquin and ER. Vail (Editors), Mesozoic and Cenozoic Sequence Stratigraphy of European Basins. Soc. Econ. Paleontol. Mineral. Spec. Publ., 60:561-581. Jacquin, T. and De Graciansky, E-C., 1998. Major transgressive/regressive cycles: the stratigraphic signature of European basin development. In: E-C. De Graciansky, J. Hardenbol, T. Jacquin and ER. Vail (Editors), Mesozoic and Cenozoic Sequence Stratigraphy of European Basins. Soc. Econ. Paleontol. Mineral. Spec. Publ., 60: 15-29. Jenkyns, H.C., 1988. The early Toarcian (Jurassic) anoxic event: stratigraphic, sedimentary and geochemical evidence. Am. J. Sci., 288: 101-151. Johannessen, E.E, MjOs, R., Renshaw, D., Dalland, A. and Jacobsen, T., 1995. Northern limit of the "Brent delta" at the Tampen Spur a sequence stratigraphic approach for sandstone prediction. In: R.J. Steel, V. Felt, E.E Johannessen and C. Mathieu (Editors), Sequence Stratigraphy of the Northwest European Margin. Norwegian Petroleum Society (NPF), Special Publication 5. Elsevier, Amsterdam, pp. 213-256. Livbjerg, F. and MjOs, R., 1989. The Cook Formation, an offshore sand ridge in the Oseberg area, northern North Sea. In: J.D. Collinson (Editor), Correlation in Hydrocarbon Exploration. Norwegian Petroleum Society. Graham and Trotman, London, pp. 299-312. Lord, A.R., 1978. The Jurassic Part 1 (Hettangian-Toarcian). In: R.H. Bate and J.E. Robinson (Editors), A Stratigraphic Atlas of British Ostracoda. See House Press, Liverpool, pp. 189-212. Lord, A.R., 1982. Metacopine ostracods in the Lower Jurassic. In: F.T. Banner and A.R. Lord (Editors), Aspects of Micropalaeontology. George Allen and Unwin, London, pp. 62-277. Marjanac, T., 1995. Architecture and sequence stratigraphic perspectives of the Dunlin Group formations and proposal for new typeand reference-wells. In: R.J. Steel, V.L. Felt, E.E Johannessen and C. Mathieu (Editors), Sequence Stratigraphy of the Northwest European Margin. Norwegian Petroleum Society (NPF), Special Publication 5. Elsevier, Amsterdam, pp. 143-165. Marjanac, T. and Steel, R.J., 1997. Dunlin Group sequence stratigraphy in the Northern North Sea: a model for Cook sandstone deposition. Am. Assoc. Pet. Geol., 81: 276-292.

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Mitchener, B.C., Lawrence, D.A., Partington, M.A., Bowman, M.B.J. and Gluyas, J., 1992. Brent Group sequence stratigraphy and regional implications. In: A.C. Morton, R.S. Haszeldine, M.R. Giles and S. Brown (Editors), Geology of the Brent Group. Geol. Soc., London, Spec. Publ., 61:45-80. Parkinson, D.N. and Hines, F.M., 1995. The Lower Jurassic of the North Viking Graben in the context of western European Lower Jurassic stratigraphy. In: R.J. Steel, V.L. Felt, E.R Johannessen and C. Mathieu (Editors), Sequence Stratigraphy of the Northwest European Margin. Norwegian Petroleum Society (NPF), Special Publication 5. Elsevier, Amsterdam, pp. 97-107. Partington, M.A., Copestake, P., Mitchener, B.C. and Underhill, J.R., 1993. Biostratigraphic calibration of genetic stratigraphic sequences in the Jurassic-lowermost Cretaceous (HettangianRyazanian) of the North Sea and adjacent areas. In: J.R. Parker (Editor), Petroleum Geology of North-West Europe: Proceedings of the 4th Conference. Geological Society, London, pp. 371-386. Rattey, R.P. and Hayward, A.B., 1993. Sequence stratigraphy of a failed rift system: The Middle Jurassic to Early Cretaceous basin evolution of the central and northern North Sea. In: J.R. Parker (Editor), Petroleum Geology of North-West Europe: Proceedings of the 4th Conference. Geological Society, London, pp. 215-249. Rider, M., 1996. Sequence stratigraphy and stratigraphy. The geological interpretation of well logs. Whittles, Caithness, 280 pp. Riding, J.B., 1987. Dinoflagellate cyst stratigraphy of the Nettleton Bottom Borehole (Jurassic: Hettangian to Kimmeridgian), Lincolnshire, England. Proc. Yorkshire Geol. Soc., 46:231-266. Ryseth, A., 2000. Sedimentology and palaeogeography of the Statfjord Formation (Rhaetian-Sinemurian), North Sea. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway - - Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 67-85 (this volume). Ryseth, A. and Ramm, M., 1996. Alluvial architecture and differential subsidence in the Statfjord Formation, North Sea: prediction of reservoir potential. Pet. Geosci., 2: 271-287. Steel, R.J., 1993. Triassic-Jurassic megasequence stratigraphy in the Northern North Sea: rift to post rift evolution. In: J.R. Parker (Editor), Petroleum Geology of North-West Europe: Proceedings of the 4th Conference. Geological Society, London, pp. 299-315. Steel, R.J. and Ryseth, A., 1990. The Triassic-Early Jurassic succession in the northern North Sea: megasequence stratigraphy and intra-Triassic tectonics. In: R.F.E Hardman and J. Brooks (Editors), Tectonic Events Responsible for Britain's Oil and Gas Reserves. Geol. Soc., London, Spec. Publ., 55: 139-168. Underhill, J.R. and Partington, M.A., 1993. Jurassic thermal doming and deflation in the North Sea: implications of the sequence stratigraphic evidence. In: J.R. Parker (Editor), Petroleum Geology of North-West Europe: Proceedings of the 4th Conference. Geological Society, London, pp. 337-345. Van Buchem, F.S.E andKnox, W.O'B., 1998. Lower and Middle Liassic depositional sequences of Yorkshire (U.K.). In: E-C. De Graciansky, J. Hardenbol, T. Jacquin and P.R. Vail (Editors), Mesozoic and Cenozoic Sequence Stratigraphy of European Basins. Soc. Econ. Paleontol. Mineral. Spec. Publ., 60: 545-559. Van Wagoner, J.C., Posamentier, H.W., Mitchum, R.M., Vail, ER., Sarg, J.F., Loutit, T.S. and Hardenbol, J., 1987. An overview of the fundamentals of sequence stratigraphy and key definitions. In: C.K. Wilgus, B.S. Hastings, C.G.St.C. Kendall, H.W. Posamentier, C.A. Ross and J.C. Van Wagoner (Editors), Sea-Level Changes: An Integrated Approach. Soc. Econ. Paleontol. Mineral. Spec. Publ., 42: 40-45. Vollset, J. and Dor6, A.G., 1984. A revised Triassic and Jurassic lithostratigraphic nomenclature for the Norwegian North Sea. Nor. Pet. Directorate, Bull., 3, 53 pp. Woolam, R. and Riding, J.B., 1983. Dinoflagellate cyst zonation of the English Jurassic. Institute of Geological Sciences Report 83/2, pp. 1-42. Yielding, G., Badley, M.E. and Roberts, A.M., 1992. The structural

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evolution of the Brent Province. In: A.C. Morton, R.S. Haszeldine, M.R. Giles and S. Brown (Editors), Geology of the Brent Group. Geol. Soc., London, Spec. Publ., 61: 27-43.

M.A. CHARNOCK I.L. KRISTIANSEN A. RYSETH LEG. FENTON

Ziegler, EA., 1990. Geological Atlas of Western and Central Europe (2nd edition). Shell International Petroleum Maatschappij, The Hague, 239 pp.

Norsk Hydro Research Centre, P.O. Box 7190, N-5020 Bergen, Norway Norsk Hydro Exploration, N-0246 Oslo, Norway Norsk Hydro Research Centre, P.O. Box 7190, N-5020 Bergen, Norway Present address: Norsk Hydro Harstad, Storakern 11, Kanebogen, N-9401 Harstad, Norway Robertson Research International Ltd., Llandudno, North Wales LL30 1SA, UK

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Divergent development of two neighbouring basins following the Jurassic North Sea doming event: the Danish Central Graben and the Norwegian-Danish Basin Jan Andsbjerg, Lars Henrik Nielsen, Peter N. Johannessen and Karen Dybkja~r

The two neighbouring basins, the Danish Central Graben and the Norwegian-Danish Basin were both affected by the regional uplift of the North Sea and adjacent areas in the early Middle Jurassic that caused the formation of a regionally extensive unconformity. The uplifted area was not a simple dome structure but of a more irregular shape with an east-west oriented branch that included the Ringk~bing-Fyn High and much of the Norwegian-Danish Basin and the Fennoscandian Border Zone. Late Aalenian-Bajocian deposition was confined to fault-controlled depocentres in both the Danish Central Graben and the Norwegian-Danish Basin. An initial southward slope in the Danish Central Graben changed to a north- to eastward slope before the end of the Middle Jurassic, and the change possibly coincides with the formation of a conspicuous sequence boundary in the Bathonian. The depositional area began to expand in the late Middle Jurassic as a result of a regional sea-level rise. In the Danish Central Graben, accelerating half-graben subsidence during the Callovian-Early Kimmeridgian enhanced the sea-level rise. Several periods of rapid subsidence during the Callovian-Volgian (mainly in the Oxfordian-Early Kimmeridgian and latest Kimmeridgian-Middle Volgian) gave accommodation space to more than four kilometres of marine mud. A break in subsidence in the late Kimmeridgian, probably related to a change of fault directions, resulted in deposition of shallow marine sandstones on platforms and hanging-wall slopes. The Norwegian-Danish Basin was characterised by a small rate of subsidence and continuous expansion of the depositional area throughout the Late Jurassic. The slow subsidence and a large supply of sediment from the Fennoscandian Border Zone caused repeated progradational events from the northeast. Hydrocarbon discoveries are known only from the Danish Central Graben where Middle Jurassic and Upper Jurassic reservoirs have been charged from Upper and to a smaller degree Middle Jurassic source rocks. Within the Norwegian-Danish Basin, reservoir rocks are abundant in the Upper Triassic-lowermost Jurassic, the Middle Jurassic and Upper Jurassic successions. The presence of mature source rocks, however, is the main risk factor as they most likely only occur within Lower Jurassic mudstones deeply buried in rim-synclines and in local grabens.

Introduction

The Danish Central Graben (Fig. l a,b) is a mature hydrocarbon province, and though the principal production comes from Upper Cretaceous-Danian chalk, exploration of the Jurassic rift succession has shown several encouraging discoveries both in the Middle and Upper Jurassic. The Jurassic of the Danish part of the Norwegian-Danish Basin (Fig. l c) likewise contains both reservoirs and potential source rocks, but commercial hydrocarbon accumulation has not been found yet. The two neighbouring basins, the Danish Central Graben and the Norwegian-Danish Basin are located in an area that was affected by extensional movements during the Jurassic. Both the Danish Central Graben and the eastern part of the NorwegianDanish Basin show a relatively complete MiddleUpper Jurassic succession. However, whereas the Middle Jurassic succession in the two basins shows a

remarkable similarity, the Upper Jurassic succession differs strongly in thickness, distribution and type of sedimentary facies, and distribution and quality of reservoir and source rocks (Figs. 2 and 3). The variables that affected these patterns may include tectonic events, climate changes, eustasy and sediment supply. Among the tectonic events that had a pronounced influence on both basins was the formation of the "mid-Cimmerian" unconformity at the base of the Middle Jurassic. This regionally extensive unconformity evolved as a response to regional uplift that has been described by various authors including Eynon (1981) and Ziegler (1982, 1990). Recently, Underhill and Partington (1993) have presented a model for the early Middle Jurassic uplift in the North Sea that has been frequently cited. This model depicts the results of the uplift as a broad domal structure, that within its concentric periphery includes the Danish Central Graben, the Ringk~bing-Fyn High and much of the Norwegian-Danish Basin (Fig. 4). However,

Sedimentary Environments Offshore Norway - Palaeozoic to Recent edited by O.J. Martinsen and T. Dreyer. NPF Special Publication 10, pp. 175-197, Published by Elsevier Science B.V., Amsterdam. 9 Norwegian Petroleum Society (NPF), 2001.

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Fig. 1. (a) The North Sea area. The study area including the Danish Central Graben and the Norwegian-Danish Basin is outlined by a dashed frame. (b) Danish Central Graben with well locations; structures active in the Late Jurassic are shown. (c) Eastern part of the Norwegian-Danish Basin (Danish Basin) with well locations; structural elements in the Jurassic are shown.

both pre-Middle Jurassic subcrop and Middle-Upper Jurassic onlap data presented by Underhill and Partington (1993) differ significantly from the results of the present study (Fig. 5). This paper attempts to compare the two neighbouring basins in order to investigate the relative influence of the variables that determined the depositional pattern and distribution of reservoir and source rocks in the two basins. The comparison demonstrates that the uplift of the North Sea Basin included the large west-northwest-east-southeast trending RingkObingFyn High and the uplifted area formed a large arch rather than a concentric dome as postulated by Underhill and Partington (1993, 1994). Recently, the Danish Central Graben and the eastern part of the Norwegian-Danish Basin have been studied extensively by the Geological Survey of Denmark and Greenland (GEUS) and this work forms the basis for this paper. Based on well logs and palynological data from more than 50 wells, a sequence stratigraphic framework was established for the Jurassic in the Danish Central Graben by Andsbjerg and Dybkja~r (2001). This work was further supported by regional seismic lines and by sedimentological studies of all available Middle Jurassic and Upper Jurassic cores from the area. The core studies are presented in detail by Johannessen (2001), who mainly undertook studies of the Upper Jurassic and by Andsbjerg (1997,

2001) for the Middle Jurassic. For the eastern part of the Norwegian-Danish Basin, Nielsen (1995, 2001) has presented sedimentological and sequence stratigraphic studies based on well logs, palynological data and detailed core studies of more than 40 wells. Other papers relevant to this work include Johannessen and Andsbjerg (1993), Johannessen et al. (1996), Johannessen (1997), Ineson et al. (2001), Michelsen et al. (2001) and MNler and Rasmussen (2001) on the Danish Central Graben, and Michelsen (1989a,b) and Poulsen (1996) on the Norwegian-Danish Basin. The lithostratigraphic subdivision of the two basins is summarised in Fig. 6.

Geological background The Danish Central Graben and the NorwegianDanish Basin were formed as a result of plate reorganisations in Late Carboniferous-Early Permian time, and have both undergone a long complex history of differential subsidence (Ziegler, 1982; Vejb~ek, 1989). In Early Jurassic time, tectonic quiescence prevailed. A marine shelf covered both basins, and thick laterally consistent sequences of homogeneous mudstones were deposited under strong influence of eustatic changes (Michelsen, 1978; Pedersen, 1985; Nielsen, 1995, 2001). These conditions lasted until regional uplift occurred in earliest Middle Jurassic due to crustal pro-

Divergent development of two neighbouring basins following the Jurassic North Sea doming event

Fig. 1 (continued).

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Fig. 2. Chronostratigraphic summary diagram of the Danish Central Graben and the eastern part of the Norwegian-Danish Basin showing distribution of lithologies and interpreted depositional environments (modified from Nielsen, 1995, 2000; Andsbjerg and Dybkjam 2000; Johannessen, 2000). The chronostratigraphic scale is from Gradstein et al. (1995).

cesses related to the updoming of the central North Sea as described by Hallam and Sellwood (1976), Eynon (1981), Ziegler (1982, 1990) and Underhill and Partington (1993, 1994). The regional uplift and subsequent rifting caused development of grabens, which determined the depositional style, and thus the distribution of the Middle and Upper Jurassic sediments. In the Danish Central Graben, the thickest succession of Middle Jurassic deposits occurs in the SCgne

Basin and the Tail End Graben, along the main boundary fault at the western margin of the RingkCbing-Fyn High (Figs. lb, 2 and 3). A thin Middle Jurassic succession occurs across a wider area in the southern and southeastern part of the Danish Central Graben. Progressively younger Upper Jurassic deposits are found updip on the western hanging-wall slope of the Danish Central Graben. Middle Jurassic deposits extend across a large part

179

Divergent development of two neighbouring basins following the Jurassic North Sea doming event

ENE

WSW

1

450 km

45 km

T Mid North Sea

H~gh

Norwegian - Danish Basin

Danish Central Graben

Ringk~bing- Fyn

H~gh

Onlapping Upper Jurassic

I

FennoscandianBorder Zone

/lll~ i ~1~ ,~

I Baltic Shield

~ Middle Jurassic

sediments

sediments Sorgenfrei - Tornquist Zone

Onlapping Upper Jurassic

sediments

Middle Jurassic

1000 m

sediments Fig. 3. Schematic diagram of the two basins displaying the situation after the completion of Jurassic deposition. Whereas the Middle Jurassic succession differs very little between the two basins, the Upper Jurassic exhibits a significant difference in thickness.

of the Norwegian-Danish Basin and the Fennoscandian Border Zone. However, a thick Middle Jurassic succession is found only in the Sorgenfrei-Tornquist Zone, a narrow northwest-southeast faulted part of the Fennoscandian Border Zone that formed the transition from the Baltic Shield to the NorwegianDanish Basin (Figs. l c, 2 and 3). Progressively younger Upper Jurassic sediments occur updip along the southern margin of the basin.

Early-Middle Jurassic uplift The depositional conditions that prevailed during the Early Jurassic changed abruptly in the earliestMiddle Jurassic due to regional uplift. The restricted and oxygen-poor shelf environment established in the Norwegian-Danish Basin during the Early Toarcian continued throughout the Late Toarcian-Early Aalenian as a response to the early phase of the uplift (Michelsen, 1978; Michelsen and Nielsen, 1991; Nielsen, 1995,2001). Due to extensive and continued uplift, a highly erosive regional unconformity formed in the North Sea, in large parts of the NorwegianDanish Basin, on the RingkObing-Fyn High and in the Fennoscandian Border Zone (Fig. 2). In the southern part of the Danish Central Graben the resulting hiatus comprises the upper Pliensbachian, Toarcian

and parts of the Aalenian. The time span of the hiatus increases to the north, where Middle-Upper Jurassic strata onlap Triassic, Permian and Carboniferous rocks (Fig. 5; Johannessen et al., 1996). A few tens of kilometres southwest of the Danish Central Graben, in the Dutch Central Graben, continuous deposition of marine mud took place across the LowerMiddle Jurassic boundary. Deposition in that area was terminated by a later uplift that resulted in the development of a late Bathonian-Callovian unconformity (Van Adrichem Boogaert and Kouwe, 1993). The lowermost part of the Middle Jurassic succession that overlies the unconformity in the Danish Central Graben has been assigned an Aalenian?-Bajocian age based on palynological evidence (Figs. 6 and 7; Andsbjerg, 1997; Andsbjerg and Dybkja~r, 2001). In the fault-bounded Sorgenfrei-Tornquist Zone, slow subsidence occurred while the rest of the area was uplifted (Nielsen, 1995, 2001). In this area the uplift-related unconformity is replaced by a basinward shift in facies from offshore marine mudstones (Fjerritslev Formation) to shoreface sandstones (Haldager Sand Formation) overlying a regressive surface of marine erosion (Fig. 8, Haldager-1 and Vedsted-1 wells). The surface is dated to the Aalenian (top Opalinum Zone; Poulsen, 1996) similar to the age of the facies change that occurs at the marginal

1 80

J. Andsbjerg et al.

Fig. 4. "Mid-Cimmerian" unconformity subcrop patterns suggested by Underhill and Partington. Modified from Underhill and Partington (1993, 1994).

zone of the uplifted area in the northern North Sea (Underhill and Partington, 1993). On ramps dipping towards the Sorgenfrei-Tornquist Zone, the surface is developed as a marked erosional unconformity, which shows truncation of progressively older strata toward the RingkCbing-Fyn High (Fig. 5a and Fig. 8). Onlap of the overlying strata shows younging in the same direction (Fig. 5b and Fig. 8). On the shallowest parts of the Ringk0bing-Fyn High, the Lower JurassicTriassic successions were eroded. In southern Sweden, the early-Middle Jurassic uplift was accompanied by faulting, erosion and

volcanism with basalts intruding along northwestsoutheast trending faults and fracture zones (Norling and Bergstr6m, 1987; Erlstr6m et al., 1997). The oldest basalt is palaeomagnetically dated to a Toarcian-Aalenian age, radiometrically dated to the Bajocian (167 Ma), while related tuffites are dated to the Aalenian by palynology (Printzlau and Larsen, 1972; Tralau, 1973; Klingspor, 1976; Norling and Bergstr6m, 1987; Bylund and Halvorsen, 1993). Sedimentation changed from marine muds in the Early Jurassic to continental-paralic deposits in the Middle Jurassic which became confined to fault-bounded

Divergent development of two neighbouring basins following the Jurassic North Sea doming event

181

Fig. 5. (a) Map showing subcrop to the "base Middle Jurassic unconformity" in the Norwegian-Danish Basin and the Danish Central Graben. (b) Map showing the Upper Jurassic marine onlap to the "base Middle Jurassic unconformity". Note the deepening of the truncation and the younging of the onlap toward the Ringkobing-Fyn High.

areas (Norling and Bergstr6m, 1987; Norling et al., 1993).

Fault-controlled deposition: Aalenian-Bajocian-early Bathonian The Aalenian-Bajocian deposits that overlie the intra-Aalenian unconformity are confined mainly to

narrow fault-controlled depocentres, where they attain a thickness of 150 to 250 m (Figs. 2, 7 and 8). The major depocentres were the SOgne Basin and the Tail End Graben, including its southern extension into the Salt Dome Province in the Danish Central Graben (Fig. l b) and the deep part of the Sorgenfrei-Tornquist Zone in the Norwegian-Danish Basin (Fig. l c). These grabens subsided slowly; sed-

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Fig. 6. Jurassic lithostratigraphy of the Danish Central Graben (left) and Norwegian-Danish Basin (right).

Fig. 7. Log correlation panel of the Middle Jurassic in the Danish Central Graben. The lower boundary of the Bryne Formation corresponds to the base Middle Jurassic unconformity that truncates Triassic rocks in the West Lulu-3 well and Lower Jurassic Fjerritslev Formation mudstones in the two southernmost wells. The base of the Middle Jurassic is not penetrated in the Amalie-1 well.

iment thickness combined with palynological data indicates a subsidence rate of about 10-20 m/m.y. The continuous existence of paralic environments in the grabens suggests that subsidence was balanced

by sediment input. The sediment package thickens towards the main boundary fault of the Danish Central Graben, indicating syn-sedimentary fault activity (Damtoft et al., 1992; Korstg5rd et al., 1993).

Divergent development of two neighbouring basins following the Jurassic North Sea doming event

183

Fig. 8. Log correlation panel of the Middle Jurassic and lower Upper Jurassic in the Norwegian-Danish Basin. In the Sorgenfrei-Tornquist Zone, a relatively continuous Lower-Middle Jurassic succession occurs, and the base Middle Jurassic unconformity corresponds here to an intra-Aalenian marine regressive surface of erosion at the base of the first shoreface sandstones (Vedsted-1 and Haldager-1 wells). On the ramps in Fars~-1 and Frederikshavn-1 wells, a major hiatus separates Toarcian mudstones of the Fjerritslev Formation from Bathonian? fluvial sandstones of the Haldager Sand Formation. Further to the southwest in the Vemb-1 well, Lower Pliensbachian mudstones are overlain by fluvial sandstones presumably of Callovian age.

In the Danish Central Graben, deposition of fluvial channel sands, overbank deposits and extensive lacustrine muds of the Bryne Formation dominated during the late Aalenian?-early Bathonian (Fig. 7). The presence of thin marine mudstones in the lower part of the succession in the southern part of the Danish Central Graben and contemporaneous marine mudstones in the Dutch Central Graben suggests a regional slope towards the south. In the northern part of the Danish Central Graben, the Bryne Formation is dominated by up to 10 m thick packages of channel sandstones that are separated by up to 50 m thick suites of floodplain deposits (Fig. 7; Andsbjerg, 2001). Individual chan-

nel units can be traced between most wells in the area suggesting that they represent laterally migrating channels. In the wells closest to the basin axis, lacustrine mudstones frequently dominate the floodplain deposits in the upper part of the succession. In the southern and southwestern part of the Danish Central Graben, floodplain and lacustrine deposits are more prominent, and channel sandstones are generally thinner and less correlatable. The presence of thin interbedded marine mudstones in this area, and more extensive marine mudstones in the Dutch sector a few tens of kilometres to the south, may suggest that the channel sandstones were deposited in minor distributary channels on a coastal plain.

184 In the eastern part of the Norwegian-Danish Basin, the deposition was confined to the deepest part of the narrow Sorgenfrei-Tornquist Zone bounded by the Fjerritslev and BCrglum Faults and their southeastward continuation in Kattegat, Oresund and SkSne (Figs. 2 and 8; Nielsen, 1995, 2001). The initial deposition of shoreface sands, 18-20 m thick, was succeeded by the deposition of transgressive muds and thin sands during the Aalenian. The marine mudstones were later incised and overlain by the fill of an estuary that was confined to the deepest part of the Sorgenfrei-Tornquist Zone. The estuarine fill is about 45 m thick and comprises fluvial sandstones that are overlain by lagoonal sandstones and mudstones with thin coaly seams and topped by barrier sandstones (Fig. 8). Transgressive marine mudstones and sandstones overlie the valley-fill. Based on a top occurrence of cysts of the dinoflagellate species Nannoceratopsis gracilis, these deposits are not younger than the lower Bajocian (Poulsen and Riding, 2001). After a sea-level fall accompanied by deep incision, rising sea level resumed the generation of accommodation space. Estuarine valley-fills, up to 25 m thick, were deposited consisting of fine-grained, muddy sandstones, thin mudstones and possibly coal seams (Fig. 8). The fills contain a mixed assemblage of marine and freshwater palynomorphs indicating an early Bathonian age. Deposition overstepped the Sorgenfrei-Tornquist Zone as accommodation space was created on the lower part of the basin-ward dipping ramps and braided fluvial channel sandstones were deposited (Fig. 9). The sandstones are sharply topped by lacustrine and lagoonal mudstones reflecting generation of further accommodation space. In general, however, accommodation space was limited during the Aalenian-Callovian as indicated by the absence of highstand and uppermost transgressive systems tracts, which presumably were cannibalised due to erosion during the subsequent sea-level falls (Nielsen, 1995, 2001).

Ramp to basin deposition: Bathonian-Callovian A pronounced and extensive intra-Bathonian sequence boundary occurs in both the Central Graben and the Norwegian-Danish Basin (Figs. 2, 7 and 8). Its formation was associated by changes in depositional patterns and sedimentary environments (Nielsen, 1995, 2001; Andsbjerg, 1997, 2001). Extensive channel deposits initially dominated sedimentation above the sequence boundary. On the ramps dipping towards the deep basins, braided-river deposits have been encountered above the sequence boundary in both the Danish Central Graben and in the Norwegian-Danish Basin (Figs. 9 and 10). In

J. Andsbjerg et al. more basin-ward settings, interbedded estuarine and fluvial deposits in the Danish Central Graben and interbedded estuarine and shallow marine deposits in the Norwegian-Danish Basin were formed. Braided-river gravel beds referred to the Bryne Formation are found above the intra-Bathonian unconformity in the Elly-3 well in the southwestern part of the Danish Central Graben (Figs. l b and 10). They occur as a 7.5 m thick unit of cross-bedded, mainly clast-supported, pebble conglomerate. On the gamma-ray log, this unit exhibits a pronounced blocky pattern, which can also be seen in the U-1 well, 35 km to the southeast (Fig. lb; Andsbjerg, 1997, 2001). In the northern part of the Danish Central Graben, an up to 40 m thick succession of incised valley-fill deposits belonging to the Bryne Formation occurs in several wells (Fig. 7; Andsbjerg, 1997, 2001). The incised valley is bounded at the base and laterally by the intra-Bathonian unconformity. The valley fill consists of cross-bedded sandstones with abundant double mud-drapes, mud-flasers and mudstone- and coal clasts (Fig. 10). Pebble layers, interpreted as basal channel lags, occur at the base of amalgamated sandstone units, and thinner mudstone units separate the sandstone units. Most of the succession was deposited within an estuary channel environment (Andsbjerg, 1997, 2001). In the Danish part of the Norwegian-Danish Basin, biostratigraphic evidence of the Bathonian-Callovian is poor (Michelsen, 1978; Poulsen, 1996). Fluvial erosion may have prevailed for a long period in large parts of the basin, and deep erosion into the Lower Jurassic succession occurs locally southwest of the Sorgenfrei-Tornquist Zone (Fig. 2; Nielsen, 1995, 2001). Also in the deep part of the SorgenfreiTornquist Zone, erosion appears to have been significant as upper transgressive to highstand systems tracts deposits of the underlying Bajocian-Bathonian sequence seem to be absent. On the ramps dipping toward the Sorgenfrei-Tornquist Zone, the intra-Bathonian sequence boundary is overlain by 520 m of mainly cross-bedded, medium-grained sandstones (Fig. 9). The sandstones were deposited by braided rivers as an initial response to the formation of accommodation space presumably in the late Bathonian-Callovian. In the more basin-ward parts of the Sorgenfrei-Tornquist Zone, the initial deposition on the sequence boundary was 5-10 m of transgressive shoreface sands (Figs. 2 and 8; Nielsen, 1995, 2001).

Regional transgression: Callovian-Oxfordian A regional Callovian-Oxfordian transgression caused the formation of interfingering paralic and

Divergent development of two neighbouring basins following the Jurassic North Sea doming event

~85

Fig. 9. Cores from the FarsO-1 and Ars-1 wells in the Norwegian-Danish Basin; both wells are located in a ramp setting. The lower cores in both wells show two units of braided-river sandstones separated by lacustrine mudstones. The lower sandstone unit, presumably of Bathonian age is overlying Toarcian offshore mudstones of the Fjerritslev Formation; the boundary corresponds to the "base Middle Jurassic unconformity". The upper sandstone unit is sitting on an intra-Bathonian unconformity. The fluvial sandstones are capped by transgressive lagoonal beach deposits of the Flyvbjerg Formation. The late Callovian-earliest Oxfordian lagoonal transgressive surface is erosive and overlain by claystone clasts (Ars-1). The upper core in Ars-1 represents the extensive lagoonal deposits that developed in the Early Oxfordian as a response to the regional transgression. The lagoonal deposits are overlain by marine mudstones and thin sandstones that constitute the upper part of the Flyvbjerg Formation. For location see Fig. lb. For legend on sedimentary structures see Fig. 14.

shoreface deposits in both the SOgne Basin and the northern part of the Tail End Graben. A series of east- and northward prograding shoreface units of the Callovian Lulu Formation stepped back towards the west and southwest (Michelsen et al., 2001). This took place on the lower part of the hanging-wall ramp of the evolving half-graben (Figs. 2, 7 and 11). The final

transgression of this area was achieved at the end of the Callovian. Further south in the Danish Central Graben an extensive, low-energy coastal plain dominated by lagoons and coastal swamps was transgressed in the earliest Oxfordian (Fig. 7; Andsbjerg, 1997, 2001). In the northern part of the Danish Central Graben, the transgressive succession consists of an up to 4 m

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Divergent development of two neighbouring basins following the Jurassic North Sea doming event

thick, regionally extensive coal bed overlain by two, approximately 10 m thick, sandstone units. Each unit consists of thin transgressive shelf sandstones overlain by thin offshore mudstones and by up to 8 m thick upward-coarsening shoreface to beach and back-barrier sandstones (Fig. 11). Less than 5 km updip on the hanging-wall slope the sandstone units wedge out to 2 m thick, upward-coarsening sandstones interbedded with the deposits of estuary channels, bay-head and tidal deltas, tidal flats and lagoons (Andsbjerg, 1997,2001). At the top of the succession, a transgressive conglomerate occurs, which is overlain by thoroughly bioturbated shallow marine siltand sandstones that are succeeded by fully marine mudstones of the Oxfordian to early Kimmeridgian Lola Formation. In the southern part of the Danish Central Graben, the transgressive succession is dominated by lagoonal and lacustrine mudstones of the Middle Graben Shale Formation. In this area, abundant coal beds at the base of the succession may locally attain a thickness of 7 m (Andsbjerg, 1997, 2001). The initial sea-level rise that created accommodation space over the intra-Bathonian sequence boundary in the Sorgenfrei-Tornquist Zone and on the lower parts of the flanking ramps continued in the Oxfordian (Fig. 2; Nielsen, 1995, 2001). In the Sorgenfrei-Tornquist Zone, up to 40 m of transgressive marine muds and thin sands belonging to the Flyvbjerg Formation were deposited on the shoreface sandstones that constitute the upper part of the Haldager Sand Formation (Fig. 8). Palynological data suggest that the change from deposition of shoreface sands to offshore muds occurred in the late Callovian (Poulsen, 1992, 1996). On the ramps on both sides of the Sorgenfrei-Tornquist Zone, deposition of braided fluvial sands was succeeded by deposition of lagoonal muds and thin sands occasionally with abundant rootlets (Figs. 2, 8 and 9). The deepening continued and the lagoonal deposits were overlain by transgressive shoreface sands and offshore muds, which show younging toward the northeast (Frederikshavn-2, Fig. 8). To the southwest the deepening is recorded by a few metres of lagoonal or marine mudstones (Vemb-1, Fig. 8). The lower part of the Flyvbjerg Formation thus shows backstepping up the ramps toward the Baltic Shield and toward the southwestern part of the basin testifying that these areas again became part of the depositional basin. Dur-

187

ing this time the Ringkc~bing-Fyn High functioned as a low-relief hinterland, that only supplied minor amounts of sediments to the basin, as indicated by the absence of significant sand at this stratigraphic level (Nielsen, 1995,2001).

Differential subsidence and transgression: Oxfordian-Kimmeridgian A significant difference in the rate of subsidence between the Danish Central Graben and the Norwegian-Danish Basin began in the Oxfordian. During the Oxfordian, a high rate of subsidence along the main boundary fault of the Danish Central Graben initiated the deposition of a 900 m thick succession of marine mudstones of the Lola Formation (Fig. 12). The rate of subsidence decreased and came to a temporary halt before the end of the early Kimmeridgian (Andsbjerg and Dybkja~r, 2001; Mc~ller and Rasmussen, 2001). The marine mudstones onlap the pre-Jurassic and Middle Jurassic on the hanging-wall slopes (Fig. 12; Mc~ller, 1986). Turbidite sands forming only a few metres thick unit, were deposited in the axial parts of the basin, and locally marginal marine sands accumulated on the hanging-wall slope (Andsbjerg, 1997; Andsbjerg and Dybkja~r, 2001). A low rate of subsidence still characterised the Norwegian-Danish Basin. The southwestern part of that basin was finally transgressed in the middle-Late Oxfordian and a thin, up to 15 m thick succession of marine offshore mudstones, that shows thinning toward the Ringk~bing-Fyn High, was deposited. In the latest Oxfordian, a short-lived regressive event of coastal progradation from the Baltic Shield deposited 5-15 m of fluvial sands that continued seaward into to 2-10 m of shell-bearing shoreface sandstones and siltstones constituting the upper part of the Flyvbjerg Formation (Figs. 2 and 8). The regression was followed by a renewed transgression that caused deposition of backstepping thin coastal sands that was succeeded by deposition of more widespread marine mud of the Bc~rglum Formation in the early Kimmeridgian (Figs. 2 and 8; Michelsen, 1978; Nielsen, 1995, 2001). Fully marine conditions with deposition of offshore mud probably dit not reach the Fennoscandian Border Zone before the late Kimmeridgian (Poulsen, 1992, 1996).

Fig. 10. Core and gamma-ray logs from the Middle Jurassic of the Danish Central Graben. In the Elly-3 well, the intra-Bathonian unconformity separates braided-river conglomerates above from a floodplain dominated succession below the unconformity. A thin marine mudstone is present near the base of the succession. In the West Lulu-3 and Amalie-1 wells the unconformity bounds an incised valley with a valley fill dominated by estuarine channel sandstones. For legend on sedimentary structures see Fig. 14.

188

J. Andsbjerg et al.

Divergent development of two neighbouring basins following the Jurassic North Sea doming event

189

Fig. 12. Log correlation panel from the Oxfordian-Kimmeridgian in the Danish Central Graben. The asymmetry of the subsiding half-graben is evident from the dip line Elly-3-Nora-1. Onlap of the Heno Plateau by marine mudstones is seen in the north in the Feda Graben (Gert-1, Jeppe-1, Gwen-2 wells) and in the south (Ravn-1, Falk-1, Elly-3).

Pause in subsidence: late Kimmeridgian In the Danish Central Graben, rift-related subsidence ceased or slowed down significantly for a period in late Kimmeridgian time, possibly in relation to a shift in activity from north-south trending to northwest-southeast trending faults (Johannessen et al., 1996; M011er and Rasmussen, 2001). The cessation of rift-related subsidence and the associated decrease in accommodation-space generation, and possibly an increase in sediment supply, caused the progradation of shallow marine sands of the late Kimmeridgian Heno Formation in the northwestern parts of the Danish Central Graben (Fig. 12). Sands sourced from the Mid North Sea High prograded towards the east on the Heno Plateau (Fig. 13), while sands sourced from the Mandal High prograded towards the west on the Gertrud Plateau and in the Feda Graben (Fig. 12; Johannessen and Andsbjerg, 1993; Johannessen et al., 1996; Johannessen, 2001). In the southern part of the Feda Graben, syn-depos-

itional subsidence balanced by a large sediment supply caused the development of an up to 90 m thick succession of aggradational back-barrier sandstones (Johannessen and Andsbjerg, 1993; Johannessen et al., 1996; Johannessen, 2001; Andsbjerg and Dybkjam 2001). The back-barrier sandstones are interbedded with mudstones; strong bioturbation and abundant water-escape structures have destroyed primary sedimentary structures. Thin coals and abundant rootlets also occur. The sediments are typically arranged in 3-8 m thick upward-coarsening to upward-fining units (Fig. 14A). The back-barrier sandstones are separated from an overlying shoreface succession by a ravinement surface. On the Heno Plateau, which formed a part of the hanging-wall slope of the Danish Central Graben, a coarsening- to fining-upward succession of mainly very fine- to fine-grained sandstones was deposited (Fig. 2). The succession is in places more than 100 m thick and contains one or two distinct pebble conglomerate beds. The sandstones, which normally are

Fig. 11. Cores from the middle Callovian-earliest Oxfordian in the SCgne Basin. Lulita-1 well wedge out in the wells further updip, where estuarine deposits are overlain by Oxfordian offshore mudstones of the Lola Formation. The complete transgression of the Danish Central Graben. For legend on sedimentary structures

The upward-coarsening shoreface-dominated succession in the dominant. The Callovian shallow marine-paralic succession is succession represents a stepwise Callovian-earliest Oxfordian see Fig. 14.

190

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Divergent development of two neighbouring basins following the Jurassic North Sea doming event

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Fig. 14. (A) Core and gamma-ray log from the Heno Formation in the Gert-1 well. The succession consists of stacked back-barrier sandstones with thin coal beds and lagoonal mudstones. Above the cored interval is a smaller succession of shoreface sandstones 9 (B) Core and gamma-ray log from the Heno Formation in the Ravn-1 well. Mainly shallow marine sandstones and conglomerates occur 9 The conglomerates represent amalgamated sequence boundaries and ravinement surfaces 9

Fig. 13. Dip section 9 Log correlation panel from the Upper Jurassic of the Danish Central Graben. Half-graben-induced asymmetry is very well developed in both the Oxfordian-early Kimmeridgian Lola Formation and in the latest Kimmeridgian-early Ryazanian Farsund Formation. The Middle Jurassic Bryne and Lulu Formations represent deposition during the early half-graben subsidence 9 The Lola and Farsund Formations represent deposition during two main subsidence phases 9

192

completely bioturbated, occur as 2-8 m thick upward-coarsening units that were deposited on the lower to middle shoreface (Fig. 14B). The conglomerate beds represent fluvial gravels that were formed during a fall in sea level and later reworked during a subsequent sea-level rise. They thus represent amalgamated sequence boundaries/ravinement surfaces (Johannessen, 2001). A regressive event similar to that of the Heno Plateau in Kimmeridgian time seems not to have affected the Norwegian-Danish Basin (Fig. 2). Here the deposition of relatively uniform offshore marine mud of the BOrglum Formation continued throughout the Kimmeridgian and into the middle Volgian, amounting up to ca. 125 m centrally in the basin (Poulsen, 1992, 1996; Nielsen, 2001). Closer to the basin margins the thickness of the marine claystones decreases to a few metres as seen on the SkagerrakKattegat Platform, where nearshore conditions lasted until late Kimmeridgian time (Fig. 2). A regressive event with coastal progradation began in the early Volgian with deposition of shallow marine sands of the Frederikshavn Formation (Fig. 15; Poulsen, 1996; Nielsen, 2001; Michelsen et al., 2001). The thickness

J. Andsbjerg et al.

of the marine claystones likewise decreases toward the southwestern part of the basin, possibly due to a limited sediment input from the Ringk0bing-Fyn High, that probably was submerged at this time owing to the general sea-level rise. At the basin margin further to the southeast in southern Sweden lagoonal and lacustrine variegated claystones and siltstones were deposited (Norling et al., 1993).

Contrasting development in the late Kimmeridgian-Ryazanian The structural evolution of the Danish Central Graben and the Norwegian-Danish Basin differed strongly during the late Kimmeridgian-Ryazanian. In the former basin rates of subsidence increased dramatically to a maximum in latest Kimmeridgianmiddle Volgian times while the subsidence in the Norwegian-Danish Basin showed a slow increase (MNler, 1986; Nielsen, 1995, 2001; Andsbjerg, 1997; Andsbjerg and Dybkja~r, 2001). In the Danish Central Graben transgression to the west and southwest of the hanging-wall slope continued (Fig. 2). A more than 3000 m thick, strongly asym-

Fig. 15. Log correlation panel from the Kimmeridgian-Ryazanian of the Norwegian-Danish Basin. The marine mudstones are interrupted by two major prograding wedges of shallow marine and fluvial-estuarine sandstones in the northeastern part of the basin. Notice the scale difference from Fig. 11.

Divergent development of two neighbouring basins following the Jurassic North Sea doming event

metric wedge of marine mudstones of the Farsund Formation was deposited reflecting an average rate of subsidence of approximately 300 m/m.y, centrally in the basin (Fig. 13; MNler, 1986; Britze et al., 1995). Organic-rich shales are associated with the latest Kimmeridgian-Early Volgian maximum flooding surfaces on the western ramp of the Danish Central Graben. These shales probably reflect an updip trapping of clastic sediments and a high rate of organic production in the relatively shallow, nutrition-rich water on the ramp following transgression of the late Kimmeridgian coastal plain. A wedge of shallow marine sands further updip on the hanging-wall slope in the Outer Rough Basin in the westernmost part of the Danish Central Graben, prograded during the early to middle Volgian (Damtoft et al., 1992; Mackertich, 1996). In Late Volgian-Ryazanian time, the Danish Central Graben broke up into smaller sub-basins. Cessation of movements along segments of the eastern boundary fault allowed drainage of new sediment source areas. Turbidite sands and other gravity flow deposits occur both in fans along the eastern boundary fault of the basin and centrally on the basin floor (Figs. 2 and 16; Damtoft et al., 1992). Intermittent transgressive phases during the otherwise regressive trend caused the deposition of another organic-rich shale, the widespread Bo Member (Fig. 2; the former "hot unit" equivalent to the most organic-rich part of the Mandal Formation; Dybkj~er, 1998; Michelsen et al., 2001) which is the most important source rock in the Danish Central Graben (Ineson et al., 2001). During this period subsidence in the NorwegianDanish Basin did not exceed ca. 15 m/m.y. The depositional environment was dominantly a shallow shelf with recurrent episodes of delta and coastal plain progradation from the Baltic Shield depositing marine silts and minor sands, interbedded with thin fluvial sands and a few coaly beds to the northeast (Fig. 2; Michelsen, 1978; Nielsen, 2001). In the middle of the basin and toward the west deposition of marine mud dominated (Michelsen, 1989a). However, in the southwestern part of the basin, the occurrence of fine-grained sand in the marine mudstones suggests that the Ringk0bing-Fyn High was emerged at least in part at this time. The resulting deposits form a wedge of mudstones and sandstones, up to 150 m thick that display a general thickening toward the northeast (Fig. 15). Discussion

Early-Middle Jurassic upfift The regional "mid-Cimmerian" unconformity in the North Sea area has been interpreted to reflect

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]Slump structure

Fig. 16. (A) Core and gamma-ray log from the Jeppe-1 well in the northern part of the Danish Central Graben. The core is dominated by debris flow deposits and thin-bedded turbidites. (B) Core and gamma-ray log from the Iris-1 well in the northeastern part of the Danish Central Graben. The core consists of turbidite deposits interbedded with hemipelagic mudstones. Both cores are of latest Volgian-Early Kimmeridgian age. For legend on sedimentary structures see Fig. 14.

uplift and erosion related to emplacement of a mantle plume (Hallam and Sellwood, 1976; Eynon, 1981; Ziegler, 1982, 1990; Leeder, 1983; Underhill and Partington, 1993, 1994), a major eustatic sea-level fall (Haq et al., 1988) or local response to footwall uplift (Barr, 1987; Badley et al., 1988). Underhill and Partington (1993, 1994) portrayed a concentric pattern, with a diameter of more than 800 km centred on the North Sea triple junction. Based on this pattern they interpreted domal uplift and erosion followed

194

by deflation and gradual onlap of the uplifted area. However, Ziegler (1982, 1990) described the uplifted structure as a broad arch that extended from the southern Viking Graben into the southern North Sea, where it was transected by the Central Graben, and eastwards to Denmark. Uplift of the east-west trending Ringkobing-Fyn High across the Danish area in the latest-Early Jurassic-early-Middle Jurassic was described by Michelsen (1978) and Koch (1983). When the subcrop contours depicted by Underhill and Partington (1993, 1994) are compared with stratigraphic data from the Danish area and the Dutch Central Graben, certain discrepancies are evident (Figs. 4 and 5). The southeastern quarter of the general concentric dome structure depicted by Underhill and Partington (1994) does not match available data. The contours showing the outer limit to Toarcian, Pliensbachian and S inemurian subcrop are drawn in a roughly north-south direction across the RingkObing-Fyn High and in the western part of the Norwegian-Danish Basin (Fig. 4). Underhill and Partington (1994) assumed that the lack of Early Jurassic deposits on the Ringk0bing-Fyn High was caused by later rift-shoulder uplift and erosion. However, Lower Jurassic subcrop patterns and Middle-Upper Jurassic onlap patterns in the Norwegian-Danish Basin run roughly east-west, parallel to the northern margin of the RingkCbing-Fyn High (Fig. 5). This indicates that the RingkCbingFyn High and most of the Norwegian-Danish Basin took part in the early Middle Jurassic uplift (Koch, 1983; Nielsen, 1995, 2001). In the Central Graben, the subcrop contours of Underhill and Partington (1994) show gradually younging of subcrops from the northern part of the Danish Central Graben to the southern limit of the Dutch Central Graben (Fig. 4). In their litho-chronostratigraphic chart, Van Adrichem Boogaert and Kouwe (1993) depict progressively older subcrops from the northern part of the Dutch Central Graben towards the south. In the northern part of the Dutch Central Graben where Van Adrichem Boogaert and Kouwe (1993) indicate the presence of lower-Middle Jurassic deposits and thus the absence of a major hiatus, Underhill and Partington (1994) have drawn their Sinemurian subcrop contour (Fig. 4). A few tens of kilometres to the north in the southern part of the Danish Central Graben, Pliensbachian deposits subcrop below the Aalenian unconformity in the Sinemurian subcrop field of Underhill and Partington (1994). The occurrence of Middle Jurassic basalts in southern Sweden in addition to the volcanic centres in the North Sea, e.g. the Fisher Bank, Glenn, Puffin and Egersund Basins (Smith and Ritchie, 1993) further indicate that the mechanism behind the uplift is not

J. Andsbjerg et al.

one large simple mantle plume. The shape of the uplifted arch described by Ziegler (1982, 1990), which includes the Ringk0bing-Fyn High, and which is transected by the Central Graben, is more in accordance with the available data. Ziegler (1982, 1990) describes Bajocian-Bathonian non-marine deposition in the Central Graben after the uplift event followed by Callovian-Oxfordian transgression. This is in accordance with data from the Danish Central Graben that show marine transgression of non-marine Middle Jurassic rocks during the Callovian-Early Oxfordian (Andsbjerg, 1997; Andsbjerg and Dybkja~r, 2001). On their map of post-uplift marine onlap, Underhill and Partington (1994) show Late Oxfordian onlap in the Danish Central Graben. The role of the long-lived Sorgenfrei-Tornquist Zone should also be considered, as this fundamental fracture zone forms a transition or buffer zone between the Danish Basin and the Baltic Shield (Sorgenfrei and Buch, 1964; Baartman and Christensen, 1975; Eugeno-S Working Group, 1988; Michelsen and Nielsen, 1991, 1993; Mogensen, 1994, 1996). This fault zone accommodated late PalaeozoicMesozoic transtension, phases of volcanic activity (Late Carboniferous-Early Permian, Middle Jurassic and Cretaceous) and Late Cretaceous-Palaeogene transpression resulting in tectonic inversion. Contemporaneously with the regional Aalenian-Bajocian uplift, slow subsidence occurred in the SorgenfreiTornquist Zone and this zone of crustal weakness probably functioned as a hinge zone between the northeastward-tilting basin and the Baltic Shield (Nielsen, 1995,2001).

Implications for hydrocarbon exploration Whereas oil and gas accumulations of commercial value have not been encountered in the Jurassic succession of the Norwegian-Danish Basin, the oil and gas fields Harald and Lulita in the northeastern Danish Central Graben produce from the Middle Jurassic Bryne Formation and Lulu Formation sandstones. The differences in the history of uplift and subsidence across the studied basins had a pronounced influence on the location and possible occurrence of the oil and gas fields. Potential reservoir rocks are abundant in most of the area. In the Norwegian-Danish Basin widespread fluvial, paralic and shoreface sandstones were deposited during Rhaetian to early Sinemurian times (Gassum Formation; Bertelsen, 1978; Hamberg and Nielsen, 2000; Nielsen, 2001). During the early Middle Jurassic uplift, older sandstones were exposed and eroded and deposited as potential reservoir sandstones in accommodation space generated by incipient rift-

Divergent development of two neighbouring basins following the Jurassic North Sea doming event

195

related subsidence in both the Norwegian-Danish Basin and the Danish Central Graben. In the northern part of the Danish Central Graben intra-formational, non-marine source rocks that were generated during this phase are locally important (Petersen et al., 2000). However, the generation of the most important source rocks in the Danish Central Graben, the widespread Upper Jurassic marine mudstones of the Farsund Formation, is a result of rapid subsidence during the main rift phase combined with under-supply of coarse clastics. The position of the NorwegianDanish Basin close to the major sediment source areas of the Baltic Shield and the low rate of subsidence in the Late Jurassic prevented that this basin developed into a starved or under-supplied basin. As a result, conditions in general were unfavourable for deposition of organic-rich muds during the Late Jurassic, although locally both marine and lagoonal mudstones may possess some potential. However, the smaller amount of uplift of the basin during earlyMiddle Jurassic time compared to the central North Sea allowed a much better preservation of the Lower Jurassic mudstones of the Fjerritslev Formation below the base Middle Jurassic unconformity. The upper part of these mudstones shows in places excellent source rock characteristics with TOC of 3-5% and HI of 400-500 (Michelsen, 1989b). The relatively slow regional subsidence in the Norwegian-Danish Basin after the Middle Jurassic uplift did not bury the Fjerritslev Formation mudstones sufficiently for maturation before Neogene regional uplift stopped the maturation process, and in general, the mudstones are immature to marginal mature (Thomsen et al., 1987; Japsen and Bidstrup, 1999). However, in places such as rim-synclines around salt structures and some local grabens, the Fjerritslev Formation occur at depths that may have brought the mudstones into the oil-generating zone.

earliest Oxfordian in the Norwegian-Danish Basin. Transgressive lagoonal and shoreface deposits backstepped up the flanking ramps of the deep areas. These coastal plain environments were drowned during Early-Middle Oxfordian time due to continued sea-level rise and basin expansion. During the Oxfordian-Ryazanian both basins were dominantly marine. As a result of half-graben development of the Danish Central Graben initiated in the late Callovian-Oxfordian, the rate of subsidence increased significantly during Oxfordian-Ryazanian times, up to twenty times that of the NorwegianDanish Basin. The combination of a much slower rate of subsidence and a closer location to sediment sources of the Baltic Shield caused a slow transgression during the Oxfordian-early Volgian followed by recurrent coastal progradation in the NorwegianDanish Basin during middle Volgian-Ryazanian time. This prevented formation of widespread organic-rich shales similar to the Bo Member in the Danish Central Graben. In contrast only one significant regressive event, deposition of the late Kimmeridgian Heno Formation, took place on the western hanging-wall slope of the Danish Central Graben, presumably as a result of tectonic quiescence during a period of change of regional fault directions.

Summary of depositional history

Andsbjerg, J., 1997. Sedimentology and sequence stratigraphy of Middle Jurassic deposits, Danish and Norwegian Central Graben. Unpublished Ph.D Thesis, University of Copenhagen, Parts 1, 2, 3 and 4, 165 pp. and 111 figures and tables. Andsbjerg, J., 2001. Sedimentology and sequence stratigraphy of the Bryne and Lulu Formations, Middle Jurassic, northern Danish Central Graben. In: F. Surlyk and J.R. Ineson (Editors), The Jurassic of Denmark and Greenland. Geol. Den. Surv. Bull. (in press). Andsbjerg, J. and Dybkja~r, K., 2001. Jurassic sequence stratigraphy of the Danish Central Graben. In: F. Surlyk and J.R. Ineson (Editors), The Jurassic of Denmark and Greenland. Geol. Den. Surv. Bull. (in press). Baartman, J.C. and Christensen, O.B., 1975. Contribution to the interpretation of the Fennoscandian Border Zone. Danmarks Geologiske UndersOgelse II R~ekke 102, 47 pp. Badley, M.E., Price, J.D., Dahl, C.R. and Agdestein, T., 1988. The structural evolution of the northern Viking Graben and its bearing upon extensional modes of basin formation. J. Geol. Soc., London, 145: 455-472.

From an initially rather uniform situation in the Middle Jurassic, the Danish Central Graben and the eastern part of the Norwegian-Danish Basin developed significant differences during the Late Jurassic (Figs. 2 and 3). From mid-Aalenian to Bajocian times deposition occurred within the deepest graben areas only. The Late Jurassic expansion of the depositional areas may have started already in the early Bathonian in the eastern part of the NorwegianDanish Basin with deposition of fluvial and lacustrine sediments on the lower ramps. However, the significant expansion of the depositional areas occurred in the latest Bathonian-Callovian time in the Danish Central Graben and in the late Callovian-

Acknowledgements The Energy Research Programme at the Danish Energy Agency, the Danish Research Academy and GEUS financially supported the studies. The referees A. Folkestad and J. Gelbjerg and editor O. Martinsen are thanked for valuable suggestions. Gurli Hansen, Eva Melskens and Alice Rosenstand made the drawings.

References

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J. Andsbjerg et al. Johannessen, EN., 2001. Depositional environments and sequence stratigraphy of paralic and shallow marine Upper Jurassic reservoir sandstones in the northern part of the Danish Central Graben. In: E Surlyk and J.R. Ineson (Editors), The Jurassic of Denmark and Greenland. Geol. Den. Surv. Bull. (in press). Johannessen, EN. and Andsbjerg, J., 1993. Middle to Late Jurassic basin evolution and sandstone reservoir distribution in the Danish Central Trough. In: J.R. Parker (Editor), Petroleum Geology of Northwest Europe: Proceedings of the 4th Conference. Geological Society, London, pp. 271-298. Johannessen, EN., Dybkjam K. and Rasmussen, E.S., 1996. Sequence stratigraphy of Upper Jurassic sandstones in the northern part of the Danish Central Trough, North Sea. Mar. Pet. Geol., 13: 755-770. Klingspor, I., 1976. Radiometric age-determinations of basalts, dolerites and related syenite in Sk~ne, southern Sweden. Geol. F6ren. Stockholm F6rh., 98:195-216. Koch, J.O., 1983. Sedimentology of Middle and Upper Jurassic sandstone reservoirs in Denmark. In: J.EH. Kaasschieter and T.J.A. Reijers (Editors), Petroleum Geology of the Southeastern North Sea and the Adjacent Onshore Areas. Geol. Mijnbouw, 62: 115-129. Korstgfird, J.A., Lerche, I., Mogensen, T.E. and Thomsen, R.O., 1993. Salt and fault interactions in the northeastern Danish Central Graben: observations and inferences. Bull. Geol. Soc. Den., 40: 197-255. Leeder, M.R., 1983. Lithospheric stretching and North Sea Jurassic clastic sourcelands. Nature, 305:510-514. Mackertich, D., 1996. The Fife Field, UK Central North Sea. Pet. Geosci., 2: 373-380. Michelsen, O., 1978. Stratigraphy and distribution of Jurassic deposits of the Norwegian-Danish Basin. Geological Survey of Denmark, Series B, 2, 28 pp. Michelsen, O., 1989a. Revision of the Jurassic lithostratigraphy of the Danish Subbasin. Danmarks Geologiske Unders0gelse Serie A, 24, 21 pp. Michelsen, O., 1989b. Log-sequence analysis and environmental aspects of the Lower Jurassic Fjerritslev Formation in the Danish Subbasin. Geological Survey of Denmark, Series A, 25, 23 pp. Michelsen, O. and Nielsen, L.H., 1991. Well records on the Phanerozoic stratigraphy in the Fennoscandian Border Zone, Denmark; Hans-l, Saeby-1 and Terne-1 wells. Geological Survey of Denmark, Series A, 29, 37 pp. Michelsen, O., Nielsen, L.H., Johannessen, EN., Andsbjerg, J. and Surlyk, F., 2001. Jurassic lithostratigraphy and depositional development onshore and offshore Denmark. In: E Surlyk and J.R. Ineson (Editors), The Jurassic of Denmark and Greenland. Geol. Den. Surv. Bull. (in press). Mogensen, T.E., 1994. Palaeozoic structural development along the Tornquist Zone, Kattegat area, Denmark. Tectonophysics, 240: 191-214. Mogensen, T.E., 1996. Triassic and Jurassic structural development along the Tornquist Zone, Kattegat, Denmark. Tectonophysics, 252: 197-220. Mr J.J., 1986. Seismic structural mapping of the Middle and Upper Jurassic in the Danish Central Trough. Geological Survey of Denmark, Series A, 13, 37 pp. Mr J.J. and Rasmussen, E.S., 2001. Middle Jurassic-Early Cretaceous rifting of the Danish Central Graben. In: F. Surlyk and J.R. Ineson (Editors), The Jurassic of Denmark and Greenland. Geol. Den. Surv. Bull. (in press). Nielsen, L.H., 1995. Genetic stratigraphy of the Upper TriassicMiddle Jurassic deposits of the Danish Basin and Fennoscandian Border Zone. Unpublished Ph.D Thesis, University of Copenhagen, Parts 2 and 3, 162 pp. and 109 figs. and tables. Nielsen, L.H., 2001. Late Triassic-Jurassic development of the Danish Basin and Fennoscandian Border Zone, southern Scandinavia. In: F. Surlyk and J.R. Ineson (Editors), The Jurassic of Denmark and Greenland. Geol. Den. Surv. Bull. (in press).

Divergent development of two neighbouring basins following the Jurassic North Sea doming event Norling, E. and Bergstr6m, J., 1987. Mesozoic and Cenozoic tectonic evolution of Scania, southern Sweden. In: EA. Ziegler (Editor), Compressional Intra-Plate Deformations in the Alpine Foreland. Tectonophysics, 137: 7-19. Norling, E., Ahlberg, A., Erlstr6m, M. and Sivhed, U., 1993. Guide to the Upper Triassic and Jurassic Geology of Sweden. Sveriges Geologiska Unders6kning Ca 82, 71 pp. Pedersen, G.K., 1985. Thin, fine-grained storm layers in a muddy shelf sequence: an example from the Lower Jurassic in the Stenlille-1 well. J. Geol. Soc., London, 142: 357-374. Petersen, H.I., Andsbjerg, J., Bojesen-Koefoed, J.A. and Nytoft, H.E, 2001. Coal-generated oil: source rock evaluation and petroleum geochemistry of the Lulita oilfield, Danish North Sea. J. Pet. Geol., 23: 55-90. Poulsen, N.E., 1992. Jurassic dinoflagellate cyst biostratigraphy of the Danish Subbasin in relation to sequences in England and Poland; a preliminary review. Rev. Palaeobot. Palynol., 75:20 pp. Poulsen, N.E., 1996. Dinoflagellate cysts from marine Jurassic deposits of Denmark and Poland. American Association of Stratigraphic Palynologists, Contribution Series, 31,227 pp. Poulsen, N.E. and Riding, J.B., 2001. The Jurassic dinoflagellate cyst zonation of Subboreal northwest Europe. In: F. Surlyk and J.R. Ineson (Editors), The Jurassic of Denmark and Greenland. Geol. Den. Surv. Bull. (in press). Printzlau, I. and Larsen, O., 1972. K/At age determinations on alkaline olivine basalts from Skfine, south Sweden. Geol. F6ren. Stockholm F6rh., 94: 259-269. Smith, K. and Ritchie, J.D., 1993. Jurassic volcanic centres in the Central North Sea. In: J.R. Parker (Editor), Petroleum Geology of Northwest Europe: Proceedings of the 4th Conference. Geological Society, London, pp. 519-531.

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Sorgenfrei, T. and Buch, A., 1964. Deep Tests in Denmark 19351959. Danmarks Geologiske Undersogelse III R~ekke, 36, 146 pp. Thomsen, E., Damtofte, K. and Andersen, C., 1987. Hydrocarbon plays in Denmark outside the Central Trough. In: J. Brooks and K. Glennie (Editors), Petroleum Geology of North West Europe. Graham and Trotman, London, pp. 497-508. Tralau, H., 1973. En palynologisk fildersbestfimning av vulkanisk aktivitet i Sk~ne. Fauna Flora, 4:121-125. Underhill, J.R. and Partington, M.A., 1993. Jurassic thermal doming and deflation in the North Sea: implications of the sequence stratigraphic evidence. In: J.R. Parker (Editor), Petroleum Geology of Northwest Europe: Proceedings of the 4th Conference. The Geological Society, London, pp. 337-345. Underhill, J.R. and Partington, M.A., 1994. Use of Genetic sequence stratigraphy in defining and determining a regional tectonic control on the "Mid Cimmerian Unconformity": implications for North Sea basin development and the global sea-level chart. In: E Weimer and H.W. Posamentier (Editors), Siliciclastic Sequence Stratigraphy: Recent Developments and Applications. Am. Assoc. Pet. Geol., Mem., 58: 449-484. Van Adrichem Boogaert, H.A. and Kouwe, W.F.E, 1993. Stratigraphic nomenclature of the Netherlands, revision and update by RGD and NOGEPA. Mededelingen Rijks Geologische Dienst, 50. Vejb~ek, O.V., 1989. Effects of asthenosperic heat flow in basin modelling exemplified with the Danish Basin. Earth Planet. Sci. Lett., 95: 97-114. Ziegler, EA., 1982. Geological Atlas of Western and Central Europe. Shell International Petroleum Maatschappij B.V., The Hague. Ziegler, EA., 1990. Tectonic and palaeogeographic development of the North Sea rift system. In: D.J. Blundell and A.D. Gibbs (Editors), Tectonic Evolution of the North Sea Rifts. pp. 1-36.

Geological Survey of Denmark and Greenland (GEUS), Thoravej 8, DK-2400 Copenhagen NV, Denmark Geological Survey of Denmark and Greenland (GEUS), Thoravej 8, DK-2400 Copenhagen NV, Denmark Geological Survey of Denmark and Greenland (GEUS), Thoravej 8, DK-2400 Copenhagen NV, Denmark Geological Survey of Denmark and Greenland (GEUS), Thoravej 8, DK-2400 Copenhagen NV, Denmark

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An integrated study of the Garn and Melke formations (Middle to Upper Jurassic) of the Smorbukk area, Halten Terrace, mid-Norway Stephen Corfield, lan Sharp, KjelI-Owe H&ger, Tom Dreyer and John Underhill

This contribution presents a re-evaluation of the shallow marine reservoir sandstones of the Garn Formation and their relationship to heterolithic shelf facies of the Not and Melke formations in the Sm~rbukk area, Halten Terrace, mid-Norway. Most of the existing interpretations of the Garn Formation suggest development as a relatively homogeneous and simple sheet-like sandstone, with a lateral extent of 10s of kin. Deposition has been previously interpreted to have occurred during a period of tectonic quiescence. This interpretation is largely based on the lithostratigraphic correlation of remarkably uniform wireline log profiles supported by regional biostratigraphic correlations. In contrast, an integrated study of seismic, seismic attribute, core and biostratigraphic data in the Sm~rbukkSmcrbukk South area indicate that the Garn Formation is strongly age and facies diachronous, often over short distances (5-10 km). Furthermore, it can be demonstrated that sediment dispersal and stratigraphic architecture were influenced by fault-created physiography. The lower part of the Garn Formation comprises up to 50 m of aggradationally stacked shallow marine sandstones which are in disconformable contact with underlying heterolithic shelf facies of the Not Formation. These lower sands onlap onto tectonically active structural highs (Sm~rbukk fault block and Sm~rbukk South field) and are of limited geographical extent, only being present in down-dip depocentres. The overlying upper Garn sandstones comprise a series of aggradational to progradational shelf-shoreface cycles, with individual cycles progressively backstepping and onlapping onto structural highs. Collectively, these cycles define a retrogradational stacking pattern. Seismic and well log correlation indicate that these upper sands are coeval with Melke Formation shelf/offshore mudstone facies in down-dip depocentres. There is a clear and hitherto largely overlooked relationship between fault-controlled basin physiography, facies distribution, sediment dispersal and stacking patterns. These new data have important implications for understanding palaeogeographic development and reservoir geometry of the Garn Formation.

Introduction

The base of the Jurassic syn-rift sequence on the Halten Terrace is currently interpreted as the contact between the pre-rift Middle Jurassic shallow marine sandstones of the Fangst Group (Ile, Not and Garn formations) and overlying shelf mudstones of the Upper Jurassic Viking Group (Melke and Spekk formations, Fig. 1; Dalland et al., 1988; Ehrenberg et al., 1992; Koch and Heum, 1995). In particular, existing sedimentological and reservoir interpretations of the Garn Formation suggest a relatively homogeneous and simple sheet-like sandstone with lateral extent of 10s of km deposited during "an interval of nearly total tectonic quiescence" (e.g. Ehrenberg et al., 1992; see also discussion of the Middle Jurassic in Dor6, 1992). This interpretation is largely based on lithostratigraphic correlation of remarkably uniform wireline log profiles supported by regional biostratigraphic correlations. However, regional-scale facies diachroneity of the Garn and Melke formations was demonstrated early on in the exploration history of

the Halten Terrace by Gjelberg et al. (1987). In particular, Gjelberg et al. (1987) concluded that the Garn Formation comprised a series of "backstepping clastic wedges", and implied a facies diachroneity between Garn Formation sandstones and Melke Formation mudstones. These data appear to have been largely overlooked by later workers. The aim of the present study was thus to readdress the depositional architecture of the Garn Formation, and its relationship to the under- and overlying Not and Melke formations. We focus on the Sm~rbukkSm~rbukk South area of the Halten Terrace, where there is a wealth of good-quality 3D seismic, wireline log, core and biostratigraphic data. An integrated well-tied seismic stratigraphic approach was used. These new data clearly indicate facies and age diachroneity of the Garn Formation over short distances (5-10 km) and support the original regional "backstepping wedges" model of Gjelberg et al. (1987). There is also clear evidence of fault-related topography during deposition of the Garn and older formations. These new data have major implications

Sedimentary Environments Offshore Norway - Palaeozoic to Recent edited by O.J. Martinsen and T. Dreyer. NPF Special Publication 10, pp. 199-210, Published by Elsevier Science B.V., Amsterdam. 9 Norwegian Petroleum Society (NPF), 2001.

200

S. Corfield et al.

Fig. 1. Lithostratigraphy of the Halten Terrace region. Modified from Dalland et al. (1988).

for understanding palaeogeographic development and sandbody connectivity within the Garn Formation.

Regional setting and structural style A full regional and structural review of the Halten Terrace and SmCrbukk/SmOrbukk South region is beyond the scope of this paper, and the reader is referred to Koch and Heum (1995) and Corfield and Sharp (2000) for more details. The key structural

observations pertinent to this paper are summarised below and on Fig. 2. The main SmCrbukk field consists of an easterly tilted fault block bounded to the west by a major NNE-SSW trending basement-involved normal fault (SmCrbukk Fault, Fig. 2; Aasheim et al., 1986; Ehrenberg et al., 1992; Koch and Heum, 1995; Corfield and Sharp, 2000). The Sm0rbukk Fault loses displacement towards both the north and south, with displacement transferred to the Revfallet and Trestakk-Sm0rbukk

An integrated study of the Garn and Melke formations (Middle to Upper Jurassic) of the SmOrbukk area

South faults, respectively. It is important to note that, whilst the present structural crest of the SmOrbukk field is located in the northwest, the area of maximum footwall uplift during Middle-Late Jurassic extension was located further south (well 6506/11-2 region). This is clear from the amount of Upper Jurassic erosion evident in this area (Figs. 2 and 4). The northward shift of the structural crest is interpreted as the product of renewed extension in the Cretaceous that increased in magnitude northward toward the Nordland Ridge. This younger episode was of greater magnitude than the Jurassic rifting (cf. Pascoe et al., 1999, in their discussion of the Nordland Ridge), a feature that has largely been overlooked in previous studies. It is important to bear this change of structural elevation in mind when addressing deposition and distribution of the Middle and Upper Jurassic sequences. In contrast to the SmOrbukk field, Sm0rbukk South is a simple dome-like structure overlying a Triassic salt pillow, bounded along its western flank by a major, N-S trending, basement-involved fault (Trestakk-SmOrbukk South Fault, Fig. 2). Displacement on the Trestakk-SmOrbukk South Fault decreases northwards to a well-defined tip-point. Associated with this northwards loss of displacement the geometry of the Jurassic interval evolves from a "classic" wedge thickening towards the TrestakkSmOrbukk South Fault in the south, to a westerly verging monoclinal flexure close to the northerly tip of the fault (Fig. 2, sections 1 to 3). The Middle and Upper Jurassic sequences thicken northwestwards over this monoclinal flexure into a synclinal depocentre located between SmOrbukk and Sm0rbukk South. This south to north variation in geometry is clear evidence that, during the deposition of the Lower, Middle and Upper Jurassic sequences, the TrestakkSmOrbukk South Fault was associated with a surface break in the south (with a well-developed half-graben geometry) while to the north the fault was buried and overlain by a fault-propagation/fault-tip fold (Corfield and Sharp, 2000). This structural style, and in particular the broad folding, is interpreted to have developed due to decoupling of the Jurassic sequences from underlying thick Triassic salt (cf. Withjack et al., 1989; Pascoe et al., 1999). Comparisons can be made to geometries recently described by Gawthorpe et al. (1997), Gupta et al. (1999) and Sharp et al. (2000) for marine syn-rift sequences adjacent to propagating faults and folds from the Suez Rift, and also to alluvial deposits developed parallel to the Tobin fault in the Western USA extensional province (Jackson and Leeder, 1994). These structural observations indicate that faultcreated basin physiography was more significant in the SmOrbukk region in the Lower and Middle Juras-

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sic than previously thought. In particular, structural lows/depocentres can be identified in the hanging wall of the Trestakk-SmOrbukk South Fault and in the vicinity of wells 6506/12-7 and 6506/12-10 in the northwestern part of the study area. Structural highs are evident in the footwall of the SmOrbukk and Trestakk-SmOrbukk South faults.

Well ties and seismic stratigraphy During the course of this study it became apparent that, despite the good-quality 3D seismic data, the "Top Garn" Formation lithostratigraphic well pick could not be tied to a single seismic reflector. Similarly, biostratigraphic data revealed a range of ages for the "Top Garn" Formation pick. An initial attempt to solve this problem was to map the underlying Not-Ile Formation contact, which represents one of the most regionally correlatable bioand lithostratigraphic picks on the Halten Terrace at the top Aalenian-base early Bajocian contact. This event consistently tied with a trough on the seismic data and is interpreted as a regionally extensive MFS, probably equivalent to the base Rannoch Formation and MFS J24 in the northern North Sea (Partington et al., 1993). The top Garn picks were then mapped in the overlying section above the Ile/Not datum, resulting in the identification of two (locally three) distinct ages and distributions of Garn Formation sandstones. Seismic data indicate that the oldest Garn sandstone is restricted to the northern part of the synclinal depocentre between Sm0rbukk and Sm0rbukk South. This sand is penetrated in wells 6506/12-7, 6506/12-9 and 6506/12-10 and is overlain by the thickest Melke section. Biostratigraphic data indicate an early Bajocian to lowest late Bajocian age (Fig. 6). The "Top Garn" seismic expression in these wells is a weak, discontinuous peak that onlaps both the monoclinal flexure of SmOrbukk South and the hanging wall dipslope of the SmOrbukk fault block to the south and west (Figs. 3 and 4). The mapped extent of the peak indicates a N E - S W elongate body restricted to the deepest part of the depocentre. An age-equivalent Garn sandstone interval is interpreted to be present in the immediate hanging wall of the Trestakk-SmOrbukk South Fault. This sand increases in thickness southwards as a product of increased displacement on the fault (e.g. Fig. 2, section 1). A sequence of younger (late Bajocian to early Bathonian) Garn sandstones are penetrated in wells located on the structural highs of SmOrbukk South and the crestal area of SmCrbukk (Fig. 3). Seismic data indicate that this sequence thickens westwards from SmOrbukk South and onlaps the tilted hanging wall dipslope of the SmOrbukk fault block. Seismic,

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An integrated study of the Garn and Melke formations (Middle to Upper Jurassic) of the SmOrbukk area

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Fig. 3. Cartoon of the geometry of the Not, Garn and Melke formations between SmOrbukk and SmOrbukk South as evident from well ties and seismic. The key wells are located in their approximate structural positions. The colours of the time lines correspond to the seismic reflectors in Fig. 4.

wireline log and biostratigraphic data also clearly indicate that sands on the structural highs are coeval with Melke Formation offshore/shelf distal shoreface and mudstone facies in the lows. Well picks for this younger "Top Garn" sandstone on SmOrbukk South tie with two laterally consistent, high-amplitude seismic peaks. The lower of the two peaks (red/purple reflector, Fig. 4) ties with the "Top Garn" sandstone on SmOrbukk South (6506/12-5, Figs. 3 and 4) and Melke Formation offshore mudstones/distal shoreface facies in wells 6506/12-7 and 6506/12-10 in the synclinal depocentre to the northwest of SmOrbukk South. Westwards, the reflector onlaps the hanging wall dipslope of SmOrbukk downdip of wells 6506/11-2 and 6506/11-4 (Fig. 3). The upper reflector (white, Fig. 4) ties with younger "Top Garn" sandstone on SmOrbukk South in wells 6506/12-3, 6506/12-8 and 6406/3-3 and Melke Formation mudstones in wells 6506/12-7 and 6506/12-10. However, this reflector onlaps higher onto the hanging wall dipslope of Sm0rbukk and ties with the "Top Garn" in wells 6506/11-2 and 6506/11-4. Updip of these wells, the reflector "dims" dramatically (Fig. 4), indicating possible erosion of the Garn sands (cf. Ehrenberg et al., 1992, their fig. 3). This relationship is also evident on seismic attribute maps. Fig. 5 illustrates the excellent correlation of the youngest "Top Garn" pick from wells, with the seismic am-

plitude of the white reflector. The red colour on the amplitude map essentially correlates with Melke Formation mudstones while the yellow/green colour correlates with Garn Formation sandstone. The map indicates that the youngest Garn sands are concentrated in a linear belt high on the hanging wall dipslope of the Sm0rbukk fault block whilst Melke Formation mudstones predominate in the northern part of the synclinal depocentre in the region of wells 6506/12-7 and 6506/12-10. Similarly, the presence of mudstones at this horizon in well 6506/12-5 on Sm0rbukk South and sands in wells 6506/12-3, 6506/12-8 and 6406/3-3 correlate well with the attribute map. The use of these relationships enable the prediction of Garn sandstones at this level in the undrilled immediate hanging wall and northern tip of the Trestakk-SmOrbukk South Fault (Fig. 5). In this area a linear N-S trending yellow/green body is evident surrounded by higher amplitudes (red). Seismic sections across this body indicate that it has a subtle but positive structural relief and is draped by overlying reflectors. Such relationships are positive for the presence of sand. The northern tip of this body appears to curve round the tip of the TrestakkSmOrbukk Fault to connect with low-amplitude reflectors (?sand) located northwest of well 6506/12-3. In summary, at least two ages and distributions of Garn sandstones can be mapped in the SmOrbukk area;

Fig. 2. (a) Location of the study area in relation to mainland Norway and the main structural elements of the mid-Norway region. (b) Base Cretaceous time map of the Sm0rbukk area showing the key structural elements and wells used in this study. Note the overall southerly plunge from SmCrbukk towards the Trestakk-SmOrbukk Fault. (c) South to north seismic sections across the Trestakk-SmCrbukk Fault, showing the strike variation in structural geometry and mapped Jurassic seismic packages. Cross-section location shown in b. Note that the Top Garn Formation pick is equivalent to the white reflector in Figs. 3 and 4, and that in up-dip wells the reflector ties to sand, whilst it ties to mudstone in down-dip wells.

PO 0

--X

Fig. 4. Dip line from the footwall crest of SmCrbukk to well 6506/12-5 on SmCrbukk South. Note the onlap of the red/purple and white reflectors onto the hanging wall dipslope of Sm~arbukk.

An integrated study of the Garn and Melke formations (Middle to Upper Jurassic) of the SmOrbukk area

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Fig. 5. Amplitude map of the younger "Top Garn" (white reflector). Note the excellent correlation with the lithologies penetrated in the wells (sandstone = yellow/green; mudstone = red). The geometry of the interpreted sandbodies appears to parallel basin physiography. Also note that the overall southerly dip, and the exaggerated relief of the SmOrbukkSouth field is due to Cretaceous uplift in the north. This overall southerly dip was not present in the Jurassic, with a southern and northern depocentre identifiable. one occurring in depocentres in the vicinity of well 6506/12-10 and the immediate hanging wall of the Trestakk-Sm0rbukk Fault, and a younger Garn sand generally occurring on the structurally elevated areas.

Implications for depositional models and palaeogeographic development The Not, Garn and Melke formations are fully marine in origin, with the Garn Formation typically comprising high-energy deltaic, shoreface and shelf sandstones, whilst the Not and Melke formations comprise bioturbated offshore/shelf mudstones, siltstones and minor distal lower shoreface/storm sands (Gjelberg et al., 1987; Dalland et al., 1988; Dor6, 1992). The Not Formation is of Aalenian-early Bajocian age, whilst the Garn Formation is essentially agecontemporaneous with the whole of the Brent Group of the northern North Sea, i.e. Bajocian-Bathonian. The Melke Formation is of Bajocian to Oxfordian

age (Fig. 1; Gjelberg et al., 1987; Dalland et al., 1988; Dor6, 1991, 1992). Whilst facies diachroneity of the Garn and Melke formations has previously been identified on a regional scale (e.g. Gjelberg et al., 1987), seismic well correlations presented in this paper indicate diachroneity on a local scale (over distances of 5 - 1 0 kin). These new data have implications for depositional models and understanding the palaeogeographic development. The Not to Garn contact is clearly associated with onlap observed on seismic data (Fig. 4). Examination of cored intervals in the Sm0rbukk region revealed that the basal contact of the Garn Formation with the underlying Not Formation is often sharp and erosive, with high-energy shallow marine sandstones abruptly overlying offshore to lower shoreface/shelf stormdominated mudstones and fine-grained sandstones. The contact is typically marked by a coarse pebble lag. Analysis of dip-meter data in the SmOrbukk region also indicates that the Garn to Not contact is a

206 disconformity associated with a slight dip change. It is thus reasonable to assume that the contact represents a regionally significant regressive erosion surface of intra-early Bajocian age. This regression was likely associated with structural tilting to produce the dip-discordance and seismically defined onlap. However, in other nearby Halten wells, notably 6406/2-3, 6406/3-2, 6407/4-1 and 6407/6-3, the Not to Garn contact is more gradational, associated with an apparently genetically related progradational association of facies from offshore/shelf to upper shoreface/shelf. Also, where the contact is sharp, as it is in many of the SmCrbukk wells, the shift in facies across the contact is minimal (upper shoreface/shelf sandstones overlying progradational offshore/shelf heterolithics). Biostratigraphic and lithostratigraphic data also indicate that erosion at the base of the Garn Formation into the underlying Not Formation is minimal, the Not Formation and associated key time-lines being present over almost the entire Halten Terrace. Absent section is possibly as little as 5-10 m. These data can lead to two quite differing interpretations for the genesis of the Not-Garn contact. The first is that the Not-Garn contact represents a regionally significant regressive erosion surface. "Proximal" or up-dip wells are associated with pronounced erosion, whilst "distal" or down-dip wells are associated with a more genetically related progradational signature and limited facies dislocation (cf. Gjelberg et al., 1987). In this scenario, limited erosion in distal/down-dip wells can be attributed to the development of regressive surfaces of marine erosion during forced regression (Hunt and Tucker, 1992). An explanation for the lack of extensive up-dip erosion of the Not Formation is more problematic, although it can be attributed to sediment by-pass and non-deposition across the shelf. The second scenario is to interpret the Not-Garn contact in terms of a genetic depositional relationship resulting from, for example, the migration of a tidally influenced shelf sandridge complex over, and eroding into, adjacent intra-sandridge shelf heterolithics. In such a depositional setting local erosion can occur at the base of the sandridge, typically in "proximal" areas or in areas of low accommodation development (e.g. structural highs). In contrast, in more "distal" settings, the contact can be more gradational with genetically related units from offshore to upper shoreface/sandwave crest preserved. This situation is comparable to that documented by Reynaud et al. (1999) for the Quaternary-Recent Celtic Sea sandbanks and to the Middle Jurassic Sortehat and VardeklOft formations of East Greenland (Heinburg and Birkelund, 1984; Surlyk, 1990a,b; Engkilde and Surlyk, 1993). On balance, however, we feel this second scenario is the less likely. This is based on fa-

S. Corfield et al.

cies data from core, the clear seismically defined onlap and dip-discordance evident across the Not-Garn contact, and also the apparently isochronous nature of this erosive contact over the Halten region. However, it is clear that increased biostratigraphic resolution of the early Bajocian interval is required to clarify the nature and duration of the depositional break. Seismic and biostratigraphic data above this basal erosion surface clearly indicate that the oldest Garn Formation sandstones are located in down-dip depocentres, and that these sands onlap towards structural highs (e.g. top early Bajocian time line in Fig. 6). Moving up stratigraphic section, successively younger sands can be demonstrated to onlap further up-dip. This up-dip migration of sand onlap is also associated with a lateral facies change, with onlapping shallow marine sandstones on structural highs being laterally time equivalent to offshore/shelf mudstone facies of the Melke Formation down-dip (Fig. 6). The Garn to Melke Formation contact is thus diachronous over relatively short distances. This relationship is most apparent when a correlation is made between wells located on SmCrbukk South (6506/12-5, 6506/12-3, 6406/3-3 and 6506/12-8) and wells in the synclinal depocentre to the north (6506-12/7, 12/6 and 12/10, Fig. 6). Shoreface/shelf sandstones on the structural high of SmCrbukk South pass laterally down-dip and "offstructure" into offshore/shelf mudstones and thin sandstones. Individual shoreface to offshore cycles are typically aggradational to progradational in both the up-dip and down-dip wells, with the tops of cycles in both settings marked by extensive bioturbation and the development of early diagenetic cements. These surfaces are interpreted as resulting from periods of marine flooding and omission colonisation (cf. Taylor and Gawthorpe, 1993), and can be correlated with reasonable confidence in the study area using core and wireline log data. Collectively, shoreface cycles in the upper part of the Garn Formation stack in a retrogradational pattern, defining a retrogradational parasequence set (i.e. successive shoreface cycles are interpreted to have been deposited in deeper water). This situation is particularly well developed in well 6506/12-5 (Fig. 6). Seismic and well log correlations also indicate that individual cycles progressively backstep and onlap onto structural highs. Comparisons can be made with the approximately age-equivalent Pelion and Fossilbjerget members (Vardekl0ft Formation) of East Greenland (Heinburg and Birkelund, 1984; Engkilde and Surlyk, 1993), and with the Tarbert and Heather formations of the northern North Sea (e.g. Graue et al., 1987; Helland-Hansen et al., 1992). In both these examples pulsed progradation occurred in an overall

~.

c~ cc

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~.

~~

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Fig. 6. Simplified correlation panel flattened on the top early Bathonian time line. Note that the older Garn 1 sands are restricted to the northern synclinal depocentre and that the younger Garn 2 sands on the highs are equivalent to Melke Formation mudstone facies off-structure. Also note retrogradational log pattern evident towards the Garn-Melke lithostratigraphic contact (e.g. 6506/12-5). Logs are gamma ray. Shading to right of logs indicates core control.

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transgressive setting resulting in the development of retrogradationally stacked shoreface cycles. The tops of the Pelion Member and Tarbert Formation are also highly time-diachronous. Using the seismic, biostratigraphic and facies correlations outlined above, a simplified three-stage palaeogeographic evolution of the Not, Garn and Melke formations can be suggested in the greater SmOrbukk area. Stage 1 - - forced regressive to Iowstand deposition

A regional regressive event of intra-early Bajocian age is associated with erosion of 5-10 m of the un-

S. Corfield et al.

derlying Not Formation on structural highs (Fig. 7), with a more progradational signature occurring in down-dip areas. Onlap observed on seismic data of the basal parts of the Garn Formation toward the structural highs and the presence of a disconformity observed in core data indicate that this contact is mainly controlled by footwall uplift and rotation of the SmCrbukk fault block. In the Sm0rbukk area, two main depocentres can be identified during this time: a northern depocentre located between SmCrbukk South and the northern, present-day structural crest of Sm0rbukk, and a southern depocentre located in the immediate hanging wall of the Trestakk-Smc~rbukk Fault. The northern depocentre is penetrated by wells.

Fig. 7. Simplified tectono-stratigraphic evolution of the Garn Formation sandstones and Melke Formation mudstones in relation to growth of the SmCrbukk and SmCrbukk South structures. Stage 1. Garn sandstone 1 (intra-early Bajocian to early-late Bajocian). Deposition of the oldest sand, possibly subsequently eroded from the structural highs. Stage 2. Garn sandstone 2 (late Bajocian to early Bathonian). The sands are restricted to structural highs with the contemporaneous deposition of Melke Formation mudstones in the depocentre. Note the possible incised geometry of the youngest sand in well 6506/12-8.

An integrated study of the Garn and Melke formations (Middle to Upper Jurassic) of the SmOrbukk area

In contrast, the interpretation of Gain sandstone in the southern depocentre is based upon seismic data. Seismic and well data in the northern depocentre indicate that the sandbody is elongated NE-SW. The area in between these depocentres and also the structural highs were undergoing active erosion or non-deposition. It is probable that the present-day distribution and geometry of this lower Garn sandstone reflects transgressive erosion by the overlying younger Garn sandstones (cf. "Depositional Remnants" of Martinsen and Krystinik, 1998).

Stage 2 m pulsed progradation in a transgressive setting Stage 2 is characterised by a marked transgression associated with backstepping of shoreline/shelfal sands onto structural highs in at least two main phases (Fig. 7). Transgressive episodes were punctuated by progradation, with successively younger parasequences retreating further up-dip to produce a retrogradational parasequence set. Regional comparisons, for example to the Tarbert Formation and Vardekloft Formation, indicate that transgression was primarily related to a regional Jurassic sea-level rise superimposed on developing rift structures. Seismic and well data from SmOrbukk indicate that the northern and southern depocentres persisted during this stage and were mudstone dominated, whilst the footwall crests of both the SmOrbukk and TrestakkSm0rbukk faults were undergoing subaerial erosion, with linear, attached shorelines developed parallel to the structural highs. Up-dip sands are coeval with down-dip mudstones. Seismic attribute mapping indicates the presence of a linear sandbody in the southern depocentre (?sandridge) and a large sandbody close to the northern tip of the Trestakk-SmOrbukk South Fault (Fig. 5). The latter sandbody has an elongate and widening external form northwards. It possibly represents a spit system or shelf sand comparable to the Upper Jurassic sands in the Draugen field (Provan, 1992). Both sandbodies could be related to focusing of tidal currents in the elongate depocentre between the main SmOrbukk field and the TrestakkSmOrbukk South Fault (cf. Vardekloft Formation in East Greenland; Engkilde and Surlyk, 1993).

209

data indicate that the Spekk Formation unconformably overlies the eroded Garn Formation (Fig. 4). This three-stage evolutionary scenario of initial lowstand sand deposition (stage 1), pulsed progradation during transgression (stage 2) and ultimately drowning (stage 3) is comparable to that suggested by Gjelberg et al. (1987) for the generation of sheet-like sandbodies over large geographical areas. However, the implications of this model for internal sandbody architecture and sandbody connectivity remain relatively poorly understood. Seismic and biostratigraphically constrained log correlations presented in this paper indicate that a simple sheet-like reservoir zonation scheme for the Garn Formation, as currently used in the Smc~rbukk, SmOrbukk South, Heidrun, Trestakk and Tyrihans hydrocarbon fields, is perhaps unreasonable. Such sheet-like schemes do not address the diachronous, onlapping nature of the lower bounding surface or the diachroneity of the upper bounding surface.

Conclusions In contrast to previous studies, which interpret the Garn Formation as a sheet-like and essentially synchronous unit, this study demonstrates that the Gain Formation is highly diachronous in nature over relatively short distances (5-10 kin). The base of the Garn Formation can still be interpreted as relatively isochronous, but the top is clearly diachronous. This relationship is supported by both seismic and biostratigraphic data and is here related to fault-related physiography and growth of the Sm0rbukk and Sm0rbukk South structures. At least three Garn sandstones with distinct ages and distributions can be identified. The oldest (early Bajocian) Garn sandstone is restricted to a synclinal depocentre between the main SmOrbukk structure and SmOrbukk South, whilst the younger Garn sandstones backstep onto the structural highs of the footwall of SmOrbukk and SmOrbukk South. The younger Garn sands on the structural highs are coeval with Melke Formation marine mudstones in down-dip depocentres. This new interpretation has implications for understanding the timing of fault growth in the Jurassic of the Halten Terrace and consequently, the sediment dispersal, stacking patterns and lateral continuity of reservoir sandstone units within the Garn Formation.

Stage 3 - - final drowning

Acknowledgements Transgression of the Garn Formation and widespread deposition of Melke facies mudstones appears to have occurred by the top early Bathonian. However, erosion of structural highs possibly persisted up-dip of the well penetrations on SmOrbukk, where seismic

The authors wish to acknowledge the constructive reviews of Tony Dor6 and Jim Steidtmann. Discussions on the geology of mid-Norway and East Greenland with John Gjelberg, Trond Lien, Ian Carr and Carl

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Watkins also helped the formulation of ideas presented in this paper. References Aasheim, S.M., Dalland, A., Netland, A. and Thon, A., 1986. The SmCrbukk gas/condensate discovery, Halten-banken. In: A.M. Spencer (Editor), Petroleum Geology of the Northern European Margin. Norsk Petroleums forening, Graham and Trotman, London, pp. 299-305. Corfield, S. and Sharp, I.R., 2000. Structural style and stratigraphic architecture of fault propagation folding in extensional settings: a seismic example from the SmCrbukk area, Halten Terrace, Mid-Norway. Basin Res., 12: 329-341. Dalland, A., Worsley, D. and Ofstad, K., 1988. A lithostratigraphic scheme for the Mesozoic and Cenozoic succession offshore midand northern Norway. Norw. Pet. Direct. Bull., 4. Dor6, A.G., 1991. The structural foundation and evolution of Mesozoic seaways between Europe and the Arctic. Palaeogeogr., Palaeoclimatol., Palaeoecol., 87: 441-492. Dor6, A.G., 1992. Synoptic palaeogeography of the Northeast Atlantic Seaway: late Permian to Cretaceous. In: J. Parnell (Editor), Basins on the Atlantic Seaboard: Petroleum Geology, Sedimentology and Basin Evolution. Geol. Soc. London, Spec. Publ., 62: 421-446. Ehrenberg, S.N., Gjerstad, H.M. and Hadler-Jacobsen, E, 1992. SmCrbukk Field: a gas condensate trap in the Haltenbanken province, offshore mid-Norway. In: M.T. Halbouty (Editor), Giant Oil and Gas Fields of the Decade 1978-1988. Am. Assoc. Pet. Geol., Mem., 54: 323-348. Engkilde, M. and Surlyk, E, 1993. The Middle Jurassic VardeklCft Formation of East Greenland analogue for reservoir units of the Norwegian shelf and Northern North Sea. In: J.R. Parker (Editor), Petroleum Geology of Northwest Europe: Proceedings of the 4th Conference. Geological Society, London, pp. 533-542. Gawthorpe, R.L., Sharp, I.R., Underhill, J.R. and Gupta, S., 1997. Linked sequence stratigraphic and structural evolution of propagating normal faults. Geology, 25: 795-798. Gjelberg, J., Dreyer, T., Hr A., Tjelland, T. and Lilleng, T., 1987. Late Triassic to Mid-Jurassic development on the Barents and Mid-Norwegian shelf. In: J. Brooks and K. Glennie (Editors), Petroleum Geology of North West Europe. Graham and Trotman, London, pp. 1105-1129. Graue, E., Helland-Hansen, W., Johnsen, J.R., LCmo, L., N~ttvedt, A., Rcnning, K., Ryseth, A. and Steel, R., 1987. Advance and retreat of the Brent Delta system, Norwegian North Sea. In: J. Brooks and K. Glennie (Editors), Petroleum Geology of North West Europe. Graham and Trotman, London, pp. 915-937. Gupta, S., Underhill, J.R., Sharp, I.R. and Gawthorpe, R.L., 1999. Role of fault interactions in controlling synrift sediment dispersal patterns: Miocene, Abu Alaqa Group, Suez rift, Sinai, Egypt. Basin Res., 11: 167-189. Heinburg, C. and Birkelund, T., 1984. Trace-fossil assemblages and basin evolution of the Vardekl~ft Formation (Middle Jurassic), Central East Greenland. J. Palaeontol., 58: 362-397. Helland-Hansen, W., Ashton, M., L~mo, L. and Steel, R., 1992.

S. CORFIELD I. SHARP K.-O. HAGER T. DREYER J. UNDERHILL

Advance and retreat of the Brent Delta: recent contributions to the depositional model. In: A.C. Morton et al. (Editors), Geology of the Brent Group. Geol. Soc. London, Spec. Publ., 61: 109-127. Hunt, D. and Tucker, M.E., 1992. Stranded parasequences and the forced regressive wedge systems tract: deposition during base level fall. Sediment. Geol., 81: 1-9. Jackson, J.S. and Leeder, M.R., 1994. Drainage systems and the development of normal faults: an example from the Pleasant Valley, Nevada. J. Struct. Geol., 16:1041-1059. Koch, J.-O. and Heum, O.R., 1995. Exploration trends of the Halten Terrace. In: S. Hanslien (Editor), Petroleum Exploration and Exploitation in Norway. Norwegian Petroleum Society (NPF), Special Publication 4. Elsevier, Amsterdam, pp. 235-251. Martinsen, R.S. and Krystinik, L.E, 1998. Depositional Remnants: Products of the Interplay Between Synsedimentary Tectonics and Changes in Relative Sea Level. American Association of Petroleum Geologists, Annual Convention Abstract, Salt Lake City, UT, May, 1998. Partington, M.A., Copestake, E, Mitchener, B.C. and Underhill, J.R., 1993. Biostratigraphic calibration of genetic stratigraphic sequences in the Jurassic of the North Sea and adjacent areas. In: J.R. Parker (Editor), Petroleum Geology of Northwest Europe: Proceedings of the 4th Conference. Geological Society, London, pp. 371-386. Pascoe, R., Hooper, R., Storhaug, K. and Harper, H., 1999. Evolution of extensional styles at the southern termination of the Nordland Ridge, Mid-Norway: a response to variations in coupling above Triassic salt. In: A.J. Fleet and S.A.R. Boldy (Editors), Petroleum Geology of Northwest Europe: Proceedings of the 5th Conference. Geological Society, London, pp. 83-90. Provan, D.M.J., 1992. Draugen Oil Field, Haltenbanken province, offshore Norway. In: M.T. Halbouty (Editor), Giant Oil and Gas Fields of the Decade 1978-1988. Am. Assoc. Pet. Geol., Mem., 54: 371-382. Reynaud, J.-Y., Tessier, B., Proust, J.-N., Dalrymple, R., Marssets, T., Batist, M., Bourillet, J.-E and Lericolais, G., 1999. Eustatic and hydrodynamic controls on the architecture of a deep shelf sand bank (Celtic Sea). Sedimentology, 46: 703-721. Sharp, I.R., Underhill, J.R., Gawthorpe, R. and Gupta, S., 2000. Fault-Propagation Folding in Extensional Settings: Examples of Structural Style and Syn-Rift Sedimentary Response from the Suez Rift, Sinai, Egypt. Bull. Geol. Soc. Am., October issue. Surlyk, F., 1990a. A Jurassic sea-level curve for East Greenland. Palaeogeogr., Palaeoclimatol., Palaeoecol., 78: 71-85. Surlyk, F., 1990b. Timing, style and sedimentary evolution of late Palaeozoic-Mesozoic extensional basins of East Greenfand. In: R.EE Hardman and J. Brooks (Editors), Tectonic Events Responsible for Britain's Oil and Gas Reserves. Geol. Soc. London, Spec. Publ., 55: 107-125. Taylor, A.M. and Gawthorpe, R.L., 1993. In: J.R. Parker (Editor), Petroleum Geology of Northwest Europe. Proceedings of the 4th Conference. Geological Society, London, pp. 317-335. Withjack, M.O., Meisling, K.E. and Russell, L.R., 1989. Forced folding and basement-detached normal faulting in the Haltenbanken area, offshore Norway. In: A.J. Tankard and H.R. Balkwill (Editors), Extensional Tectonics and Stratigraphy of the North Atlantic Margins. Am. Assoc. Pet. Geol., Mem., 46: 567-575.

Department of Earth Sciences, University of Manchester, Oxford Road, Manchester M13 9PL, UK Norsk Hydro Research Centre, P.O. Box 7190, N-5020 Bergen, Norway Norsk Hydro Research Centre, P.O. Box 7190, N-5020 Bergen, Norway Norsk Hydro Research Centre, P.O. Box 7190, N-5020 Bergen, Norway Department of Geology and Geophysics, University of Edinburgh, Edinburgh EH9 3JW, UK

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Middle Jurassic-Lower Cretaceous transgressive-regressive sequences and facies distribution off northern Nordland and Troms, Norway Morten Smelror, Atle Mork, Mai Britt E. Mork, Hermann M. Weiss and Helge Loseth

Application of sequence stratigraphic principles, detailed facies analyses, mineralogy, organic geochemistry and high-resolution biostratigraphic correlations to continuously cored units from the Middle Jurassic-Lower Cretaceous succession of the Troms III and Nordland VII areas provide new information on the sedimentary history offshore northern Norway. Two new lithostratigraphical units, the M~snykan Formation and Rau~te Member, are formally described. In the Troms III area, the lowermost sequence boundary of late Bajocian-early Bathonian age occurs in the upper part of the Sto Formation, while at Andoya and in the Nordland VII area it occurs on top of weathered basement. The Bathonian to Oxfordian succession is dominated by the transgressive systems tracts formed by the proximal sandstones of the M~tsnykan (new unit in the Nordland VII area) and Rams~ (And~ya) formations which distally grade into the Fuglen Formation in the offshore Troms area. The Hekkingen Formation occurs in all areas above the late Oxfordian sequence boundary. In the Nordland VII area, siltstones of the Rau~tte Member (new) represent the most proximal deposits of the Hekkingen Formation. Basinward the Rau~te Member grades into the very organic-rich claystone of the Alge Member. We explain the deposition of the Alge Member by a modified "expanding puddle" model. Overlying the Alge Member a thick (Volgian-Berriasian), highstand systems tract forms the dark claystones and siltstones of the Krill Member in all areas. Sediments reflecting tectonic activity in the earliest Cretaceous are overlain by a regional early Valanginian sequence boundary that marks a clear shift in sedimentary regime. Slow sedimentation rate and high bioclastic input resulted in deposition of a condensed marly succession on shallow shelf areas (Klippfisk and Nybrua formations). The late Hauterivian-earliest Barremian sequence boundary is followed by renewed marine dark claystone deposits (Kolje Formation). The change in depositional regime across the condensed unit is associated with mineralogy changes of the claystones, which suggest increased influence from erosion of feldspar-rich rocks towards the end of this period.

Introduction

Exploration activities during the past 20 years have provided an extensive documentation of the Mesozoic and Cenozoic strata and depositional history offshore mid- and northern Norway, but major less known areas still exist between the main petroleum provinces in the Norwegian and Barents seas (Fig. 1). On And0ya, where the only onshore exposures of Jurassic and Cretaceous sediments are found, investigations of the Jurassic-Lower Cretaceous succession started as early as 1867, and deep wells were drilled in 1960-1970. The first modern account of the Jurassic and Lower Cretaceous strata was through the detailed work of Dalland (1975, 1981). Seismic studies from offshore Vester~len-AndCya (Fig. 1) have documented a thick Mesozoic succession (Brekke and Riis, 1987; Gabrielsen et al., 1990; L0seth and Tveten, 1996). In 1991 IKU Petroleum Research carried out shallow stratigraphic drilling

in the Nordland VI and VII areas with the objective to record Mesozoic and Cenozoic deposits off Vester~len and Lofoten. Two of the coreholes (6814/04-U-01 and -/04-U-02) penetrated the Lower Cretaceous to Middle Jurassic sediments (and terminated in the crystalline basement) in the northern Ribban Basin (Nordland VII area). One year prior to the Nordland VII drillings, IKU carried out a similar shallow drilling project in the Troms III area. Here three cores (7018/05-U-01,-/05-U-02 and -/05-U-06) were drilled through the Lower Cretaceous-Middle Jurassic succession at the northeastern margin of the Harstad Basin. In this paper the sedimentary successions of these cores are described and compared with the time-equivalent succession on AndCya. The excellent recovery of continuous cores through the Middle Jurassic to Lower Cretaceous strata in the Nordland VII and Troms III areas provides a good framework for correlating the sedimentary se-

Sedimentary Environments Offshore Norway - Palaeozoic to Recent edited by O.J. Martinsen and T. Dreyer. NPF Special Publication 10,

pp. 211-232, Published by Elsevier Science B.V., Amsterdam. 9 Norwegian Petroleum Society (NPF), 2001.

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Fig. 1. Location map for cores included in the study offshore Nordland and Troms and onshore AndCya.

quences and interpreting the depositional history in these two areas. High-resolution biostratigraphy and detailed facies analyses of the cored sedimentary successions enable interpretations of transgressiveregressive units and lateral facies changes, and provide new information which we use to define a new depositional model for the Upper Oxfordian-Kimmeridgian organic-rich sediments off Troms.

The Lofoten-Vesterfilen area is located on a thinned continental margin (LCseth and Tveten, 1996) that experienced several phases of late- and post-Caledonian uplift and erosion (LOseth and Stiberg, 1998). The boundary between basement and sediments in the Nordland VII area coincides with an angular unconformity in the offshore seismic sections, and has been correlated to regional uplift in Early to early-Mid-

Middle Jurassic-Lower Cretaceous transgressive-regressive sequences and facies distribution dle Jurassic times (LOseth and Stiberg, 1998; LOseth, 1999). During Bajocian times, faulting and subsidence initiated on AndOya as well as in Vesterfilen and the northern Ribban Basin. Mesozoic sedimentation took place in an active tectonic regime with ongoing fault activity throughout the succession. The Kimmerian rift phase is evidenced by faulting west of the core location in the Nordland VII area in Middle and Late Jurassic time. During the Early Cretaceous (late Barremian) a major rift phase was initiated and the Lofoten Ridge developed as a horst relative to the Ribban and Vestfjorden halfgraben basins (L0seth and Tveten, 1996). These two basins, together with the Vester~len and AndOya areas, experienced rapid subsidence, and the Lofoten Ridge was an important local clastic source area during the Early Cretaceous (LOseth and Stiberg, 1998). The last rift phase initiated in mid-Cretaceous and terminated by continental break-up and sea-floor spreading during latest Paleocene-earliest Eocene time.

Core descriptions: sedimentary units and depositional environments Lithostratigraphically the cored sediments have been correlated with the St0, Fuglen, Hekkingen, Klippfisk and Kolje formations in the western Barents Sea Shelf (Worsley et al., 1988; Smelror et al., 1998). A new formation (M~snykan Formation) and a new member (Raufite Member) are defined in the Nordland VII area (Appendix A). B iostratigraphic datings are summarised in Appendix B. Core data are presented in Figs. 2-4 (for legend see Fig. 3), and a lithostratigraphic correlation to AndOya is suggested in Fig. 5. Troms III area

Shallow cores in the Troms III area (Weiss et al., 1991) penetrated a 830 m thick succession of upper Toarcian to lower Albian sediments in four cores (Fig. 2a). A fifth core penetrated Upper PaleoceneEocene sediments. S t o Formation Approximately 100 m of shallow marine sandstone of Toarcian-Bajocian age referable to the StO Formation were penetrated in core 7018/05-U-06 (Fig. 2b). The lowermost part of the core is clayey and includes lenticular bedded and highly bioturbated siltstones. Both trace fossils (Planolites and Helminthopsis) and palynomorphs suggest a marine depositional environment for this lower part, possibly with upward increasing restriction (?lagoonal). The main part of the formation is a sandstone unit. The lower part is fully bioturbated, very fine- to fine-

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grained and moderately to well sorted. These sediments show a high ichnodiversity (Thalassinoides, Teichichnus, Helminthopsis and Skolithos). The depositional environment is interpreted to be marine on the basis of palynomorphs and trace fossils, with deposition just below the fair-weather wave base in a lower shoreface-offshore transition zone. The main sandstone comprises alternations of well sorted fine-grained sandstone and variably sorted mediumgrained sandstone. B ioturbation is almost absent, except minor Skolithos in the upper part. Primary sedimentary structures such as current ripple lamination, small trough cross-beds and horizontal lamination are common. Fining-upward successions often show clean sand at the tops. Such beds may result from wave action by progradation from lower/middle to upper shoreface. The upper part of the sandstone section is richer in mud and strongly bioturbated by Thalassinoides, Diplocraterion, Planolites and Helminthopsis composing a mature Cruziana ichnofacies assemblage. The increase in bioturbation may reflect marine flooding. This upper succession indicates slow progradation/aggradation within a lower shoreface setting. Fuglen Formation The Fuglen Formation (cores 7018/05-U-06 and -U-02, Fig. 2b), shows a gradual transition from the underlying sands of the Str Formation. This contrasts with the sharp transition from sandstone to mudstone in the Hammerfest and TromsO basins (Worsley et al., 1988). The formation consists of micaand pyrite-bearing sandy mudstone and clayey siltstone, with an overall fining to silty claystone. This may represent a continued deepening trend from the St0 Formation, as shown by increased ichnodiversity with abundant Terebellina and Helminthopsis (core -U-06). Increasing distance to shore upward in the section is supported also by the palynofacies. Ammonites, belemnites and bivalves are common in the finer intervals. The fossils and the thorough bioturbation, including Thalassinoides in core -U-02, suggest deposition in open oxic marine environments. The mudstones of the Fuglen Formation typically contain between 1 and 2 wt% of total organic carbon (TOC), except for bitumen-stained intervals and a thin carbonaceous siltstone layer that have higher TOC contents. Results from bulk pyrolysis (Rock-Eval), pyrolysis-gas chromatography and visual kerogen analysis suggest that the (thermally immature) organic matter is predominantly derived from poorly preserved land plants and has no potential for generation of liquid hydrocarbons (type-III and type-IV kerogen, hydrogen index typically 50-150 mg/g TOC). A slight increase in the hydrogen index

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Fig. 3. Legend to lithological logs and correlation figures.

towards the top of the formation corresponds with a decrease in bioturbation intensity.

Hekkingen Formation The Hekkingen Formation overlies the Fuglen Formation with a sharp contact. The base of the formation is recognised by a sudden increase in gamma readings (Fig. 2b) at the transition from the grey silty mudstones of the Fuglen Formation to the overlying dark claystones of the Hekkingen Formation. The Hekkingen Formation was penetrated in cores 7018/05-U-02 and 7018/05-U-01 (Fig. 2a,b), and the two cores show a stratigraphic overlap of 51 m. The composite thickness of 260 m is comparable to the thicknesses in the type well in the Hammerfest Basin (Worsley et al., 1988). The sediments consist of finely laminated, non-bioturbated, organic-rich dark-grey to black claystones. The lower 50 m of the Hekkingen Formation (Fig. 2b) are characterised by high gamma readings and high TOC contents (6-16 wt%) and are referred to as the Alge Member. Nektonic fossils (dispersed by pelagic larvae) such as ammonites and belemnites occur, and benthic bivalves are sporadically present. The organic matter typically contains 10-20% algal material and 10-20% amorphous matter, the remainder consisting mainly of terrigenous material of variable composition. Yield and composition of the pyrolysates from these thermally immature rocks also indicate a mixed marine-terrigenous organic matter type (predominantly type-II/III kerogen) with rood-

erate potential for generation of liquid hydrocarbons (hydrogen index 141-321 mg/g TOC, typically 200300 mg/g TOC). Generally, the lack of evidence of any current activity and the high organic content indicate deposition in "deep" shelf conditions, possibly in a somewhat restricted basin, for the entire formation. The few occurrences of benthic bivalves and foraminifera suggest sporadic ventilation of the sea bottom, although deposition took place mainly under anoxic to dysoxic conditions (see discussion below). The upper 210 m of the cored Hekkingen Formation are referred to as the Krill Member, and can be distinguished from the Alge Member by significantly lower gamma readings (Fig. 2), as also reported in the type area (Worsley et al., 1988). Total organic carbon content is also significantly lower (2.2-5.8 wt%) than in the Alge Member, while Rock-Eval hydrogen index values are similar and partly somewhat higher (115-399 mg/g TOC, typically 200-350 mg/g). Values of more than 350 mg/g TOC are typically associated with bitumen-stained intervals and therefore not indicative of a higher generation potential or different kerogen quality. As in the Alge Member, the organic matter represents a mixture of marine planktonic and terrigenous material. The fossil content is sparse in the lower part of the Krill Member, but relatively high in the upper 50 m where ammonites and benthic bivalves are present. Carbonate-cemented beds with siderite, ankerite and dolomite (less common) occur in both members and

Fig. 2. (a) Interpreted seismic section showing location of the Lower/Middle Jurassic-Lower Cretaceous cores in the Troms III area (location of the section is plotted in Fig. 1). (b) Lithological logs compiled from four cores in the Troms III area (a) which overlap each other stratigraphically. Note lithological and compositional differences across sequence boundaries shown by various logs. Legend to logs is shown in Fig. 3. Kerogen type colour code is in part based on extracted samples to eliminate the effect of hydrocarbon staining on the hydrogen index.

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are most abundant in the middle part of the Krill Member. Regarding thermal maturity, organic richness and kerogen type, the claystones of the Hekkingen Formation in this area appear to be broadly similar to those in Troms I (well 7120/12-1; Bjor0y et al., 1983). Higher TOC contents in the Alge Member at Troms III and poorer kerogen quality in the upper part of the Krill Member at Troms I seem to be the main differences.

Lower Cretaceous condensed beds (Klippfisk Formation) The uppermost dark claystones of the Hekkingen Formation are sharply overlain by a condensed unit (1.25 m thick) of grey-green claystone overlain by fossiliferous, nodular limestone (10 cm thick) in core 7018/05-U-01 (Fig. 2b). This condensed unit is assigned to the Klippfisk Formation described by Smelror et al. (1998). The boundary between the Hekkingen Formation and the condensed Klippfisk Formation is defined by a sudden decrease in gamma readings. The Lower Cretaceous claystone is bioturbated by Chondrites and contains macrofossils dominated by the thick-shelled bivalve Inoceramus. The nodular limestone includes echinoderm fragments. The condensed nature is demonstrated by presence of palynomorphs of Valanginian age and nannoplankton (coccolithophorids) of Hauterivian age. Missing sections are indicated by the biostratigraphic data at the boundary with the underlying Hekkingen Formation where the upper Berriasian is missing, and on top of the nodular bed.

Kolje Formation The Klippfisk Formation is overlain by 30 m of dark claystones in core 7018/05-U-01, and 213 m in core 7018/07-U-01 ascribed to the Kolje Formation (Fig. 2b). Palynomorphs and foraminifera show a succession of middle Barremian to early Albian age (Appendix B). The formation consists of laminated claystones with thin carbonate beds in the lower and upper parts, separated by turbiditic siltstones and mudstones of middle Aptian to early Albian age. Marine fossils, scarcity of burrows, lack of wave structures, laminae with gradational bases, and lack of any traction-generated structures suggest deposition below storm wave base in a slope or basinal setting. The claystones contain 0.9-5.5 (typically 1.5-3.0) wt% TOC. A significant fraction of this (up to 1.5 wt%) represents bitumen staining, which caused increased hydrogen index values, particularly in the lower half of the formation (Fig. 2b). Analysis of solvent-extracted samples and microscopic examination, however, shows that the organic matter consists

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of largely to predominantly terrigenous, partly oxidised material with no liquid hydrocarbon generation potential (type-IV and some type-III kerogen). The middle part of core 7018/07-U-01 corresponds to two wedge-shaped units recognised on seismic data (Weiss et al., 1991). These intervals include turbiditic beds (sharp or erosional base, often graded) with trace fossils restricted to Chondrites and Planolites. The wedge shape of the middle intervals was taken as evidence that the deposition was syntectonic (Weiss et al., 1991). The organic richness of these deposits is similar to that of the surrounding claystones (0.5-3.5 wt%). However, the kerogen is more strongly dominated by terrigenous material with a considerable content of opaque, "coaly" particles. It has no hydrocarbon generation potential (kerogen type IV with hydrogen index values of <50 mg/g TOC).

Nordland VII area The two cores drilled in the Nordland VII area are located on the same seismic line (IKU-LO4787-90) with a gap of about 80 ms (~ 120 m, Fig. 4a) between the top of core 6814/04-U-01 and the base of core 6814/04-U-02. The cored succession ranges from Middle Jurassic (overlying crystalline basement) to Early Cretaceous (early Barremian) (Hansen et al., 1992).

M~snykan Formation The lower part of core 6814/04-U-01 includes a sandstone unit of Bathonian to middle Callovian age that is time-equivalent to the more distal claystones of the Fuglen Formation described in the Troms III area (Fig. 4b, Fig. 5). This unit is herein defined as the Mfisnykan Formation (see Appendix A). A thin unit (7 m) of muddy debris flow deposits occurs below the sandstones, on top of weathered basement gneiss. The debris flow unit includes conglomeratic beds, mudstones, feldspathic sandstones and kaolinitic weathering products from underlying basement. Coal fragments and coalified wood also occur. Palynofacies shows a dominance of terrestrial woody palynomacerals. Pollen, spores and the fresh-water algae Botryococcus are present, altogether suggesting deposition in a lacustrine or possibly restricted marine environment. The overlying sandstone unit comprises a 114 m thick succession of fine- to coarse-grained, quartzrich (sublitharenitic) sandstone (Fig. 4). Micaceous laminae are abundant in the lower part. Several stacked coarsening-upward successions, separated by thinner fining-upward units, are identified. Sedimentary structures, such as low-angle and trough

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Fig. 4. (a) Interpreted seismic section showing location of shallow stratigraphic cores in the Nordland VII area (location of the section is plotted in Fig. 1). (b) Lithological logs for the two cores in the Nordland VII area which are separated by 80 ms or ~120 m (a). Note lithological and compositional differences across sequence boundaries shown by various logs. Legend to logs is shown in Fig. 3.

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Fig. 5. Lithostratigraphic and sequence stratigraphic correlation between cored successions at Troms III and Nordland VII and Dalland's (1981) lithostratigraphy from AndCya. Legend is shown in Fig. 3.

cross-beds, current ripple lamination, flaser beds associated with coal fragments, and more rarely hummocky beds suggest deposition above wave base. A marine environment is clearly indicated by the types of organic debris present such as marine palynomorphs (dinoflagellates and acritarchs above 156.2 m). Bioturbation is common and includes highenergy forms such as Skolithos and Macaronichnus segregatis (cf. Bromley, 1996) and clear marine indicators such as Ophiomorpha, Palaeophycus and Helminthopsis. The combined sedimentological and biostratigraphic data suggest deposition in shoreface

to foreshore environments. The coarsening-upward sequences may represent a regressive development from an offshore transition zone-lower shoreface to an upper shoreface-foreshore setting, and the erosively based fining-upward sequences may represent channels. A few thin, but very organic-rich (5-27 wt% TOC) mudstones occur in the lower part of the Mgtsnykan Formation. They contain thermally immature to early mature organic matter with a considerable liquid hydrocarbon potential (type-II/III kerogen, hydrogen index 230-409 mg/g TOC).

Middle Jurassic-Lower Cretaceous transgressive-regressive sequences and facies distribution

Hekkingen Formation The Hekkingen Formation was penetrated in both cores and includes the new Raugtte Member and the Alge and Krill members. In the Nordland VII area (core 6814/04-U-01) the muddy siltstones of the formation unconformably overlie the sandstones of the Mgtsnykan Formation (Fig. 4b). At this proximal location the lower 20 m differ significantly from the Hekkingen Formation elsewhere, and we introduce the name Raufite Member for this part (Appendix A). The Raufite Member consists of dark-grey, calcareous and micaceous muddy siltstone and is of Late Oxfordian-Kimmeridgian age. The main difference from the Hekkingen Formation elsewhere is the higher silt content and the extensive bioturbation. A diverse ichnofauna of Terebellina, Helminthopsis, Planolites, Zoophycos and Chondrites suggests deposition in a well oxygenated shelf setting. Palynofacies is dominated by marine forms and amorphous materials. One belemnite was recognised. The Raugtte Member becomes increasingly muddy upwards, defining a gradual transition to the organicrich, dark claystones typical of the overlying Alge Member. The transition is expressed also by increases in gamma ray and neutron density readings at the boundary (Fig. 4b). The TOC content in the Raugtte Member varies between 1.6 and 5.0 wt% (Fig. 4). The organic matter is composed mainly of strongly degraded amorphous material and opaque, "coaly" particles, but also contains some planktonic, possibly algal material. This composition and hydrogen index values of approximately 70-220 mg/g TOC suggest some gas potential but no significant liquid hydrocarbon generation potential for these rocks. The Alge Member (core 6814/04-U-01) consists of organic-rich, finely laminated, pyritic and micaceous, dark- to very dark-grey silty claystone. It is mainly non-bioturbated, except for sporadic occurrences of Terebellina in the lower part and Chondrites and Helminthopsis in the upper part. Ammonites and bivalves are common. These rocks are considerably richer in organic matter (4-9 wt% TOC) than those of the Raugtte Member, and the highest TOC contents were recorded in the lower part of the Alge Member. Hydrogen index values ranging from about 200 to about 330 mg/g TOC suggest the presence of a type-II/III kerogen with a moderate liquid hydrocarbon generation potential. The values are rather constant over the whole interval and show no relation to the TOC content. The organic material is dominated by amorphous matter and contains about 15% algal material. The presence of both tasmanitids and Botryococcus-type algae is con-

Pl 9

sistent with a relatively marginal marine depositional environment. The Hekkingen Formation in core 6814/04-U-02 is assigned to the Krill Member. The Krill Member shows a higher silt content than the Alge Member, and a slight decrease in gamma values in the upper 100 m (Fig. 4b). Thin carbonate beds and nodules are abundant especially in the middle part. Trace and body fossils are lacking in the lower part, but in the upper part Planolites and Helminthopsis are present and in the uppermost part Zoophycos occurs. The content of wood fragments, cephalopod hooks and ammonites increases upward where belemnites, bivalves, coquina beds, Buchia and Dorsoplanites are occasionally present. The claystones of the Krill Member show considerably lower TOC contents (typical values of 2.5-4 wt%) than in the Alge Member in 6814/04-U-01. However, the hydrogen index values are similar to those of the Alge Member claystones (typically between 200 and 350 mg/g TOC) and show a general decrease from the pure claystones at the bottom to the siltier and slightly bioturbated lithologies at the top of the Krill Member (Fig. 4b). The organic matter is dominantly amorphous material, marine algae and dinoflagellate cysts. It has a moderate liquid hydrocarbon generation potential. Layers of less organic-rich rocks with TOC contents of 1.3-2.7 wt% and hydrogen index values of less than 150 mg/g TOC (type-III and type-IV kerogen) occur in the upper part of the Krill Member, suggesting periods of oxic depositional conditions. The depositional environment for the Raugtte Member is interpreted as open marine and well oxygenated, with indications of upward deepening. The Alge and Krill members were deposited in shelf environments below wave base, partly at anoxic conditions with subordinate dysoxic incursions (see later discussion). Dysoxic/oxic incursions may have been more frequent during deposition of the Krill Member and eventually culminated in the establishment of oxic conditions.

Klippfisk Formation The Klippfisk Formation overlies the Hekkingen Formation in core 6814/04-U-02, with a gradual lithological transition (Fig. 4). It consists of a ~30 m thick unit of dark-grey and grey-green, calcareous siltstone with nodules and beds of limestone. This resembles the platform sediments of the Klippfisk and Lyr formations rather than the time-equivalent, more basinal Knurr and Lange formations (Dalland et al., 1988; Worsley et al., 1988; Smelror et al., 1998). The development in the Nordland VII area is quite similar to the development at And~ya where

220 alternating siltstones and carbonate beds form the Leira Member of the Nybrua Formation (Dalland, 1975, 1981; Fig. 5). The lower boundary is defined where calcite cement is abundant and bioturbation is thorough. Fossils are abundant, including bivalves as fragments and coquina beds, Inoceramus shells and prisms and molluscs with sponge borings. Palynofacies of woody debris, marine algae, dinoflagellate cysts exist and in some beds abundant black, woody phytoclasts are present. The marine fauna and palynomorphs, as well as the lack of primary sedimentary structures, imply that deposition occurred below wave base. The formation is totally burrowed by Zoophycos, suggesting deposition in well oxygenated, open starved shelf environments. The formation is distinctly thicker than in the Troms III area, and the biostratigraphy data indicate that the deposition occurred over a long time span (Appendix B). The high proportion of calcareous mud in the sediment may be explained by low clastic input. As micritisation is absent in the limestone beds, we suggest that deposition took place below the photic zone. The core is located at the crest of a rotated fault block that was active during deposition. The formation increases to approximately three times the cored thickness in the graben 6 km to the west of the core location. A possible interpretation for the deposits may be an attenuated succession at the upper part of the fault block that laterally grades into basinal sediments resembling the Knurr Formation. The claystones from the lower part of the Klipprisk Formation contain significant concentrations of organic matter (1.1-3.4 wt% TOC) with a poor hydrocarbon generation potential (type-III and type-IV kerogen). The overlying claystones, marls and siltstones are organic-poor.

Kot]o Formation The calcareous sediments of the Klippfisk Formation are overlain by ~33 m of dark-grey claystones of late Hauterivian-?early Barremian age with a sharp base in core 6814/04-U-02 (Fig. 4). These are correlated with the Kolje Formation in the Hammerfest Basin (Worsley et al., 1988) and the Skja~rmyrbekken Member of the Nybrua Formation and the lower Nordelva Member of the Skarstein Formation on AndOya (Dalland, 1975, 1981). The formation consists here of dark-grey pyritic mudstone with thin carbonate beds of siderite and ankerite in the lower part. Siltstone beds in the upper part may be ripple laminated and carbonate cemented. B ioturbation is only locally observed. The rocks contain appreciable concentrations of organic matter (1.3-4.4 wt% TOC), but have no significant hydrocarbon generation potential (hydrogen

M. Smelror et al.

index 20-90 mg/g TOC). The kerogen contains abundant structured and well preserved woody phytoclasts and common degraded wood. It is richer in spores and leaf cuticles compared to the underlying Hekkingen and Klippfisk formations. Presence of dinoflagellates shows the depositional environment to have been marine, whereas absence of microfossils and only sparse bioturbation indicate a restricted environment. A possible shelf environment of deposition is suggested beneath water masses affected by restricted circulation. The core is located at the crest of the rotated fault block, and the thickness of this formation increases to approximately three times the cored thickness in the graben 6 km to the west of the core location.

Sequence stratigraphic framework and correlation Sequence stratigraphy is used to divide sedimentary successions into genetic units for stratigraphic correlations. It also enables facies mapping and interpretations of the sedimentary history of a basin. The definition of sequences has been intensively discussed in the literature, and different sequence stratigraphic models have been proposed. The depositional sequence was defined and discussed at length by Exxon geologists (Vail et al., 1977; Vail, 1987; Posamentier et al., 1988; Van Wagoner et al., 1990), the genetic stratigraphic sequence was defined by Galloway (1989), and the T-R (transgressive-regressive) cycle was discussed by Johnson et al. (1985) and Embry (1993, 1995). Each of these models has its merits, and the practical application of the models will depend of how easy it is to recognise the key stratigraphic surfaces used to define the sequence boundaries. Often, and especially in areas with few wells, only T-R trends are recognised. Unconformities, transgressive surfaces, and maximum flooding surfaces need to be recognised over wider areas and not only in single wells. Main emphasis is put on recognising and dating the key stratigraphic surfaces used to delineate and subdivide the sequences. The key surfaces are (1) the unconformities and correlative conformities (used as sequence boundaries in the Exxon model), (2) the transgressive surfaces (which together with the correlative unconformities define the boundary of T-R sequences), and (3) the maximum flooding surfaces (or condensed units, which are used as boundaries for the genetic depositional sequences or systems tracts separators in the other models). The present paper pragmatically follows the principles of T-R sequences because the unconformities and correlative transgressive surfaces are seismically, lithologically

Middle Jurassic-Lower Cretaceous transgressive-regressive sequences and facies distribution

and biostratigraphically the most easy recognisable stratigraphic surfaces in the studied successions. Systems tracts are contemporaneous depositional systems active during a specific time interval of sequence development (Posamentier et al., 1988). They are the building blocks used in sequence analysis. In a T - R sequence two systems tracts are recognised: a transgressive systems tract comprising all strata that occur between the basal sequence boundary (i.e. the unconformity and correlative transgressive surface) and the maximum flooding surface; and a regressive systems tract which includes all the strata that occur between the maximum flooding surface and the upper sequence boundary, i.e. the next unconformity and correlative transgressive surface (Embry, 1993). These systems tracts differ somewhat from those defined for a depositional sequence (Posamentier et al., 1988). As outlined by Embry (1993), the transgressive systems tracts are the same for both types of sequences, but the regressive systems tract of a T - R sequence is equivalent to the highstand systems tract of one depositional sequence combined with the lowstand or shelf margin system tract of the overlying depositional sequence (sensu Posamentier

224

et al., 1988). As demonstrated by Embry and Johannessen (1993) and Surlyk (1991), facies analyses should be done separately for each systems tract within each sequence because facies associations commonly differ significantly from one systems tract to the next. This is because factors which strongly influence facies development, such as sediment supply, climate, tectonic influence and subsidence, change across sequence boundaries and maximum flooding surfaces (Embry, 1993). In the present study four major sequence boundaries are recognised (Fig. 6). The lower two are represented by major regional transgressive surfaces in the upper Bajocian and the Upper Oxfordian delineating two 2nd-order sequences (order concept of Embry, 1995). The upper two sequence boundaries underlie and overlie the condensed unit of ValanginianHauterivian age. A subsequent regional transgression in the lower Barremian initiated the deposition of a new transgressive dark shale unit during BarremianAptian time. All these sequence boundaries can be traced throughout the Barents Sea Shelf and also further to other areas of the Arctic (MOrk and Smelror,

2001).

Fig. 6. Sequence stratigraphic correlation between Troms III, Nordland VII and Andoya giving a space-time framework for interpretation of variations in the depositional environments off Nordland and Troms. Legend is shown in Fig. 3. Time-scale from Gradstein et al. (1999).

222

M. Smelror et al.

Upper Bajocian-Lower Oxfordian T-R sequence In the late Bajocian-earliest Bathonian a major transgressive event led to major facies changes along the basins and shelf margins off Norway and Greenland (Surlyk, 1991). In the Troms III area, nearshore (lower shoreface-foreshore) sandy deposits assigned to the St0 Formation were drowned and replaced by offshore muddy siltstones of the Fuglen Formation. The transgressive surface is most likely located in the upper St0 Formation, but can not be picked exactly in the present core (7018/05-U-06). In the Nordland VII area (and And0ya, Fig. 5), the sea transgressed major areas of peneplained crystalline basement. At base a thin unit of continental or restricted marine deposits occurs in structurally controlled embayments. A stratigraphic break (unconformity) between the fluvial and/or restricted marine deposits and the overlying

transgressive shoreface sands in core 6814/04-U-01 is possible, but not documented. On And0ya the oldest Mesozoic deposits (Ramsgt Formation) of sandstone with beds of kaolinitic shale and channel-coal (Hestberget Member) are interpreted as braided river deposits draining into a lagoonal environment (Dalland, 1975, 1981). These are overlain by lagoonal, bituminous shale deposits (Kullgr0fta Member) and fluviatile sandstone (lower Bonteigen Member). As shown in Fig. 6, we have tentatively correlated the lower boundary of the upper Bajocian-lower Oxfordian T - R sequence with the base of the Mesozoic succession on And0ya. A schematic depositional model is presented in Fig. 7. The timing of the initiation of this transgression is not clearly defined, but dinoflagellate assemblages found in the lower part of the transgressive unit in Troms III, core 7018/05-U-06 (Fig. 2) indicate a late Bajocian age. This appears contemporaneous

Schematic depositional models off Nordland-Troms Oxfordian - Kirnmeridgian

The Late Oxfordian transgression resulted in changing sedimentary patterns. In proximal areas the Rauc~te Member (Nordland VII) and Breisand Member (Andoya) represent renewed sediment input while very slow sedimentation took place in the distal areas reaching from the Troms III area throughout the Barents Shelf. Stratified water masses, with input of fresh water, resulted in high organic produ~on and algal blooms. At the mostly anoxic sea bottom very organic-rich sediments were deposited, resulting in one of the best hydrocarbon source rocks of the boreal areas.

Early Oxfordian

In the Early Oxfordian, very slow sedimentation in the Troms III area resulted in a condensed section. In the Nordland VII area the late Callovian - Early Oxfordian was characterised by non-deposit3on and~or erosion.

Bathonian - Callovian

Following the early Bathonian transgression proximal sediments were deposited as the Mc~snykan Formation in Nordland VII and Ramsc~Fm in Andoya. These sediments grade into a more silty development with strongly bioturbated proximal shelf sediments of the Fuglen Formation in the Troms III area and more distal shelf conditions of the same formation in the Hammerfest Basin.

Fig. 7. Schematic depositional models off Nordland and Troms. The drillship indicates drilling localities.

Middle Jurassic-Lower Cretaceous transgressive-regressive sequences and facies distribution with the transgression initiating the deposition of the Fossilbjerget Member within the Jameson Land embayment in East Greenland, where this rapid transgression falls into the late Bajocian Cranocephalites pompeckji Zone (Surlyk, 1991 ). Manum et al. (1991) also suggested a Bajocian age for the oldest Mesozoic deposits on And~ya. The present evidence suggests that the lower boundary of the upper Bajocian-lower Oxfordian unit is diachronous across the western Barents Sea Shelf. The overall transgression continued through the early to mid-Bathonian, and by late Bathonian times fully marine conditions were established over the whole western Barents Sea Shelf area. On Spitsbergen and Kong Karls Land ammonite faunas assignable to the mid-Bathonian Arcticoceras ishmae Zone are found near the base of the Agardhfjellet Formation (L~faldli and Nagy, 1980; Nagy and Basov, 1998). Similar faunas are recovered in the lowermost part of the Fuglen Formation off Troms III. A distinct acme of Dichadogonyaulax sellwoodii near the base of the lower massive sandstone of the Mgtsnykan Formation in the Nordland VII core 6814/04-U-01, and the incoming of Sirmiodinium grossii in the overlying beds, allow a correlation to the Arcticoceras ishmae Zone in cores 7018/05-U-01 and -02 (where the same events are recognised). The period of relative sea-level rise culminated in the mid-Callovian. In Nordland VII, upper Callovian-lower Oxfordian deposits are missing. In the Troms III core 7018/05-U-02 the upper boundary of the upper Bajocian-middle Callovian transgressive unit is represented by a thin carbonate bed immediately overlying thin claystone, giving a small kick on the gamma ray readings. The presence of very small, scattered Chondrites in this claystone may indicate dysaerobic conditions (cf. Bromley and Ekdale, 1984). This unit is interpreted as a maximum flooding surface and dated as mid-Callovian, based on the LAD (last appearance datum) of common Chytroeisphaeridia hyalina a few metres below the gamma spike. The succeeding upper Callovian-lower Oxfordian highstand unit of dark claystone is only about 8 m thick in core 7018/05-U-02. The youngest sediments of this unit are dated by the co-occurrence of Trichodinium scarburghense and Wanaea fimbriata, suggesting a calibration to the early Oxfordian Quenstedoceras mariae or Cordioceras cordatum ammonite zones.

Upper Oxfordian-Berriasian T-R sequence Middle Oxfordian strata are missing over large parts of the western Barents Sea Shelf, including

223

the Troms III and Nordland VII areas. In the Nordland VII core 6814/04-U-01, the hiatus below the upper Oxfordian-lower Kimmeridgian transgressive deposits corresponds to the late Callovian-middle Oxfordian time interval, while in the Troms III core 7018/05-U-02 (Fig. 2) middle Oxfordian strata are missing. The profound regional transgression taking place in the late Oxfordian initiated the deposition of Upper Jurassic dark shales all over the western Barents Sea (Hekkingen Formation) and Svalbard (Agardhfjellet Formation; M~ark et al., 1999). The recovery of ammonites in the lower part of these transgressive shales in the Troms III and Nordland VII areas dates the start of this transgression to the Amoeboceras serratum ammonite Zone. The corresponding sequence boundary is presently not well documented on Ande~ya. The dinoflagellate record in the shallow marine sands from the upper part of the Bonteigen Member (Ramsfi Formation; Vigran and Thusu, 1975) points to an Oxfordian-Kimmeridgian age at this level. An early Kimmeridgian age for the overlying Breisanden Member of the Dragneset Formation is confirmed by the presence of ammonites related to the Rosenia cymadoce Zone (Dalland, 1981). Apparently, the sequence boundary coincides with a disconformity within the Bonteigen Member (i.e. 25 m above the base of the member), probably corresponding to the transition between the fluvial sediments and the overlying beach and shallow marine deposits (as indicated in fig. 4b in Dalland, 1981, as the "Callovian transgression?"). The flooding of the western Barents Sea Shelf reached its maximum in the late Kimmeridgian (sensu gallico) (possibly in the Aulacostephanus eudoxus Zone), when the transgressive deposits of the lower Alge Member (Hekkingen Formation) were followed by dysoxic to oxic shales of the overlying Krill Member. The transition to highstand and regressive deposits is recognised by increasingly oxic depositional conditions and lowering of the total organic carbon content of the shales. Following our interpretation, the time of maximum flooding does not correspond to the time when the most radioactive shales of the Alge Member were deposited. We interpret the time of maximum flooding to coincide with the time when the sedimentation rate was lowest (Fig. 8), i.e. just above the transition to the Krill Member within the latest Kimmeridgian (Aulacostephanus eudoxus ammonite Zone). At the Janusfjellet and Agardhfjellet sections on Spitsbergen a similar shift in depositional rate is noted in latest Kimmeridgian time, i.e. from high rates of more than 30 m/Ma in the latest early Kimmeridgian (Aulacostephanoides mutabilis ammonite Zone) to around 10 m/Ma in the late

224 Age in Ma 144.0

M. Smelror et al. Stage Boreal Berriasian

Ammonite zones L

H. kochi R ma)mci

~~

~m~~

Troms III

.......

Uncored interval (Troms III)

+ + +~

U R tenuicostatus

Volgian

M

L

150.7 Kimmeridgian sensu gallico

154.1 Oxfordian

M L

159.4

U

Callovian

M L

164.4 Bathonian

C. tenuiserratum C. densiplicatum c. cordatum Q. mariae 62. lamoertt R athleta E. coronatum K. jason S. calloviense c. nordenskjoeldi C. apertum c. calyx C. variabile A. cranocephaloides

A. ishmae A. greenlandicus A. arcticus

169.2 Bajocian

E. vogulicus L. groenlandicus C. anguinus E. pseudapertum D. gracilis D. liostracus R communis R rugosa R iatriensis D. primus R pectinatus R hudlestoni wheatleyensis 9 R ele.qans A. autissiodorensis A. eudoxus A. mutabilis R. cymodoce R ba]/lei A. rosenkrantzi A. regulare A. serratum

U

Condensed ~/ deposits and/or I1v hiatus /

9 Krill Mb ,,, , ; I , , ~ P " ..- . . . , , . m., ,J" "" "" q r + § +

" mfs Condensed deposits /

,,*, .. + + + § .%

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I

+.

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Legend: Nordland VII

t errorbarof

biostratigraphic datum

! ?

Condensed deposits and/or hiatus

~'

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~' M&snykan ~' Fm ?

Fuglen Fm ......-.'"" ,nunutuuunuuuumu ~tnnumn uunnumnuUuueunuuuuuuuu

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mu nu in ul uu

C. pompeckji C. indistinctus c. borealis

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6'0;0~'0~'0

' ' ' ' ' ' ' ' ' ' 200 ' 210 ' 220 ' 230 ' 240 ' 250 ' 100 110 120 130 140 150 160 170 180 190

Thickness (m)

Fig. 8. Time-thickness diagrams for the Fuglen, Mfisnykan and Hekkingen formations in the Nordland VII and Troms III areas. Note the condensed deposition during the uppermost Oxfordian-Kimmeridgian (Autissiodorensis-Elegans ammonite zones). Time scale from Gradstein et al. (1999).

Kimmeridgian (Nagy and Basov, 1998). As discussed by Surlyk (1991), continued sea-level rise is difficult to document in shale-dominated successions because of the absence of clear depth indicators. However, Surlyk (1991) suggested, based on shifts in kerogen composition and development of fine lamination, that in East Greenland the maximum flooding surface of the upper Oxfordian-Volgian sequence falls within the late Kimmeridgian Aulacostephanus eudoxus Zone. The Aulacostephanus eudoxus Zone also marks the peak extent of black shale deposition in the British onshore Jurassic (Wignall and Hallam, 1991). On the Barents Sea Shelf, black shale deposition continued during the Volgian-Berriasian. A lowering of relative sea level by mid-Volgian time (dated by Dorsoplanites) is inferred from a slight increase of silt in the claystone deposits of core 6814/04-U-02 (Fig. 4b), and from the incoming of more abundant carbonate beds and bivalves. Signals of lowering sea-level are more difficult to trace at the more distal Troms III site in core 7018/05-U-01 (Fig. 2b). On Spitsbergen, Nagy and Basov (1998) noted a shift from low sedimentation rates in the lower Volgian (10-12 m/Ma) to significantly higher rates (21 m/Ma) in the middle part of the middle Volgian (Dorsoplanites beds).

A new transgression in the late Volgian is recognised by a slight increase in gamma ray readings in the Nordland VII core 6814/04-U-02 (Fig. 4b), and by an increase of the TOC content and hydrogen index in the dark offshore shales in core 7018/05-U-01 (Fig. 2b). The presence of Laugeites cf. groenlandicus a few metres below the facies change in the latter core dates this event to some time after the latest middle Volgian. The recovery of Craspedites below and above the gamma spike in core 6814/04-U-02 suggests that this flooding event terminated in the late Volgian Craspedites taimyrensis ammonite Zone. From the time-thickness diagrams for the Troms III and Nordland VII areas (Fig. 8) it appears that this transgressive episode was associated with a period of slow sedimentation and/or condensation. In Spitsbergen, this time interval is characterised by slow sedimentation rates of around 4 m/Ma (Nagy and Basov, 1998). A new regressive phase, starting in the late Volgian (approximately middle C. taimyrensis Zone) and continuing into the Berriasian (possibly interrupted by a smaller relative sea-level rise in the latest Berriasian), is recognised by decreased gamma ray readings (increased silt content), a decrease in TOC content and increased incoming of kerogen type III

Middle Jurassic-Lower Cretaceous transgressive-regressive sequences and facies distribution

in core 6814/04-U-02 (Fig. 4b). At the Troms III site 7018/05-U-01 more uniform sedimentation prevailed, although a slight decrease in the TOC content and a slightly more terrigenous organic matter type is recognised. The age of these sediments is determined by the presence of Subcraspedites (Borealites) suprasubditus in core 6814/04-U-02, which allows a correlation to the Hectoroceras kochi ammonite Zone. The recovery of specimens with affinity to the bivalve Buchia unschensis in core 7018/05-U-01 also suggests an age around the Jurassic-Cretaceous boundary (and a possible correlation to the Hectoroceras kochi Zone).

Valanginian-Hauterivian condensed unit In both the Troms III and Nordland VII areas the Klippfisk Formation is represented by condensed units which in core 7018/05-U-01 (Fig. 2) consists of lower Valanginian transgressive claystone overlain by Hauterivian limestone. The contemporaneous Valanginian-Hauterivian deposits in the Nordland VII core 6814/04-U-01 consist of calcite-cemented siltstone that shows facies similarities to the Nybrua Formation on AndOya (Fig. 5). The early Valanginian age at the base of the unit in core 6814/04-U-01 is based on the occurrence of Buchia ex. gr. keyserlingiinflata, whereas calcareous nannofossils recovered from the upper part of the unit indicate a correlation to the late Valanginian-early Hauterivian (Jakubowski, 1987). On AndOya the lower boundary of the Valanginian-Hauterivian condensed unit corresponds to the boundary between the Dragneset and Nybrua formations. The Skjaermyrbekken Member of the Nybrua Formation includes red siltstones and mudstones. These may be facies analogues to the contemporaneous R~adryggen Member on East Greenland (Zakharov et al., 1981), and can likewise be interpreted as deposited in an oxidised environment on a submarine high on the crest of a tilted fault block.

(Uppermost Hauterivian ?-) lower Barremian-Aptian transgressive deposits In the Troms III area the condensed ValanginianHauterivian deposits are overlain by Barremian darkgrey shales (Fig. 2b). The transition between the units is sharp and represents an unconformity. A similar sharp transition between the condensed carbonates of the Klippfisk Formation and the overlying dark shales of the Kolje Formation is previously documented over large parts of the western Barents Sea Shelf (Arhus, 1991; Smelror et al., 1998). In the Nordland VII area the condensed Valanginian-Hauterivian unit

225

is also sharply overlain by Barremian transgressive dark-grey shales (Fig. 4b). Following the establishment of Barremian dark shale deposition over most of the western Barents Sea Shelf, deposition of dark open marine shales continued to the Aptian-Albian in both the Troms III and Nordland VII areas. A major detrital mineral change is seen from the Hekkingen Formation to the Kolje Formation in the Troms III area, marked by a distinct increase in feldspar, kaolinite and chlorite and a decrease in quartz. A similar but more gradual change is seen in the Nordland VII area. The increase in feldspar (Figs. 2 and 4) above the unconformity may be explained by a change in the provenance to include new erosion products from non-weathered igneous or metamorphic rocks. It has been suggested elsewhere that the Lofoten Ridge became an important clastic source in the earliest Cretaceous (Le~seth and Tveten, 1996). On And0ya (Figs. 5 and 6) the transgression initiating the Barremian-Aptian deposits is marked by the transition from the upper Skjaermyrbekken Member to the ?upper Hauterivian-lower Barremian Nordelva Member of the Skarstein Formation. Dalland (1981) suggested a continued deepening from the Nordelva Member to the overlying Aptian Hellneset Member that was probably deposited in deep water, in part by turbidity currents. According to L~aseth and Tveten (1996) the base of the Hellneset Member corresponds to a seismic reflection that is located above the cored interval in 6814/04-U-02.

A depositional model for the organic-rich Alge Member (Hekkingen Formation) Several models have been proposed for the extensive deposition of Upper Jurassic organic-rich shales extending from Arctic Canada and the Barents Sea Sheik along the margins/shelves of the present Norwegian-Greenland Sea, and south to the English Channel (e.g. Hallam, 1975; Hallam and Bradshaw, 1979; Tyson et al., 1979; Oschmann, 1988, 1991; Miller, 1990; Wignall and Hallam, 1991; Leith et al., 1993). A summary and discussion of some of these models can be found in Wignall and Hallam (1991). Over the greater part of the Barents Sea Shelf the upper Oxfordian-Boreal Berriasian sequence of organic-rich shale does not exceed a thickness of 100 m and exhibits a marked thinning east towards Novaya Zemlya (Leith et al., 1993). In the Hekkingen Formation of the western Barents Sea Shelf, the more organic-rich radioactive horizons typically occur in the Alge Member of late Oxfordian-Kimmeridgian age (Worsley et al., 1988; Leith et al., 1993). In contrast, on the eastern Barents Sea Shelf the most

226

organic-rich black shales occur in the late Kimmeridgian and Volgian, typically showing a thickness of 20-30 m (Leith et al., 1993). None of the existing models fully accounts for the observed distribution pattern of the Upper Jurassic organic-rich deposits of the Barents Sea Shelf. However, some of the principal mechanisms applied to the "stratified basin model" presented by Tyson et al. (1979) and the "expanding puddle model" introduced by Wignall and Hallam (1991) appear valid for the deposition of the organicrich Alge Member offshore Troms, and are discussed below. The "stratified basin model" (Tyson et al., 1979) explains the accumulation of dark, organic-rich sediments beneath a pycnocline (density interface) which isolates bottom waters in the deeper parts of the basins from oxygenated surface waters. In (landlocked) epeiric seas with poor horizontal circulation, supply of oxygen to the sea bottom depends on vertical circulation. If this is prevented by a density interface (due to thermal or salinity gradients), bottom conditions will become favourable for accumulation and preservation of quantities of organic matter. The influx of abundant terrestrial organic matter and common fresh-water algae (i.e. Botryococcus) in the Alge Member may indicate the presence of a brackishwater surface wedge during deposition. Progradation of shallow marine sandstones on And0ya (Breisanden and TaumhNet members) and shallow shelf deposits in the Nordland VII area (Raufite Member) may support this interpretation. Sedimentation rates determined for the Alge Member in the Troms IlI area are quite high (approx. 13 m/Ma, Fig. 8), and in fact even higher than the rates estimated for the contemporaneous and more silty Raufite Member in the Nordland VII area. This suggests that in addition to relatively high input of fine-grained clastics, an exceptional favourable preservation potential for such organic-rich sediments may have existed off Troms in the late Oxfordian-Kimmeridgian. An effective pycnocline that isolated the bottom waters in the deeper parts of the basin from the oxygenated surface waters may have been a main factor controlling this depositional environment. However, one should note that the macrofauna and microfossil assemblages found in the Alge Member do not indicate reduced salinity in the surface water masses. The macrofauna contains ammonites throughout the Alge Member and belemnites in the upper part. Calcareous microfossils are absent, and the dinoflagellate assemblages are of low diversity (1-8 taxa per sample). However, the Alge Member does contain common to abundant radiolaria at some levels. The presence of abundant radiolaria may suggest an oceanic connection (Dyer and Copestake, 1989), and this apparently

M. Smelror et al.

contradicts the presence of a brackish surface water wedge during deposition of the Alge Member. The Alge Member contains arenaceous foraminifera at several levels, and a few bivalves (Buchia and undetermined forms of Pelecypoda) and echinoderms are also recorded. This suggests dysaerobic depositional conditions, rather than anoxic bottom conditions during deposition of larger parts of the Alge Member. On the other hand, the lack of bioturbation in the Alge Member may suggest that the O2/H2S interface was situated close to the sediment surface and prevented the development of an in-fauna of deep-burrowing and deposit-feeding bivalves and worms. Hallam and Bradshaw (1979) proposed a model involving irregular bottom topography to explain the occurrence of dark organic-rich shales in areas of high subsidence rates. This model, termed the "puddle model", is essentially an extension of the stratified basin model that incorporates a structural/subsidence component. Basically, the puddle model fails to predict the several examples of transgressive organic-rich dark shales that are found resting on shallow-water sediments in basin margin settings (Wignall and Hallam, 1991). To account for this, Wignall and Hallam (1991) proposed the "expanding puddle model" for these transgressive black shales. During early transgression, deep-water conditions will become proportionally more extensive, due to the combined effects of subsidence, a rapid rate of sea-level rise and sediment starvation. In epeiric seas, with mainly vertical and limited horizontal advection of the water masses, large areas will be covered with sufficiently deep water for accumulation of organic-rich dark shales. The expanding puddle model predicts the expansion of deep-water organic-rich shale depositional conditions into basin margins rather than the development of a shallow-water black shale facies (Wignall and Hallam, 1991). Following our interpretations, the Alge Member (and the contemporaneous Raugtte Member) falls within the transgressive systems tract of the upper Oxfordian-Berriasian T-R sequence. In the Troms III area the dark shales of the Alge Member rest with a hiatus (corresponding to the Middle Oxfordian) on silty mudstones of the Fuglen Formation. In the Nordland VII area the Raufite Member unconformably overlies the Mfisnykan Formation. Here the hiatus spans the late Callovian to middle Oxfordian. This pattern agrees with the principles of the "expanding puddle model". This model may also account for the presence of radiolaria in the Alge Member, as these "oceanic" microfaunas may have been brought by oceanic water masses into the shelf areas during the extensive late Oxfordian transgression. However, the model does not fully explain the high accumulation

Middle Jurassic-Lower Cretaceous transgressive-regressive sequences and facies distribution

rate of organic-rich deposits of the Alge Member, as sediment starvation is to be expected during the early transgression. One important aspect when applying the "expanding puddle model" to the offshore northern Nordland and Troms areas, is that the model predicts the presence of older Upper Jurassic dark, organic-rich, shales in the central basinal areas. To our knowledge, middle Oxfordian strata are not documented from offshore northern Norway. Generally, the middle Oxfordian corresponds to a hiatus, or it is too condensed to be recognised. However, the data recovery may be biased since most exploration wells are drilled on structural highs, and the model of middle Oxfordian basinal organic-rich shales is still to be tested. Oschmann (1988) suggested seasonal upwelling in the Upper Jurassic Boreal Ocean. He suggested that a wind-driven surface current flowed from Tethys to the Boreal Ocean and that a bottom current in reverse direction brought deep water south to the English Channel to complete the circulation. This bottom water received no oxygen during its southward flow and became dysoxic and eventually anoxic. Oschmann's (1988) model predicts that dark organic-rich sediments should be present along the entire passage of the bottom waters from the Arctic Ocean to the English Channel, and that the sediments should show evidence of seasonal variations. Wignall and Hallam (1991) argued that during most of the Late Jurassic, organic-rich deposition did not extend over most of England, but rather was restricted to the principal depocentres. They also questioned the evidence for seasonal variations as the presence of very fine parallel lamination is only seen in a few organic-rich shales. Miller (1990) pointed out that, following the principles of Ekman transport (Coriolis force), the proposed southern wind should have generated an eastward-flowing surface current rather than a surface current flowing from the Tethys to the Boreal Ocean. Miller (1990) further argued that the basically Boreal nature of the Kimmeridge Clay fauna contradicts Oschmann's (1988) proposed current systems. As an alternative he suggested a current model with an overall flow from north to south, forming a "Boreal current", with no countercurrent. The Coriolis force would concentrate this current to the western side of the Late Jurassic ocean, i.e. areas off Greenland and Canada. According to Miller's (1990) model, deposition of the Upper Jurassic organic-rich shales apparently occurred in a stratified sea. An upper layer flowed southward from the Boreal Ocean to Tethys, on top of warm saline bottom water. Further, according to Miller (1990), much evidence exists for a negative water balance, but none for unusual organic productivity.

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Miller's (1990) model predicts several of the observations made on the organic-rich shales of the upper part of the Hekkingen Formation (i.e. the VolgianBerriasian Krill Member), but is less plausible for the Alge Member, which was influenced by fresh-water influx and high terrestrial input (algal Botryococcus and liptinitic materials). Whereas the model of Miller (1990) generally requires an overall arid climate, the late OxfordianKimmeridgian Alge Member (and the contemporaneous lower Kimmeridge Clay Formation) was deposited in an overall humid climate with significant terrestrial runoff to the depositional areas. In conclusion, none of the existing models fully predicts the observations made on the organic-rich deposits of the Alge Member. Any model for these upper Oxfordian-Kimmeridgian sediments has to take into account a complex interplay between the extensive sea-level rise, local tectonics related to Kimmerian rifting phases, and an overall humid climate and relatively high continental runoff to the sub-basins along the ocean margins. The "expanding puddle" model of Wignall and Hallam (1991), modified to account for relative high accumulation rates and possibly a brackish surface water wedge during periods of deposition, may most closely mirror the depositional conditions of the organic-rich, dark shales of the Alge Member.

Summary and conclusions Four well defined sequence boundaries (late Bajocian-early Bathonian, late Oxfordian, early Valanginian and close to the Hauterivian-Barremian boundary) have been identified in the studied Middle Jurassic-Lower Cretaceous section off Troms and Nordland. The corresponding systems tracts enabled an improved correlation of proximal and distal facies developments within the studied successions. (1) The lowermost regional sequence boundary of late Bajocian-early Bathonian age in the Troms III area occurs in the upper part of the St0 Formation. At Ande~ya and in the Nordland VII area it occurs on top of weathered basement that may be overlain by a very thin continental succession. (2) The Bathonian to Oxfordian succession is dominated by the transgressive systems tracts formed by the proximal sandstones of the Mfisnykan (Nordland VII) and Ramsgt (AndOya) formations which distally grade into the Fuglen Formation in the area offshore Troms. (3) The Hekkingen Formation occurs in all areas above the late Oxfordian sequence boundary. Siltstones of the Raufite (Nordland VII) and Breisand (AndOya) members occur in the most landward areas.

228 Stratigraphically and regionally these units grade into the very organic-rich claystone of the Alge Member, and form together the transgressive systems tract. A thick regressive systems tract forms the organic-rich claystones and siltstones of the Krill Member in all areas. (4) The observations made on the Alge Member may be explained by deposition according to an "expanding puddle" model (Wignall and Hallam, 1991), modified to account for relatively high sediment accumulation rates and for the possible existing wedge of brackish surface water during periods of deposition. High organic productivity in the surface-near water accompanied by anoxic sea bottom conditions resulted in very organic-rich sediments. The organic richness is higher at Troms III (6-16 wt% T o e ) than at Nordland VII (4-9 wt%), but the liquid hydrocarbon generation potential is moderate in both areas. (5) Sediments reflecting tectonic activity in the earliest Cretaceous are succeeded by a regional sequence boundary in the early Valanginian that marks a clear shift in the sedimentary regime. Slow sedimentation rate and high bioclastic input resulted in deposition of a condensed marly succession on shallow shelf and platform areas (Klippfisk and Nybrua formations), while calcareous shales were deposited in basinal areas further north (Knurr Formation). (6) The late Hauterivian-earliest Barremian sequence boundary is followed by renewed marine claystone deposits. The change in depositional regime across the condensed unit is associated with mineralogical changes of the claystones, which suggest increased influence from erosion of feldspar-rich rocks towards the end of this period. The compositional change can be a response to a new rift phase that was initiated during the late Barremian-Aptian. (7) Application of sequence stratigraphic principles supported by detailed biostratigraphic observations enables the recognition of systems tracts where proximal and distal sedimentary units form stratigraphic hydrocarbon plays. Such models may be applicable in wider areas and not only along the Norwegian coast.

Acknowledgements The present study is based on cores obtained during the "IKU Shallow Drilling Troms III 1990" and "IKU Shallow Drilling Nordland VI and VII" programs. We would like to thank a number of colleagues who contributed with field- and laboratory work to these projects. Richard G. Bromley (Denmark) studied the trace fossils, Andrzej Wierzbowski (Poland) and Natasha Shulgina (Russia) studied the

M. Smelror et al.

macrofossils for the early industry reports. Simon R.A. Kelly (England) commented on an early version of the manuscript. The paper was reviewed by Ashton E Embry and Rob Gawthorpe. They are all acknowledged for their contributions. Berit Fossum helped us to improve the figures.

Appendix A Definition of the M~snykan Formation and the Rau~te Member A formal lithostratigraphic scheme for the Mesozoic and Cenozoic succession offshore northern Norway was published by Dalland et al. (1988) and Worsley et al. (1988), replacing the informal units used by Olaussen et al. (1984), Berglund et al. (1986) and Gjelberg et al. (1987). In the Hammerfest Basin the transition from the sands of the StO Formation to the mudstones of the overlying Fuglen Formation is sharp. In the Troms I area the Fuglen Formation is of late Callovian to Oxfordian age (Worsley et al., 1988). The shallow stratigraphic cores 7018/05-U-02 and 7018/05-U-06 in the Troms III area (Fig. 2) document a more extensive development of the Fuglen Formation, with a wider time range of deposition and more gradual transition from the underlying StO Formation. Shallow coreholes drilled in the Nordland VI and VII areas offshore Lofoten and Vesterfilen (Hansen et al., 1992) have penetrated a new unit of sandstones younger than the Str Formation and time-equivalent with the Fuglen Formation in the Troms III and I areas. This sandstone unit is here formally introduced as the Mfisnykan Formation. As for the Fuglen Formation and the overlying Hekkingen Formation (Oxfordian-Berriasian), the formation name for the new Mg~snykan Formation is taken from a lighthouse located on an island adjacent to the type locality of the formation. Corehole 6814/04-U-01 (Fig. 4) in the Nordland VII area documents an undescribed unit of upper Oxfordian (to lowermost Kimmeridgian) muddy siltstones. This unit is here formally introduced as a new member of the Hekkingen Formation, i.e. the Raufite Member.

M~snykan Formation (new) (M~snykanformasjonen) Name. Named after the Mfisnykan lighthouse located on a small island at the western tip of LangCya in BO, Vestergtlen (68~ and 14~ particularly known for its large colonies of sea birds.

Middle Jurassic-Lower Cretaceous transgressive-regressive sequences and facies distribution

Type section. Corehole 6814/04-U-01, 49.17-163.40 m (fully cored); coordinates 68~ and 14~ 11'08.9"E (Fig. 4b). Reference section. Corehole 6710/03-U-01, 132.40148.65 m (fully cored); coordinates 67~ 10~ Thickness. The formation is 126.23 m thick in the type corehole (Fig. 4b) and 16.25 m in the reference corehole. Lithology. The formation consists of fine to coarse sublitharenitic sandstones, which generally are mineralogically mature and well sorted. Minor kaolinitic mudstones may occur in the lower part of the formation. Basal stratotype. In the type section the lower boundary is defined by an abrupt change from underlying crystalline basement gneiss (subjected to intensive kaolinite weathering at the top) to basal polymict conglomerates and arkosic sandstones. This boundary is recognised by an abrupt change from block to serrate sonic log pattern. Characteristics of the upper boundary. In the type section the upper boundary occurs by an abrupt change from sandstones of the Mfisnykan Formation to overlying Upper Jurassic dark-grey siltstones. In the reference section the upper boundary is recognised by a change from sandstones to overlying Lower Cretaceous claystones. On the electrical logs these boundaries are recognised by marked increases in gamma log intensity and in velocity readings. In both the type and reference coreholes a significant stratigraphic break separates the Mfisnykan Formation from the overlying strata. Distribution. The formation is present in the Nordland VI and VII areas. Age. Bathonian-(?middle) Callovian. Depositional environment. Shallow marine (shoreface to foreshore-backshore). Remarks. The Mfisnykan Formation is time-equivalent to the lower part of the Fuglen Formation in the Troms III area. The formation is also partly timeequivalent to the Ramsfi Formation on AndOya, but due to uncertainties in the dating of the Ramsfi Formation sandstones, the relationship is still obscure. The formation is also time-equivalent to the Melke Formation as defined in the Tra~nabanken area (Dalland et al., 1988). Worsley et al. (1988) defined the Fuglen Formation to include upper Callovian to Oxfordian mudstones deposited on local block structures in the Hammerfest Basin. It was deposited during a highstand with ongoing tectonic movements. The present information from Nordland VII and Troms III shows

229

that closer to the palaeo-coastline thick sequences of mudstones/shales assignable to the Fuglen Formation also accumulated during Bathonian and early Callovian time. By the mid-Callovian, the offshore Troms areas were characterised by sediment starvation or non-deposition, and possibly local minor erosion on structural highs. During the syntectonic period, both tectonics and sea-level fluctuations controlled the sedimentation pattern. By the late Callovian-early Oxfordian, previous areas/blocks of non-deposition were flooded and became sites of slow sedimentation.

Definition of the Rau~te Member (included in the Hekkingen Formation) (Rau~teleddet) Name. Named after the copepod Calanus finmarchicus (in Norwegian called raufite) which is common in the North Atlantic, along the Norwegian coast, and in Atlantic water masses in the Barents Sea. Type section. Corehole 6814/04-U-01, 29.80-49.17 m (fully cored); coordinates 68~ and 14~ (Fig. 4b). Thickness. The member is 19.37 m thick in the type corehole. Lithology. The Raufite Member consists of dark-grey, calcareous and micaceous muddy siltstone, which becomes muddier upwards. Basal boundary. The basal boundary is sharp and marked by an unconformity where the upper Oxfordian dark grey siltstone of the Raufite Member overlies ?middle Callovian sandstones of the Mfisnykan Formation. Characteristics of the upper boundary. The upper boundary is recognised by a change from dark-grey muddy siltstones to overlying dark-grey, finely laminated, pyritic and slightly calcareous, claystones (assigned to the Krill Member). On the gamma log this boundary is recognised by a change to higher gamma readings. Distribution. The Raufite Member is presently only documented from the Nordland VII area. Age. Late Oxfordian-early Kimmeridgian. Depositional environment. Shelf, below wave base in a well oxygenated environment. Remarks. The Raufite Member is time-equivalent to the Alge Member offshore Troms. However, the Rau~te Member represents, in contrast to the Alge Member, deposition in an oxic environment, possibly relatively closer to the palaeo-coast than the Alge Member. The Rau~tte Member is included within the Hekkingen Formation and represents the most coarse-grained and proximal facies of this formation.

M. Smelror et al.

230

Appendix B Selected biostratigraphic events of the Nordland VII and Troms III cores:

Core 6814/04-U-01: Aulacostephanus sp. Amoeboceras kochi Amoeboceras cf. subkitchini Amoeboceras ex gr. bauhini-bayi Amoeboceras ernesti Amoeboceras cf. serratum Evansia dalei Mendicodinium groenlandicum Lithodinia callomonii Valvaeodinium theresae FAD Chlamydophorella ectotabulata Acme Ctenidodinium sellwoodii Core 6814/04-U-02: LAD Buchia cf. keyserlingi LAD Buchia ex gr. keyserlingi-inflata Bojarkia cf. mesezhnikov Subcraspedites suprasubditus LAD Praetollia sp. ?Craspedites sp. LAD Virgatosphinctes cf. exoticus Craspedites sp. LAD Buchia unschensis FAD Buchia unschensis LAD Dorsoplanites spp. FAD Buchia fischeriana ?Dorsoplanites sp. Dorsoplanites cf. gracilis ?Pectinatites sp. FAD Pectinatites sp.

Ammonite Ammonite Ammonite Ammonite Ammonite Ammonite Dinoflagellate Dinoflagellate Dinoflagellate Dinoflagellate Dinoflagellate Dinoflagellate

10.26 12.20 24.70 28.25 28.60 47.89 68.72 68.72 76.70 76.70 108.55 156.20

Bivalve Bivalve Ammonite Ammonite Ammonite Ammonite Ammonite Ammonite Bivalve Bivalve Ammonite Bivalve Ammonite Ammonite Bivalve Bivalve

50.06 62.04 71.04 80.64 85.33 97.63 103.10 106.57 119.79 123.95 136.21 140.85 151.16 154.97 168.90 189.90

Dinoflagellate Dinoflagellate Bivalve Foraminifer Ammonite Ammonite Bivalve Ammonite Ammonite Ammonite Ammonite Foraminifer Ammonite Foraminifer Bivalve Bivalve Dinoflagellate Ammonite Ammonite Ammonite

Dinoflagellate Foraminifer Ammonite Ammonite Bivalve Dinoflagellate Dinoflagellate Dinoflagellate Dinoflagellate Dinoflagellate Dinoflagellate

160.24 158.30 145.80 142.25 141.90 122.61 101.78 101.78 93.80 93.80 61.95 14.30

m m m m m m m m m m m m

Kimmeridgian Kimmeridgian Kimmeridgian Early Kimmeridgian Early Kimmeridgian Late Oxfordian L. Bath.-Callovian L. Bath. or younger L. Bath.-Callovian L. Bath.-Callovian Late Bathonian Late Bathonian

m m m m m m m m m m m m m m m m

421.69 409.71 400.71 391.11 386.42 374.12 368.65 365.18 351.96 347.80 335.54 330.90 320.59 316.78 302.85 281.85

m m m m m m m m m m m m m m m m

Valanginian Early Valanginian Berriasian Early Berriasian Early Berriasian Late Volgian Late Volgian Late Volgian Late Volgian Late Volgian Middle Volgian Middle Volgian Middle Volgian Middle Volgian Early Volgian Early Volgian

48.50 48.50 55.12 58.16 84.03 82.55 83.60 85.86 102.11 107.80 149.56 167.30 185.05 243.63 244.10 246.15 256.38 262.98 264.50 270.96

m m m m m m m m m m m m m m m m m m m m

579.67 579.67 573.05 570.01 544.14 545.62 544.57 542.31 526.06 520.37 478.61 460.87 443.12 384.54 384.07 382.02 371.79 365.19 363.67 357.21

m m m m m m m m m m m m m m m m m m m m

E. Valanginian E. Valanginian Early Berriasian Early Berriasian Early Berriasian Early Berriasian Early Berriasian Early Berriasian Middle Volgian Middle Volgian Middle Volgian Middle Volgian Middle Volgian Early Volgian Early Volgian Late Kimmeridgian E. Kimmeridgian E. Kimmeridgian E. Kimmeridgian Late Oxfordian

20.12 48.49 74.93 99.27 102.19 107.89 107.89 111.47 115.80 119.65 146.45

m

404.05 375.68 349.24 324.90 321.98 316.28 316.28 312.70 308.37 304.52 277.72

Core 7018/05-U-01: FAD Nelchinopsis kostromiensis FAD Oligosphaeridium complex LAD Buchia unschensis FAD Gaudryina gerkei LAD Borealites sp. LAD Surites sp.

Buchia okensis FAD Surites sp. LAD Laugeites cf. groenlandicus FAD Laugeites cf. groenlandicus Dorsoplanitinae FAD Bulbobaculites vetustus ?Dorsoplanitinae

Evolutinella emaljanzevi LAD Buchia tenuistriata FAD Buchia tenuistriata FAD Oligosphaeridium patulum Amoeboceras ex gr. kitchini ?Rasenia sp. indet Amoeboceras (?Prionodoceras) sp.

Core 7018/05-U-02: LAD Cribroperidinium longicorne FAD Evolutinella emeljanzevi LAD Amoeboceras ex. gr. alternans FAD Amoeboceras ex. gr. alternans FAD Buchia concentrica Acme Trichodinium scarburghense LAD Wanaea fimbriata FAD Wanaea fimbriata LAD Chytroeisphaeridia hyalina LAD Lacrymodinium warenii FAD Endoscrinium galeritum

m m m m m m m m m m

Early Volgian Late Kimmeridgian Late Oxfordian Late Oxfordian Late Oxfordian Early Oxfordian Early Oxfordian Early Oxfordian Middle Callovian Early Callovian Early Callovian

Middle Jurassic-Lower Cretaceous transgressive-regressive sequences and facies distribution

231

Appendix B (continued) Core 6814/04-U-01: FAD Gonyaulacysta pectinigera Arcticoceras sp. (cf. k~lowi) FAD Evansia cerebraloides LAD Riyadhella sibirica Pseudoceras sp. (cf. mudum) Acme Ctenidodinium sellwoodii FAD Cassiculosphaeridia dictydia Arcticoceras sp. (ex. gr. ishmae) A rctocephalites? sp.

Dinoflagellate Ammonite Dinoflagellate Foraminifer Ammonite Dinoflagellate Dinoflagellate Ammonite Ammonite

151.59 158.96 174.43 183.60 202.06 202.90 220.36 260.30 261.03

m m m m m m m m m

272.58 265.21 249.74 240.57 222.11 221.27 206.81 163.87 163.14

m m m m m m m m m

Late Bathonian Late Bathonian Late Bathonian Late Bathonian Late Bathonian Late Bathonian Late Bathonian "Mid" Bathonian "Mid" Bathonian

Dinoflagellate Ammonite Ammonite Ammonite Dinoflagellate Dinoflagellate

52.74 70.78 78.93 83.60 109.91 121.23

m m m m m m

192.43 174.39 166.24 161.57 135.26 123.94

m m m m m m

Late Bathonian Late Bathonian "Mid" Bathonian "Mid" Bathonian Late Bajocian Early Bajocian

Core 7018/05-U-06" FAD Cassiculosphaeridia dictydia ?Arcticoceras sp. LAD Arctocephalites (?) sp. FAD Arctocephalites (?) sp. FAD Ctenidodinium sellwoodii LAD Phallocysta eumekes

(FAD -- first appearance datum/oldest occurrence; LAD = last appearance datum/youngest occurrence; acme = abundance peak; first value = depth in the core; second value = height above base in Figs. 2 and 4)

References Arhus, N., 1991. The transition from deposition of condensed carbonates to dark claystone in the Lower Cretaceous succession of the southwestern Barents Sea. Nor. Geol. Tidsskr., 71: 259-263. Berglund, L.T., Augustson, J., Faerseth, R., Gjelberg, J. and Ramberg-Moe, H., 1986. The evolution of the Hammerfest Basin. In: A.M. Spencer et al. (Editors), Habitat of Hydrocarbons on the Norwegian Continental Shelf. Norwegian Petroleum Society. Graham and Trotman, London, pp. 319-338. Bjomy, M., Bue, B. and Elvsborg, A., 1983. Organic geochemical analysis of the first two wells in the Troms I area (Barents Sea). In: M. Bjomy et al. (Editors), Advances in Organic Geochemistry 1981. Wiley, Chechester, pp. 16-27. Brekke, H. and Riis, F., 1987. Tectonics and basin evolution of the Norwegian Shelf between 62~ and 72~ Nor. Geol. Tidsskr., 67: 295-322. Bromley, R.G., 1996. Trace Fossils: Biology, Taphonomy and Application (2nd ed.). Chapman and Hall, London, 361 pp. Bromley, R.G. and Ekdale, A.A., 1984. Chondrites: a trace fossil indicator of anoxia in sediments. Science, 224: 872-874. Dalland, A., 1975. The Mesozoic rocks of And0ya, northern Norway. Nor. Geol. Unders., 316: 271-287. Dalland, A., 1981. Mesozoic sedimentary succession at AndOya, northern Norway and relation to structural development of the North Atlantic area. In: J.W. Kerr and A.J. Ferguson (Editors), Geology of the North Atlantic Boderlands. Can. Soc. Pet. Geol., Mem., 7: 563-584. Dalland, A., Worsley, D. and Ofstad, K., 1988. A lithostratigraphic scheme for the Mesozoic and Cenozoic succession offshore midand northern Norway. Norwegian Petroleum Directorate Bulletin, 4, 65 pp. Dyer, R. and Copestake, E, 1989. A review of Late Jurassic to earliest Cretaceous radiolaria and their biostratigraphic potential in the North Sea. In: D.J. Batten and M.C. Keen (Editors), Northwest European Micropalaeontology and Palynology. Ellis Harwood, Chichester, pp. 214-235. Embry, A., 1993. Transgressive-regressive (T-R) sequence analysis of the Jurassic succession of the Sverdrup Basin, Canadian Arctic Archipelago. Can. J. Earth Sci., 30: 301-320. Embry, A., 1995. Sequence boundaries and sequence hierarchies: problems and proposals. In: R.J. Steel, V. Felt, E.E Johannessen and C. Mathieu (Editors), Sequence Stratigraphy of the Northwest

European Margin. Norwegian Petroleum Society (NPF), Special Publication 5. Elsevier, Amsterdam, pp. 1-11. Embry, A.F. and Johannessen, E.E, 1993. T-R sequence stratigraphy, facies analysis and reservoir distribution in the uppermost Triassic-Lower Jurassic succession, western Sverdrup Basin, Arctic Canada. In: T.O. Vorren, E. Bergsager, O.A. Dahl-Stamnes, E. Holter, B. Johansen, E. Lie and T.B. Lurid (Editors), Arctic Geology and Petroleum Potential. Norwegian Petroleum Society (NPF), Special Publication 2. Elsevier, Amsterdam, pp. 121-146. Gabrielsen, R., Faerseth, R.B., Jensen, L.N., Kalheim, J.E. and Riis, F., 1990. Structural elements of the Norwegian continental shelf. Part I: The Barents Sea region. Norwegian Petroleum Directorate Bulletin, 6, 33 pp. Galloway, W.E., 1989. Genetic stratigraphic sequences in basin analysis: architecture and genesis of flooding-surface bounded depositional units. Am. Assoc. Pet. Geol. Bull., 73: 125-142. Gjelberg, J., Dreyer, T., H0ie, A., Tjelland, T. and Lilleng, T., 1987. Late Jurassic to Mid-Jurassic sandbody development on the Barents and Mid-Norwegian shelf. In: J. Brooks and K. Glennie (Editors), Petroleum Geology of Northwest Europe. Graham and Trotman, London, pp. 1105-1129. Gradstein, F.M., Agterberg, F.P., Ogg, J.G., Hardenbol, J. and B~ickstr6m, S., 1999. On the Cretaceous time scale. Neues Jahrb. Geol. Pal~iontol. Abh., 212: 3-14. Hallam, A., 1975. Jurassic Environments. Cambridge University Press, Cambridge, 269 pp. Hallam, A. and Bradshaw, M.J., 1979. Bituminous shales and oolitic ironstones as indicators of transgressions and regressions. J. Geol. Soc., London, 136:157-164. Hansen, J.W., Bakke, S., Fanavoll, S., L0seth, H., Mork, A., MOrk, M.B.E., Rise, L., Smelror, M., Verdenius, J.G., Vigran, J.O. and Weiss, H.M., 1992. Shallow drilling Nordland VI and VII 1991. Main report. IKU Report 23.1594.00/02/92 390 pp. Restricted. Jakubowski, M., 1987. A proposed Lower Cretaceous calcareous nannofossil zonation scheme for the Moray Firth area of the North Sea. Abh. Geol. Bundesanst., 39: 99-119. Johnson, J.G., Klapper, G. and Sandberg, C.A., 1985. Devonian eustatic fluctuations in Euramerica. Geol. Soc. Am., Bull., 96: 567-587. Leith, T.L., Weiss, H.M., MOrk, A., Arhus, N., Elvebakk, G., Embry, A.F., Brooks, EW., Stewart, K.R., Pchelina, T.M., Bro, E.G., Verba, M.L., Danyushevskaya, A. and Borisov, A.V., 1993. Mesozoic hydrocarbon source-rocks of the Arctic region. In: T.O.

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232 Vorren, E. Bergsager, O.A. Dahl-Stamnes, E. Holter, B. Johansen, E. Lie and T.B. Lund (Editors), Arctic Geology and Petroleum Potential. Norwegian Petroleum Society (NPF), Special Publication 2. Elsevier, Amsterdam, pp. 1-25. Lcfaldli, M. and Nagy, J., 1980. Foraminiferal stratigraphy of Jurassic deposits on KongsCya, Svalbard. Nor. Polarinst. Skr., 172: 6395. LCseth, H., 1999. Paleogeographical evolution of the Lofoten and Vesterfilen onshore and offshore area. Geonytt, Abstract Volume 1999, pp. 71-72. L0seth, H. and Tveten, E., 1996. Post-Caledonian structural evolution of the Lofoten and Vester~len offshore and onshore areas. Nor. Geol. Tidsskr., 76: 215-230. LCseth, H. and Stiberg, J.E, 1998. Apatite and zircon fission track data from Lofoten and Vesterhlen. EAGE Conference and Technical Exhibition, Leipzig. Extended Abstract Vol. 2, 3-02. Manum, S.B., Bose, M.N. and Vigran, J.O., 1991. The Jurassic flora of AndCya. Rev. Palaeobot. Palynol., 68: 233-256. Miller, R.G., 1990. A paleogeographic approach to the Kimmeridge Clay Formation. In: A.Y. Huc (Editor), Deposition of Organic Facies. Am. Assoc. Pet. Geol., Stud. Geol., 30:13-26. MCrk, A., Dallmann, W., Dypvik, H., Johannessen, E.E, Larssen, G.B., Nagy, J., NCttvedt, A., Olaussen, S., Pchelina, T.M. and Worsley, D., 1999. Mesozoic lithostratigraphy. In: W.K. Dallmann (Editor), Lithostratigraphic Lexicon of Svalbard. Upper Palaeozoic to Quaternary bedrock: Review and recommendation for nomenclature use. Norsk Polarinstitutt, Tromsr pp. 127-214. MCrk, A. and Smelror, M., 2001. Correlation and non-correlation of high order circum-Arctic Mesozoic sequences. Proceedings from International Conference on Arctic Margins 1998, Celle. Polarforschung, Bremerhaven (in press). Nagy, J. and Basov, V.A., 1998. Revised foraminiferal stratigraphy of Bathonian to Ryazanian deposits in Spitsbergen. Micropaleontology, 44:217-255. Olaussen, S., Dalland, A., Gloppen, T.G. and Johannessen, E., 1984. Depositional environment and diagenesis of Jurassic reservoir sandstones in the eastern part of the Troms I area. In: A.M. Spencer, S.O. Johnsen, A. MOrk, E. Nys~ether, E Songstad and A. Spinnanger (Editors), Petroleum Geology of the North European Margin. Graham and Trotman, London, pp. 61-80. Oschmann, W., 1988. Kimmeridge Clay sedimentation - - a new cyclic model. Palaeogeogr., Palaeoclimatol., Palaeoecol., 65:217251. Oschmann, W., 1991. Distribution, dynamics and palaeoecology of Kimmeridgian (Upper Jurassic) shelf anoxia in western Europe. In: R.V. Tyson and T.H. Pearson (Editors), Modern and Ancient Continental Shelf Anoxia. Geol. Soc. London, Spec. Publ., 58: 381-395. Posamentier, H.W., Jervey, M.T. and Vail, ER., 1988. Eustatic

M. SMELROR A. MORK M.B.E. MORK H.M. WEISS H. LOSETH

controls on clastic deposition I - - conceptual framework. In: C.K. Wilgus, B.S. Hastings, C.G.St.C. Kendall, H.W. Posamentier, C.A. Ross and J.C. Van Wagoner (Editors), Sea-Level Changes: An Integrated Approach. Soc. Econ. Paleontol. Mineral., Spec. Publ., 42: 109-124. Smelror, M., MCrk, A., Monteil, E., Rutledge, D. and Leereveld, H., 1998. The Klippfisk Formation - - a new lithostratigraphic unit of Lower Cretaceous platform carbonates on the Western Barents Shelf. Polar Res., 17:181-202. Surlyk, E, 1991. Sequence stratigraphy of the Jurassic-lowermost Cretaceous of East Greenland. Am. Assoc. Pet. Geol. Bull., 75: 1468-1488. Tyson, R.V., Wilson, R.C. and Downie, C., 1979. A stratified water column environmental model for the type Kimmeridge Clay. Nature, 277: 377-380. Vail, ER., Mitchum, R.M., Jr., Todd, R.G., Widmier, J.M., Thompson, S., III, Sangree, J.B., Bubb, J.N. and Hatelid, W.G., 1977. Seismic stratigraphy and global changes of sea level. In: C.E. Payton (Editor), Seismic S t r a t i g r a p h y - Applications to Hydrocarbon Exploration. Am. Assoc. Pet. Geol., Mem., 26:49-212. Vail, ER., 1987. Seismic stratigraphy interpretation using sequence stratigraphy. In: A.W. Bally (Editor), Atlas of Seismic Stratigraphy. Am. Assoc. Pet. Geol., Stud. Geol., 27: 1-10. Van Wagoner, J.C., Mitchum, R.M., Campion, K.M. and Rahmanian, V.D., 1990. Siliciclastic sequence stratigraphy in well logs, cores, and outcrops. Am. Assoc. Pet. Geol., Methods Explor. Ser., 7: 1-55. Vigran, J.O. and Thusu, B., 1975. Illustration and distribution of Jurassic palynomorphs of Norway. NTNF's Continental Shelf Project, Publication no. 65, 56 pp. Weiss, H.M., Bakke, S., Bugge, T., Hansen, J.W., LCseth, H., MCrk, A., MCrk, M.B.E., Ritter, U., Smelror, M., Verdenius, J.G., Vigran, J.O. and ,~rhus, N., 1991. Shallow drilling Troms III 1990. Main report. IKU Report 23.1433.00/00-/02/91, 346 pp., 15 apps. Restricted. Wignall, EB. and Hallam, A., 1991. Biofacies, stratigraphic distribution and depositional models of British onshore Jurassic black shales. In: R.V. Tyson and T.H. Pearson (Editors), Modern and Ancient Continental Shelf Anoxia. Geol. Soc. London, Spec. Publ., 58: 291-309. Worsley, D., Johansen, R. and Kristensen, S.E., 1988. The Mesozoic and Cenozoic succession of Tromsr In: A. Dalland, D. Worsley and K. Ofstad (Editors), A lithostratigraphic scheme for the Mesozoic and Cenozoic succession offshore mid- and northern Norway. Norw. Pet. Direct. Bull., 4: 42-65. Zakharov, V.A., Surlyk, F. and Dalland, A., 1981. Upper JurassicLower Cretaceous Buchia from AndCya, northern Norway. Nor. Geol. Tidsskr., 61: 261-269.

Geological Survey of Norway, N- 7491 Trondheim, Norway SINTEF Petroleum Research, N-7465 Trondheim, Norway SINTEF Petroleum Research, N-7465 Trondheim, Norway SINTEF Petroleum Research, N-7465 Trondheim, Norway Statoil Research Centre, N-7005 Trondheim, Norway

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Outcrop studies of tidal sandstones for reservoir

characterization (Lower Cretaceous Vectis Formation, Isle of Wight, southern England) Shuji Yoshida, Matthew D. Jackson, Howard D. Johnson, Ann H. Muggeridge and Allard W. Martinius

The Vectis Formation (Lower Cretaceous) in the Isle of Wight, southern England serves as an excellent outcrop analogue for tidal sandstone reservoirs. It consists of two juxtaposed marginal-marine depositional systems separated by an erosional, regionally extensive sequence boundary. The lower system is characterized by a muddy succession which overlies fluvial strata and lacks tidal indicators. This is interpreted as a low-energy, broad lagoonal or embayment complex. The upper, sandstone-dominated system is characterized by a wide range of heterolithic deposits, with abundant evidence of tidal processes, as indicated by sedimentary structures and palaeocurrent patterns. It is interpreted as a laterally migrating tidal bar and channel complex within the inner part of a mixed-energy estuary (sensu Dalrymple et al., 1992). The Vectis Formation sandstones contain a highly variable array of sedimentary structures and heterogeneities, ranging from the regional scale (tens of kilometres wide, tens to hundreds of metres in thickness) to individual laminae scale (mm in thickness). These have been described and quantified in order to provide a basis for evaluating their impact on fluid flow and hydrocarbon recovery. The large-scale heterogeneities and facies variations are explained within a sequence stratigraphic framework and reflect the transgressive infill of a mixed-energy estuarine complex. The intermediate- to small-scale heterogeneities ( m m - m in thickness, representing the scale of individual bedforms and laminae) reflect autocyclic tidal processes and have been examined in detail by (1) constructing a detailed 2-D facies model directly from the outcrop (measuring 400 m x 6 m), (2) photographing and digitizing small, representative areas (ca. 1 m x 1 m) of 2-D outcrops, and (3) using serial sectioning techniques to reconstruct the 3-D geometry of tidal sedimentary structures (wavy-, lenticular- and flaser-bedding) directly from large (ca. 60 c m x 60 cm x 20 cm) rock specimens. These reconstructed rock models have been used to investigate the validity of using core-plug and well-log derived permeability data to represent heterolithic facies in tidal sandstone reservoir models.

Introduction

Tidal sandstone reservoirs host major hydrocarbon accumulations in the North Sea, including the Lower Jurassic Cook Formation of the Gullfaks field (e.g. Marjanac and Steel, 1997) and the Middle Jurassic Beryl Formation in the Beryl and Bruce fields (e.g. Robertson, 1997; Dixon et al., 1997). They are also important hydrocarbon producers elsewhere, including the well-documented Eocene Misoa Formation in the Lagunillas and LL-652 fields, Venezuela (e.g. Maguregui and Tyler, 1991; Mellere et al., 1999). However, the characterization and modelling of tidal sandstone reservoirs is difficult. This is because tidal sandstones contain a complex array of sedimentary heterogeneities at length scales ranging from millimetres to kilometres, and our understanding of the effect that these heterogeneities have on the flow of fluids during recovery is poor. To properly represent the heterogeneities in a reser-

voir model, the reservoir must be subdivided into a logical hierarchy of units, which can be modelled individually, and the smaller-scale heterogeneities upscaled so that their effect on flow can be represented in larger-scale simulation models (e.g. Haldorsen, 1986; Pickup et al., 1994; Peijs-van Hilten et al., 1998). For modelling and simulation purposes, we have subdivided the reservoir/outcrop heterogeneities within a tidal estuary system into a threefold hierarchy (Fig. 1). (1) Large-scale heterogeneity refers to the size, shape and distribution of tidal bars and channels within a tidal estuarine incised valley system (ca. tens to hundreds of metres in thickness). (2) Intermediate-scale heterogeneity refers to the facies distribution within individual bars and channels (ca. tens of centimetres to tens of metres in thickness). (3) Small-scale heterogeneity is concerned with single bedform architectures (ca. millimetres to tens of centimetres in thickness) within individual facies. Ideally, quantitative information on the geometry

Sedimentary Environments Offshore Norway - Palaeozoic to Recent edited by O.J. Martinsen and T. Dreyer. NPF Special Publication 10, pp. 233-257, Published by Elsevier Science B.V., Amsterdam. 9 Norwegian Petroleum Society (NPF), 2001.

234

S. Yoshida et al.

Fig. 1. Hierarchy of heterogeneities within an estuarine depositional system (tide-dominated or the inner part of mixed-energy estuaries of Dalrymple et al., 1992). Our large-, intermediate- and small-scale heterogeneities used to characterize the Lower Cretaceous Vectis Formation in the Isle of Wight and the Lower Jurassic Tilje Formation, offshore mid-Norway, are genetically correlatable to the macro-, meso- and micro-scale heterogeneities proposed by Peijs-van Hilten et al. (1998) for the Lower Cretaceous Glauconite Channel in Alberta, Canada. The time and physical scales associated with these three-fold hierarchical terms, however, varies from one example to another.

and dimensions of tidal sandstone reservoir heterogeneities (e.g. distribution of facies and mud drapes) should be derived from outcrop analogues at each of these length scales. With the rapid development of sequence strafi' graphy, numerous outcrop analogue studies have been conducted over the last decade, with the aim of characterizing subsurface tidal sandstone reservoirs. These include the Eocene Roda Sandstone (e.g. Nio and Yang, 1991) and the Amettla Formation (e.g. Dreyer and Ffilt, 1993; Dreyer, 1994) in the Spanish Pyrenees, and the Cretaceous Western Interior Basin in the USA, notably the Tocito Sandstone in New Mexico (e.g. Jennette and Jones, 1995), the Sego-Upper Castlegate Sandstones in the Book Cliffs area, Utah (e.g. Van Wagoner et al., 1991; Yoshida et al., 1996; Yoshida, 2000; Willis, 2000), the Fall River Formation in Wyoming (e.g. Willis, 1997) and the Frontier Formation in Wyoming (e.g. Willis et al., 1999). The majority of these studies focus on the scale of stratigraphic sequences and systems tracts (our large-scale heterogeneities; Fig. 1), whereas a limited number focus on the intermediate scale (e.g. White and Barton, 1997; Wonham et al., 1998; Willis et al., 1999; Jackson and Muggeridge, 1999; Fig. 1). However, few studies focus on small-scale heterogeneities (ca. centimetres

to tens of centimetres in thickness; the scale of flaser, wavy, and lenticular bedding; Fig. 1), yet small-scale structures may have a significant impact on fluid flow and hydrocarbon recovery (e.g. Kortekaas, 1985; Lasseter et al., 1986; Ringrose et al., 1993; Jones et al., 1995; Jackson et al., 1999). Consequently, quantitative data is required to characterize tidal sandstone reservoirs at the small- to intermediate-scales, within the large-scale framework. We have used the outcrop of the Vectis Formation (Lower Cretaceous) in the Isle of Wight, southern England (Fig. 2), to characterize the small- to large-scale sedimentary heterogeneities found within a succession of tidal deposits. Furthermore, we have quantified the heterogeneities at each scale, with emphasis on the hitherto poorly documented small- and intermediate-scale heterogeneities. The Vectis Formation is a broadly transgressive succession of muddominated coastal and nearshore deposits, which sharply overlies the fluvial Wessex Formation and is succeeded by the shallow-marine Lower Greensand Group (Ruffell, 1988; Wach and Ruffell, 1991; Fig. 2). Within this succession there is a prominent and well-exposed sandstone body (the Barnes High Sandstone Member; Fig. 2), which contains a wide array of features indicative of deposition in an

Outcrop studies of tidal sandstones for reservoir characterization

935

Fig. 2. Geology of the Isle of Wight, southern England. The Vectis Formation is the youngest formation within the Wealden Group, and records the first transgressive event within the non-marine Wealden Group. The Vectis Formation is transgressively overlain by the open-shallow-marine Lower Greensand. The outcrops of the Barnes High Sandstone Member of the Vectis Formation occur in three areas on the island: Compton Bay, Sandown Bay and Brighstone Bay (Barnes High-Shepherd's Chine). The outcrop widths of the member in the first two localities are less than 100 m due to high structural tilt, whereas in the Brighstone Bay area the member is continuously exposed for 1.7 km. The Bembridge-St. Valery Line served as the northern margin of the E-W-trending extensional rift basin (Portland-South Wight Basin) during the Jurassic-Cretaceous, and then changed to a compressional fault line with some strike-slip deformation due to Tertiary inversion related to the Alpine orogeny (Hawkes et al., 1998). The Wealden (Biscay cycle) sequence follows the terminology of Hawkes et al. (1998).

inshore tide-dominated (meso- to low macro-tidal) environment. The member is characterized by several vertically stacked, upward-coarsening units, and is interpreted as a composite tidal sand bar formed within an incised estuarine valley. The facies and facies associations (sensu Reading, 1986) are analogous to those found in many subsurface tidal sandstone reservoirs, partly because it contains an almost complete range of heterogeneity types. Furthermore, the facies and facies associations are comparable to those forming important hydrocarbon-bearing reservoirs in the Lower Jurassic Tilje Formation (outcrop facies calibrated to Tilje 1C interval; see Martinius et al., 2001), which hosts several oil and gas accumulations in the Haltenbanken area, offshore mid-Norway, including the Heidrun and SmCrbukk (part of Asgard) fields (e.g. Whitley, 1992; Olsen et al., 1999).

This paper focuses on the outcrop studies of the Barnes High Sandstone Member. The main objectives of this paper are as follows: (1) to place the Vectis Formation within a sequence stratigraphic framework in order to better understand the regional facies distributions and large-scale heterogeneities; (2) to demonstrate the detailed level of field/laboratory work, modelling and visualization required to properly characterize intermediate- and small-scale heterogeneities in tidal sandstone reservoirs; (3) to provide quantitative information on these sedimentological heterogeneities for subsurface facies studies; and (4) to compare conventional upscaling approaches applied to small-scale heterogeneities and to evaluate their problems.

236

Sequence stratigraphy of the Vectis Formation The Isle of Wight is located along the northern margin of the Portland-South Wight Basin, as defined by Hawkes et al. (1998) (Fig. 3), which is a subbasin of the more extensive Wessex Basin of southern England (Stoneley, 1982). Both are E-W-trending extensional rift basins, which originated during the Permian and were re-activated during Jurassic-Early Cretaceous times (e.g. Chadwick et al., 1983). The Vectis Formation comprises the uppermost part of the Lower Cretaceous Wealden Group (Fig. 2), which unconformably overlies Jurassic strata. Based on detailed regional studies integrating seismic, outcrop and well data of southern England, Hawkes et al. (1998) interpreted the Lower Cretaceous succession, including the Wealden Group and the overlying Lower Greensand Group, as a syn-rift 'megasequence' (termed the Wealden sequence or Biscay cycle) bounded by regional unconformities (Fig. 2). The lower part of the mega-sequence is the Wessex Formation, which is characterized by a reddish-brown mudstone-dominated succession containing freshwater fossils (including well-preserved dinosaur bones and footprints), abundant pedogenic features and subordinate fluvial channel sandstone bodies. It is interpreted as a succession of floodplain deposits dissected by meandering rivers flowing to the east, parallel to the rift trend (e.g. Stewart, 1981, 1983; Insole and Hutt, 1994). The Vectis Formation conformably overlies the Wessex Formation, and is characterized by pale grey mudstones containing fresh- and brackishwater fossils. It is divided into three members: in ascending order, Cowleaze Chine, Barnes High Sand-

S. Yoshida et al.

stone and Shepherd's Chine Members (Daley and Stewart, 1979; Stewart, 1981; Fig. 2). The Cowleaze Chine Member records the first brackish transgression within the mega-sequence exposed on the Isle of Wight (e.g. Insole et al., 1997). The Vectis Formation is erosionally overlain by the fully marine Lower Greensand Group (e.g. Wach and Ruffell, 1991; Ruffell and Wack, 1998). Due to the absence or scarcity of marine fauna, dating of the Wealden Group is poorly constrained (Barremian to early Aptian). The Vectis Formation was interpreted by earlier workers as a conformable, mudstone-dominated succession deposited within a broad lagoonal complex, containing a possible "open marine" interval in terms of salinity (sensu Pritchard, 1967; cf. Dalrymple et al., 1992) in the upper part of the Shepherd's Chine Member (e.g. Stewart et al., 1991). The Barnes High Sandstone Member within the formation was interpreted as a lagoonal delta complex debouching from the mouths of meandering river systems, as represented by the Wessex Formation. These deltas prograded eastwards into low-salinity brackish water, as represented by the muddy Cowleaze Chine and Shepherd's Chine members (e.g. Wach, 1991; Wach and Ruffell, 1991). The apparent absence of a basinward facies shift and coarse lag deposits (e.g. Van Wagoner et al., 1990; Zaitlin et al., 1994) at the base of the upward-coarsening Barnes High Sandstone Member in the Brighstone Bay area may, at first, appear as a conformable succession as suggested by the earlier workers. However, based on detailed outcrop mapping, utilizing sedimentary facies and architectural element analysis methodology, we postulate that the lower part of the Vectis Formation consists of at least two juxtaposed marginal-marine de-

Fig. 3. Gross isopach map for the Wealden sequence in Wessex Basin (taken from Hawkes et al., 1998 with minor modifications), indicating the control of active extensional faulting during Wealden deposition.

Outcrop studies of tidal sandstones for reservoir characterization

237

Fig. 4. Sequence stratigraphy and facies architecture of the Vectis Formation in the Isle of Wight. TS = transgressive surface; SB/TS = composite surface of a sequence boundary and a transgressive surface, following the definitions of these terms by Allen and Posamentier (1993). The overall thickness of the Barnes High Sandstone Member increases to the north within the southern half of the island, both to the northwest along the coast as shown in this figure, and to the Sandown Bay area in the northeast (e.g. Wach, 1991). The Barnes High Sandstone is thinnest at Barnes High, where it erosively overlies the palaeosol at the top of the Cowleaze Chine Member interpreted as an intra-valley high. See Fig. 2 for the locations of vertical sections, and Table 1 for the explanation of lithofacies assemblages.

positional systems separated by a regionally extensive erosional hiatus surface (Fig. 4). The main features of the two separate depositional systems are as follows. (1) The lower marginal-marine depositional system (Cowleaze Chine Member) is characterized by a muddy succession containing fresh- and brackishwater fossils, with few lateral facies variations. Sedimentary structures indicative of tidal process (flaser, wavy, and lenticular bedding, double mud drapes, etc.) are virtually absent. It sharply overlies fluvial strata of the Wessex Formation and is interpreted as a broad transgressive lagoonal or estuarine embayment complex in a low-energy, micro-tidal setting. (2) The upper marginal-marine depositional system (Barnes High Sandstone Member) is a sandstonedominated unit characterized by a broadly coarsening-upward succession. This passes from muddominated to sand-dominated heterolithic facies with abundant evidence of tidal processes (bi-directional cross-bedding, wavy, lenticular, and flaser bedding, double mud drapes, etc.). It also contains a mixture of fresh- to brackish-water faunas. The system is interpreted to have been formed in the inner part of a meso- to low macro-tidal, mixed-energy estuary (sensu Dalrymple et al., 1992), where high-energy tidal processes dominate.

We suggest that the erosion surface at the base of the upper system represents a composite surface of a sequence boundary and a transgressive surface (sensu Allen and Posamentier, 1993; Fig. 4), with the overlying Barnes High Sandstone Member and most of the overlying Shepherd's Chine Member representing a transgressive incised valley system because of the following five reasons. (1) There is a marked facies contrast across the erosion surface (Fig. 4), with bioturbated mudstones below and finely laminated, lenticular bedded mudstones above (Fig. 5). In addition, there is an abundance of tidal indicators above the surface, but none below (Table 1). (2) The erosion surface is regionally extensive. This surface represents an erosive mud-mud contact in the Brighstone Bay area, where it is physically traceable for the whole outcrop length for 1.7 km from Barnes High to Shepherd's Chine (Figs. 4 and 6). It also strongly correlates to the erosion surfaces at the base of the Barnes High Sandstone Member in Compton Bay (mud-sand contact; Fig. 7) and Sandown Bay (mud-mud contact; reported by Wach and Ruffell, 1991), thereby covering the whole island. (3) The erosion surface displays a regionally discordant geometry. This is reflected in the regional

TABLE 1 Lithofacies assemblage of the Vectis and Wessex Formations in the Isle of Wight, southern England

bO Cad

r~ ~,~~

Outcrop studies of tidal sandstones for reservoir characterization

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Fig. 5. Erosional contact between the Cowleaze Chine Member and the overlying Barnes High Sandstone Member (marked as SB/TB). The Cowleaze Chine Member (Assemblage V8) is intensively bioturbated (see white dots representing horizontal burrows) immediately below the erosion surface. This surface is laterally traceable for the whole outcrop length (1.7 km) in the area (Fig. 4), and is interpreted as a composite surface of a sequence boundary and a transgressive surface (labelled as SB/TS; sensu Allen and Posamentier, 1993). Distal tidal bar facies (Assemblages V5 and V6) occur above the surface, and constitutes the lowermost part of the Barnes High Sandstone Member in the area. Gentle dipping on the outcrop is due to tectonic tilt, and the depositional dip is sub-horizontal.

thickness distribution of the Barnes High Sandstone Member, which exhibits a lenticular geometry; the member is thinnest in the southern part of the Isle of Wight (ca. 6 m), it thickens northwards towards the E-W-trending axis of the island (Fig. 2) and along the coast both to the northwest (Fig. 4) and southeast (Wach, 1991), and it is thickest (ca. 15 m) at the west and east ends of the island. The Barnes High Sandstone Member also wedges out to the north and northwest and is absent on the nearby UK mainland. This suggests a valley origin to the erosion surface within the overall E-W-trending thalweg extending from Compton Bay to Sandown Bay. (4) In Compton Bay (Fig. 2), where the Barnes High Sandstone Member is the thickest, the erosion surface is floored by abundant conglomerates. These are interpreted as lag deposits associated with the fluvial incision during lowstand and close to the valley thalweg (Assemblage V 1 on Figs. 4 and 7 and Table 1). (5) At Barnes High (Fig. 2), where the member is thinner, a 10-15 cm thick zone of tightly cemented sideritic mudstone occurs immediately below the erosion surface (Fig. 4). This zone grades downwards into a 1-2 m thick zone of mudstones containing abundant vertical fractures, some of which are filled with coarse-grained sands. This interval is interpreted as a palaeosol formed on an intra-valley high.

Facies characteristics of the Barnes High Sandstone Member The Vectis Formation has been divided into twelve lithofacies assemblages (Table 1). This paper focuses on the Barnes High Sandstone Member (the upper marginal-marine depositional system), which exhibits considerable lateral facies variations.

Tidal bar complex In the central and eastern parts of the island (Barnes High to Sandown Bay; Fig. 2), the basal part of the upper tidal system is characterized by vertically stacked coarsening-upward units containing abundant tidal sedimentary structures (Assemblages V2 to V6 on Fig. 4 and Table 1). These units are particularly well-exposed between Barnes High and Cowleaze Chine (Fig. 6) and are interpreted as being part of a composite tidal bar formed within a broad incised estuarine valley (Fig. 4). Mud-dominated facies within the Barnes High Sandstone Member (Assemblage V9; Table 1)contain lenticular bedding made of very fine sands and silts, which occur either between the tidal bars (Figs. 6 and 8) or in the basal part of the coarsening-upward succession immediately above the sequence boundary (Fig. 8). This facies locally thickens in a landward direction (to the northwest), and is interpreted as subtidal interbar mud deposits (Figs. 6 and 8).

0

Fig. 6. Barnes High Sandstone Member exposed at the mouth of Cowleaze Chine, Isle of Wight (outcrop AJ on the 2-D outcrop facies panel shown in Fig. 8). Here it comprises three superimposed upward-coarsening units above an extensive erosional surface interpreted as a sequence boundary. It is interpreted as a composite tidal sand bar. The gentle dip on the outcrop (to the right) is due to tectonic tilt. Abbreviations used for key surfaces: S B / T S = composite surface of a sequence boundary and a transgressive surface (sensu Allen and Posamentier, 1993); BFS = bay flooding surface; DS -- discontinuous scour surfaces of shallow wave/tidal reworking; BAS -- bar abandonment surface; BB -- bar boundary (other than BAS, BFS or SB/TS); LAS = large-scale accretion surfaces. Abbreviations used for sedimentary structures: L = lenticular bedding; W = wavy bedding; F = flaser bedding; TXB = trough cross-bedding; CR = current ripples; WR = wave ripples; MR = mega-ripples. See Table 1 for details of lithofacies assemblages. See Fig. 8 for detailed facies distribution. The outcrop belt in the Brighstone Bay area (Barnes High-Shepherd's Chine; Fig. 2) is virtually 2-D and trends N W - S E .

t...,

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Fig. 7. Outcrop of the Barnes High Sandstone Member in Compton Bay, Isle of Wight. Abbreviations for bounding surfaces: S B / T S -= composite surface of a sequence boundary and a transgressive surface (sensu Allen and Posamentier, 1993)" B F S - bay flooding surface; S B ? = candidate sequence boundary; C H B -- channel base (corresponding to fifth-order surface of Miall, 1988); C H T - - channel top (corresponding to fourth-order surface of Miall, 1988). Abbreviations for architectural elements (in ascending order)" L - C H -- lower channel; M - C H = middle channel; U - C H -- upper channel; H e - C H -- heterolithic channel. See Fig. 4 for vertical section. V1, V8, VIO, V l l = lithofacies assemblage ID (see Table 1 for details). Strata in this locality are tilted at 22 ~ towards the north-northeast, and the BFS represents the stratigraphic horizontal. Outcrop trends NW-SE.

The five sandstone-dominated facies within the Barnes High Sandstone Member (Assemblages V 3 V6) occur in a continuum, both laterally and vertically, within each upward-coarsening and minor upward-fining succession (Figs. 5, 6 and 8 and Table 1). They are interpreted as representing different parts of tidal sand bars (proximal to distal in relation to the bar centre; Figs. 4 and 8). These facies (Assemblages V3-V6), as well as muddy facies (Assemblage V9), all lack evidence of subaerial exposure such as coals, desiccation cracks, pedogenic features or evaporites, and are interpreted to have been formed in a subtidal setting. The uppermost part of the Barnes High Sandstone forms a thin, laterally extensive cemented sand-

stone sheet (Assemblage V2; Figs. 6 and 8 and Table 1). This unit contains evidence of shallowwater processes, including discontinuous scour surfaces (Figs. 6 and 8), current- and wave-ripples, 1020 cm high subaqueous dunes, and fresh- and brackish-water shells and pebbles that are concentrated in the troughs between the ripple crests (Fig. 9). The most abundant trace fossils include Diplocraterion, Beaconites (Fig. 9D), and Ophiomorpha. Significantly, dinosaur tracks are locally preserved on top of the member (Radley et al., 1998). This cemented sandstone is very similar to the emergent part of modern tidal sand bars, including those in The Wash of eastern England (Ke et al., 1996; Fig. 10) and the Oosterscheldt Estuary in The Netherlands (Nio

~.,.~

Fig. 8 . 2 - D outcrop facies panel of the Barnes High Sandstone Member (composite tidal sand bar in the inner, tide-dominated part of a mixed-energy estuary) in the Cowleaze Chine-Shepherd's Chine area, Isle of Wight. Vertical exaggeration 6 times. See Table 1 for details of lithofacies.

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Fig. 9. Tidal bar top facies (Assemblage V2) of the Barnes High Sandstone Member. (A) Mega ripples on top of the Barnes High Sandstone Member. Scale 15 cm. (B) Symmetrical ripples on top of the Barnes High Ss Mb. Scale 10 cm. (C) Plan view of bioclasts (Unio and rare Filosina, indicating fresh-water to slightly blackish-water conditions; e.g. Stewart et al., 1991; Radley and Barker, 1998) on top of the Barnes High Ss Mb. Scale 10 cm. (D) Plan view of Beaconites (non-marine indicator; e.g. Goldring and Pollard, 1995), on top of the Barnes High Ss Mb. Scale 15 cm.

et al., 1980), and probably represents a combination of both wave and tidal reworking in a mixed-energy estuary setting. More specifically it is interpreted as a partly condensed, winnowed lag deposit in which the morphology of migrating ripples and dunes have been preserved; it may have formed in either an intertidal (e.g. bar top) to partly shallow subtidal environment. The sharp contact between this sandstone facies and the overlying muddy facies (Shepherd's Chine Member; Assemblage V 10) is interpreted as a bay flooding surface (Figs. 4, 6, 8 and 9).

(3.5 m high) tangential cross-beds (Middle Channel on Fig. 7). These sandy channel fills are interpreted as alternate bars or side bars (e.g. Collinson and Thompson, 1989) within a straight or low-sinuosity, high-energy, tidally influenced channel at the head of a mixed-energy estuary (Dalrymple et al., 1992). A distinct, but thinner (1.2 m) channel occurs at the top of the member (Heterolithic Channel on Fig. 7). This channel is filled with heteroliths made of mud, silt and fine-grained sands with low-angle lateral accretion surfaces, and is interpreted as a meandering tidal channel fill at/near the head of the estuary.

Tidal-fluvial channel complex Around 7 km to the northwest (Compton Bay; Fig. 2), the system is characterized by medium- to very coarse-grained sandstones with blocky to finingupward grain-size profiles (Assemblage V 1 on Figs. 4 and 7; Table 1). Wood fragments and disseminated shell debris are important secondary constituents, while bioturbation is rare to absent. Palaeocurrent patterns are bidirectional but with dominant sediment transport to the south to southeast, which is in a broadly offshore direction. The sandy channel fills in the lower to middle part of the member lack clear lateral accretion surfaces (Lower and Upper Channels on Fig. 7), and may comprise single sets of large

Bayfill deposits Both the tidal channel and tidal bar complex are sharply overlain by a muddy succession containing fresh- and brackish-water fossils, with evidence of wave and storm action (Assemblage V 10; Figs. 4 and 6-8 and Table 1). This 50 m thick succession displays an overall upward-increase in salinity (e.g. Stewart et al., 1991). This is interpreted as a low-energy bayfill succession deposited within a broad transgressive coastal embayment, which in part may correspond to the low-energy 'central basin' of a mixed-energy estuary of Dalrymple et al. (1992).

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Fig. 10. Modern intertidal deposits (The Wash embayment in eastern England). (A) 3-D dunes (mega-ripples) and current/ladder ripples on top of the tidal sand bar exposed during low tide. (B) Lag concentrations (shells and bioclasts) in the troughs between current ripples. Compare the similarity with Fig. 9.

Palaeogeography of the Barnes High Sandstone Member The Barnes High Sandstone Member is interpreted as an amalgamation of bar and channel complexes formed within the inner part of a broad incised, mainly subtidal, estuarine valley. It displays features that may correspond to a 'bayhead delta' zone of a mixed-energy estuary (Dalrymple et al., 1992), or an 'estuary delta' at the head of modem, mixedenergy estuaries, such as the Gironde estuary, where tidal processes dominate (e.g. Allen, 1991; Allen and Posamentier, 1993; Fig. 11). The inner part of the 'bayhead delta' zone, as inferred for the Compton Bay area, was dominated by fluvial and ebb-dominated tidal processes, which resulted in extensive channelization. The outer, more distal, tide-dominated part of the bayhead-delta zone, as observed in the Cowleaze Chine area, contains a series of coalesced, elongate, NW-SE-oriented tidal sand bars, which are inferred to have been aligned broadly parallel to the estuary axis. The exact location of the estuary margin is uncertain. However, the Barnes High Sandstone Member thickens to the north (from 6 m to 15 m), within the southern half of the island, and pinches out between the island and onshore England, some 10-20 km to the west.and northwest (e.g. Wach, 1991). The Lower Cretaceous succession, which is up to 1000 m thick in the southern part of the island, is completely absent to the north of a ma-

jor E-W-trending, basin margin fault system between Sandown Bay and Compton Bay (Fig. 2). This trend is part of the regional Purbeck-Isle of Wight disturbance zone (Underhill and Paterson, 1998), which connects with the regionally extensive Bembridge-St. Valery Line (Hawkes et al., 1998; Figs. 2, 3 and 11). This fault system underwent active extension during the Early Cretaceous as a result of rifting associated with the opening of the North Atlantic (e.g. Stoneley, 1982; Chadwick et al., 1983; Hawkes et al., 1998; Ruffell, 1998). Hence, this estuary was probably partly fault bounded, as was the underlying fluvial Wessex Formation (Selley and Stoneley, 1987). The upward-coarsening units of the member are unlikely to have been formed as prograding/regressive sandbodies into an open-marine environment because of (1) abundant fresh- and brackish-water fauna indicative of very low salinity (Stewart et al., 1991; Radley and Barker, 1998), (2) absence or scarcity of fully marine indicators in terms of sedimentary structures, lithology, fossils and palaeocurrents (Table 1), (3) presence of an angular, regional erosion surface at the base of the member implying its estuarine valley origin rather than prograding tidal delta origin (Dalrymple, 1999), and (4) the gradual upward increase in salinities towards the top of the Shepherd's Chine Member, indicating an overall transgressive history (Stewart et al., 1991; Radley and Barker, 1998). The 2-D geometry of the outcrop belt of the Barnes High Sandstone Member in the Brighstone Bay area,

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Fig. 11. (A) Palaeogeographic reconstruction of the upper marginal-marine depositional system within the Vectis Formation (Barnes High Sandstone Member) in the Isle of Wight. The exact position of the sub-environment boundary is not clear due to the scattered occurrence of the outcrops and the absence of subsurface data. (B) Comparison with the modern meso- to low macro-tidal Gironde estuary in France (modified from Allen, 1991). The three-fold zonation and terminology of depositional environments within the Gironde estuary is after Dalrymple et al. (1992) who selected the Gironde estuary as an example of a mixed-energy estuary. At present (in the post-Holocene stillstand phase of sea level since 4000 BP; Fairbanks, 1989) tidal flats are building out from the estuary margins; however, previous tidal fiats were eroded away during the preceding Holocene transgression (Allen and Posamentier, 1993). The upper depositional system of the Vectis Formation may be more comparable to the Holocene transgressive stage of the Gironde estuary, though it is not clear whether the scale of the Vectis estuary may be comparable in size to the Holocene Gironde estuary or, possibly, larger.

here interpreted as a tidal 'bar' complex, render it difficult to precisely interpret the 3-D geometry of the upward-coarsening unit(s). However, it is not likely to have been formed as prograding deltas/lobes at an estuary mouth, such as a flood/ebb-tide delta because of (1) the very low salinity indicators, as mentioned above, (2) lack of tidal inlet channels feeding flood/ebb delta(s) (e.g. Nio and Yang, 1991), and (3) high-angle relationship between palaeocurrent trend and bar accretion directions. Within each tidal bar there are large-scale accretion surfaces dipping northwards, whose strike trend approximately parallel to the weakly bimodal NE-SE-oriented palaeocurrent direction (Fig. 6). The high-angle relationship be-

tween the accretion direction and the weakly bimodal palaeoflow trend suggests a dominance of lateral accretion/migration of the tidal bars (Dalrymple and Rhodes, 1995). The preferred environment is that of a poorly confined or broad channel located in the inner, tide-dominated part of a mixed-energy estuary, such as the modern Gironde estuary. The lack of any unequivocal tidal fiat deposits probably reflects a combination of their limited development and low preservation potential due to continuous lateral combing by migrating tidal channels, and/or the landward migration of a tidal prism, as occurred during the transgressive evolution of the Lower Jurassic Cook Formation (Marjanac and Steel,

246 1997). The 'tidal ravinement surface' (sensu Allen and Posamentier, 1993) associated with the landward migration of tidal inlets has been attributed as the cause of the low preservation potential of tidal flat deposits within the Gironde estuary (Allen and Posamentier, 1993). However, there is no evidence indicative for the presence of outer estuary or estuary mouth deposits within the Barnes High Sandstone Member as discussed above. If they are preserved, they would be located further basinward (eastward) of the Isle of Wight outcrops (Fig. 11). It is not clear whether the estuary complex represented by the Barnes High Sandstone Member was partly protected by some barrier bars or spits (wavedominated or mixed-energy estuary), as suggested above, or whether it was part of an open mouth tidedominated estuary, such as the modem Broad Sound in Australia (Cook and Mayo, 1977). The very low salinity and the presence of inferred central basin facies supports a partly closed estuary mouth system rather than a fully open one. The confinement of the lowsalinity brackish water and the high input of fresh water may have been further enhanced by regional-scale tectonic confinement (Wach and Ruffell, 1991).

Quantification of outcrop heterogeneities for subsurface studies This quantitative study has focused on intermediate- and small-scale heterogeneities, especially on the length distribution of mud drapes in the small scale.

Heterogeneities associated with each hierarchy The sequence stratigraphic framework of the Vectis Formation (Fig. 4) places the large-scale heterogeneities and facies variations into their regional context. The basal sequence boundary and the bay flooding surface at the boundary between the Barnes High Sandstone and Shepherd's Chine Members are both of regional extent. At this scale muddy zones occur either as (1) laterally extensive bayfill muds (Assemblage V 10) above a bay flooding surface (Figs. 4 and 8), or (2) mud plugs of inner estuarine channels (Assemblage V 1 l; Fig. 4) in the updip part of a bayhead-delta zone. Hence, recognition of sequence stratigraphic surfaces is crucial to correctly identifying large-scale reservoir architecture (e.g. vertical partitioning of individual genetic elements) and associated heterogeneity patterns. Basement faulting is also likely to have influenced regional thickness and lateral facies changes within the Vectis Formation as discussed in the previous section. Intermediate-scale heterogeneities (tens of centimetres to metres in thickness, the scale of individual

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channels and bars) were examined using the architectural element analysis methodology developed originally for fluvial successions by Allen (1983) and Miall (1985, 1988). We mapped facies distributions within the continuously exposed 2-D outcrop (400 m x 6 m) of the Barnes High Sandstone Member in the Cowleaze Chine-Shepherd's Chine area (Fig. 8). Within the member three separate phases of tidal sand bar deposition are bounded by laterally extensive (hundreds of metres) flooding surfaces (bar boundaries and bar abandonment surfaces on Fig. 8), which are overlain by muddy facies. Within the tidal bar complex muddier facies increase downwards as well as locally landwards to the northwest. At this scale, interbar muds (Assemblage V9) are the most significant mud barrier, which occur either (1) between tidal bars 1 and 2, or (2) as interfingering muds at the distal edge of tidal bar 1, above the sequence boundary. Within tidal bars 2 and 3, large accretion surfaces that extend for tens of metres across individual sandbodies serve as a focus for mud drape deposition in the proximal part of each bar (Figs. 4 and 8). In the distal part of the tidal bar the accretion surface is draped by thin zones of lenticular bedding (Assemblage V4B; Figs. 5 and 8), which merge laterally into thick interbar muds (Assemblage V9; Figs. 5 and 8). Small-scale heterogeneities occur as follows: (1) mud drapes on foresets/toesets and set boundaries of small-scale (5-50 cm thick) cross-bedding (Fig. 12); (2) rip-up clasts in similar locations to (1), above; (3) flaser bedding or thin (0.1-1 cm thick), wavy mud layers with variable continuity on top of ripples; and (4) as a muddy matrix within wavy bedding (sandstone 1-4 cm thick) and lenticular bedding (sandstone and siltstone 0.5-2 cm thick; Fig. 7).

Methodology of data acquisition and analysis for quantifying small-scale heterogeneities Small-scale heterogeneities were examined in two ways: (1) photographing and digitizing small areas (ca. 5 m x 20 m to 20 cm x 20 cm) of 2-D outcrops; and (2) using serial sectioning techniques to reconstruct the 3-D geometry of tidal sedimentary structures directly from large rock specimens (see the next section). In this section we concentrate on the quantification of the 2-D outcrops. We cleaned up the outcrop face and photographed it using a 500 mm telephoto lens from a distal point normal to the outcrop face to minimize the distortion effects (Fig. 12). The photograph was digitized at a high resolution, enhanced where necessary, and a large image file (TIF: ca. 1-5 MB) was made for each photo. The geological features of interest (mud drape, sedimen-

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Fig. 12. Mud drapes along the foresets and set boundaries of tangential to low-angle cross-bedding within the proximal tidal bar facies (Assemblage V4A) of tidal bar 1, Barnes High Sandstone Member (Fig. 8), Cowleaze Chine, Isle of Wight.

tary structures, etc.) on the image were traced using graphical software, and the validity of the traced image was re-checked in the field. We found that this combined approach of digitizing and field checking is more accurate and time-efficient than conventional outcrop mapping using binoculars, photomontage and graphical sketching/overlays in the field. A photograph (200 ASA film) which covers an area of 30 cm high and 40 cm wide typically gives a maximum resolution of 0.5 mm using this methodology. On the representative outcrop we placed a grid with dimensions such that most of the geologic features, particularly mud drapes, have both ends within the grid (Fig. 13). Then we measured the actual length of all the geologic features which occur, at least in part, within the grid. We found that the grid must contain more than 50 geologic features under study to obtain a reliable frequency analysis. To examine spatial variation of the length occurrence (e.g. upward change in the mean shale layer length), the grid must contain several hundred geologic features (ca. N = 200-300). For each mud drape type we assumed a simple geometry, or an idealized line shape which can be applied for both dip and strike sections of a particular sedimentary structure and its associated mud drape (Fig. 13). Both single and double mud drapes are assumed to be a single continuous line as long as at

least one component of the mud drape occurs continuously along one particular stratigraphic surface (e.g. set boundary or cross-bedding). For example, in documenting the dimension of individual mud drapes within the trough cross-bedding, we measured on the 'mud map' the vertical and horizontal coordinates of the left, right, top and the bottom points of the mud drape (Fig. 13). The horizontal and vertical extent of the mud drapes were calculated, and their frequency of length occurrence was analyzed. The coordinates of the contact points between mud drapes were also tabulated. This new methodology, which we have invented, though time-consuming, provides more geologically realistic values for shale length distribution than conventional methods in terms of the following points: (1) the geometry of the object can be calculated from the measured data; (2) the measured values of the coordinates can be used for both stochastic and deterministic models; (3) spatial variation of the length occurrence can be easily calculated.

Results of 2-D quantification of small-scale heterogeneities We analysed the length frequency of each type of mud drape, most of which show close to log-normal distributions. Fig. 14 shows the length distribution of

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Fig. 13. (A) Idealized geometry of mud drapes which occur along the trough cross-bedding. (B) Size of grid used to document length occurrence of the mud drapes within trough cross-bedding.

different types of mud drapes within the Barnes High Sandstone Member (Assemblages V3 and V4A). On each plot, the dashed-line curve denotes a log-normal distribution, with the mean and standard deviation corresponding to that of the measured data. The continuous-line curve denotes a log-normal distribution with the mean equal to the mode of the measured data, and the standard deviation adjusted where necessary to improve the fit as if more measurements of small lengths had been made. It is interesting to note that in all three cases continuous lines fit better to the frequency distribution. Traditional methods of quantifying length distributions in outcrop truncate both limbs of the frequency distribution, as both large length (> grid dimensions) and small length (< grid resolution) measurements are discarded. Using this methodology, all large length measurements are included, which may explain why the frequency distributions are slightly skewed towards large lengths. If so, the adjusted frequency curves (continuous line) accommodate both large and small lengths beyond the grid resolution and, hence, may be closer to the true length distribution. We also analyzed spatial changes of mud drape length distribution. Fig. 15A to C shows graphs in which the horizontal extent of mud drapes (corresponding to Fig. 14A to C, respectively) are crossplotted against their vertical coordinates. There is

a clear trend of upward-decrease in the length of mud drapes. Such features should be taken into account when constructing geologically realistic reservoir models for flow simulation. Stochastic reservoir models based solely on the length distribution of the mud drapes within the bulk volume of rock (Fig. 14A-C) would be misleading. For example, within tidal bar 1 (Fig. 14A,B), fluid migration may occur preferentially in the upper part of the bar where the horizontal extent of the mud drapes is smaller. Most mud drapes associated with cross-bedding (e.g. Assemblage V4 within tidal bar 1; Fig. 12) are isolated in two dimensions, which implies that their connectivity in three dimensions is extremely low. In certain instances, however, clustering of geologic features occurs, notably (1) lenticular bedding, and (2) thin wavy mud drapes. In both of these cases, as well as in the case of cross-bedding, clustering or high connectivity of objects commonly occurs periodically, and is constrained to several stratigraphic horizons.

Three-dimensional reconstruction of small-scale heterogeneities The three-dimensional architecture and lateral continuity of small-scale tidal sedimentary structures is poorly understood, because the existing models are based either on two-dimensional outcrop data (e.g.

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Fig. 14. Frequency analysis for the horizontal extent of mud drapes: (A) along set boundaries (assemblages V3 and V4 in tidal bar 1 (outcrops U-S) on Fig. 8; grid 3.5 m high x 17.2 m wide, see Fig. 13); (B) along the foreset of trough cross-bedding (assemblages V3 and V4 in tidal bar 1 (outcrops V-U); grid 3 m high x 4.9 m wide); (C) associated with the strike section of trough cross-bed ("mega" flaser bedding) within Assemblage V3 in tidal bar 3 (outcrops H-E; grid 0.8 m high x 13.4 m wide).

Fig. 15 (right). Cross-plot between stratigraphic height against the horizontal extent of mud drapes: (A) along set boundaries (assemblages V3 and V4 in tidal bar 1 (outcrops U-S) on Fig. 8; grid 3.5 m high x 17.2 m wide, see Fig. 13); (B) along the foreset of trough cross-bedding (assemblages V3 and V4 in tidal bar 1 (outcrops V-U); grid 3 m high x 4.9 m wide); (C) associated with the strike section of trough cross-bed ("mega" flaser bedding) within Assemblage V3 in tidal bar 3 (outcrops H-E; grid 0.8 m high x 13.4 m wide).

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Fig. 16. Visualization of the reconstructed models of the rock specimens. Light colours denote sand; dark colours denote mud. Note that the dark colour on the top face of specimen two is caused by the 3-D perspective shading.

Reineck and Wunderlich, 1968), or on modern analogues in which the preservation potential is uncertain (e.g. Terwindt, 1981). We used serial sectioning techniques to reconstruct the 3-D architecture of smallscale tidal sedimentary structures directly from large rock specimens. The resulting rock models are a close representation of the rock specimens, and are not based upon synthetic data or stochastic modelling techniques. The rock specimens, measuring approximately 60 cm x 60 cm x 20 cm, were taken from Assemblages V5 and V6 (distal tidal bar facies) of the Vectis Formation, and are representative of facies found throughout the Tilje Formation and other heterogeneous tidal sandstone reservoirs. Specimen one predominantly contains lenticular sand lenses connected both horizontally and vertically; subordinate, but more continuous, wavy beds are also present. Specimen two predominantly contains isolated mud flasers embedded in sand. Both specimens were sectioned at ca. 20 mm intervals, and the section faces photographed. Each section photograph was digitized, and the boundaries between sand and mud traced. These boundaries represent the intersections between bedding surfaces and the section faces, and by manually correlating the traced boundaries between each 2-D section, the 3-D architecture of the bedding surfaces was reconstructed.

The reconstructed rock specimen models may be visualized using any standard reservoir modelling package. Fig. 16 shows the reconstructed models visualized using IRAP RMS (Smedvig Technologies, 1998). The advantage of the reconstructed models is that a range of visualization options is available. For example, the models may be sectioned in any orientation. Fig. 17A shows two orthogonal sections through specimen one; note that the lenticular sandbodies often "stack up" and are connected vertically. Fig. 17B shows two orthogonal sections through specimen two; note the variation in the lateral continuity and vertical connectivity of the mud flasers. Alternatively, individual blocks may be sampled from the whole model. Fig. 18 shows several core-plug-sized blocks sampled from specimen one; note that the lateral continuity of the sand and mud layers is likely to be higher at the scale of a core-plug than at the scale of the whole model. Finally, individual sand and mud layers may be viewed independently of the rest of the model. Fig. 19 shows sand and mud layers within specimen one; note that many of the lenticular sandbodies are connected horizontally in 3-D, although they may appear to be isolated in 2-D (Fig. 17A). Note also that the mud layers are not entirely laterally continuous where the sandbodies connect vertically. The rock specimens from which models one and two were reconstructed sample only a small volume

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Fig. 17. Visualization of the reconstructed models of the rock specimens: (A) two orthogonal sections through specimen one; (B) two orthogonal sections through specimen two. Light colours denote sand; dark colours denote mud.

of heterolithic, small-scale tidal sedimentary structures. It is therefore possible that they are not generally representative of those structures. Nevertheless, the models provide valuable information about the 3-D distribution of sand and mud, which would not be available from 2-D observations. For example, 3-D visualization of specimen two (lenticular-wavy bedding) indicates that lenticular sandbodies may appear isolated in 2-D yet be connected in 3-D. Moreover, mud layers may appear continuous in 2-D, yet be discontinuous in 3-D where sandbodies connect vertically. These observations indicate that a proper understanding of the 3-D architecture of heterolithic, small-scale tidal sedimentary structures has a significant impact on our assessment of their reservoir quality, particularly reservoir connectivity.

The representation of small-scale heterogeneities in reservoir-scale models The representation of small-scale heterogeneities in reservoir-scale models is problematic: as yet, they cannot be represented explicitly, because the resolution of a reservoir-scale model grid is insufficient

(Haldorsen, 1986). Typically, they are grouped within facies types, and their "effect on flow" is represented using "averaged" petrophysical data obtained from core-plugs and well-logs. This approach is valid only if the averaged petrophysical data, measured at the scale of the core-plugs and well-logs, properly represents the effect on flow of the small-scale heterogeneities at the scale of a reservoir model grid-block (Haldorsen, 1986). We used the reconstructed model of specimen one to investigate numerically the effect of sample volume on the measured single-phase permeability of lenticular- wavy-bedded facies. The effect of sample volume was investigated by calculating the effective permeability of the rock model for a range of sample volumes, using the "sealed side" pressure solver of Warren and Price (1961). This is the numerical equivalent of taking cores of different sizes from the rock specimens, and measuring their permeability in the laboratory. The permeability of the sand and mud facies in the fine-grid model (sand (ks) and mud (km)) w e r e assumed to be uniform and isotropic, and permeability values were assigned on a facies basis. For

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Fig. 18. Individual blocks sampled from rock specimen one. Each block measures 8 x 8 x 2 cm. Light colours denote sand; dark colours denote mud, with some variation in the light colour of sand due to 3-D perspective shading.

each sample volume, effective (measured) permeability values (ke) w e r e derived for the x, y (horizontal) and z (vertical) directions, and expressed in terms of a dimensionless (normalized) permeability k, where k for a given direction is given by: ke -- km

k=

ks - km

Expressed in this way, the normalized measured permeability values for a facies are independent of the dimensional permeability values (ks and km) of the sand and mud; rather, they depend only upon the sand/mud permeability ratio: ks X/~ =

km

Fig. 20A, B shows the variation in the (arithmetic) mean measured permeability as a function of sam-

ple volume, for the case X~ = 10, 000. The mean horizontal permeability (k~) generally decreases with increasing volume. In contrast, the mean vertical

permeability (kz)oscillates with increasing sample volume, and no clear trend is present. Note the large contrast between the horizontal (k~) and vertical (kz) mean permeabilities, which is a characteristic feature of these heterolithic facies. For comparison, Fig. 20C, D shows the variation in the modal (most frequently measured) permeability as a function of sample volume. For the horizontal permeabilities (k~ and ky), the modal value oscillates as the sample size increases (Fig. 20C), and no clear trend is present. In contrast, for the vertical permeability (kz), the modal value generally increases with increasing sample size (Fig. 20D), and varies by almost an order of magnitude between the smallest and largest sample volumes used. These results demonstrate that the measured permeability varies with sample volume, and more significantly, that the "averaged" (modal and mean) measured permeability also varies with sample volume. This indicates that permeability distributions obtained

Outcrop studies of tidal sandstones for reservoir characterization

253

Fig. 19. Individual sand and mud layers within rock specimen model one. Images A and B show sand and mud layers towards the top of the model; images C and D show sand and mud layers towards the base of the model. Light colours denote sand; dark colours denote mud.

from numerous core-plug and well-log measurements of complex bedform-scale sedimentary structures do not properly represent their effective permeability in reservoir-scale models. Conclusions Tidal deposits in the Lower Cretaceous of southern England (Vectis Formation) have been used as outcrop analogues for heterolithic tidal sandstone reservoirs. Quantitative data have been collected to constrain numerical flow simulations at a variety of scales. The key results are as follows. (i) The outcrops comprise meso- to low macrotidal estuarine deposits. The main outcrop area is characterized by vertically stacked, coarseningupward units and is interpreted as a composite tidal bar formed within the inner part of a broad, mixedenergy, estuary (incised valley). (ii) Reservoir heterogeneities have been characterized at the following scales: (1) large-scale (external

geometry defined by key stratal surfaces-sequence boundaries and flooding surfaces); (2) intermediatescale (internal geometry defined by sand bar accretion surfaces, bar abandonment surfaces and facies boundaries); and (3) small-scale (internal facies variability covering lenticular-wavy-flaser bedding). (iii) Quantification of small- to intermediate-scale heterogeneities, notably mud drapes along the set boundaries and on cross-bed foresets, has been achieved by placing a grid on small, representative areas of the 2-D outcrop and measuring the coordinates of mud drapes. The length of most mud drape types exhibits a close to log-normal distribution. In addition, mud drape length shows a progressive upwarddecreasing trend within each upward-coarsening unit. (iv) Reservoir characterization of small-scale heterogeneities (lenticular-wavy-flaser bedding) has been achieved through the use of serial sectioning techniques to reconstruct their 3-D architecture directly from large rock specimens. (v) The reconstructed rock specimen models have

S. Yoshida et al.

254

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Fig. 20. The variation in measured (normalized) permeability with sample volume, for specimen one with a sand/mud permeability ratio of Xk -- 10, 000. Plot A shows the horizontal mean measured permeability (kx); plot B shows the vertical mean measured permeability (kz). In both plots, the error bars denote one standard deviation; note that the decrease in size of the error bars with increasing sample volume reflects in part the decrease in the number of samples measured owing to the finite size of the rock specimen model. Plot C shows the horizontal modal measured permeability (kx and ky); plot D shows the vertical modal measured permeability (kz). In both plots, the error bars denote the bin sizes used to obtain the modal permeability.

been used to investigate representative reservoir properties for individual facies types. For lenticularwavy-bedded facies, both the mean and modal measured single-phase permeability varied with sample volume, which indicates that "averaged" well-log and core-piug measurements do not yield representative permeability data for these facies. Consequently,

it may not be valid to assign rock properties to heterolithic facies in tidal reservoir models, using conventional well data. (vi) Quantitative outcrop studies offer one approach to obtain more representative reservoir properties for heterolithic tidal facies.

Outcrop studies of tidal sandstones for reservoir characterization

Acknowledgements This study is part of the FORCE project (Forum of Reservoir Characterization and Reservoir Engineering) on tidal sandstone reservoir characterization, funded by BP-Amoco, Norske Conoco A/S, Fortum Petroleum A/S, Saga Petroleum (now Norsk Hydro), and Statoil. The authors gratefully acknowledge these sponsors, and the Norwegian Petroleum Directorate for helping to coordinate the project. Smedvig Technologies are thanked for providing the IRAP RMS software. We appreciate our discussions with Roland Goldring on the fauna and ichnofacies of the Wealden Group. Robert Dalrymple, Brian Willis, Duna Mellere and Steve Flint, Duna Mellere and Lars-Magnus F~ilt, as well as other geoscientists at Statoil, are thanked for their comments and constructive suggestions. Field and technical assistance was provided by John Dennis, Tom Huang, Zoe Hansen, Francis Longworth, Richard Evans, Susan Nuttall, Innocent Ofoma, Nigel Huggins and Mark Buckley. The paper has greatly benefited from the reviews of Brian Zaitlin and Phillip Ringrose. Nick Lee proofread and improved our final manuscript. We thank the editors Ole Martinsen and Tom Dreyer for helping with preparation of the final manuscript.

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S. YOSHIDA

M.D. JACKSON H.D. JOHNSON A.H. MUGGERIDGE A.W. MARTINIUS

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Centre for Petroleum Studies, T.H. Huxley School of Environment, Earth Sciences and Engineering, Imperial College, Prince Consort Road, London SW7 2BP, UK Present address: Surface Processes and Modern Environments Research Group, Department of Geology, Royal Holloway, University of London, Egham, Surrey TW20 OEX, UK Centre for Petroleum Studies, T.H. Huxley School of Environment, Earth Sciences and Engineering, Imperial College, Prince Consort Road, London SW7 2BP, UK Centre for Petroleum Studies, T.H. Huxley School of Environment, Earth Sciences and Engineering, Imperial College, Prince Consort Road, London SW7 2BP, UK Centre for Petroleum Studies, T.H. Huxley School of Environment, Earth Sciences and Engineering, Imperial College, Prince Consort Road, London SW7 2BP, UK Statoil Research Centre, Arkitekt Ebbellsveg 10, N-7005 Trondheim, Norway Present address: c/o Statoil Venezuela - Sincor Project, N-4035 Stavanger, Norway

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259

Lower Cretaceous (Barremian-Albian) deltaic and shallow marine sandstones in N o r t h - E a s t G r e e n l a n d

-- sedimentology,

sequence stratigraphy and regional implications Michael Larsen, Tor Nedkvitne and Snorre Olaussen

Lower Cretaceous sandstones of the Steensby Bjerg Formation are exposed in northern Hold with Hope where they form an unusual coarse-grained clastic wedge in the otherwise mudstone-dominated 'Cretaceous succession of North-East Greenland. The sandstones were deposited as early post-rift basin fill, in a deltaic and shallow marine environment during overall sea-level rise following intense faulting, block rotation and subaerial erosion in the Late Jurassic-earliest Cretaceous. Seven major facies associations are identified and the succession is divided into unconformity-bounded units. Deposition started with transgressive shoreface sandstones in the early Barremian followed by at least two phases of southwards delta progradation and valley incision during the Barremian-early Aptian. In the late Aptian-early Albian the deltaic system backstepped and progressively deeper-water facies were deposited. The coarse-grained clastic system was drowned in the early Albian and marine mudstones dominated from this time and onwards. A tectonic phase with renewed uplift and submarine erosion in the middle Albian is represented by an angular unconformity between mudstones of the Steensby Bjerg Formation and mudstones with intercalated sandy turbidites of the overlying Home Forland Formation. The position of the Hold with Hope clastic wedge at a relay ramp in the western bounding fault of the Mesozoic rift basins implies that Lower Cretaceous, shallow marine sandstone wedges may be predicted to form a new reservoir play model along steps in older Mesozoic lineaments in the Mesozoic basins offshore Norway and the West Shetland-Faeroe Basin.

Introduction

The Upper Permian-Jurassic sedimentary succession in East Greenland has long been recognised to be an important analogue for the petroliferous basins offshore Norway and in the northern North Sea (e.g. Surlyk et al., 1986; Stemmerik et al., 1993; Price and Whitham, 1997). The recent move of exploration to the western, deep-water areas has, however, changed the interest towards the Cretaceous-Palaeogene succession. The rift basins of North-East Greenland (north of 72~ include well-exposed CretaceousPalaeogene sediments with a pre-drift position approximately 100-150 km northwest of the Gjallar Ridge (Fig. 1). Data from the onshore sections may help to derive new play models for the offshore basins on the western Norwegian Shelf. The Cretaceous succession of North-East Greenland is more than 2 km thick, and consists of siliciclastic, mainly marine sediments deposited following a major rift phase and reorganisation of the basin in the latest Jurassic-earliest Cretaceous (Surlyk, 1978). During the Barremian-Aptian, the rift topography became submerged and offshore mudstones

were deposited in most of the East Greenland Basin (Donovan, 1957) (Fig. 2). Fine-grained marine sediments onlap Triassic and Jurassic sediments on the footwall of rotated fault blocks at Traill 0, Geographical Society 0 and Wollaston Forland. Only locally, coarse-grained deltaic and shallow marine sandstones were deposited in this overall transgressive stage. The best-exposed example of these Lower Cretaceous sandstones is the Steensby Bjerg Formation of northern Hold with Hope (Figs. 2 and 3). The Upper Cretaceous succession in East Greenland consists mainly of offshore mudstones with subordinate turbidite sandstones and fault-scarp derived conglomerates (Donovan, 1957; Stemmerik et al., 1993, 1997; Surlyk and Noe-Nygaard, 2001) (Fig. 2). The palaeocurrent directions changed from mainly coast-parallel (north-south) during the Early Cretaceous to offshore-directed flows (eastwards) in the Late Cretaceous (Stemmerik et al., 1997; Whitham et al., 1999). Koch (1931) and Maync (1949) described the succession at Hold with Hope, but not until recently (Kelly et al., 1998) a formal lithostratigraphy for the Barremian-lower Albian sandstones and mudstones

Sedimentary Environments Offshore Norway - Palaeozoic to Recent edited by O.J. Martinsen and T. Dreyer. NPF Special Publication 10, pp. 259-278, Published by Elsevier Science B.V., Amsterdam. 9 Norwegian Petroleum Society (NPF), 2001.

260

M. Larsen et al.

Fig. 1. Pre-drift reconstruction of the northern North Atlantic showing Cretaceous outcrops in East Greenland, major faults and the position of the study area relative to the Vcring Basin. Note the relatively short distance between the study area in North-East Greenland and the Gjallar Ridge on the Norwegian Shelf, which is the area of recent drillings (1999) for hydrocarbons. Pre-drift reconstruction by courtesy of SAGA Petroleum.

Fig. 2. Stratigraphic scheme of the Barremian-Maastrichtian succession in North-East Greenland. Based on Donovan (1957), NChr-Hansen (1993), Kelly et al. (1998), Surlyk and Noe-Nygaard (2001) and H. NChr-Hansen (pers. commun., 1999).

Lower Cretaceous (Barremian-Albian) deltaic and shallow marine sandstones in North-East Greenland

261

Fig. 3. Lower Cretaceous (Barremian-lower Albian) sediments (LC) of the Steensby Bjerg Formation are exposed along the north coast of Hold with Hope in North-East Greenland. The sandstone-dominated Lower Cretaceous succession unconformably overlies rotated Lower Triassic (LT) and Middle-Upper Jurassic (MJ) strata. The succession is overlain by a thin horizon of sandstones and conglomerates of Paleocene? age overlain by Paleocene and Eocene flood basalts (P). View towards the east. The Lower Cretaceous succession (LC) is approximately 300 m thick.

of the Steensby Bjerg Formation and the overlying middle Albian-Santonian mudstones of the Home Forland Formation was established (Fig. 4). A general sedimentological interpretation dividing the succession into coarse shallow marine, fine shallow marine and coarse basinal facies associations was presented by Whitham et al. (1999). They furthermore identified six coarsening-upward sequences and two sequence boundaries within the Lower Cretaceous sandstones. In this paper, we present a sedimentological facies analysis and sequence stratigraphic interpretation of Steensby Bjerg Formation based on fieldwork in 1996-1998. The formation is up to 300 m thick and exposed in the coastal cliffs of northern Hold with Hope for a distance of approximately 18 km (Fig. 5). The studied succession is deposited as the result of a sea-level rise followed by at least two episodes of southward delta progradation and valley incision. The sandstone-dominated depositional system was drowned during Aptian marine flooding and the delta stepped back towards the north. From the late-early Albian mudstone deposition prevailed. Reconstruction of the Lower Cretaceous palaeogeography in the North-East Greenland Basin shows that the sand-rich deltaic system on Hold with Hope is located at a major relay zone in the western boundary fault of the Mesozoic rift basins (Whitham et al., 1999). Similarly located Lower Cretaceous, shallow

marine sandstone wedges may be present in similar structures in the older Mesozoic lineaments in the Norwegian Sea and Faeroe-Shetland Basin and may there form important new exploration targets.

Regional setting A series of north-south elongate sedimentary basins of late Palaeozoic-Palaeogene age is exposed along the East Greenland margin due to Neogene uplift. The basins formed due to extensional collapse following the Caledonian orogeny and later episodes of rifting (Surlyk, 1990). The basins are the result of a complex series of tectonic events caused by plate movements and reorganisations and thermal contraction. Mesozoic rifting culminated in North-East Greenland in Late Jurassic-earliest Cretaceous time (Vischer, 1943; Surlyk, 1990). The Lower Cretaceous succession overlies the degraded rift topography of this tectonic phase. The Mesozoic rift basins of East Greenland are mainly bounded by eastward-dipping normal faults (Vischer, 1943; Haller, 1970). Between the Traill O-Geographical Society 0 area and the Wollaston Forland area the major north-south-trending fault zone steps approximately 50 km towards the east forming a major southward-facing relay zone across Hold with Hope (Whitham et al., 1999) (Fig. 1).

262

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Fig. 4. Litho- and biostratigraphy of the Lower Cretaceous succession in northern Hold with Hope. Lithostratigraphy based on Kelly et al. (1998). Biostratigraphy based on Kelly et al. (1998), NChr-Hansen (1993) and unpublished work by H. Nchr-Hansen (pers. commun., 1999). See Fig. 5 for location.

The western part of Hold with Hope probably represents the southwards extension of the fault block forming a structural high exposing westerly- and southwesterly-dipping crystalline basement in Clavering 0 (Vischer, 1943; Stemmerik et al., 1993). The Clavering 0 high is bounded to the east by an eastward-dipping normal fault, which probably has its southwards extension in Fosdalen at Hold with Hope (Fig. 5). Kelly et al. (1998) suggested that the Fosdalen fault was active in Early Cretaceous time and thereby influenced deposition of the Lower Cretaceous sediments. A number of smaller synthetic faults present in the footwall block of the Fosdalen fault seem also to have been active in the Early Cretaceous. The Lower Cretaceous succession rests with an angular unconformity (3~ ~ on Lower Triassic (Wordie Creek Formation) (Koch, 1931; Nielsen, 1935; Maync, 1949) and Middle-Upper Jurassic strata (Pelion, Payer Dal and Bernbjerg Formations) (Stemmerik et al., 1997; Kelly et al., 1998; Larsen et al., 1998). The unconformity, Ukl of Kelly et al. (1998), was formed following Late Jurassic-earliest Cretaceous rifting associated with block faulting and rotation (Fig. 4). The Lower Triassic and Jurassic strata are, thus, rotated in 0.5 to 2 km wide fault

blocks and dip 5~ 14~ towards the southwest, whereas the overlying Lower Cretaceous strata dip approximately 3~ ~ towards the southwest. In Steensby Bjerg, a number of north-south-striking normal faults of Late Jurassic-Early Cretaceous age can be seen to terminate against the unconformity at the base of the Cretaceous succession (Larsen et al., 1998). Most of the Mesozoic faults in northern Hold with Hope, however, were reactivated during the Palaeogene and a number of new, southwest-northeast-striking faults were formed which cross-cut the Cretaceous sediments and the overlying basalts. The Cretaceous succession is truncated by an erosional unconformity, Ut of Kelly et al. (1998), suggested to be of Paleocene age (Maync, 1949; Upton et al., 1980). In northern Hold with Hope, up to 24 m of sandstones and conglomerates with subordinate mudstones overlie the unconformity. These sediments of possible Paleocene age are overlain by up to 900 m of Paleocene-Eocene flood basalts (Upton et al., 1980).

Stratigraphy The Cretaceous succession in northern Hold with Hope consists of the sandstone-dominated Barremian-

Lower Cretaceous (Barremian-Albian) deltaic and shallow marine sandstones in North-East Greenland

263

Fig. 5. Geological map of the studied area in northern Hold with Hope. Geographical names mentioned in the text and the position of the geological cross-sections shown in Figs. 6, 7 and 13 are indicated.

lower Albian Steensby Bjerg Formation, up to 300 m thick, unconformably overlain by the mudstone-dominated middle Albian-Santonian Home Forland Formation, up to 1000 m thick (Kelly et al., 1998) (Fig. 4). In the studied sections only the lowest part of the Home Forland Formation, the Fosdalen Member, is exposed. The sediments crop out in two areas separated by the wide valley of the river Gulelv (Fig. 5). East of the river at Steensby Bjerg-Diener Bjerg, the Steensby Bjerg Formation is divided into the Diener Bjerg (lower Barremian), the Gulelv (Barremian-Aptian) and the Rc~delv (lower Aptian) Members (Kelly et al., 1998) (Fig. 4). At Diener Bjerg and in the eastern part of Steensby Bjerg the formation is unconformably overlain by the Fosdalen Member (middle Albianupper Turonian). At Stensi6 Plateau to the west of Gulelv, the Steensby Bjerg Formation consists of the Stribedal (lower Barremian), Blfielv (upper Barremian?), Stensi6 Plateau (upper Barremian-lower Aptian), Gulelv (Aptian-lower Albian) and the ROdelv (lower A1bian) Members (Kelly et al., 1998) (Fig. 4). The

litho- and biostratigraphy was thoroughly discussed by Kelly et al. (1998) and ages were assigned to the lithostratigraphic units based on dinoflagellate cysts and ammonites. Our observations roughly confirm the established stratigraphy; however, a detailed palynological study (H. NOhr-Hansen, pets. commun., 1999) suggests, that the age of the Stribedal Member should be earliest Barremian Nelchiopsis kostromiensis Subzone (I1) of N0hr-Hansen (1993) and not the Pseudoceratium anaphrissum Subzone (I2) as stated by Kelly et al. (1998). This shifts the timing of the inundation of the Hold with Hope block down towards the Hauterivian-Barremian boundary. The Cretaceous succession is bounded by two major angular unconformities (Ukl and Ut of Kelly. et al., 1998) as described above. In addition, four unconformities are present within the Cretaceous succession. They are of late Barremian (SB1), early Aptian (SB2), late Aptian (SB3) and late-early A1bian (SB4) age (Fig. 4). SB 1-SB3 are characterised by erosional surfaces, lag deposits and abrupt facies changes across the unconformity surface, whereas

264 an angular unconformity is present across SB4. SB4 separates the Steensby Bjerg and Home Forland Formation and was originally described by Kelly et al. (1998) under the heading Uk2. Across two of the unconformities (SB 1, late-early to late Barremian, and SB4, late-early Albian) a hiatus is present based on the dinoflagellate cysts stratigraphy (H. NOhr-Hansen, pers. commun., 1999) (Fig. 4).

Sedimentary facies associations A number of sedimentary facies are identified based on lithology, sedimentary structures and body and trace fossils. For the purpose of description, the facies are grouped into seven major facies associations, of which each characterises a depositional environment. The facies associations allow a better and slightly different subdivision of the succession than the lithostratigraphic units defined by Kelly et al. (1998).

Shoreface sandstones and conglomerates The shoreface association consists of interbedded pebble conglomerates, fine- to medium-grained sandstones and mica-rich mudstones. The conglomerates can be followed laterally for a few tens of metres and vary from a few cm to 40 cm of thickness over short distances. The bases of the conglomerate beds are erosional, whereas the upper boundary may be either sharp or shows a gradual transition into the overlying sandstones. The conglomerates are mainly clast-supported with well-rounded spherical to discoidal quartz pebbles (grain size around 3 cm) and locally cobbles (up to 25 cm). The conglomerates are commonly structureless, but locally show trough cross-bedding with foresets dipping towards northeast. Rare belemnites are present throughout the association and reworked Middle Jurassic ammonites are present in the conglomerate bed overlying the base Cretaceous unconformity (see also Kelly et al., 1998). The sandstones are generally well-sorted and form beds of 1-2 m thick, with a maximum of up to 5 m thick. In the fine-grained sandstones the dominant structure is wave-ripple cross-lamination, but locally hummocky and swaley cross-stratification are present. Scour-and-fill and trough cross-bedding dominate in the more coarse-grained beds. Foreset dip azimuths are bi-directional, but with a clear dominance towards the northeast (mean 39~ The finegrained sandstone beds are strongly burrowed showing the trace fossil Curvolithos multiplex, whereas Skolithos isp. occur scattered in the medium-grained beds. Ophiomorpha nodosa locally occur in great numbers extending from bedding surfaces at the top

M. Larsen et al.

of the sandstone succession (upper boundary of the Gulelv Member). Laminae and thin beds of mica-rich and carbonaceous mudstones occur throughout the association. They either form lenticular units draping troughs of cross-sets or they form laterally continuous beds, up to a few centimetres thick. Dinoflagellate cysts and disseminated plant fragments are present in the mudstones. The shoreface association is present at the base of the Cretaceous succession. It reaches a maximum thickness of 37 m in the eastern part of Stensi6 Plateau (Stribedal Member) (Fig. 6). At Steensby Bjerg it forms the lower part, up to 24 m thick, of the Gulelv Member. At Diener Bjerg a conglomeratic lag deposit up to 1 m thick forms the base of the Diener Bjerg Member. The shoreface association is furthermore present at the top of the Gulelv Member representing the wave-reworked top of the Lower Cretaceous sandstone succession (Figs. 6 and 7). The sedimentary structures and the strong marine burrowing suggest that the association was deposited in a shoreface environment (see also Whitham et al., 1999). The abundant conglomerates with scoured bases and local occurrence of swaley and hummocky cross-stratification suggest periodic highenergy, wave-dominated conditions during storms.

Cross-bedded delta front sandstones Coarse-grained locally pebbly sandstones dominate sediments of this association. The sandstones are planar and trough cross-stratified with set thicknesses between 0.5 and 2 m, locally up to 5 m, and forming cosets, up to 40 m thick, bounded by coarse-grained lag deposits (Figs. 7 and 8a). Locally, the cross-sets can be seen to form intrasets in large-scale compound cross-beds with low-angle master-surfaces dipping a few degrees towards the southwest. Foreset dip directions indicate that the cross-bedded units prograded towards the south and southwest (mean 222~ In the eastern part of Steensby Bjerg and around Diener Bjerg a single large-scale foreset bed with foresets up to 50 m high occurs (Figs. 7 and 8b). The largescale foresets dip towards the southeast (mean 150~ The sandstones are generally unfossiliferous although Kelly et al. (1998) reported a single ammonite specimen from the basal part of the large-scale cross-bedded unit at Diener Bjerg. Locally, sinuous burrows occur along the bedding planes of the large-scale foresets. The sandstones grade downdip into toesets of laminated silty and sandy mudstones with intercalated thin sandstone beds. The grain-size and the number of sandstone beds increase upward and the toesets form upward-coarsening units up to 5 m thick. The

Lower Cretaceous (Barremian-Albian) deltaic and shallow marine sandstones in North-East Greenland

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265

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266

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Lower Cretaceous (Barremian-Albian) deltaic and shallow marine sandstones in North-East Greenland

267

Fig. 8. (a) Cross-bedded pebbly sandstones of the Gulelv Member exposed at Steensby Bjerg. The cross-beds are compound consisting of a master-bedding dipping a few degrees to the southwest with internal cross-sets 2-4 m thick. Person for scale. (b) Coarse-grained sandstones of the Gulelv Member exposed north of Diener Bjerg. The sandstones consist of a single, large-scale foreset unit (stippled) 45 m high. The sandstones overlie poorly exposed mudstones of the Diener Bjerg Member. (c) Tidally influenced cross-strata in the upper part of the Gulelv Member at Steensby Bjerg. Note tidal bundles and double mud-drapes (see detail in d). Person for scale. (d) Mudstone-rich bottomset of a tidally influenced, planar cross-set. Note double mud-drapes and reactivation surface (arrow). Pencil, 10 cm long, for scale. (e) Well-rounded vein quartz clast, long axis 24 cm (above pencil) embedded in wave-rippled coarse-grained sandstones forming the top of the Gulelv Member at Steensby Bjerg. The outsized clasts mark the lower Aptian sequence boundary SB2. Pencil (encircled) for scale. (f) Outcrop of the angular unconformity SB4 southwest of the river Gulelv. The unconformity separates siltstones of the lower Albian ROdelv Member from black mudstones with thin sandy turbidites of the middle Albian Fosdalen Member. Person for scale.

mudstones are rich in disseminated organic material of terrigenous origin (plant fragments) and are locally strongly bioturbated. A few burned and abraded dinoflagellate cysts were found in the fine-grained toeset sediments. The delta front association forms the bulk of the Gulelv Member in Steensby Bjerg and Diener Bjerg (Fig. 7).

The coarse-grained, locally pebbly lithology, the dominant unimodal, planar and trough cross-bedded structure and the lack of wave-generated structures suggest that the association was deposited in a fluvial-dominated mouthbar or delta front environment. This is supported by the apparent lack of marine fossils including dinoflagellates and the fact that trace

268

fossils are very rare in the sandstones. The mudstones forming the toeset were deposited from suspension in front of the prograding deltaic systems. The distinct large-scale foreset bed present in the eastern part of the area suggests deposition in a delta lobe (Gilberttype) building into a marine body of water. The height of the foresets may give a rough indication of water depths reaching up to 50 m. The southeasterly palaeocurrent direction in this unit that is almost 90 ~ offset from the general palaeocurrent direction is ambiguous. The large-scale foresets and the direction of progradation, however, may have been controlled by a sudden increase in the water depth across eastwarddipping faults synthetic to the Fosdalen Fault. The delta interpretation of the cross-bedded sandstones given here differs somewhat from the interpretation given by Whitham et al. (1999), which suggested a tidally influenced shallow marine environment for the Gulelv Member. Tidally influenced cross-beds, however, are based on our observations, restricted to the uppermost part of the Gulelv Member (Figs. 6 and 7) and are described below in the tidal channel/tidal bar association.

M. Larsen et al.

Massive channel sandstones The association is characterised by massive and cross-bedded sandstones forming lenticular units up to 23 m thick and several kilometres wide (Fig. 6). The lower boundary is strongly erosional and concave-up, whereas the upper boundary is nearly horizontal (Figs. 6 and 9a). The sandstones are mediumto coarse-grained, locally pebbly with a distinct coarse-tail grading in the lower few metres. Large rip-up mudstone and fine-grained sandstone clasts are common (Fig. 9b,c). The thickest middle part of the sandstone units consists of well-sorted fineto medium-grained, massive sandstones. In the upper 2-3 m parallel-laminated fine-grained sandstones with primary current lamination occur in one of the units. Otherwise recumbent cross-beds and waterescape structures are present in the upper part of the sandstone units (Fig. 9d). Palaeoflow was directed towards the southwest as indicated by the geometry of the sandstone bodies, direction of parting lineation and foreset dip in the cross-sets. The thickest developed unit of the association is present in the B lfielv Member of Kelly et al. (1998). It is up to 23 m

Fig. 9. (a) Lenticular unit of massive sandstones cut into mudstones of the Stensi6 Plateau Member. The sandstone unit is 12 m thick and represents one of several channelized sandstones (including the Blgtelv Member) present at Stensi6 Plateau. Person for scale. (b) Erosional contact between the Bl~elv Member (above) and the Stribedal Member (below). Note the massive appearance of the sandstones of the Blgtelv Member and the large rip-up clast of mudstone (M). Hammer encircled for scale. (c) Large-scale slump structure in the basal part of the Blfielv Member. Note large-scale mudstone clasts, contorted heterolithic mudstones and units of pebbly sandstones. Hammer encircled for scale. (d) Recumbent cross-sets (arrow) associated with the channelized massive sandstones. Person for scale.

Lower Cretaceous (Barremian-Albian) deltaic and shallow marine sandstones in North-East Greenland

thick and shows a clear tri-partite division. The lower part is dominated by large clasts, up to 4 m across, of contorted and slumped, internally laminated, heterolithic mudstones and fine-grained sandstones interbedded with medium- to coarse-grained, locally pebbly sandstones (Fig. 9c). Well-rounded quartz pebbles, fragments of coalified wood, bivalves and reworked Hauterivian belemnites are common (Kelly et al., 1998). The middle part, up to 18 m thick, consists of massive well-sorted sandstones, with clear normal grading. In some sections faint large-scale cross-bedding with a set thickness of up to 5 m is developed. The upper part, up to 3 m thick, shows parallel lamination with primary current lamination. The association also occurs in the lower part of the Gulelv Member at Stensi6 Plateau and in the uppermost part of the Gulelv Member at Steensby Bjerg (Figs. 6 and 7). The presence of faint large-scale cross-bedding in the otherwise massive sandstones and the recumbent cross-bedding in the upper part of the units may point towards deposition by traction currents in a river or tidal channel. Recumbent deformed cross-bedding is, thus, said to be characteristic for fluvial deposits impending a strong traction drag on the sediments (Allen and Banks, 1972). The large, slumped mudstone and sandstone clasts, however, could not have survived long-distance transport in a traction regime. The large folded and contorted clasts may be interpreted as representing partly consolidated sediment deposited at the base of the channel following collapse of the channel margins. Massive sandstones associated with bank failure have been described from large rivers (see Martin and Turner, 1998) and in the lower part of incised valley fills (Martinsen, 1994). We suggest that the massive channel sandstones can be related to channel incision into unstable partly consolidated sediments leading to high rates of sedimentation associated with episodic bank failure. The presence of reworked Hauterivian belemnites indicates that significant erosion of older sediments occurred at this time and thus may lend some credence to the interpretation of the sandstones as incised valley-fill deposits.

Tidal channel/tidal bar sandstones The association is characterised by medium- and coarse-grained, mica-rich sandstones, which are generally well-sorted, although thin laminae of mudstones and disseminated organic detritus commonly drape foresets and troughs (Fig. 8c,d). The dominant structure is planar cross-bedding and locally double mud-drapes and tidal bundles are present. Set thickness varies between 10 and 80 cm. Foreset dip directions indicate a dominant palaeocurrent direc-

269

tion towards the southwest (mean 220~ B ioturbation including Ophiomorpha isp. is common. The association is present in the upper part of the Gulelv Member at Stensi6 Plateau and Steensby Bjerg (Figs. 6 and 7). The cross-bedded sandstones are interpreted as deposited in tidally influenced sandbars in a shallow marine environment or within tidal channels (see also Whitham et al., 1999). The presence of double mud-drapes and low-angle reactivation surfaces indicates that deposition took place within the intertidal zone. The tidal currents were southwesterly (ebbdominated) and with pronounced slack water periods suggesting an in-channel depositional environment (Nio and Yang, 1991).

Tidal flat sandstones The association consists of a monotonous succession, up to 45 m thick, of well-sorted, fineto medium-grained sandstones with abundant muddrapes rich in disseminated organic material (Fig. 6). The sandstones show parallel lamination, low-angle cross-stratification and locally wave-ripple cross-lamination. No macrofossils were found, but dinoflagellates are common. The sandstones show a moderate degree of bioturbation. The association is present at Stensi6 Plateau (Gulelv Member), where it forms a large-scale coarsening-upward succession overlain by medium-grained sandstones of the tidal channel/tidal bar association. The fine-grained sandstones are interpreted as deposited in a shallow marine environment. The lack of scouting and general plane-parallel lamination and small-scale structures suggest a relatively low-energy environment. However, none of the structures are diagnostic of a certain depositional environment. Based on association with the massive channel sandstones and the tidal channel/channel bar association we tentatively suggest that the sandstones were deposited in a tidal flat environment. Offshore silty mudstones Mica-rich, silty mudstones and fine-grained muddy sandstones dominate the association. The mudstones are generally poorly exposed, but locally show plane lamination. The mudstones contain ammonites, bivalves and gastropods preserved in concretionary horizons. Dinoflagellate cysts are abundant and diverse. Trace fossils are common and includes Planolites isp., Chondrites isp. and Munsteria isp. Kelly et al. (1998) also reported Zoophycos isp. from laminated mudstones at Stensi6 Plateau. Offshore silty mudstones dominate the Diener Bjerg, Stensi6 Plateau and Re~delv Members (Figs. 6 and 7). The general parallel bedding planes and fine grain size suggest a low-energy environment with deposi-

270

M. Larsen et al.

tion from suspension. The diverse dinoflagellate and trace fossil assemblages indicate fully marine conditions. The fine-grained mudstones are interpreted as deposited below storm-wave base in an offshore marine environment.

Offshore mudstones and graded sandstones The association consists of thick successions of laminated dark mudstones intercalated with laterally consistent, fine-grained sandstone beds up to 15 cm thick. The sandstones are graded and show parallel lamination and small-scale cross-lamination. The mudstones are generally mica-poor compared with the offshore silty mudstone association. B ioturbation is rare and included Planolites isp., Chondrites isp. and Munsteria isp. In Fosdalen, east of Diener Bjerg, the association contains large slumped blocks of Middle Jurassic? sandstones. The association is characteristic for the lower part of the Fosdalen Member (Fig. 8f). The mudstones were deposited from suspension in a low-energy marine environment. The lack of wavegenerated structures suggests that deposition occurred below storm-wave base. The intercalated sandstone beds are interpreted as formed from suspension fallout from sediment gravity flows. They may represent distal storm beds or low-density turbidites. The association is interpreted as representing a generally deeper water environment than the silty mudstone association and probably formed in an outer shelf environment.

Facies successions The basal Steensby Bjerg Formation forms an overall backstepping facies succession following exhumation and erosion in the earliest Cretaceous. The initial transgression is punctuated by progradation of sand-dominated deltaic and tidal successions separated by marine flooding events. The progradation was probably controlled by falling relative sea level and part of the succession represents incised valley fills. The coarse-grained elastic system was drowned in the early Albian. In order to illustrate the lateral differences along.the outcrop, three localities are described separately below.

Stensi6 Plateau The basal unit at Stensi6 Plateau consists of characteristic white, mica-rich sandstone up to 12 m thick, which lies with angular unconformity on the Middle Jurassic Pelion Formation (Ukl). It is well-sorted, fine- to medium-grained, shows faint cross-lamination and is tentatively included in the shoreface association marking the first transgressive deposits at this locality. The shoreface association (Stribedal

Member) reaches a maximum thickness of 37 m in the eastern part and a minimum thickness of 10 m towards the western end of the outcrop (Fig. 6). It consists of stacked shoreface units 2-5 m thick. The Stribedal Member is erosionally truncated by a more than 2 km wide and 23 m deep channel scour (SB 1) (Figs. 6 and 10). SB 1 forms the basal surface of an incised channel, filled in by massive and tidally influenced recumbent, cross-bedded sandstones of the B lfielv Member (Fig. 10). The channel fill has a channel-like geometry with a flat upper boundary to horizontally laminated mudstones approximately 20 m thick (Stensi6 Plateau Member). The offshore mudstones grade upward into fine-grained tidal flat sandstones of the Gulelv Member. Lenticular units of the massive channel sandstone association (SB2) occur in the thick tidal flat sandstone succession in a similar way as seen in the B lfielv Member below. However, the massive part of the sandstones are less prominent and the channel fill is dominated by recumbent cross-beds. The channel-fill sandstones are overlain by a monotonous succession of tidal flat sandstones forming a largescale coarsening-upward succession of 45 m thick. It is overlain by a 30 m thick unit of planar cross-bedded sandstones of the tidal bar association. A major Palaeogene volcanic sill disrupts the sandstones, but the succession probably continues without stratigraphic break above the sill where similar tidally influenced sandstones occur. A marked erosional surface (SB3) marked by large, rounded pebble and locally cobble-sized quartz is present in the middle of the tidally influenced sandstones and marks the turn-around point from the coarsening-upward to fining-upward grain-size trend. A thin, 1-2 m thick wave-bedded sandstone unit tops the succession. The shoreface sandstones are overlain by offshore mudstones of the ROdelv Member.

Steensby Bjerg The unconformity (Ukl) forming the base of the Gulelv Member is overlain by pebble conglomerates and coarse-grained sandstones of the shoreface association (Fig. 7). The conglomerate either forms a single fining-upward lag deposit 0.5-1 m thick or a unit up to 8 m thick of several amalgamated conglomerate beds. The conglomerates grade upward into a 24 m thick overall upward-fining succession of trough cross-bedded shoreface sandstones. The transgressive sandstones are erosionally overlain by the first of two progradational pulses of large-scale cross-stratified sandstone units of the delta plain and delta front association (Fig. 11). The first progradational pulse is represented by compound planar cross-beds and

Lower Cretaceous (Barremian-Albian) deltaic and shallow marine sandstones in North-East Greenland

i?_71

Fig. 10. At Stensi6 Plateau, Middle Jurassic sandstones are unconformably overlain by sandstones and conglomerates of the shoreface association (Shsst) (Stribedal Member). Massive channel sandstones (Msst) (Blfielv Member) erosionally truncate the up to 27 m stacked shoreface association. The erosional surface represents a distinct upper Barremian unconformity SB 1 interpreted as the base of an incised valley system. Person for scale.

trough cross-bedded sandstones, 15-25 m thick. The sandstones thin towards the east and are probably not present in the Diener Bjerg area (Fig. 7). The second progradational pulse consists of a 30-50 m thick delta unit locally with large-scale foreset beds (Fig. 7). It is topped by a pebble conglomerate (SB 1) interpreted as the most proximal facies in the succession. The deltaic units prograded towards the south and southwest. The upper delta unit prograded furthest into the basin and is probably the only deltaic unit present to the east at Diener Bjerg (see below). SB 1 is overlain by a fining-upward deltaic sandstone unit of trough and planar bedded sandstones. The content of organic

material increases upward and in the upper 10 m tidally influenced cross-beds with double mud-drapes occur for the first time in the succession. A 30-40 m thick unit of the tidal bar and tidal channel associations overlies the deltaic sandstones. The dominant structure is tidally influenced planar cross-bedding, but locally massive sandstone beds associated with recumbent deformed cross-bedding occur (Fig. 7). A wave-bedded shoreface sandstone unit, 5 m thick, containing large pebbles and cobble-sized clasts (SB2) truncates the tidal sandstones (Fig. 8e). The wave-bedded shoreface sandstones form the uppermost unit of the Gulelv Member. Laminated off-

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Fig. 11. At Steensby Bjerg coarse-grained sandstones of the Gulelv Member unconformably overlie Middle Jurassic sandstones (Ukl). The succession shown consists of shoreface sandstones (SHsst) overlain by two prograding units of the delta plain and delta front association (DFsst).

shore mudstones and siltstones of the ROdelv Member overlie the shoreface sandstones.

Diener Bjerg At Diener Bjerg Ukl is overlain by a lag conglomerate up to 50 cm thick forming the basal part of the Diener Bjerg Member. It is overlain by black, laminated mudstones of the offshore marine association, which forms an 18 m thick, slightly coarsening-upward succession (Fig. 7). A coarse-grained locally pebbly unit of the delta front association rather abruptly overlies the mudstones of the Diener Bjerg Member. The deltaic sandstones belong to the Gulelv Member and locally form a single large-scale foreset bed up to 45 m thick overlain by tidally influenced, planar cross-bedded sandstones up to 15 m thick (Fig. 12). At the easternmost outcrop the large-scale foreset bed contains a low-angle reactivation surface and may be divided into three progradational phases. The boundary to the tidally influenced cross-bedded unit is tentatively correlated with SB 1. The planar cross-bedded unit belongs to the tidal channel/bar association and is topped by a wave-reworked sandstone bed. The sandstones are overlain by offshore marine mudstones of the ROdelv Member. The boundary between the Gulelv andR0delv Members at Diener Bjerg is tentatively correlated with SB2 at Steensby Bjerg (Fig. 12), although the surface has not been walked out in the field.

Sequence stratigraphy and basin evolution A sequence stratigraphic interpretation of the Steensby Bjerg Formation can only be preliminary due to low biostratigraphic resolution in the coarse-grained facies associations and the fact that the sequence stratigraphic model is based on one, largely two-dimensional, outcrop. In the following we describe the basin evolution and discuss the sequence stratigraphic model (Figs. 13 and 14).

Early Barremian The angular unconformity forming the base Cretaceous in northern Hold with Hope has already been thoroughly discussed in the regional setting (see also Larsen et al., 1998). The unconformity spans the Late Jurassic (Volgian)-earliest Cretaceous (Hauterivian) and was named Ukl by Kelly et al. (1998). The unconformity probably formed by subaerial erosion following a major rifting event accompanied by faulting and block rotation. The unconformity was reworked by wave ravinement during marine transgression in earliest Barremian time. A lag with clast-size up to 25 cm and reworked Middle Jurassic ammonites formed by winnowing of older marine and probably fluvial deposits. The oldest dated deposits above Ukl are shoreface sandstones (Stribedal Member) of earliest

Lower Cretaceous (Barremian-Albian) deltaic and shallow marine sandstones in North-East Greenland

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Fig. 12. At the eastern Steensby Bjerg and Diener Bjerg the Gulelv Member unconformably overlies Lower Triassic strata (Ukl). The Gulelv Member consists of large-scale cross-bedded sandstones of the delta plain and delta front sandstone association (DFsst) overlain by tidally influenced sandstones. The sandstones are overlain by fine-grained offshore mudstones (Omst) of the R0delv Member.

Barremian (Nelchiopsis kostromiensis Subzone (I1)) age exposed to the west at Stensi6 Plateau (Fig. 4) (H. N0hr-Hansen, pers. commun., 1999). In Diener Bjerg to the east the unconformity is overlain by a basal conglomerate followed by mid-Barremian (Pseudoceratium anaphrissum Subzone, I2) mudstones (Diener Bjerg Member) representing the transgressive systems tract (Fig. 13). This may suggest that the transgression progressed from the southwest towards the northeast across a westward-dipping and possibly segmented fault block (Fig. 14).

Late early Barremian-late Barremian The rising sea level created accommodation for a thick sedimentary wedge along the ramp and axis of the old Mesozoic lineament. In late-early Barremian time the overall backstepping of the shallow marine system was succeeded by southward progradation of coarse-grained deltaic units of the highstand systems tract. During the first phase of progradation the lower of two deltaic units was deposited in the Steensby Bjerg area (Fig. 13). Following a minor sea-level rise, which shifted the delta back towards the north, the second deltaic unit prograded. This time the delta reached further towards the south and deltaic sediments were deposited across the former crest

of the fault block and into the Diener Bjerg area (Fig. 14). The relatively sharp base of the deltaic sandstones in this area may suggest that progradation occurred during a forced regression and the upper deltaic unit may thus represent a late highstandfalling stage systems tract of this sequence. The maximum southward progradation of the delta is marked by a pebble lag (SB1) at the top of the upper cross-stratified deltaic unit at Steensby Bjerg (Fig. 13). The lag consists of an up to 1.5 m thick gravel and pebble-rich structureless conglomerate with rare intercalated cross-bedded sandstones. Direct correlation across the river Gulelv is not possible, but we tentatively suggest that SB 1 correlates to a distinct unconformity truncating shoreface deposits of the Stribedal Member at Stensi6 Plateau. At Stensi6 Plateau, the surface can be followed laterally along the entire outcrop forming a deep channel interpreted as part of an incised valley system. The age of the unconformity (SB1) at Steensby Bjerg is uncertain, constrained only by the age of the oldest deposits above. These fining-upward delta sandstones and tidally influenced sandstones contain dinoflagellates of the early Aptian Pseudoceratium nudum Zone (II) (H. N0hr-Hansen, pers. commun., 1999). At Stensi6 Plateau the sediments below the unconformity (Stribedal Member) are of earliest Bar-

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Fig. 13. Geological cross-sections of the Lower Cretaceous succession. (a) The Steensby Bjerg-Diener Bjerg area. (b) The Stensi6 Plateau area. Colour coding indicates major sedimentary facies associations/depositional environments. Positions of the two cross-sections are shown in Fig. 5.

Lower Cretaceous (Barremian-Albian) deltaic and shallow marine sandstones in North-East Greenland

275

Fig. 14. Geological cross-section through the Hold with Hope area depicting the depositional history of the Lower Cretaceous sandstone-dominated system. The model shows the spatial distribution of the main sedimentary facies associations through time. Note that the main sediment transport direction was towards the south and southwest (towards the viewer). See text for further explanation.

remian (Nelchiopsis kostromiensis Subzone (I1)) age and the oldest fossiliferous sediments above the unconformity (Stensi6 Plateau Member) are of late Barremian-early Aptian age suggesting that part of the Barremian is missing (H. NOhr-Hansen, pers. commun., 1999) (Fig. 4). SB 1 is tentatively suggested to be formed in the late Barremian.

Late Barremian-early Aptian The massive sandstones of the B lfielv Member are interpreted as the lower part of the lowstand systems tract and overlie the unconformity SB1. The sandstones grade up into offshore marine mudstones of the Stensi6 Plateau Member representing the trans-

276 gressive systems tract and possibly include the maximum flooding surface. The offshore mudstones of the Stensi6 Plateau Member indicate a relative sea-level rise leading to a flooding of the sandstone-dominated system at Hold with Hope. The differences in water depth from below wave-base at Stensi6 Plateau to intertidal at Steensby Bjerg may have been governed by the topography created by valley incision (SB 1) (Fig. 14), or by a late Barremian-early Aptian westwards tilting of the fault block (Whitham et al., 1999). At the top of the Gulelv Member in Steensby Bjerg, well-rounded quartz cobbles up to 20 cm long are embedded in shoreface sandstones. Cobbles of this size are not present anywhere else in the deltaic succession and we suggest that they were originally deposited in a fluvial environment, reworked by marine processes during a following transgression. Sequence boundary SB2 is thus characterised by indirect evidence of removal of coarse-grained, most likely fluvial deposits of the falling stage and lowstand systems tract. It is dated as early Aptian by the age of the overlying mudstones of the R~delv Member (Kelly et al., 1998). At Stensi6 Plateau a candidate for the surface is the base of a massive, tidal channel fill in the lower part of the Gulelv Member reflecting a basinwards shift in facies. These channel-fill sandstones represent the lowstand systems tract and are overlain by a thick coarsening-upward succession of tidal flat sandstones.

Late Aptian-early Albian The final progradation of the sandstone-dominated system in Hold with Hope is represented by the large-scale coarsening-upward succession at Stensi6 Plateau representing the transgressive-highstand systems tract. Well-rounded quartz pebbles and cobbles up to 15 cm long are present in a lag deposit above an erosional surface near the top of the succession. This surface is a candidate for a sequence boundary, SB3. The unconformity formed in the late Aptian as indicated by dinoflagellate cysts in the underlying tidal flat sandstones and by the age of the overlying mudstones of the ROdelv Member (Kelly et al., 1998). The lag is overlain by a fining-upward succession of tidal channel/bar sandstones overlain by a thin shoreface sandstone unit of the transgressive systems tract. The sandstone-dominated deposits correlate with offshore silty mudstones (ROdelv Member) in the Steensby Bjerg-Diener Bjerg area, indicating that the depocentre for the tidal sandstones was situated in the west (Fig. 14). This implies a westward shift in the position of the depocentre from the Barremian to the Aptian. Lack of stratigraphic control and truncation

M. Larsen et al.

by SB4 and Ut in the eastern part of the outcrop area, however, hinders conclusive interpretation of the palaeogeography and distribution of the depositional systems at this stage. In early Albian time, the coarse-grained depositional system of the Hold with Hope area was drowned and offshore marine conditions prevailed.

Late-early Albian-Cenomanian SB4 forms a tectonically controlled, erosional surface described and interpreted in detail by Kelly et al. (1998, their U~:2). The unconformity surface represents a marine surface of erosion truncating strata of Early Triassic, Middle Jurassic, and Early Cretaceous age. The youngest sediments below the SB4 are earliest-middle Albian mudstones of the RCdelv Member (Kelly et al., 1998). The surface is overlain by marine mudstones with thin-bedded sandstones of middle Albian age (Fosdalen Member). The unconformity can be traced south along the Gulelv where the offshore marine mudstones and siltstones of the ROdelv Member are overlain by mudstones and thin-bedded turbidites of the Fosdalen Member with an angular unconformity (Fig. 8f). Kelly et al. (1998) interpreted SB4 as the eastward-facing slope of a degraded fault block that was gradually transgressed during the middle Albian-early Cenomanian (Fig. 14). The onlapping succession consists of offshore mudstones with thin sandstone turbidites of the Home Forland Formation indicating a change to a relatively deep, mud-dominated shelf environment.

Regional implications The Lower Cretaceous sandstones form a new reservoir unit in the North-East Greenland stratigraphy and may in certain areas show to be equally important as Middle Jurassic reservoir units well known from the Norwegian Shelf and the northern North Sea. In Hold with Hope, the Lower Cretaceous sandstones to a large extent originated from extensive reworking and erosion of the Pelion Formation, thus locally destroying the classic Middle Jurassic sandstone play. The presence of a thick proximal succession at Hold with Hope suggests that the area acted as a major sediment entry point to the East Greenland shelf during the Early Cretaceous. It is important to stress that sediments in the Lower Cretaceous were transported in southerly and southwesterly directions parallel to the faults bounding the Mesozoic rift basins (Fig. 15). It is probably not until the middle Albian-Turonian that the Jurassic-Lower Cretaceous rift topography was filled in, opening for sediment

Lower Cretaceous (Barremian-Albian) deltaic and shallow marine sandstones in North-East Greenland

HoM with Hope----

I

[

50km

I-/-8-'

Fluvial conglomerates (inferred)

[::ii:::i~i:i,:::iiDeltaic :i:i] and shallow marine sandstones (outcrop) L---~ Offshore mudstones (inferred) ~

deep water sandstones of the Agat field west of the Norwegian coast (Gulbrandsen, 1987) and the transgressive shallow marine sandstones of Victory field west of the Shetland Islands (Goodchild et al., 1999). Onshore paralic to shallow marine equivalents, although without hydrocarbon potential, are known from the Kangerlussuaq area in southern East Greenland (Larsen et al., 1996, 1999) and in Svalbard (Helvetiafjellet Formation) (Gjelberg and Steel, 1995). Both the onshore and offshore successions have been interpreted as reflecting overall transgression following uplift and erosion of the basin margins in the Hauterivian-Barremian. However, offshore Norway the transgressive, paralic to shallow marine sandstone play of the Lower Cretaceous is a hitherto unproven play concept.

] Erosion/non-deposition

~ ) 1 Offshore mudstones (outcrop) ~

277

Normal fault

Direction of delta progradation Fig. 15. Palaeogeographic reconstruction (Barremian-Aptian) of the North-East Greenland margin. Note the position of the coarsegrained deltaic and shallow marine system of the Steensby Bjerg Formation south of a major relay ramp in the Mesozoic fault system. The dominant sediment transport direction was towards the south, parallel to the major faults delineating the Mesozoic rift basins.

transport across the East Greenland shelf and into basinal areas to the east (Whitham et al., 1999). This can be seen by a change in depositional systems towards deeper water facies and a 90 ~ shift in palaeocurrent directions from the Lower (south and southwesterly) and into the Upper Cretaceous (east) (Stemmerik et al., 1997; Surlyk and Noe-Nygaard, 2001). Based on comparison with the Hold with Hope succession, Lower Cretaceous sandstone wedges may be predicted to occur along older Mesozoic lineaments in the deeply buried basins offshore Norway and the Faeroe-Shetland Basin. These sandstone wedges represent a stratigraphic play concept along the structural highs, which can open a new exploration playground. Attention should be focused in areas with major translation of fault zones that are linked to a large updip fluvial catchment area. Barremian-Albian, hydrocarbon-bearing sandstone wedges are found at several places in the northern North Atlantic both offshore, such as the

Acknowledgements The study of the Cretaceous sediments in NorthEast Greenland forms part of the project "Resources of the sedimentary basins of North and East Greenland" supported by the Danish Research Councils. We also acknowledge financial support by SAGA Petroleum a.s.a. We are grateful to H. NOhr-Hansen (GEUS) for providing the essential stratigraphical control based on dinoflagellate cysts. H. Amundsen and S. Prosser (Saga Petroleum) and J. Therkelsen and H. Vosgerau (GEUS) are thanked for good companionship and inspiring discussions during field work in East Greenland. We wish to thank reviewers D. Leckie and A. Ryseth for their comments and suggestions to improve an early version of the manuscript and A.G. Whitham for discussions on sedimentology and stratigraphy. The paper is published with permission of the Geological Survey of Denmark and Greenland.

References Allen, J.R.L. and Banks, N.L., 1972. An interpretation and analysis of recumbent-folded deformed cross-bedding. Sedimentology, 19: 257-283. Donovan, D.T., 1957. The Jurassic and Cretaceous systems in East Greenland. Medd. GrOnl., 155(4), 214 pp. Gjelberg, J. and Steel, R.J., 1995. Helvetiafjellet Formation (Barremian-Aptian), Spitsbergen: characteristics of a transgressive succession. In: R.J. Steel, V. Felt, E.E Johannessen and C. Mathieu (Editors), Sequence Stratigraphy on the Northwest European Margin. Norwegian Petroleum Society (NPF), Special Publication 5. Elsevier, Amsterdam, pp. 571-593. Goodchild, M.W., Henry, K.L., Hinkley, R.J. and Imbus, S.W., 1999. The Victory gas field, west of Shetland. In: A.J. Fleet and S.A.R. Boldy (Editors), Petroleum Geology of Northwest Europe: Proceedings of the 5th Conference. Geological Society, London, pp. 713-724. Gulbrandsen, A., 1987. Agat. In: A.M. Spencer. C.J. Campbell, S.H. Hanslien, E. Holter, EH.H. Nelson, E. Nys;ether and G. Ormaasen (Editors), Geology of the Norwegian Oil and Gas Fields. Graham and Trotman, London, pp. 363-370.

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278 Haller, J., 1970. Tectonic Map of East Greenland (1:500,000). Medd. GrCnl., 171(5), 287 pp. Kelly, S.R.A., Whitham, A.G., Koraini, A.M. and Price, S.M., 1998. Lithostratigraphy of the Cretaceous (Barremian-Santonian) Hold with Hope Group, NE Greenland. J. Geol. Soc. London, I55: 993-1008. Koch, L., 1931. Carboniferous and Triassic Stratigraphy of East Greenland. Medd. GrCnl., 83(2), 99 pp. Larsen, M., Hamberg, L., Olaussen, S. and Stemmerik, L., 1996. Cretaceous-Tertiary pre-drift sediments of the Kangerlussuaq area, southern East Greenland. Bull. GrCnl. Geol. Unders., 172: 37-41. Larsen, M., Piasecki, S., Preuss, T., Seidler, L., Stemmerik, L., Therkelsen, J. and Vosgerau, H., 1998. Petroleum geological activities onshore East Greenland in 1997. Geol. Greenl. Surv. Bull., 180: 35-42. Larsen, M., Hamberg, L., Olaussen, S., NCrgaard-Pedersen, N. and Stemmerik, L., 1999. Basin evolution in Southern East Greenland: an outcrop analogue for Cretaceous-Paleogene basins on the North Atlantic volcanic margins. Bull. Am. Assoc. Pet. Geol., 83: 1236-1261. Martin, C.A.L. and Turner, B.R., 1998. Origins of massive-type sandstones in braided river systems. Earth Sci. Rev., 44: 15-38. Martinsen, O.J., 1994. Evolution of an incised-valley fill, the Pine Ridge sandstone of southeastern Wyoming, USA: systematic sedimentary response to relative sea-level change. In: R.W. Dalrymple, R. Boyd and B.A. Zaitlin (Editors), Incised-Valley Systems: Origin and Sedimentary Sequences. Soc. Econ. Paleontol. Mineral. Spec. Publ., 51: 109-128. Maync, W., 1949. The Cretaceous beds between Kuhn Island and Cape Franklin (Gauss Peninsula), northern East Greenland. Medd. Gr~nl., 133(3), 291 pp. Stemmerik, L., Christiansen, EG., Piasecki, S., Jordt, B., Marcussen, C. and Nchr-Hansen, H., 1993. Depositional history and petroleum geology of Carboniferous to Cretaceous sediments in the northern part of East Greenland. In: T.O. Vorren, E. Bergsager, ~.A. Dahl-Stamnes, E. Holter, B. Johansen, E. Lie and T.B. Lund (Editors), Arctic Geology and Petroleum Potential. Norwegian Petroleum Society (NPF), Special Publication 2. Elsevier, Amsterdam, pp. 67-87. Nielsen, E., 1935. The Permian and Eotriassic vertebrate bearing beds at Godthaab Gulf (East Greenland). Medd. GrCnl., 98(1), 111 pp. Nio, S.D. and Yang, C.S., 1991. Diagnostic attributes of clastic tidal

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deposits: a review. In: D.G. Smith, G.E. Reinson, B.A. Zaitlin and R.A. Rahmani (Editors), Clastic Tidal Sedimentology. Can. Soc. Pet. Geol., Mere., 16: 3-27. NChr-Hansen, H., 1993. Dinoflagellate cyst stratigraphy of the Barremian to Albian, Lower Cretaceous, North-East Greenland. Bull. GrCnl. Geol. Unders., 166, 171 pp. Price, S.R and Whitham, A.G., 1997. Exhumed hydrocarbon traps in East Greenland: analogs for the Lower-Middle Jurassic play of NW-Europe. Bull. Am. Assoc. Pet. Geol., 81: 196-221. Stemmerik, L., Clausen, O.R., Korstggtrd, J., Larsen, M., Piasecki, S., Seidler, L., Surlyk, F. and Therkelsen, J., 1997. Petroleum geological investigations in East Greenland: project "Resources of the sedimentary basins of North and East Greenland". Geol. Greenl. Surv. Bull., 176: 29-38. Surlyk, E, 1978. Submarine fan sedimentation along fault scarps on tilted fault blocks (Jurassic-Cretaceous boundary, East Greenland. Bull. GrCnl. Geol. Unders., 128, 109 pp. Surlyk, E, 1990. Timing, style and sedimentary evolution of Late Palaeozoic-Mesozoic extensional basins of East Greenland. In: R.EP Hardman and J. Brooks (Editors), Tectonic Events Responsible for Britain's Oil and Gas Reserves. Geol. Soc. London Spec. Publ., 55: 107-125. Surlyk, F. and Noe-Nygaard, N., 2001. Cretaceous faulting and associated coarse-grained marine gravity flow sedimentation, Traill O, East Greenland. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 293-319 (this volume). Surlyk, E, Hurst, J.M., Piasecki, S., Rolle, E Scholle, EA., Stemmerik, L. and Thomsen, E., 1986. The Permian of the western margin of the Greenland Sea - - a future exploration target. In: M.T. Halbouty (Editor), Future Petroleum Provinces of the World. Am. Assoc. Pet. Geol., Mem., 40: 629-659. Upton, B.G.J., Emeleus, C.H. and Hald, N., 1980. Tertiary volcanism in northern E Greenland: Gauss Halve and Hold with Hope. J. Geol. Soc. London, 137: 491-508. Vischer, A., 1943. Die Post-Devonische tektonik von Ostgr6nland zwischen 74~ to 75~ Kuhn ~, Wollaston Forland, Clavering 0 und angrenzende gebiete. Medd. GrCnl., 133(1): 1-194. Whitham, A.G., Price, S.P., Koraini, A.M. and Kelly, S.R.A., 1999. Cretaceous (post-Valanginian) sedimentation and rift events in NE Greenland (71-77~ In: A.J. Fleet and S.A.R. Boldy (Editors), Petroleum Geology of Northwest Europe: Proceedings of the 5th Conference. Geological Society of London, pp. 325-336.

Geological Survey of Denmark and Greenland (GEUS), Thoravej 8, DK-2400 Copenhagen NV, Denmark Norsk Hydro Exploration, N-0246 Oslo, Norway Norsk Hydro Research Centre, P.O. Box 7190, N-5020 Bergen, Norway

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The depositional history of the Cretaceous in the northeastern North Sea Tom Bugge, Bjarne Tveiten and Sven B&ckstr6m

The Cretaceous of northwest Europe seems to have had periods of significant tectonic activity, and was not only a period of passive, post-rift infilling as earlier was the dominant view. Late Jurassic rifting probably continued into the Cretaceous and lasted until the late Barremian. During Hauterivian-Barremian times some of the local structural highs and the eastern basin margin of the northeastern North Sea were uplifted, and a relative lowering of sea level in the latest Barremian resulted in widespread and locally deep erosion. The resulting relief was therefore more important in some places for the Cretaceous depositional history than the Base Cretaceous relief. The Barremian landscape was largely filled in by the end of the Early Cretaceous, and all structural highs and most of the eastern basin margin were flooded in the Santonian. A marine environment prevailed during the entire Cretaceous period with deposition of hemipelagic clay, interrupted by sand deposition in certain periods. The sand was deposited in a slope-to-basin-floor setting by gravity mass flows, primarily debris flows and turbidity currents, and was probably sourced from shallow marine sands in the east. The sand transport was concentrated in certain fairways and followed topographic lows. Some of the Albian sands in the Agat area seem to be deposited in local slide scars, thus explaining why there is no pressure communication between large sandstone bodies. Upper Turonian-Coniacian sands were deposited after the topography was filled in and have a slope-fan geometry. Structural closures of Cretaceous strata are few and limited in extent in the northeastern North Sea, and any hydrocarbon prospectivity will depend purely or partly on stratigraphic closure.

Introduction This paper presents results obtained in a regional study of the Cretaceous from the northeastern North Sea. The study area comprises the northern part of the North Viking Graben, the Sogn Graben and the terraces and slope area located to the south and east (Fig. 1). The Fram and Gj~a Fields are located in the southern and central part of the area, with oil discoveries in Jurassic strata. The Agat Field is situated in the north with gas discoveries in the Lower Cretaceous, while one of the recent wells in the Gj~a Field recorded gas and oil shows in the Upper and Lower Cretaceous, respectively. Structural closures of Cretaceous horizons are rare in this part of the North Sea and any economic hydrocarbon discovery relies on stratigraphic trapping. This poses a challenge that can only be solved through enhanced understanding of the regional geological development of the area, as well as of the depositional history and sedimentary processes. There is a common view that the Cretaceous in the northern North Sea was characterised by passive post-rift infilling of the topography caused by the Late Jurassic rifting. Gabrielsen et al. (1999) concluded that extension was initiated later and ter-

minated earlier in the northern North Sea than in the M~re Basin. The present study area is in the northernmost North Sea and touches the southern part of the M~re Basin. It is suggested here that rifting probably continued into the Early Cretaceous and that smaller extensional and vertical tectonic episodes with uplift of basin flanks and local highs and basin floor subsidence occurred throughout the Cretaceous period. Gravity-driven mass flows seem to have been the dominant process to transport sand into the northeastern North Sea Basin, and it is suggested that many of these can be linked to periods of increased tectonic activity. About 30 exploration wells have been drilled to date within this part of the North Sea, and there is a significant amount of 2D seismic data and fairly good coverage of 3D seismic data (Fig. 2). All of these data have been integrated in this study.

Methodology (a) A consistent biostratigraphic zonation was established for the area, and this was used to subdivide each well in a chronostratigraphic manner. Palaeowater depth was interpreted where the wells have suitable micropalaeontological coverage.

Sedimentary Environments Offshore Norway - Palaeozoic to Recent edited by O.J. Martinsen and T. Dreyer. NPF Special Publication 10, pp. 279-291, Published by Elsevier Science B.V., Amsterdam. 9 Norwegian Petroleum Society (NPF), 2001.

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Fig. 1. The study area is in the northeastern part of the North Sea with the North Viking Graben and Sogn Graben in the west and comprises the terraces and slope area to the north of the Troll Field. The Fram and Gj0a Fields contain oil in the Jurassic, while gas and oil are proven in the Cretaceous in the Agat and Gjca areas. Contours in two-way-timeto Base Cretaceous.

T. Bugge et al. considered, i.e. the deep-water sandstones of presumed turbiditic origin were not included in the analysis. The cycles identified are assumed to represent second-order transgressive and regressive events, although in a deep-water environment such an assumption must be regarded as tentative. (c) An initial well-to-well log correlation was performed using the biostratigraphic data as constraints. (d) A number of seismic units and seismic markers were mapped over parts of, or over the entire, area. This enabled a quality control of the log correlation, and in several cases it was necessary to modify the log correlation to reach a consistent interpretation. Log and seismic interpretation was thus done as an iterative process. (e) Regional data were studied in order to improve the exact timing of different events. The work of De Graciansky et al. (1998) and Jacquin et al. (1998) proved to be especially valuable due to its documentation of a detailed compilation of data from all European basins. Our data set was examined in order to assess the uncertainty of the age definition of the second-order cycle boundaries. The ages of the cycle boundaries were adjusted to match those of De Graciansky et al. (1998) if they fell within the range of uncertainty of the data set available for this study. This was relevant for the Aptian-Cenomanian and the Santonian intervals.

Geological setting

Fig. 2. The database comprises about 30 exploration wells, a dense grid of 2D seismic lines and a fairly good coverage of 3D seismic data (shaded area). (b) A detailed log analysis was performed in each well, identifying upward fining and coarsening cycles. In doing this, only the background sedimentation was

The Cretaceous lithostratigraphy has been described by different workers and informal and formal units are defined. Deegan and Scull (1977) compiled the first lithostratigraphic scheme for the North Sea; Hesjedal and Hamar (1983) revised this scheme and proposed some new informal units in the Lower Cretaceous. They also described the palaeogeography and palaeoenvironment in the southern Norwegian North Sea. Isaksen and Tonstad (1989) defined a formal lithostratigraphy for the entire Norwegian North Sea, which is still valid. Relatively few other papers dealing with the Cretaceous in the northern North Sea have been published so far. Gulbrandsen (1987) described the hydrocarbon discoveries in the Agat Field, where gas and condensate were found in the Lower Cretaceous intervals in the 1970s. There is no pressure communication between the wells over distances of a few kilometres, and the reservoirs are interpreted as sands deposited by turbidity currents. Alternative interpretations were presented by Shanmugam et al. (1995), Shanmugam and Moiola (1995) and Skibeli et al. (1995), who interpreted the Agat sandstone intervals as deposits from local slumping and massive, sandy debris flows.

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Fig. 3. Seismic lines illustrating the geological setting. (A) Line from the Sogn Graben through the Agat area. (B) Line from the North Viking Graben in the west through the Fram Field towards the coast. Note the thick sequence of .~sgard Fm. sediments (Ryazanian-Barremian) in the grabens and as erosional remnants in the east on both lines. The overlying Cretaceous sediments show an overall onlapping geometry on the local structural highs and the slope area to the east.

282 This latter interpretation has been disputed by several authors (AAPG Bulletin Volume 81, with reply by Shanmugam and Moiola, 1997, in the same volume). Skibeli et al. (1995) interpreted the Lower Cretaceous interval in a sequence stratigraphic context and divided the succession into sequences and supersequences. Jacquin and Thomson (1999) took this further and integrated the whole of the Cretaceous into a structural and sequence stratigraphic framework constructed for the Triassic to Tertiary. De Graciansky et al. (1998) presented a stratigraphic scheme of second-order facies cycles in Europe. Physiographically, the study area is characterised by the North Viking Graben and the Sogn Graben in the west, and terraces/structural highs and slopes in the south and east (Fig. 3). The North Viking Graben is a half-graben fault-bounded to the west, while the Sogn Graben is a true graben. Fig. 3 shows that both basins were infilled during the Early Cretaceous, probably from the late Ryazanian to the end of the Barremian. In both the southeast and the northeast there are erosional remnants of sediments that were deposited during the same period. The deeper basins of the M0re/V0ting area to the north also received great quantities of sediment during this period. On the shallower platform areas in the North Sea and offshore mid-Norway, transgression and reduced clastic input led to deposition of shallow marine carbonates during the same time interval. In addition, the separation between the platform areas and the terraces offshore mid-Norway became more pronounced. Blystad et al. (1995) regarded this to be the response to the last phase of the Late Jurassic rifting event. The Aptian/Albian was characterised by deposition of clay with some significant sand pulses. It is interpreted as a tectonically quiet period, but it is apparent that some of the boundary faults on the northern part of the mid-Norwegian shelf became active. The Upper Cretaceous succession shows an overall onlapping geometry in the study area (Fig. 3). Onlap features on some internal horizons (see below) are taken as evidence of tilting due to tectonic activity at certain time intervals. In contrast, the Late Cretaceous was a tectonically quiet period in the central and southern North Sea, with reduced clastic input and deposition of pure carbonate strata (Isaksen and Tonstad, 1989). The northern North Sea and the margin offshore mid-Norway were characterised by rapid deposition and the accumulation of a tremendous thicknesses of hemipelagic, argillaceous and calcareous marine sediment, interrupted by episodic sand deposition (Dalland et al., 1988; Isaksen and Tonstad, 1989). The rate of sand input was particularly high in the Vcring Basin during the Santonian/Campanian (HCgseth et al., 1999).

T. Bugge et al.

Depositional history of the Cretaceous succession The depositional history during the Cretaceous in the northern North Sea is thought to have been influenced by the Late Jurassic rifting and the structural relief formed during the rifting. A combination of the remnant relief and the regional basin trends means that the depositional environment in the early/middle Ryazanian was characterised by structurally controlled, restricted basins with anoxic conditions, leading to the deposition of the organicrich Draupne Formation shales. In the late Ryazanian the basins became less isolated and more oxygenated (e.g. Hesjedal and Hamar, 1983). This is interpreted to represent the time corresponding to "The Base Cretaceous" reflector. A deep marine setting developed across the study area following the Barremian and remained as such for the remaining part of the Cretaceous; no sediments representing the shallow marine and coastal systems seem to be preserved.

The Asgard Formation (late Ryazanian-Barremian) Up to 400 m of sediments of late Ryazanian to Barremian age are preserved close to the Norwegian coast. It is interpreted from seismic data that sediments of the same age constitute the fill in the Sogn Graben (> 800 m) and the North Viking Graben (> 300 m). Elsewhere, condensed units of this age are occasionally present, as proven by drilling. The preserved sediments near the present-day coast are seen as spectacular highs with well-developed erosional edges (Fig. 3). Internal reflection patterns are parallel and often conform to the underlying Upper Jurassic shallow marine strata. The Lower Cretaceous sediments in these highs were drilled in wells 36/7-2 and 36/1-1 and are also interpreted to be of shallow marine origin. There is no structural break across the Base Cretaceous unconformity. Along the axes of the Sogn and North Viking Grabens the Cretaceous reflections are also parallel with the underlying Jurassic strata, while they onlap the Base Cretaceous on the flanks of the basin. The environment in the grabens is assumed to be deep marine, based on palaeogeographic reconstruction from seismic data. It is proposed that sedimentation during the earliest Cretaceous took place in two basin systems: one following the Late Jurassic rift axis in the Viking/Sogn Grabens and the other along the present-day coast of Norway, possibly related to the Oygarden Fault Zone. The Oygarden Fault Zone consists of a series of down-to-the-east faults and is probably still active (F~erseth et al., 1995). The area between the two basin

The depositional history of the Cretaceous in the northeastern North Sea

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uplift during Hauterivian-Barremian and subsequent erosion is well known from most of the north Atlantic region, e.g. west of Shetland, the mid-Norway area, Barents shelf and Greenland (e.g. Dor6, 1992). Only a few of the drilled wells contain sediments of late Ryazanian/early Valanginian age, i.e. the oldest sediments above "The Base Cretaceous" boundary (Fig. 6). These wells were situated in a basin or slope position, while the wells missing this section seem to be located on structurally higher areas. A greater number of wells record Hauterivian-aged deposits, while most wells have sediments of Barremian age. This is interpreted as reflecting an overall transgressive trend lasting until the late Barremian. This was followed by a drastic lowering of relative sea level during the latest Barremian, which was related to the above described basin flank uplift and structuring. The Sola Formation (Aptian)

Fig. 4. Geological setting during deposition of late RyazanianBarremian (Asgard Fm.) sediments. Shallow marine sediments were deposited in "inner basins" and deep marine sediments in the North Viking Graben and the Sogn Graben. The area between was partly bypassed and partly covered by relatively thin deposits.

systems is interpreted as a region that received little sediment and was mainly a zone of bypass (Fig. 4). It is proposed that strong inversion of the eastern basin occurred during the late Barremian, and by the end of Barremian these uplifted sediments were strongly eroded, as seen in Fig. 5. A similar history with 36/1-1

b)

a) bypass/ thin deposits

SognNiking Graben

Inner basin

Fig. 5. Model sketch of Early Cretaceous inversion. (a) The late Ryazanian-Barremian (Asgard Fm.) sediments deposited in the "inner basins" (see Fig. 4) were (b) probably inverted when the basin flanks were uplifted and basin floor subsided in the Hauterivian-Barremian. This probably represented the end of the Late Jurassic rifting.

The Sola Formation is recognised in well logs as a highly radioactive shale of Aptian age, and is defined as a formation as such (Isaksen and Tonstad, 1989). We interpret the erosional products resulting from the sea level lowering in the latest Barremian to have been deposited in the latest Barremian-earliest Aptian. Such deposits have been drilled in one of the wells in the Selje High area immediately to the north of the Agat area, but have only been observed on seismic data in the Sogn Graben within the study area. Low-amplitude "background" reflections are seen with a few high-amplitude features internally. Mapping reveals the high-amplitude events as having an elongate shape coming out of an interpreted palaeovalley running westwards across the southern part of the Agat area. This high-amplitude seismic facies can be traced across the Sogn Graben before bending northwards when reaching the Marflo Ridge (Fig. 7). The high amplitudes are interpreted as the response of possible turbidite sandstone systems encased in the "background" shales. The Aptian succession is condensed and incomplete in most of the wells drilled. Seismic data can be interpreted to suggest that Aptian sediments exist in the deeper parts of the area, including the North Viking Graben and Sogn Graben. There is 90 m of Aptian sediment in the westernmost of the Agat wells (35/3-1), but it pinches out between wells 35/3-2 and 35/3-4. Seismic onlap and pinchout are observed further south along the basin margin. The dark shales of the Sola Formation are characterised by high gamma radiation and were deposited in a deep marine environment. A high gamma-ray level in the lower Aptian is interpreted to represent maximum flooding and is correlated with deposition of

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Fig. 6. Chronogram of selected wells, based on integration of biostratigraphic zonation data, seismic data and well data. The figure shows that many sequences are incomplete in the drilled wells, particularly in the Lower Cretaceous. It is important to notice that sequences of the same age are deposited and preserved in many of the wells. This probably indicates that the drilled areas have not been subaerially exposed and eroded.

the organic-rich clays of the Fischschiefer in Germany (Kemper, 1973; Riley et al., 1992). Together with the later mentioned earliest Turonian flooding event, this is one of the major transgressions in the Cretaceous. The rapid regression in the late Barremian was thus succeeded by rapid transgression in the earliest Aptian. The lack of younger Aptian sediments in most of the wells is interpreted as the combined result of the topographic relief left by the Barremian structuring, a relative lowering of sea level and a general regressive trend until the end of the Aptian.

The Redby and Agat Formations (Albian) The Albian period was dominated by hemipelagic deposition of clay interrupted by the sandy products

of mass-flow events. Lithologically, the claystone is assigned to the RCdby Formation and the sandstone to the Agat Formation (Isaksen and Tonstad, 1989). Gamma-ray (GR) log patterns show a stePwise reduction in the radioactivity level of the background shales with time (Fig. 8). High radioactivity is taken as an indicator of a restricted environment. The microfauna does not indicate any changes in palaeowater depth during this period. The GR trend is therefore interpreted as a stepwise opening of the basin from the early Albian to a fully open marine environment by the end of the Albian. Sandstones belonging to the Agat Formation occur frequently in the sequence. With the exception of well 35/3-5, most of the sandstones seem to be of late A1bian age. The units appear as anomalous sediments

The depositional history of the Cretaceous in the northeastern North Sea

Fig. 7. Erosional products from the latest Barremian-earliest Aptian erosion were deposited in the Sogn Graben. The lobe is interpreted as deposits from turbidity currents running in a palaeovalley through the Agat area, being deflected northward in the Sogn Graben. The map shows the high-reflectivity lobe superimposed on the Base Cretaceous surface (both as depth in seismic reflection time).

in an environment characterised by hemipelagic deposition of clay and were transported into the system as gravity mass flows. This activity started in the early/middle Albian, but increased dramatically by the middle/late Albian transition and can be traced as such

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in wells over a wide area. It is worth noticing that biostratigraphic reanalysis indicates that none of the sandstones are of Aptian age. Although Shanmugam et al. (1995) suggested that many of the sandstone units were mass deposits from large-scale slumping, there is general agreement that the apparently massive sands are the result of amalgamation of thin turbidite beds or sandy debris flows (Nystuen, 1999). From well data three possible transgressive/regressive cycles can be recognised in the Albian succession (Fig. 6). These can be correlated with the sequence stratigraphic subdivisions presented for European basins by De Graciansky et al. (1998) and Jacquin et al. (1998), who classified them as secondorder facies cycles. Most of the lower and middle Albian sandstones are interpreted as part of second-order regressive cycles, while most of the upper Albian sandstones are placed within a second-order transgressive cycle (wells 35/3-1, -2, -4, 35/9-3). Because the Agat sandstones are so widespread but restricted to certain time intervals, it could be suggested that they are related to periods of higher tectonic activity and frequency of earthquakes. There is, however, no direct evidence in seismic data of increased tectonic activity such as active faulting or change of basin configuration, tilting, onlap surfaces etc. Except for the Agat wells, most of the other wells contain upper Albian sediments, but lack sediments of early to middle Albian age (Fig. 6). Being located on structural highs or on the upper slope, this suggests that they first were bypassed and then flooded by late Albian time. This could further

Fig. 8. Albian correlation between Agat wells 35/3-1 and -5 and GjOa well 35/9-3. There is a stepwise reduction in maximum gamma-ray level throughout the Albian, which is interpreted to reflect a gradual opening of the basin to more oxygenated conditions. The prevailing hemipelagic deposition of clay was interrupted by episodic deposition of sand by gravity mass flows (Agat Fm.). Most of the Agat sand was deposited in the early part of the late Albian, with the exception of well 35/3-5, which contains both older and younger sands.

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The Svarte Formation (Cenomanian)

Fig. 9. (a-c) In the Agat area sand was probably deposited by debris flows and turbidity currents in a slope setting. The turbidites are typically 10-30 cm thick and were amalgamated to thicknesses of many tens of metres. In a bypass situation on the slope this would hardly occur. Seismic data indicate, however, that slide scars from small-scale slumping and sliding formed accommodation space for the sand to be preserved in large, isolated bodies. This would also explain why there is no pressure communication between the Agat sandstone in wells 35/3-2 and -4, which are only 5.5 km apart in an updip direction.

Deposition of hemipelagic clay seems to have continued into the Cenomanian period, and sand input by gravity mass flows decreased. There are some seismic indications that the Svarte Formation sediments onlap the underlying Lower Cretaceous sediments, thus suggesting a minor tectonic event with basin floor subsidence/flank uplift. The Svarte Formation can be divided into two subunits. The lower unit has been dated to early Cenomanian and seems to infill and smoothen some of the underlying relief. The upper part is of late Cenomanian age and has a more even thickness and more widespread distribution. The two subunits seem to correspond to the facies cycles indicated in Fig. 6. There is an overall thickening trend of the Svarte Formation towards the basin centre in the west and a slight thickening to the north. Palaeobathymetric interpretation indicates a general transgression throughout the Cenomanian, which is coincident with the last phase of the overall Early to mid-Cretaceous transgression.

The Blodeks Formation (latest Cenomanian-early Turonian) indicate that the Agat sands were sourced from sandrich areas that were flooded during a period of rising relative sea level. Gas and condensate were found in the Agat sandstones in some of the wells drilled in the Agat area (35/3-1, 35/3-2, and 35/3-4), but there appears to be no communication between the wells. The sandstones are therefore interpreted to occur as isolated bodies. These wells are located on the northern flank of the earlier mentioned E-W-striking palaeovalley, probably in a palaeogeographic slope setting with the basin floor in the Sogn Graben to the west and in the Mere Basin further out to the north. This part of the Agat area was therefore generally a bypass area for the turbidity currents. The seismic data quality is not the best, but there are indications that local slumping and sliding occurred. This would have created accommodation space for thin turbidites and debris flow deposits to amalgamate to thicknesses of tens of metres, thus creating sand bodies with no internal communication (Fig. 9). Well 35/3-5 is located more in the centre of the palaeovalley and contains thicker sandstone units that tend to follow the strike of the valley. These sandstone units may have more internal communication, but seem to be isolated from the sandstone bodies encountered in the above-mentioned wells on the northern flank of the valley.

The long-term Early to mid-Cretaceous transgression is part of the North Atlantic first-order cycle described by Jacquin et al. (1998). The cycle culminated in the latest Cenomanian/early Turonian with deposition of a condensed section of organic-rich clay referred to as the B lodCks Formation in the area of study. This formation represents the most pronounced transgression in the Cretaceous and is equivalent with a known condensed interval from Svalbard to Italy (De Graciansky et al., 1984). It corresponds to a well-known oceanic anoxic event with starvation and can be correlated with the Plenus Marl Formation further south in the North Sea (Deegan and Scull, 1977). The typical thickness in the study area is only a few metres and is therefore below seismic resolution. The impedance contrast is, however, high and the B lod~ks Formation is recognised as a strong and continuous reflection on seismic data (Fig. 10). It is recognised on well logs by high gamma-ray values, and the drilled wells show that the formation is absent on some of the structural highs and the highest parts of the eastern basinal slope. This could be explained by later erosion or by non-deposition. Because lower Turonian sediments exist where B lodCks sediments are missing, any erosion must have occurred very soon after deposition. Alternatively, B lod~ks Formation sediments bypassed and were never deposited on the structural highs and on the entire eastern slope of the basin.

The depositional history of the Cretaceous in the northeastern North Sea

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Fig. 10. The BlodCks Fm. (late Cenomanian-early Turonian age) represents the most pronounced transgression in the Cretaceous. It was followed by a significant relative lowering of sea level before the Tryggvason Fm. transgressed and backstepped the Blod0ks Fm. Downlap features in the basin represent the regressive phase, while the backstepping is expressed by onlap on the eastern basin margin. The overlying Kyrre Fm. backstepped the Tryggvason Fm. and comprises two sand lobes (upper Turonian-Coniacian) represented by high-reflection amplitudes. The onlap geometries on top Blod0ks and top Tryggvason are interpreted to reflect tectonic episodes with flank uplift and basin floor subsidence.

The Tryggvason Formation (early-middle Turonian) The Tryggvason Formation can be divided into subunits of high and low internal reflectivity. These can be interpreted by well data trends as regressive and transgressive, respectively. After deposition of the transgressive B lod0ks Formation, there was a significant regression that left vague downlap features in the basin. On the eastern basin margin the Tryggvason Formation shows clear onlap on the underlying BlodOks and Svarte Formations (Fig. 10), with increasingly higher onlap to the east. The onlap is seismically one of the most pronounced features in the Cretaceous succession and could represent a significant tectonic tilting event after deposition of the B lodCks Formation in the early Turonian. The lithology of the Tryggvason Formation is dominated by shale and marly shale, while some of the high reflectivity, regressive units contain some sandstone. Oil shows were encountered in thin sand stringers in the Agat well 36/1-2. Attribute analysis shows a pattern of semi-concentric arcs that are much like the features interpreted as water escape structures or compaction features in for instance the Eocene shales in the North Sea (Cartwright, 1996) or in shales and oozes in the Oligocene/Miocene in

the V0ring Basin (internal data). The features observed in the Tryggvason Formation are asymmetric in form and occur only in the western part of the elsewhere-symmetric polygonal features (Fig. 11). This is interpreted to be the result of gravitational forces acting on the slope where the Tryggvason Formation sediments were deposited. On seismic cross-sections these features are seen as mainly intraformational small-scale faulting. Fig. 11 is an example from the uppermost subunit of the Tryggvason Formation. We suggest that these features are related to clay compaction and water escape and that presence of sand can be postulated where they are absent. This pattern or lack of pattern could thus be used as a lithology indicator.

The Kyrre Formation (late Turonian-Campanian) There are strong onlap geometries observed at the top of the Tryggvason Formation (Figs. 11 and 12), in a similar fashion to those seen at its base (Fig. 10). This indicates that another tectonic tilting event occurred with rapid basin floor subsidence and flank uplift at the transition between the Tryggvason and Kyrre Formations in late Turonian time. Sedimentation rate increased and can be interpreted to have

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Fig. 11. Seismic attribute analysis used for lithology classification. (Upper panel) Attribute analysis of the upper part of the Tryggvason Fm. demonstrates a pattern of largely concentric half-circles similar to water escape structures observed in clay- and ooze-dominated sediments elsewhere. They coincide with intraformational faults seen on seismic data. (Lower panel) Faulting is absent where wells show presence of sand in the upper Tryggvason Fm. Because the polygonal water escape structures are related to compaction of non-sandy sediments, we suggest that presence of such features can be used to distinguish between sand and clay. The asymmetric pattern observed here is due to gravitational forces induced on the westward-dipping slope.

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Fig. 12. Conceptual diagram of the interpreted sequences. The Asgard Fm. sediments in the east were deposited in shallow marine basins, uplifted in the Hauterivian-Barremian and eroded in the latest Barremian-earliest Aptian. The younger sequences onlap each other or the Barremian unconformity. Onlap surfaces at top Albian, early Turonian, middle/late Turonian and Coniacian probably represent tectonic events with tilting of the basin. Sand was deposited in certain tectonically active periods as a result of gravity mass flows.

reached a maximum around Santonian time. Palaeobathymetric interpretation indicates deep marine environments in the Sogn Graben and North Viking Graben and shallower conditions to the east. Except for local presence of upper Turonian/Coniacian sandstone (see below) the entire Kyrre Formation is dominated by deposition of hemipelagic clay. Despite the pronounced basin floor subsidence, the basins were largely infilled by the end of the Campanian, probably with only a gentle dip remaining towards the basin centres in the west. On seismic data the upper Turonian/Coniacian sandstones are observed as well-defined high-reflectivity units (see Fig. 10). The sandstone-bearing interval can be divided into two subunits, which have a distribution suggesting that they are submarine fan units with a common apex point. The lower fan seems to extend slightly more to the west than the younger one, suggesting a transgressive and backstepping trend. The distal part of this fan was drilled in exploration well 35/9-3 and proved to contain fineto medium-grained turbidite sandstone with gas. In the Selje High area immediately to the north of the study area there are sandstone intervals of the same age, but which also seem to range into the Santonian. Elsewhere in the study area, however, there are no high-amplitude reflections that might indicate the presence of sandstone in the Kyrre Formation. The lower part of the Kyrre Formation (from its base to top of the Coniacian sandstones) shows an overall uniform thickness from east to west, even

where the sand shales out into the basin. There is a clear onlap surface between this part of the succession and the overlying units. This is interpreted to reflect basin floor subsidence and basin flank uplift and is probably related to plate reorganisation and opening of the Labrador Sea in early Campanian (chron 33, Roest and Srivastava, 1989).

The Jorsalfare Formation (Maastrichtian) The Jorsalfare Formation consists of marl in the eastern region of the study area and is shale dominated further west. Limestone dominates on the Horda Platform to the south. Sedimentation rate had decreased from that of the Kyrre Formation. The formation is divided into a lower transgressive part that was followed by deposition of an upper regressive part. Eastward thinning and signs of erosion imply that basin flanks were exposed by the end of Cretaceous.

Discussion and conclusions It can be concluded that the Cretaceous in the northern North Sea was not a period of passive postrift infilling. The effects of the rifting event in the Late Jurassic seem to have continued into the Early Cretaceous, and in addition the rifting preceding the opening of the North Atlantic in the early Tertiary is expressed as a series of precursor events in the Late Cretaceous. We suggest that the strong basin flank uplift and inversion/uplift of some intrabasinal highs

290 that can be observed until the end of the Barremian should be attributed to the last phase of the "Late Jurassic" rifting. The structural setting and topography left at the end of the Barremian should thus be taken as the true situation at the onset of the post-rift phase of basin evolution. Other possible tectonic events are represented by onlap geometries on certain horizons, generally reflecting tilting with basin floor subsidence and flank uplift. They could be minor extensional events or alternatively vertical or lateral movements without any direct link to rifting. They could be related to mantle heating or to a combination of this and effects of plate reorganisation associated with the opening of the southern and middle Atlantic Ocean. Onlap geometries are observed at the following levels (Fig. 12): (1) Top RCdby (transition Early/Late Cretaceous, Albian/Cenomanian); (2) Top Blod0ks (early Turonian); (3) Top Tryggvason (middle/late Turonian); (4) Top Kyrre sandstone (Coniacian). The possible precursor of the Tertiary seafloor spreading could be represented by the rapid basin floor subsidence in the Santonian. From latest Ryazanian to end Barremian times fairly thin sequences of clay and marl were deposited on the structural highs, while thicker sequences were deposited on the basin slopes and nearly continuous sedimentation took place in the basins (Fig. 6). It is worth noticing that lowermost Cretaceous sediments present in the drilled wells represent certain stratigraphic intervals, while sediments from the intervening periods are absent throughout. We suggest that this reflects alternating and regional periods of sediment deposition and non-deposition, and because similarly aged sediments are preserved on intra-basinal structural highs it can be inferred that the structural highs were not subaerially exposed and eroded. Sediment distribution throughout the remaining part of the Cretaceous is characterised by large missing sections on the structural highs in the Aptian/ Albian and first part of the Late Cretaceous. Sedimentation seems to have been continuous in the basin areas and on the slopes. This overall setting continued until all highs were flooded around Santonian time. Hemipelagic deposition of clay and some marl dominated throughout the entire Cretaceous. Deposition of sand is broadly restricted to certain time intervals and probably linked to and following increased tectonic activity and/or changes of relative sea level. The Agat sands are all of Albian age, with a marked increase in sand input at the onset of late Albian. There are no indications in seismic data for tectonic activity directly affecting the area of study, such as change in basin topography (faulting, tilting, subsidence, etc.). The other periods of increased sand

T. Bugge et al. supply may, however, be linked to tectonic tilting events: expressed as the Tryggvason Formation in the early-middle Turonian and the Kyrre sands in the late Turonian-Coniacian. In the late Barremian/early Aptian, sand was probably also deposited in the Sogn Graben as the result of the Barremian basin flank uplift. The Ryazanian-Barremian sand seems to be derived from uplifted and eroded areas within the basin and along the basin flanks, while the Agat, Tryggvason and Kyrre sandstones were probably sourced from the east. The Agat sandstones seem to occur in palaeovalleys, as observed in wells 35/3-5 and 35/9-3 for example, and in the Sogn Graben and its extension to the south. The Kyrre sandstone is even more localised with more or less one single entry point. It can be speculated whether the sand was originally sorted and accumulated in a coastal setting where the basin morphology concentrated sand accumulation into certain depocentres. These could occur at positions where rivers entered the sea and/or where underlying structures formed embayments to accumulate sand fed by the rivers and by longshore transport. The combination of lowered relative sea level and tectonic activity could expose these areas and earthquakes could also contribute to the triggering of mass flows. The gravitational mass flows would be guided by the underlying topography and tend to follow palaeovalleys. It is not known where the coastline was situated during the Cretaceous, but the sand could possibly be transported over long distances. This complicates the reconstruction of possible source areas and leaves us with only conceptual models for the palaeogeography to the east of the area of deposition. The results described above have important implications for hydrocarbon exploration and can be summarised as follows. (1) Compared to a geological model of passive, post-rift infilling by the Cretaceous sediments, the indications of fairly active tectonism will severely influence the hydrocarbon exploration in the northern North Sea as well as other areas. Correct structural reconstruction is not possible without taking this into account. It will also have consequences for modelling of maturation and secondary migration of oil and gas. (2) The majority of the sandstones seem to have been deposited by different mass-flow processes and can probably be linked to certain periods of tectonic activity and/or change in relative sea level. This helps to identify prospective successions and areas. (3) There are almost no structural closures in the prospective Cretaceous intervals, and any potential oil and gas prospect will partly or fully rely on stratigraphic closures. This requires enhanced under-

The depositional history of the Cretaceous in the northeastern North Sea

standing of palaeogeographic setting, sedimentary environment and depositional processes.

Acknowledgements We thank Saga Petroleum who allowed us to publish this paper. We would also like to thank Bj~rn Tore Larsen, Snorre Olaussen and Sarah Prosser for reading the manuscript and giving valuable comments, and other colleagues in Saga Petroleum who have contributed throughout the work. Erik E Johannesen and Mike Talbot gave valuable referee comments.

References Blystad, P., Brekke, H., Fa~rseth, R.B., Larsen, B.T., Skogseid, J. and T~rudbakken, B., 1995. Structural elements of the Norwegian continental shelf. Part II: The Norwegian Sea Region. Norwegian Petroleum Directorate Bulletin, 8, 45 pp. Cartwright, J.A., 1996. Polygonal fault systems: a new type of fault structure revealed from 3-D seismic data from the North Sea Basin. In: E Weimer and T.L. Davies (Editors), Application of 3-D Seismic Data to Exploration and Production. AAPG Studies in Geology 42 and SEG Geophysical Development Series, 5, pp. 225-230. Dalland, A., Worsley, D. and Ofstad, K., 1988. A lithostratigraphic scheme for the Mesozoic and Cenozoic succession offshore midand northern Norway. Norwegian Petroleum Directorate Bulletin, 4, 65 pp. Deegan, C.E. and Scull, B.J. (compilers), 1977. A standard lithostratigraphic nomenclature for the central and northern North Sea. Institute of Geological Sciences Report 77/25, Norwegian Petroleum Directorate Bulletin, 1, 35 pp. Dor6, A.G., 1992. Synoptic palaeogeography of the Northeast Atlantic Seaway: late Permian to Cretaceous. In: J. Parnell (Editor), Basins on the Atlantic Seaboard: Petroleum Geology, Sedimentology and Basin Evolution. Geol. Soc., London, Spec. Publ., 62: 421-446. Fa~rseth, R.B., Gabrielsen, R.H. and Hurich, C.A., 1995. Influence of basement in structuring of the North Sea basin, offshore southwest Norway. Nor. Geol. Tidsskr., 75: 105-119. Gabrielsen, R.H., Odinsen, T. and Grunnaleite, I., 1999. Structuring of the Northern Viking Graben and the M~re Basin; the influence of basement structural grain, and the particular role of the M~reTr~ndelag Fault Complex. Mar. Pet. Geol., 16: 443-465. De Graciansky, E-C., Deroo, G., Herbin, J.-P., Montadert, L., Muller, C., Schaaf, A. and Sigal, J., 1984. Ocean-wide stagnation episode in the Late Cretaceous. Nature, 308: 346-349. De Graciansky, E-C., Hardenbol, J., Jacquin, T. and Vail, E (Editors), 1998. Mesozoic and Cenozoic Sequence Stratigraphy of European Basins. Soc. Econ. Paleontol. Mineral., Spec. PUN., 60, 786 pp. Gulbrandsen, A., 1987. Agat. In: A.M. Spencer, S.O. Johnson, A. MCrk, E. Nysa~ther, E Songstad and A. Spinnangr (Editors), Petroleum Geology of the Norwegian Oil and Gas Fields. Graham and Trotman, London, pp. 363-370. Hesjedal, A. and Hamar, G.E, 1983. Lower Cretaceous stratigraphy and tectonics of the south-southeastern Norwegian offshore. In:

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J.EH. Kaasschieter and T.J.A. Reijers (Editors), Petroleum Geology of the Southeastern North Sea and the Adjacent Onshore Areas. Geol. Mijnbouw, 62: 135-144. H~gseth, K., Vagle, G.B., Bergfjord, E., Granholm, P.G. and Skjervold, R., 1999. The Cretaceous depositional systems of the frontier V~ring Basin evidence from the Nyk High well (6707/10-1) and the Vema Dome well (6706/11-1). In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway Palaeozoic to Recent. Extended Abstracts, Norwegian Petroleum Society/NPF Conference, Bergen, Norway, May 3-5, 1999, ISBN 82-92032-00-2, pp. 199-200. Isaksen, D. and Tonstad, K., 1989. A revised Cretaceous and Tertiary lithostratigraphic nomenclature for the Norwegian North Sea. Norwegian Petroleum Directorate Bulletin, 5, 24 pp. Jacquin, T. and Thomson, M., 1999. 4-dimensional stratigraphic modeling of the northern North Sea. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway Palaeozoic to Recent. Extended Abstracts, Norwegian Petroleum Society/NPF Conference, Bergen, Norway, May 3-5, 1999, ISBN 82-92032-00-2, pp. 49-52. Jacquin, T., Rusciadelli, G., Amedro, F., De Graciansky, E-C. and Magniez-Jannin, F., 1998. The North Atlantic cycle: an overview of 2nd-order transgressive/regressive facies cycles in the Lower Cretaceous of Western Europe. In: E-C. De Graciansky, J. Hardenbol, T. Jacquin and E Vail (Editors), Mesozoic and Cenozoic Sequence Stratigraphy of European Basins. Soc. Econ. Paleontol. Mineral., Spec. Publ., 60: 397-409. Kemper, E., 1973. The Aptian and Albian stages in northwest Germany. In: R. Casey and EF. Rawson (Editors), The Boreal Lower Cretaceous. Geol. J. Spec. Issue, 5: 345-360. Nystuen, J.E, 1999. Submarine sediment gravity flow deposits and associated facies: core examples from the Agat Formation. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway Palaeozoic to Recent. Extended Abstracts, Norwegian Petroleum Society/NPF Conference, Bergen, Norway, May 3-5, 1999, ISBN 82-92032-00-2, pp. 211-215. Riley, L.A., Harker, S.D. and Green, S.C.H., 1992. Lower Cretaceous palynology and sandstone distribution in the Scapa Field, U.K. North Sea. J. Pet. Geol., 15:97-110. Roest, W.R. and Srivastava, S.E, 1989. Sea-floor spreading in the Labrador Sea: a new reconstruction. Geology, 17: 1000-1003. Shanmugam, G. and Moiola, R.J., 1995. Reinterpretation of depositional processes in a classic flysch sequence (Pennsylvanian Jackfork Group), Ouachita Mountains, Arkansas and Oklahoma. Am. Assoc. Pet. Geol. Bull., 79: 672-695. Shanmugam, G. and Moiola, R.J., 1997. Reinterpretation of depositional processes in a classic flysch sequence (Pennsylvanian Jackfork Group), Ouachita Mountains, Arkansas and Oklahoma: Reply. Am. Assoc. Pet. Geol. Bull., 81: 476-491. Shanmugam, G., Bloch, R.B., Mitchell, S.M., Beamish, G.W.J., Hodgkinson, R.J., Damuth, J.E., Straume, T., Syvertsen, S.E. and Shields, K.E., 1995. Basin-floor fans in the North Sea: sequence stratigraphic models vs. sedimentary facies. Am. Assoc. Pet. Geol. Bull., 79:477-512. Skibeli, M., Barnes, K., Straume, T., Syvertsen, S.E. and Shanmugam, G., 1995. A sequence stratigraphic study of Lower Cretaceous deposits in the northernmost North Sea. In: R.J. Steel, V. Felt, E.E Johannessen and C. Mathieu (Editors), Sequence Stratigraphy of the Northwest European Margin. Norwegian Petroleum Society (NPF), Special Publication 5. Elsevier, Amsterdam, pp. 389-400.

Norsk Hydro ASA, N-9480 Harstad, Norway; E-maih [email protected] Norsk Hydro ASA, E & P International, N-0246 Oslo, Norway Applied Biostratigraphy, Blekksoppgrenda 41, N-1352 Kols~s, Norway

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Cretaceous faulting and associated coarse-grained marine gravity flow sedimentation, Traill t21, East Greenland Finn Surlyk and Nanna Noe-Nygaard

The Cretaceous succession of the Traill 0 region in East Greenland is more than 2 km thick and is dominated by mudstones. It occurs in a system of fault blocks, about 5-10 km wide, delimited by roughly parallel, NNE-trending faults. The main faults in the area are from west to east the Post-Devonian Main Fault, the Bordbjerg Fault, the M~nedal Fault and the Mols Bjerge Fault. Cretaceous subsidence was governed by thermal contraction following the protracted Late Bajocian-Valanginian rift phase. The regional subsidence was punctuated by several possibly rift-related fault episodes marked by deposition of breccias, conglomerates and sandstones. Transport was by sediment gravity flows and clasts were derived from the uplifted footwalls and fault scarps of the main basin bounding faults. Four coarse-clastic deep-water units represent different types of depositional systems which are described to illustrate the variety of fault-associated deposits. Three of the coarse-clastic units form the basis for defining the new Rold Bjerge, Mfinedal and Vega Sund Formations. A chaotic breccia of earliest Middle Albian age forms the new Rold Bjerge Formation. It contains blocks, up to 60 m long, and has a clast assemblage characterised by Upper Permian and Lower Triassic carbonates. It is truncated to the west by the Mfinedal Fault, and the Lower Triassic Wordie Creek Formation which is exposed in the footwall, is considerably younger than the dated clasts. The breccia was thus derived from the area further west delimited by the Post-Devonian Main Fault, while the Mfinedal Fault appears not to have been active at the time of deposition. Transport was by submarine, probably hydroplaning, debris avalanches, and the estimated runout distance was about 20 kin. An allochthonous succession comprising mudstones and sandstone turbidites of Late Albian age is also exposed immediately adjacent to the M~nedal Fault. It was formed by extensive downslope sliding and slumping along discrete, densely spaced detachment planes, and the mudstones and sandstones were folded during transport. The age of the slide-slump event is not known, but a Late Turonian-Early Coniacian age is tentatively suggested, by analogy with conglomerates and pebbly sandstones of the Mfinedal Formation which are situated in a similar position along strike. Remobilisation, downslope transport and subsequent redeposition took place during an important phase of footwall uplift. A conglomerate-sandstone package forming the new Mfinedal Formation is limited to the west by the Mfinedal Fault. The age is not precisely known, but macrofossils indicate a Late Turonian-Early Coniacian age. The common presence of boulders composed of Upper Permian carbonates indicates that the Mfinedal Fault was not yet formed or played only a minor role at the time of deposition because the rocks exposed in the footwall are of Early Triassic age. The main source area was again situated in the area west of the Post-Devonian Main Fault, and the conglomerates and pebbly sandstones were transported downslope by turbidity currents and related processes towards the east-southeast at a right angle to the faults. The system shows a crude fining-upward in the top part and is interpreted as a faulted slope apron. A thick succession of planar sandstone turbidites forming the new Vega Sund Formation is exposed 6 km east of the Mfinedal Fault. The age is highly uncertain. A Cenomanian age has been suggested previously but an Early-Middle Coniacian age is very tentatively preferred. Transport direction was again towards east-southeast at a right angle to the faults. The sheet-like bed geometry, lack of channeling and scouring, and gradual fining-upward of the top part suggest deposition on the outer part of a sand-rich basin floor fan. The four systems differ greatly in grain size, sorting, downslope transport processes, geometry and lateral extent. The transverse eastward transport directions and the clast provenances indicate, however, that the Cretaceous fault blocks were wider and not yet fragmented into the narrow present-day blocks. This interpretation contrasts with previous accounts which suggest that the Mfinedal and Bordbjerg Faults were active and exerted a profound control on sedimentation during Cretaceous time. It is, however, possible that an early expression of the Mhnedal Fault was formed in post-middle Albian time. The systems provide evidence for phases of possibly rift-related faulting accommodated by the old faults and associated footwall uplift and erosion in the earliest Middle Albian, and Late Turonian-Early Coniacian. The absence of Hauterivian and in part Barremian strata virtually throughout East Greenland may reflect marginal or probably regional uplift during an earlier fault event; the scarcity of Upper Aptian deposits and the presence of a Lower Cenomanian basal conglomerate may likewise reflect rifting and marginal uplift. The coarse-clastic deep-water deposits of the Traill 0 region were emplaced by a variety of sediment gravity flows and may serve as useful field analogs for the deeply buried correlative strata of the outer Norwegian Shelf. They illustrate the variability and complexity of the depositional systems, and allow identification of several significant Cretaceous possibly rift-associated fault pulses superimposed on the long-term regional subsidence governed by thermal contraction following the protracted Late Bajocian-Valanginian rift phase.

Sedimentary Environments Offshore Norway - Palaeozoic to Recent edited by O.J. Martinsen and T. Dreyer. NPF Special Publication 10, pp. 293-319, Published by Elsevier Science B.V., Amsterdam. 9 Norwegian Petroleum Society (NPF), 2001.

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Introduction The main Mesozoic rift event in East Greenland began in Middle Jurassic time, culminated in the Late Jurassic with formation of strongly tilted fault blocks, and faded out in the earliest Cretaceous (Vischer, 1943; Maync, 1947, 1949; Surlyk et al., 1981; Surlyk, 1990, 2001; Surlyk and Noe-Nygaard, 2000). Deposits and structures associated with this event are well exposed in East Greenland and have to some extent overshadowed the importance of subsequent Cretaceous fault phases. The scattered nature of the Cretaceous outcrops and lack of seismic profiles makes it difficult to evaluate if the faulting is related to true extensional rifting. Demonstration of episodes of footwall uplift, rejuvenation of fault scarps and formation of relatively steep slopes suggest, however, that the faulting may represent actual rift events. The Cretaceous succession of East Greenland is dominated by dark basinal mudstones, and is up to about 2 km thick. Debris avalanche breccias, slide and slump deposits, turbiditic conglomerates and sandstones occur at several stratigraphic levels. Similar Cretaceous successions are at present subject of intense exploration on the Norwegian shelf and elsewhere around the northern North Atlantic Ocean. The aim of the

study is to describe and interpret the wide spectrum of coarse-grained gravity flow deposits of Traill D, to relate them to the tectonic evolution of the area, and to make a comparison with similar Cretaceous deposits known from wells on the Norwegian Shelf.

Geological setting The important Late B ajocian-Valanginian rift phase is well known from East Greenland, the North Sea and the Norwegian Shelf (e.g. Surlyk et al., 1981; Ziegler, 1988; Surlyk, 1990, 1991, 2001; Dor6, 1992; Surlyk and Noe-Nygaard, 2000). Great economic interest is associated with this event because it created the main hydrocarbon play types in the region. Extensive outcrops of correlative successions occur in East Greenland and have received much attention as analogues for the correlative deeply buried subsurface successions (e.g. Surlyk, 1978, 1990, 1991,2001; Stemmerik et al., 1992; Price and Whitham, 1997). Cretaceous faulting events and associated syn-tectonic deposits have, however, also been reported from East Greenland (Donovan, 1953, 1955, 1957; Surlyk et al., 1981; Surlyk, 1990; Whitham et al., 1999). In Jurassic-Cretaceous time a N-S-trending rifted seaway was situated between East Greenland and

Fig. 1. Pre-drift reconstruction showing the structural framework of the mid-Cretaceous seaway between East Greenland and Norway. Based on Stemmerik et al. (1998), Granholm (1999), Larsen et al. (1999), and unpublished data.

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Fig. 2. The Mesozoic rift basin of East Greenland showing outcrop of Cretaceous deposits and main faults. Based on Haller (1970), Koch and Haller (1971), GEUS maps, and own data.

Norway (Fig. 1). The western margin of the seaway was located up to a few hundred kilometres inland from the present-day coast of central and northern East Greenland. In the study area on the island Traill 0 (O means island in Danish) (Figs. 2 and 3), Middle Jurassic early-rift deposits rest unconformably on Triassic sediments, and consist of a succession of lower(?) Bajocian fluvial pebbly sandstones of the Bristol Elv Formation, marine sandstones and mudstones of the Upper BajocianCallovian Pelion and Fossilbjerget Formations, shallow marine sandstones of the Lower-Middle Oxfor-

dian Olympen Formation, and dark deep-marine Upper Oxfordian-Kimmeridgian shales of the Bernbjerg Formation (Fig. 4). Note that the rank of some units differs from older literature (Surlyk et al., 2001). Rift climax deposits of latest Jurassic-earliest Cretaceous age are not exposed, but late-rift, relatively deepwater, red, calcareous mudstones occur in isolated outliers on eastern Traill 0 (Donovan, 1953). They are included in the Valanginian ROdryggen Member and are characteristic of submerged fault block crests suggesting the presence of a buried rift-climax prism (S urlyk, 1978).

296

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Fig. 3. Map of the Traill 0 region showing faults, fault blocks and place names mentioned in the text. Based on Stemmerik et al. (1998). The studied outcrops are marked with a number. Locality *1: submarine debris avalanche, Rold Bjerge Formation. Locality *2: conglomeratesandstone slope apron, M~nedal Formation. Locality *3: submarine slide-slump complex. Locality *4: sandstone-dominated basin floor fan, Vega Sund Formation.

The irregular Jurassic rift topography was partly filled in by a mudstone-dominated Cretaceous succession, up to about 2 km thick. It is not well exposed and has not been lithostratigraphically classified. It represents post-rift deposition during thermal contraction following Middle Jurassic-earliest Cretaceous rifting, punctuated by several possibly rift-related fault events. The age relations of the Cretaceous succession in the Traill 0 region are only broadly known because macrofossils are scarce at most levels and because dinocysts are generally poorly or not preserved due to heating by Cenozoic intrusions (Donovan, 1953, 1955, 1957; NOhr-Hansen, 1993; this study). The Cretaceous mudstones are thus disturbed by numerous dolerite dykes and sills and are overlain by Cenozoic plateau basalts. The presence in the Traill 121region of Lower Aptian, Lower and Upper Albian, and possibly Lower Cenomanian strata was demonstrated on the basis of dinocysts (NChr-Hansen, 1993), whereas Upper Aptian, basal Middle Albian, Upper Turonian-Middle(?)

Coniacian, and Upper Campanian strata are documented on the basis of ammonites, belemnites and inoceramid bivalves (Donovan, 1953, 1955, 1957; this study). The East Greenland margin was uplifted in the order of 1.5-3 km in Cenozoic times (Christiansen et al., 1992), and the results of Mesozoic rifting can now be studied in the mountains along the East Greenland coast. Uplift was not uniform but was associated with faulting, partly by rejuvenation of the older Mesozoic faults, and it is not always possible to reconstruct the Mesozoic component of the fault system nor the amount of throw on the individual faults. The main faults trend NNE and are connected by NW- and NE-trending transfer faults. The faults are identified by mapping on aerial photographs and ground mapping, and the time of their activity is estimated by comparison of successions in the footwall and hanging wall, and by matching clast lithologies in syn-tectonic deposits on the hanging wall with successions exposed in the footwall. In this

Cretaceous faulting and associated coarse-grained marine gravity flow sedimentation, Traill O, East Greenland

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Fig. 4. Stratigraphic scheme of the Upper Permian-Cretaceous of East Greenland. The Jurassic lithostratigraphic nomenclature is from Surlyk et al. (2001 ).

way it is possible to identify which of the main NNEtrending faults were active during the Cretaceous. Main faults

The fundamental, N-S-trending, Post-Devonian Main Fault (Figs. 1 and 2) was recognised in the

region north of Traill O by Vischer (1943). In the southern Geographical Society O-Traill 0 the orientation changes to NE-SW and this diagonal fault links up with the N-S-trending Stauning Alper Fault which limits the Jameson Land Basin to the west (Figs. 1 and 2). The fault system downthrows Carboniferous-Lower Permian strata to the east against

298 Devonian and older rocks and was the main basin bounding feature in Late Permian-Mesozoic times. Another fault runs parallel to the Post-Devonian Main Fault 4 km to the east (Fig. 3). It downthrows Upper Permian-Lower Triassic against Upper Carboniferous strata and is probably a late feature as it cuts Cenozoic sills. The next fault to the east, the Bordbjerg Fault, downthrows Triassic against Upper Carboniferous strata (Figs. 1 and 3). The fault has several splays, dips 50-60 ~ towards the east, post-dates the earliest Triassic and is mainly of Cenozoic age. The Mgmedal Fault downfaults Cretaceous against Lower Triassic strata, dips 50-60 ~ to the east, and is probably mainly of Cenozoic age (Figs. 1 and 3). Faults A and B are splays at the southern end of the Mfinedal Fault (Fig. 3). Fault A downfaults Middle Jurassic-Cretaceous against Triassic strata, whereas Fault B downfaults Cretaceous against Middle Jurassic-Cretaceous strata. Coarse-clastic deep-water deposits

Cretaceous coarse-clastic, deep-water deposits are exposed in dip sections in several E-W-trending valleys which cut across the N-trending Mfinedal Fault (Fig. 3). An early expression of this fault may have existed in Cretaceous time, but its present appearance is due to Cenozoic faulting. Four different types of coarse-grained deep-water deposits are described below, starting with the most chaotic and ending with the most organised system. Their general setting and age is presented, and the facies are described, followed by interpretation of transport processes and depositional environment. The coarse-clastic deposits are in most cases rather chaotic without showing any pronounced organisation in the form of vertical facies trends. Clasts in the breccias and conglomerates vary in size from pebbles to blocks up to 60 m across, and in composition from carbonate over sandstone to metamorphic quartzite, granite and gneiss. Several of the clast lithologies are sufficiently characteristic to allow determination of provenance, and clasts composed of Upper Permian carbonates and shales, Lower Triassic carbonates and pebbly sandstone and Middle Jurassic sandstones have all been identified with more or less certainty. The clast ages allow determination of the stratigraphic units exposed in the footwall fault scarps and of the minimum height of the scarps at the time of deposition. The provenance of the clasts also indicates which faults were active during deposition. The resedimented units show significant differences and are highly characteristic. Three of the units form the basis for the definition of the new Rold

F. Surlyk and N. Noe-Nygaard

Bjerge, Mfinedal and Vega Sund Formations (see Appendix A). Submarine debris avalanche - - Rold Bjerge Formation

The studied section is located at 600 m altitude in Rold Bjerge in a mountain pass, about 3 km north of the Mfinedal valley (locality *1 in Fig. 3). The pass marks the position of the N-S-trending, eastward-downthrowing, M~nedal Fault which forms the eastern boundary of the 5 km wide Mfinedal Block limited to the west by the Bordbjerg Fault (Fig. 3). The footwall of the Bordbjerg Fault exposes a succession of Carboniferous pebbly sandstones and conglomerates overlain by Upper Permian conglomerates, sandstones, hypersaline carbonates, evaporites and mudstones, topped by sandstones and shales of the Lower Triassic Wordie Creek Formation. The Mgmedal Block exposes a thick succession of shallow marine sandstones of the Wordie Creek Formation (Figs. 3 and 4). Pebbly sandstones and sandstones occurring in or adjacent to the M~nedal Fault separating Triassic and Cretaceous strata were previously interpreted as a sliver of the Middle Jurassic Pelion Formation (Donovan, 1953; Price and Whitham, 1997). They are here considered to represent the Svinhufvuds Bjerge Member of the Wordie Creek Formation (Clemmensen, 1980). If this interpretation is correct, the M~nedal Fault is much more simple than indicated by previous studies. The succession exposed in the hanging wall immediately east of the Mfinedal Fault comprises a chaotic sedimentary breccia, at least 25 m thick (Fig. 5), sharply overlain by dark-grey to black, laminated mudstones with abundant inoceramid bivalves, including Actinoceramus sp. aft. concentricus (Parkinson), indicating an earliest Middle Albian age (K.-A. Tr6ger, written commun., 1997). Dinoflagellate cysts from shales below the breccia indicate an Early A1bian age (N~hr-Hansen, 1993). The breccia is exposed in a series of bluffs extending over a distance of several hundred metres (Fig. 5). The clast size varies from pebbles and cobbles to very large blocks. The largest block observed is 60 m by more than 10 m and is oriented with the long axis parallel to bedding. It consists of brown-weathering, well sorted, fine-grained quartzose sandstone. Another block of the same lithology is 30 x 30 m in size and blocks in the 5-10 m range are common (Fig. 6). Many clasts consist of pebbly sandstone. Angular platy boulders of white or pinkish limestone are very common and may reach a maximum length of several metres. They show irregular banding or crude lamination reminiscent of algal lamination (Fig. 7).

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Fig. 5. Submarine debris avalanche breccia sheet of the Lower or lowermost Middle Albian Rold Bjerge Formation. The breccia is overlain by black basal Middle Albian mudstones and is bounded to the west by the Mfinedal Fault. The main left part of the sheet is one big block of pebbly sandstone, 60 m long. Lower Triassic sandstones of the Wordie Creek Formation are exposed in the footwall west of the fault. Locality * 1 in Fig. 3.

Fig. 6. Large blocks of whitish limestone and sandstone with densely packed angular boulders (hammer encircled for scale). Debris avalanche breccia of the lower or lowermost Middle Albian Rold Bjerge Formation. Locality * 1 in Fig. 3.

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Fig. 7. Detail of the debris avalanche breccia of the lower or lowermost Middle Albian Rold Bjerge Formation. Note the dense packing of angular blocks composed of pebbly sandstone, sandstones and Upper Permian carbonates showing wavy, possibly algal lamination (above hammer head). The breccia is strongly cemented by several generations of cement. Hammer for scale. Locality * 1 in Fig. 3.

Light-brown, cobble-sized, equidimensional and angular limestone clasts dominate in parts of the unit. The elongate clasts show a crude orientation subparallel to bedding, but the sorting is extremely poor. Grading and clast imbrication have not been identified. Two to three highly irregular bedding planes can, however, be traced through much of the unit. They show a dip of 5-10 ~ to the east. The chaotic nature, lack of organisation of the poorly defined breccia beds, and the content of very large blocks suggest deposition from submarine debris avalanches which may have been hydroplaning (cf. Mohrig et al., 1999). The laminated limestone clasts were undoubtedly derived from the Upper Permian Karstryggen Formation, whereas others are similar to carbonates of the Lower Triassic Odepas Member. The source of the pebbly sandstone clasts is more difficult to identify but they may have been derived from the Lower Triassic S vinhufvuds Bjerge Member. The provenance of clasts shows that the adjacent Mfinedal and Bordbjerg Blocks could not have served as source areas. The Karstryggen Formation and the Svinhufvuds Bjerge and Odepas Members occur in the Mfinedal and Bordbjerg Blocks in undisturbed stratigraphic succession, well below the topographical

level of the breccia. The source thus has to be sought further to the west in the footwall of the Post-Devonian Main Fault. This gives a minimum runout distance of the debris avalanches of about 25 km. The Rold Bjerge breccia thus records earliest Middle Albian rejuvenation of the Post-Devonian Main Fault, footwall uplift, and shedding of a series of submarine debris avalanches down the faulted slope (Fig. 8). The trigger mechanism was probably earthquakes caused by movements on the fault. Submarine slide-slump complex ~ Eastern Svinhufvud Bjerge

This locality is situated in an east-west-oriented valley cutting the B fault splay of the N-S-trending Mfinedal Fault (locality *3 in Fig. 3). The section was mapped using a mosaic of overlapping polaroid photographs (Fig. 9). It exposes an impressive slideslump complex which extends over several hundred metres downdip towards the east-southeast (140~ Yellow sandstones of the Pelion Formation are exposed in the footwall and gently E-dipping, dark-grey mudstones and yellow sandstones in the hanging wall. The mudstones have yielded dinoflagellate cysts indicating a Late Albian age (NOhr-Hansen, 1993).

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Fig. 8. Depositional model for the Lower to lowermost Middle Albian submarine debris avalanche breccia of the Rold Bjerge Formation. PDMF = Post-Devonian Main Fault; MF -- possible early expression of the M~nedal Fault.

Turbidite sandstone beds, mainly 10-20 cm thick, are interbedded with the mudstones. They are dominated by Bouma Ta divisions and appear relatively sheet-like. The succession is disturbed by numerous low-angle irregular faults which are roughly parallel to overall bedding, and separate packets, up to a few metres thick (Fig. 9). The strata in the packets have orientations ranging from parallel to the faults to vertical and from gently to strongly folded. Isoclinal folds are common as are overturned or subvertical beds which may show S-like folds (Figs. 9 11). Some packets are sandstone-dominated, whereas others consist of mudstone with only a few thin sandstone beds (Figs. 9 and 11). This clearly reflects an original difference in the faulted succession with alternating sandstone- and mudstone-dominated units. Orientations of slump fold axes show a broad scatter from 0 ~ over 90 ~ to 160 ~ averaging 74 ~ and vergences range from south over east to northeast. General direction of transport seems to have been roughly towards the southeast. A succession of depositional events can now be interpreted. A mud-rich turbidite system was formed in the Late Albian. The mainly thin-bedded nature of the sandy turbidites and the low sandstone : mudstone ratio suggest a distal, relatively deep-water environ-

ment, probably a fan fringe. A levee origin cannot be excluded, but is considered less likely on the basis of the apparent sheet-like nature of the turbidites, the alternation between sandstone- and mudstone-dominated units, the dominance of Ta divisions, and the overall absence of climbing ripples in the Tc divisions. At some later time the whole succession was disturbed by slope failure and was transported by sliding and slumping down the palaeoslope resulting in the formation of a thick slide-slump complex (to the left on Fig. 12). Movements were both translational and rotational as shown by the planar to curved rupture surfaces and differences in orientation of strata between the detachment planes. The timing of the slide-slump event is not known and cannot be directly determined from the section. This can only be done if a locality is found where the disturbed succession is overlain by younger, undisturbed deposits. It is suggested that the event was triggered by the same fault episode that caused the formation of the turbiditic conglomerate-sandstone complex of the Mfinedal Formation described below, and a Late Turonian-Early Coniacian age is tentatively suggested for the event. This interpretation is based on the evidence for strong footwall uplift of the Post-Devonian

C.O 0

Fig. 9. Sketch of slide-slump complex, based on field mapping on polaroid photographs. Note the abundant subparallel detachment planes delimiting strongly folded packages of dark-grey mudstones and sandstone turbidites. Locality *3 in Fig. 3.

Cretaceous faulting and associated coarse-grained marine gravity flow sedimentation, Traill 0, East Greenland

303

Fig. 10. Slide-slump complex, eastern Svinhufvud Bjerge. Interbedded dark-grey mudstones and turbidite sandstones have undergone translational down-slope sliding and slumping. The strata were folded and faulted during transport but were not homogenised and transformed into debris flows, probably due to consolidation and early cementation. The shales and sandstones were deposited in the Late Albian (N~hr-Hansen, 1993), but the subsequent slide-slump event cannot be dated. It is tentatively suggested that it was caused by the Late Turonian-Early Coniacian fault episode that resulted in deposition of the Mfinedal Formation conglomerates. Locality *3 in Fig. 3.

Main Fault during the Late Turonian-Early Coniacian event, the possible initiation of the M~nedal Fault, and the similar position of the two systems along strike of the M~nedal Fault.

Conglomerate-sandstone slope apron - - M~nedal Formation

A relatively thick succession of pebble, cobble and boulder conglomerates was described from the low northern bank of the Mfinedal valley by Donovan

304

F. Surlyk and N. Noe-Nygaard

Fig. 11. Details from the slide-slump complex of Fig. 9. Note the sharp detachment surface at the top of the sandstone-dominated lower unit.

Fig. 12. Depositional model for the Late Turonian-Early Coniacian conglomerate-sandstone slope apron of the M~nedal Formation. The submarine slide-slump complex in eastern Svinhufvud Bjerge (locality *3 in Fig. 3) is shown in the foreground to the left. The strata are of Late Albian age and the slump event is tentatively referred to the Late Turonian-Early Coniacian. PDMF = Post-Devonian Main Fault; MF = possible early expression of the M~nedal Fault.

(1953). The succession has yielded poorly preserved ammonites, belemnites and inoceramids which together suggest a Late Turonian-Early Coniacian age

(W.J. Kennedy and W.K. Christensen, written commun., 1999). The rather poor exposure starts 500 km east of the Mfinedal Fault and is about 500 m long

Cretaceous faulting and associated coarse-grained marine gravity flow sedimentation, Traill O, East Greenland

but is interrupted several places by unexposed intervals (locality *2 in Fig. 3). The strata appear to dip between 10 and 20 ~ towards the ESE, and this gives a thickness between 90 and 180 m. Donovan (1953) estimated a thickness of 80-100 m based on similar considerations. Direct measurement of the composite section made up of the numerous small, closely spaced outcrops gives, however, a thickness of only 45 m. This thickness estimate is corroborated by independent measurement by M. Larsen (pers. commun., 1999). It cannot be excluded, however, that parts of the section are cut out by later faulting, whereas repetition of strata seems not to have taken place. A total of nine outcrops were studied down-dip from west to east covering the same stretch as localities 65-62 of Donovan (1953) and mapped on polaroid photographs. Locality 33 farthest to the west and locality 38 close to the eastern end provided the best outcrop. A late fault cuts the eastern end of locality 34, but otherwise only minor faults were observed, and some of these are syn-depositional. The sections in localities 36-39 appear to be in stratigraphic continuity in spite of the poorly exposed nature of the intervening parts of the river bank. The succession is volumetrically dominated by pebble and cobble conglomerates, and pebbly sandstones are prominent at some levels, whereas others are dominated by boulders (Figs. 13-18). Shale zones up to 40 cm thick occur at several levels and form marker beds for correlation between exposures. Five facies groups are identified and are described and interpreted below.

Boulder conglomerates This facies is highly irregular in appearance. It is dominated by large, commonly outsized boulders and several characteristic boulder lithologies occur. The most common type comprises characteristic, muddraped yellow sandstone boulders (Fig. 17). They are commonly about 0.5-1 m long, may reach lengths of almost 2 m, are subrounded, and have striations on the surface. The mud drape is black and quartz and shale pebbles are commonly pressed into the surface of the boulder. The yellow sandstone is mainly well sorted but may contain stringers of quartzite pebbles. The lithology is highly similar to the Middle Jurassic Bristol Ely and Pelion Formations, but a source in the Lower Triassic Svinhufvuds Bjerge Member is also possible. The subrounded shape, surface striations and rather poorly consolidated nature of the sandstone boulders show that the sand was only weakly cemented at the time of erosion in the source area, suggesting only shallow burial before uplift, exposure, erosion and redeposition.

305

Another characteristic, but less common, boulder type consists of grey limestone (Fig. 16A). The boulders are up to 250 cm long and may have tabular to almost spherical shapes. The largest boulder observed measures 250 x 150 x 150 cm and consists of thick-bedded dark-grey wackestone to packstone with abundant brachiopods and bivalves at the base of the original succession and with algal stromatolites at the top. The lithology and fossil content of the limestone boulders are similar to the Upper Permian Wegener HalvO Formation and there is no doubt that this unit formed the source of the boulders. A third boulder lithology consists of black shale, appearing as elongate and platy, angular to subrounded clasts. The lithology is similar to the Upper Permian Ravnefjeld Formation but other black shales may also have formed the provenance. The Upper Jurassic Bernbjerg Formation is thus another possibility but was probably too poorly consolidated to form angular clasts at the time of Late Cretaceous erosion. The boulder-dominated beds pass laterally into pebble and cobble conglomerates with only a few boulders. They can thus be considered as a variant of the pebble-cobble conglomerates described below, but some boulders may be genetically unrelated to the beds in which they occur. In cases where isolated boulders or boulder trains are found at the base of a pebble-cobble conglomerate they were probably deposited by an earlier event and were subsequently embedded in a finer-grained gravity flow deposit. The mechanisms active during transport and deposition of the isolated boulders and boulder-dominated conglomerates are thus difficult to interpret. A debris fall transport mechanism (see Nemec, 1990) is suggested for those boulder accumulations forming clast assemblages which have clearly been overridden by a finer-grained tail. Others were obviously part of a finer-grained sediment gravity flow with sufficient competence to transport outsized clasts, and some were probably emplaced from hydroplaning debris flows (cf. Mohrig et al., 1999).

Pebble-cobble conglomerate This facies is volumetrically dominant (Figs. 1316). Beds are mainly about 0.5-1.5 m thick, but may reach a thickness of 3-4 m. The lower boundary of the beds may be erosional or sharp without showing signs of marked erosion. In some cases beds rest on a pronounced scour surface and the lower boundary is in a few cases developed as a detachment surface (Fig. 18). The upper boundaries are planar or scoured by the base of the overlying bed. The beds are non-graded or show a crude coarse-tail or content grading, while basal inverse

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Cretaceous faulting and associated coarse-grained marine gravity flow sedimentation, Traill O, East Greenland

307

Fig. 14. Interbedded conglomerates and pebbly sandstones of the Mfinedal Formation. Note large rounded boulders in the top of the section. White stick is 2 m long. Locality *2 in Fig. 3.

grading is noted in some beds. The clast lithologies include quartzite, sandstone, limestone, shale and, less commonly, crystalline basement rocks. The clasts of harder lithologies are generally well rounded suggesting a polycyclic history with derivation by erosion of older conglomerates (Fig. 15). Clasts of sedimentary origin are subrounded to almost angular and commonly show slab-like shapes reflecting the original bedding of the parent rock. The conglomerates are rather poorly sorted and mainly clast-supported although matrix-supported beds or levels also occur. The matrix consists of poorly sorted sand and fine pebbles, smeared mudstone chips, and in some cases there is a significant content of mud. Scattered fossil debris, mainly calcite prisms from inoceramid bivalves, occurs. The clasts show a-axis imbrication in most beds indicating transport towards the ESE (Fig. 13). The a-axis imbrication was formed by clast collisions at high clast concentrations during late stage of flow immediately preceding deposition. Each bed represents one, possibly sustained, flow event and deposition probably took place from below by gradual upward movement of the depositional interphase. The basal inverse grading reflects development of dispersive pressure caused by clast collisions at the base of the flow. The crude normal grading and mainly erosional base

suggest transport by turbidity currents. The facies is highly similar to Volgian-Valanginian resedimented conglomerates well known from Wollaston Forland farther north in East Greenland (Surlyk, 1978, 1984).

Pebbly mudstones This facies consists of sandy mudstone or muddy sandstone with scattered pebbles and cobbles (Fig. 18). Bed boundaries are only exposed in some cases where they are developed as shear or detachment planes. The beds may be completely homogenised and massive but more commonly they show internal disturbed bedding or shear planes. The internal beds may show folding and truncation against detachment planes (Fig. 18). Scattered fossils occur and include belemnites, ammonites and inoceramid bivalves mainly represented by shell prisms. The dark-grey to black matrix is very poorly sorted and contains some clay. The pebbly mudstones were formed by remobilisation of previously deposited pebble, sand and muddominated lithologies and subsequent homogenisation during downslope mass movement. The original deposits were probably emplaced on the slope by various types of sediment gravity flows. Subsequent triggering of mass movements on the unstable slope

F. Surlyk and N. Noe-Nygaard

308

Fig. 15. Clast-supported pebble-cobble conglomerates of the M~nedal Formation. Locality *2 in Fig. 3.

resulted in sliding and slumping of lithologically inhomogeneous packages which were homogenised to various degrees depending on the length of travel and the transport processes involved. In the present case a spectrum of downslope transport processes ranging from slumping to viscous debris flow can be recognised, represented by bedded but internally folded beds and structureless beds, respectively.

Sandy mudstone A few beds of dark-grey, poorly sorted sandy mudstone occur in the succession. They are up to 40 cm thick and form local marker beds which can be used for correlation between the sections. They represent relatively short time intervals with lateral shifts in the locus of coarse-grained gravity flow deposition or temporal abandonment of the coarse-grained system.

Graded, pebbly sandstone Depositional environment This facies is a characteristic but volumetrically less important part of the succession (Fig. 13). Beds are normally wedge-shaped and 0-30 cm thick, but may form amalgamated packages up to 110 cm thick, which are also wedge-shaped. The base of the beds is commonly erosional and the beds are graded from pebbly to medium-grained sandstone. The pebbles are mainly composed of quartzite or shale, and finely comminuted shale chips are common on bedding planes. Larger clasts commonly show a-axis imbrication. The pebbly sandstone beds show a lower massive portion followed by an upper laminated part. The pebbly sandstones are interpreted as turbidites showing Ta and Tab Bouma divisions. Deposition took place in shallow scours and chutes from strongly erosional turbidity currents. Transport direction as shown by the imbricated clasts was towards the ESE (Fig. 13).

Vertical trends in facies or in grain sizes cannot be detected in the succession although this might be due to the lack of outcrop continuity. There is, however, a weak tendency for a gradual fining-upward in the upper part and the succession is overlain by mudstones with thin sandstones of the Inoceramus lamarcki beds of Donovan (1953, 1955) who also noted this trend. The succession is only well known from this locality but small exposures of similar lithologies, probably of the same age, occur 3.5-5 km due north along depositional strike (Donovan, 1953; localities 51 and 59, his plate 3). Similar, undated conglomerates and sandstones exposed 10 km east of Mfinedal on Silja 0 in Vega Sund and 15 km to the east on southeastern Geographical Society 121 due north of Kap Hovgaard possibly belong to the

Cretaceous faulting and associated coarse-grained marine gravity flow sedimentation, Traill O, East Greenland

309

Fig. 16. Clast-supported cobble-boulder conglomerates of the Mfinedal Formation. Locality *3 in Fig. 3. (A) The clasts show imbrication (succession dips to the east, to the right in the picture). (B) Limestone boulder with abundant fossils, derived from the Upper Permian Wegener HalvO Formation.

same depositional system (Fig. 2) (Donovan, 1955, 1957). The depositional environment is interpreted as a faulted slope apron (Fig. 12). This is indicated by the apparent lack of overall organisation except the crude fining-upward trend in the upper part, deposition from

a variety of sediment gravity flow types, ubiquitous scouring, evidence for syn-sedimentary sliding and slumping, and an ESE transport direction at a right angle to the dominant fault direction (to the right on Fig. 12). A similar interpretation as "a boulder-bed flanking a faulted coast-line of Turonian times" was

310

F. Surlyk and N. Noe-Nygaard

Fig. 17. Clay-draped, subrounded sandstone boulders from the M~nedal Formation, probably derived from the Middle Jurassic Bristol Elv or Pelion Formations. Locality *2 in Fig. 3.

Fig. 18. Conglomerates and pebbly sandstones separated from the underlying folded pebbly mudstone by a detachment surface. M~edal Formation, Locality *2 in Fig. 3.

already offered by Donovan (1957, p. 84). Donovan (1955, 1957) proposed that the source area was the Mfinedal and Bordbjerg Blocks because Upper Permian carbonates are exposed on these blocks (Fig. 3). This view was followed by Haller (1970) but is not supported by the present data (see below). The conglomerate-sandstone slope apron of the Mfinedal Formation seems to be of considerable dimensions. In the Mgmedal section the succession is estimated to be close to 50 m thick. It extends downdip east of the Mfinedal Fault for about 15 km and may have extended westward for almost 25 km covering the area between the Mfinedal Fault and the Post-Devonian Main Fault. This gives a maximum proximal-distal extent of about 40 km. Along strike the extent seems to be at least 10 km, but there are

numerous scattered, poorly exposed conglomerates further south along the M~nedal Fault which may belong to the same system. The system is in many ways similar to the Volgian-Ryazanian Wollaston Forland Group (Surlyk, 1978, 1984, 1989). The facies types are the same, and in Wollaston Forland very coarse-grained pebble-, cobble- and boulder-dominated gravity flows travelled eastward up to about 15 km downslope away from the scarp of the Dombjerg Fault. Flows came to a rest when they reached the axis of the halfgraben basin and were checked by the westward-dipping slope of the hanging wall. The youngest conglomerates of the Valanginian Young Sund Member were, however, emplaced by axial, southward gravity flows and may have travelled for distances up to 30 km.

Cretaceous faulting and associated coarse-grained marine gravity flow sedimentation, Traill O, East Greenland

311

with centimetre-thick mudstone interbeds (Figs. 1921), and about 90 m are exposed in the narrow valley which passes upstream into a steep-walled canyon. It is intercalated between two dolerite sills. All organic material has been burned off and dinoflagellate cysts are not preserved. Donovan (1955, p. 28) recorded a single specimen of Inoceramus crippsi Mantell suggesting a Cenomanian age. We found a specimen of Inoceramus sp. with a characteristic plicate sculpture. According to K.-A. Tr6ger (written commun., 1997)

Sandstone-dominated basin floor fan - - Vega Sund Formation

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31 2

F. Surlyk and N. Noe-Nygaard

Fig. 20. Sand-rich turbidite succession of the upper part of the Vega Sund Formation. Note the planar, sheet-like bed geometry and lack of scouring. Type section, locality *4 in Fig. 3. Person encircled for scale.

Fig. 21. Turbidite sandstones of the lower Vega Sund Formation. Type section, locality *4 in Fig. 4. Note the sheet-like geometry and the high sandstone" mudstone ratio. Pencil for scale (to the left).

Cretaceous faulting and associated coarse-grained marine gravity flow sedimentation, Traill O, East Greenland

this type of sculpture is known only from the MiddleUpper Albian boundary interval, Middle Turonian (common), and Lower-Middle Coniacian (common especially in the basal parts). K.-A. Tr6ger on the other hand excluded the lower Middle Albian, Upper Albian, Cenomanian(!), Lower and Upper Turonian, Upper Coniacian, and the Santonian-Maastrichtian. A relatively young Cretaceous age is supported by the location of the section. The Cretaceous succession generally youngs towards the east on the Mols Bjerge Block. The oldest Cretaceous rocks exposed in the Mfinedal valley are Early to earliest Middle Albian mudstones with breccias and isolated blocks occurring immediately adjacent to the Mfinedal Fault. They are followed toward the east by Cenomanian mudstones, Late Turonian-Early Coniacian conglomerates and sandstones, and farthest toward the east by the succession described here. An Early-Middle Coniacian age is thus tentatively suggested. The succession consists of normally graded, medium- to fine-grained sandstone turbidites, 10-25 cm thick, showing well developed Bouma divisions, basal flute casts, parting lineation in the Tb divisions and current ripple lamination in the Tc divisions. The succession is very sand-rich (Figs. 19-21). Bedding is parallel throughout and the beds are tabular and do not wedge out within the extent of the outcrop

313

(Figs. 20 and 21). The base of the beds is always sharp and may be wavy. The sandstones are well sorted in the lower part of the section and become less well sorted higher up, but are, however, still medium-grained with abundant feldspar, and beds are graded, but Bouma sequences are less well developed as below. The sandstones are separated by thin mudstones which in most cases seem to represent Td divisions, or by mudstone turbidites, 5-10 cm thick. The thin muddy interbeds are commonly bioturbated, and C h o n d r i t e s and T a e n i d i u m - l i k e trace fossils are identified. The mudstone content decreases from the lower to the middle part of the section where they become restricted to a few, 5 cm thick, interbeds, whereas the sandstone beds are on average 30 cm thick. The upper part of the section shows a gradual upward thinning of the sandstones, decrease in the sandstone : mudstone ratio, and increase in bioturbation with abundant burrows including P l a n o l i t e s and mantled tubes. The high sandstone : shale ratio, the tabular turbidite sandstone beds of uniform thickness, the lack of scouring, channelling and well defined stacking patterns suggest deposition on the basin floor in the outer part of a sand-rich basin floor fan (Fig. 22). A levee environment can be excluded by the very high sandstone : shale ratio, laterally extensive, tabular beds, lack of climbing ripples and notable absence of slumping.

Fig. 22. Depositional model for the sand-rich basin floor fan of the Vega Sund Formation. The age is highly uncertain, but an Early-Middle Coniacian age is tentatively suggested. PDMF -- Post-DevonianMain Fault; MF - possible early expression of the MfmedalFault.

314

Tectonic implications Donovan (1953, 1955, 1957), Haller (1970) and Price and Whitham (1997) argued that the wide Jurassic block was broken up into much narrower blocks during Cretaceous rift events and that the Mfinedal and Bordbjerg Blocks formed the source areas for the coarse clastic deposits of the strata here included in the new Rold Bjerge and Mfinedal Formations. Our data do not support this interpretation. The M~nedal Block exposes Lower Triassic strata of the Wordie Creek Formation in the footwall of the Mfinedal Fault adjacent to the coarse clastic Lower Cretaceous successions. The breccias and conglomerates contain clasts derived from Upper Permian carbonates which occur in normal stratigraphic continuity with the overlying Lower Triassic sandstones of the Wordie Creek Formation in both the Bordbjerg and M~nedal Blocks (Fig. 3). They were thus not exposed to erosion at the time of conglomerate deposition, and the two blocks can be excluded as provenance areas for at least the Upper Permian limestone clasts. This leaves the block west of the NE-SW-trending segment of the Post-Devonian Main Fault as a source area (Figs. 3 and 23). This block has never been deeply buried as shown by vitrinite reflectance data (Stemmerik et al., 1992). It was probably emerged and subject to erosion in Late Cretaceous time when the Post-Devonian Main Fault was the main basin margin fault (Fig. 23). All available evidence thus indicates that the M~medal Fault did not form the main active fault controlling conglomerate deposition. The literature contains numerous examples of very long runout distances for coarse-grained submarine gravity flows (see reviews by e.g. Picketing et al., 1989; Hampton and Lee, 1996). The inferred travel distances of 25-40 km for the Mfinedal gravity flows fall well within the known spectrum of runout distances for similar flows. The inferred eastward downdip extent of the slope, based on the palaeocurrent data, indicates that the Mols Bjerge Block was not limited to the east by the Laplace Bjerg Fault in Late Cretaceous tame, but was much wider. The Laplace Bjerg Fault was either not formed at the time deposition or did not have any marked expression. If it existed, however, it did not extend south of the E-W-trending William Smith Dal Fault proposed by Donovan (1955) (Fig. 3). In Late Cretaceous time the Mols Bjerge Block was probably limited to the east by the Mols Bjerge Fault or a fault situated offshore, east of Traill 0. The western limit of the Cretaceous block was the Post-Devonian Main Fault but it is possible that an early expression of the M~nedal Fault existed already at that time and that the block(s) between the Post-Devonian Main Fault and the Mfinedal Fault

F. Surlyk and N. Noe-Nygaard

formed a relatively high but still submerged area compared to the Mols Bjerge Block. This interpretation is at marked contrast with that of Price and Whitham (1997) and Price et al. (1997) who suggested a narrow spacing of faults in the Cretaceous with strong fragmentation of the wide Jurassic block. Their interpretation was based on the interpreted amount of erosion of pre-Cretaceous strata on the present-day footwall crests and the derived estimation of fault block widths based on the assumption of a domino-model of extension (cf. Jackson and McKenzie, 1983). Price and Whitham (1997) and Price et al. (1997) further argued that the Mfinedal Fault was the main basin margin fault during Late Bajocian-?Volgian rifting and that narrower blocks were formed during a Valanginian rift phase. The first point was based on a postulated marked thickening of Middle to lower Upper Jurassic strata towards the Mfinedal Fault. This is based essentially on two sections (sections 4 and 5 of Price and Whitham, 1997, their fig. 6). Their section 4 is situated on the BjOmedal Block, limited to the west by the Va~lddal Fault (Fig. 3). This block is tilted towards the east, opposite to almost all other Mesozoic blocks in the region. The V~elddal Fault became active already in the Late Bajocian as demonstrated by detailed studies of Carr (1998) (written commun., 1999), and eastward tilting continued through Middle Jurassic time (Vosgerau et al., 2001). Section 5 of Price and Whitham (1997, fig. 6) is situated in northern Svinhufvud Bjerge, west of the Va~lddal Fault on a westwards-tilted block. It shows thicknesses of the Bristol Elv, Pelion and Olympen Formations close to the double of the maximum thicknesses known anywhere else in East Greenland. The postulated marked thickening towards the Mfinedal Fault is thus in reality based on one section only and the great thicknesses are not confirmed by our work. The second point concerns the suggested Valanginian break-up of the wide Jurassic block into narrower blocks (Price and Whitham, 1997). Clast provenance and palaeocurrent data from the conglomerates and sandstone turbidites show that the blocks were still very wide in Late Cretaceous time, and all palaeocurrent directions reported in this study are towards the east-southeast at a right angle to the main faults. If the original wide blocks were broken up into narrower blocks such as those seen today, the sediment gravity flows would have been deflected into axial SSW-NNE directions by the westwarddipping hanging-wall slopes. The narrow present-day fault blocks seem all to have formed during extensional events associated with Cenozoic break-up. It is, however, likely that some of the faults were initiated

Cretaceous faulting and associated coarse-grained marine gravity flow sedimentation, Traill O, East Greenland

31 5

Fig. 23. Schematic illustration of the Late Permian-Cretaceous tectonic evolution of the Traill 0 area. Main rift phases in the Late PermianEarly Triassic and Middle Jurassic-earliest Cretaceous followed by possibly rift-associated fault events in the earliest Middle Albian, Late Turonian-Early Coniacian and Cenozoic (not shown, at least two phases). Major uplift events in the Early Jurassic, Hauterivian-Barremian and Late Cenozoic. Not to scale. The Pre-Devonian Main Fault was the main western boundary fault throughout Late Palaeozoic-Mesozoic time. An early expression of the M~nedal Fault may have been formed in Cretaceous time, but the fault is mainly a Cenozoic feature. The position of the eastern boundary fault of the wide Late Permian-Mesozoic block is not clear. Most likely it was situated offshore, east of Traill 121,but the Mols Bjerge Fault was probably also active during the Cretaceous.

316 as minor faults during the Volgian-Valanginian as shown for Wollaston Forland by Vischer (1943) and Surlyk (1978). The two Mesozoic rift events described by Price and Whitham (1997) and Price et al. (1997) are also open to discussion. The main Mesozoic rift event has long been known to have been initiated in the Late Bajocian (e.g. Surlyk, 1977, 1990, 2001; Surlyk et al., 1981; Surlyk and Clemmensen, 1983; Surlyk and Noe-Nygaard, 2000). It intensified through the Bathonian-Kimmeridgian to reach a climax in the Volgian, and rifting gradually faded out in the Ryazanian-Valanginian (e.g. Vischer, 1943; Surlyk, 1977, 1978, 2001; Surlyk et al., 1981). The protracted Late Bajocian-Valanginian rift event was succeeded by a number of minor, possibly rift-related fault events and associated coarse-clastic gravity flow deposition in the Cretaceous (Donovan, 1957; Surlyk et al., 1981; this study). There is no direct evidence for a thick Jurassic-earliest Cretaceous syn-rift wedge buried beneath the mid-Cretaceous mudstone succession as proposed by Price and Whitham (1997). The only evidence is the presence of condensed red mudstones of the Valanginian Rc~dryggen Member known from Wollaston Forland in the Mols Bjerge Block on eastern Traill 0. This facies is restricted to the Valanginian and characterises submerged block crests which were isolated from influx of coarser clastics. This has been interpreted as a possible evidence for the Mols Bjerge Block forming an isolated block crest in the Valanginian by analogy with Wollaston Forland (Surlyk, 1978).

Implications for the outer Norwegian Shelf The Cretaceous of the outer Norwegian Shelf has been the target of intense hydrocarbon exploration in recent years. Thick successions of deep-water sandstones have been encountered in several areas in the Veering Basin west of the Nordland Ridge and in the Agat region in the northern North Sea. The sandstones of the outer Vcring Basin are mainly thick, sheet-like turbidites (e.g. Kittilsen et al., 1999; own observations), whereas the inner Veering Basin and the Agat region are dominated by more chaotic deposits, including slides, slumps and debrites (Shanmugam et al., 1994). The deep-water gravity flow deposits of Traill 0 are in general coarser grained and of a more proximal nature than their Norwegian counterparts. However, the slide-slumps complex described here shows similarities to the contorted upper Albian sandstone and mudstone facies of the Agat Field of the Norwegian northern North Sea.

F. Surlyk and N. Noe-Nygaard

The inferred depositional processes and environment of the slide-slump complex are very similar to the model proposed for the Agat region by Shanmugam et al. (1994). The main difference is the more proximal, rift- or fault-related setting of the Traill 0 occurrence. An excellent provenance study of deep-water Cretaceous sandstones in the Norwegian sandstone was presented recently by Morton and Grant (1998). They identified and mapped three distinct sandstone types (K1, K2 and K3) on the basis of heavy-mineral data. K1 occurs on the Trcmdelag Platform, Halten Terrace, and Nordland Ridge, but does not appear to be present further offshore in the deeper Veering Basin. It was derived from the Scandinavian landmass sourced from metasediments of the Caledonian Fold Belt, intrusives of the Transcandinavian Igneous Belt, and to a smaller extent, Svecofennian basement. K3 was derived from the Western Gneiss region further south. K2 occurs in a more basinal location, the mineralogy contrasts with that of K1 and K3, and is not consistent with a source area in Lofoten or the conjugate margin in the Traill 0 region based on samples from Cretaceous turbidite sandstones, some of which are from the sections described in this study. K2 was therefore interpreted as belonging to a separate axial transport system fed by a source in northern North-East Greenland (74-76~ based on zircon ages indicating a source in Palaeoproterozoic (ca. 2000 Ma) and Archaean basement. There are, however, no Archaean zircon ages known from this area. Recent work (Kristine Thrane, written commun., 1999) in East Greenland from the inner fjord zone in the Traill 0 region gives the same zircon age groups as shown for the K2 sandstones by Morton and Grant (1998, fig. 11). Metamorphic zircons in the Krummedal and Smallefjord metasediments yield ages of 900-950 Ma, while detrital zircons are in the range of 1050-2000 Ma. The underlying crystalline basement south of 73~ is Archaean and yields ages between 2500 and 3000 Ma, while the basement north of 73~ is Palaeoproterozoic with ages of 1800-2000 Ma. In the Archaean basement there are granitic intrusives, dikes and formation of leucosome also in the 1800-2000 Ma period which could explain the peak shown in fig. 11 of Morton and Grant (1998). The lack of dates between 500 and 700 Ma and between 2000 and 2400 Ma shown by Morton and Grant (1998, fig. 11) also appears in the data from the conjugate margin in East Greenland between 72 ~ and 74~ The new data from East Greenland thus indicate that large tracts of the East Greenland craton could have served as provenance for the Cretaceous deepwater sandstones in the outer Vcring Basin. The coarse-grained deep-water deposits of the Traill 0

Cretaceous faulting and associated coarse-grained marine gravity flow sedimentation, Traill O, East Greenland

area mainly belong to relatively small systems fed by local, nearby source areas such as uplifted footwalls. They were not part of the larger systems identified in the outer V~ring Basin, but show greater similarity to the systems described from the Agat region by Shanmugam et al. (1994). The sandy basin floor fan system of the Vega Sund Formation, however, is similar in several respects to some of the larger turbidite sandstone systems of the outer Voring Basin. The examples from Traill D illustrate the great variety of gravity flow processes and facies types in the late Early-Late Cretaceous North Atlantic region and in the identification of periods with influx of coarseclastic sediments associated with possibly rift-related faulting, footwall uplift and erosion.

Summary and conclusions Cretaceous subsidence in East Greenland was mainly by thermal contraction following Late Bajocian-Valanginian rifting. The Cretaceous succession is dominated by dark-grey to black siltstones and sandy mudstones and is up to about 2 km thick. Several coarse-grained intervals occur, notably in the basal Middle and Upper Albian, Upper TuronianLower Coniacian, and possibly the Lower-Middle Coniacian. The coarse clastic units relate to Cretaceous, possibly rift-associated fault events. In the Traill D area in North-East Greenland there is a Hauterivian-Barremian hiatus and the underlying strata were deeply eroded during this time interval. Sedimentation was resumed in the Early Aptian in the deepest part of the degraded Jurassic halfgraben basin topography. Extensive drowning of the depositional area seems to have taken place in the mid Late A1bian, as fine-grained sediments of this age are widely distributed, and the transgressive development continued through Late Albian-Early Campanian time. Hiatuses may, however, be present also in this part of the succession as the age control is poor. The Cretaceous deposits occur on a system of relatively narrow fault blocks delimited by N E - to NNEstriking faults. The main faults in the area are from west to east the Post-Devonian Main Fault, and the Bordbjerg, Mfinedal, Laplace Bjerg and Mols Bjerge Faults. The Mgmedal Fault forms the present-day western limit of the Cretaceous sediments. The faults are commonly well exposed but the time of their initiation and periods of major activity is difficult to date. Previous workers have suggested that the wide Jurassic fault block was fragmented into narrower blocks limited by the Bordbjerg and Mfinedal Faults during Cretaceous, possibly rift-related fault events. Provenance of clasts in Cretaceous syn-rift breccias and conglomerates provides information on the

317

age of the succession exposed in the footwall fault scarps at the time of deposition and allow assessment of which faults were active during deposition. The breccias and conglomerates contain abundant clasts of the Upper Permian Karstryggen, Wegener HalvO, and Ravnefjeld(?) Formations and the Lower Triassic Svinhufvuds Bjerge(?) and Odepas Members (Wordie Creek Formation). The position of these units in the footwall of the Mfinedal Block is well below the topographic level of the Cretaceous resedimented deposits in the hanging wall. They furthermore occur in stratigraphic continuity in the Mfinedal and Bordbjerg Blocks and were not exposed to erosion in Cretaceous time. The provenance area of the clasts was situated further towards the west in the block bordered to the east by the Post-Devonian Main Fault. The narrow present-day fault blocks are thus of post-Cretaceous age and seem to have formed during Cenozoic break-up. The coarse-clastic depositional systems described here were not part of the larger turbidite systems of the VOring Basin. They form useful field analogues and identify Cretaceous rift phases in the NorwegianGreenland Sea region.

Acknowledgements This study was carried out under the auspices of project TUPOLAR supported by the Danish Research Councils. We direct our thanks to W.J. Kennedy, K.-A. Tr6ger and W.K. Christensen for fossil determinations, M. Vesterager for wordprocessing, O.B. Berthelsen for darkroom work, L. Hansen and C. Hagen for drafting, L. Stemmerik, S. Piasecki, M. Larsen, J.-E Nystuen and O.R. Clausen for discussion, K. Thrane for providing information on zircon ages from East Greenland, T. Nedkvitne and T. Lien for useful reviews, and O.J. Martinsen for careful editing.

Appendix A

Lithostratigraphy Three of the coarse-clastic successions described in this paper are well-defined lithological units which differ markedly from the intercalated dark mudstones. They form the basis for the erection of three new formations.

Rold Bjerge Formation New History. The sediments of this unit have not been

previously described. Name. After the mountains Rold Bjerge, northeastern Traill 0 (Figs. 1-3).

318

Type section. Pass in Rold Bjerge, 3 km north of the Mfinedal valley (locality * 1 in Fig. 3, Figs. 5-7). Reference sections. The type section is the only known, good exposure of the formation but clusters of large blocks of same type as in type section occur interbedded with dark mudstones along strike on the downthrown side of the Mfinedal Fault. Thickness. A minimum of 25 m are exposed in the type section, but the base of the formation is not seen. Lithology. Chaotic, polymict block breccia with clasts including carbonates of the Upper Permian Karstryggen Formation and Lower Triassic Odepas Member and possibly pebbly sandstones of the Lower Triassic S vinhufvuds Bjerge Member. Fossils. Inoceramid bivalves occur in the immediately overlying mudstones, and dinocysts are recorded from the underlying and overlying mudstones. Depositional environment. Submarine debris avalanches on a faulted slope (Fig. 8). Boundaries. The formation interfingers with poorly exposed Lower Albian and lowermost Middle Albian mudstones both of which may belong to the Home Forland Formation defined in northern Hold with Hope by Kelly et al. (1998). Distribution. Rold Bjerge, Traill O, along the eastern side of the Mfinedal Fault (Figs. 2 and 3). Age. Earliest Middle Albian, may extend down into the Early Albian. M~nedal Formation New Name. After the valley M~medal in southern Rold Bjerge, Traill 0 (Figs. 1-3). Type section. Northern bank of the Mfinedal valley, 500-1000 m east of the Mfinedal Fault (locality *2 in Fig. 3). Reference section. Conglomerates and sandstones exposed on Silja 0 in Vega Sund, and north of Kap Hovgaard on Geographical Society O, possibly belong to this formation, and may serve as reference sections if this can be demonstrated (Fig. 2). Thickness. A thickness of about 80-100 m was estimated by Donovan .(1953). Sections measured independently by us and M. Larsen (pers. commun., 1999) give a thickness of about 45 m. Lithology. Pebble and cobble conglomerates with boulder-dominated beds at some levels, pebbly sandstones, a few pebbly mudstone beds and thin sandy shale mudstones. Fossils. Belemnites, ammonites and inoceramid bivalves. Depositional environment. Faulted slope apron (Fig. 12). Boundaries. The formation overlies dark Cenomanian

E Surlyk and N. Noe-Nygaard

mudstones with thin, slumped sandstones and is overlain by Coniacian or younger dark mudstones both of which may belong to the Home Forland Formation defined in northern Hold with Hope by Kelly et al. (1998). Distribution. M~nedal valley on Traill 0, and possibly also Silja 0 in Vega Sund, and Kap Hovgaard on Geographical Society O. Age. None of the macrofossils give precise dates and a general Late Turonian-Early Coniacian age is suggested.

Vega Sund Formation New History. The succession at the type section was described by Donovan (1953, 1955). Name. After the sound Vega Sund (Figs. 1-3). Type section. The M~.nedal valley gorge, immediately inland of the mouth of the river M~nedal into Vega Sund (locality *4 in Fig. 3). Reference section. The formation is only known from the type section. Thickness. About 90 m is exposed in the type section but the base of the formation is not seen. Lithology. Planar decimetre-thick sandstone turbidites interbedded with thin mudstones; more mudstonerich towards the top. Fossils. Rare inoceramid bivalves, trace fossils at some levels. Depositional environment. Basin floor fan (Fig. 22). Boundaries. Not seen. Probably intercalated between Upper Cretaceous dark mudstones. Distribution. Only known from type section. Age. Equivocal. Both Cenomanian and Coniacian ages have been suggested on the basis of two inoceramid bivalve specimens and geological position. The relatively young, Coniacian, age is tentatively preferred. References Carr, I., 1998. Facies Analysis and Reservoir Characterisation of Jurassic Sandstones from BjCrnedal, Central East Greenland. Unpublished Ph.D. Thesis, University of Reading, 246 pp. Christiansen, EG., Larsen, H.C., Marcussen, C., Hansen, K., Krabbe, H., Larsen, L.M., Piasecki, S., Stemmerik, L. and Watt, W.S., 1992. Uplift study of the Jameson Land basin, East Greenland. Nor. Geol. Tidsskr., 72: 291-294. Clemmensen, L.B., 1980. Triassic lithostratigraphy of East Greenland between Scoresby Sund and Kejser Franz Josephs Fjord. GrCnl. Geol. Unders. Bull., 139: 1-56. Donovan, D.T., 1953. The Jurassic and Cretaceous stratigraphy and palaeontology of Traill 0, East Greenland. Medd. GrCnl., 149(5): 1-14. Donovan, D.T., 1955. The stratigraphy of the Jurassic and Cretaceous rocks of Geographical Society 0, East Greenland. Medd. GrCnl., 103(9): 1-60. Donovan, D.T., 1957. The Jurassic and Cretaceous systems in East Greenland. Medd. GrCnl., 155(4): 1-214.

Cretaceous faulting and associated coarse-grained marine gravity flow sedimentation, Traill O, East Greenland Dor6, A.G., 1992. Synoptic palaeogeography of the Northeast Atlantic Seaway: Late Permian to Cretaceous. In: J. Parnell (Editor), Basins of the Atlantic Seaboard: Petroleum Geology, Sedimentology and Basin Evolution. Geol. Soc. London Spec. PUN., 62: 421-446. Granholm, R-G., 1999. Gjallar mest spennende boring i 1999. Geo, April, pp. 2-3. Haller, J., 1970. Tectonic map of East Greenland (1: 500,000). An account of tectonism, plutonism, and volcanism in East Greenland. Medd. Gr~anl., 171(5): 1-286. Hampton, M.A. and Lee, H.J., 1996. Submarine landslides. Rev. Geophys., 34: 33-59. Jackson, J.A. and McKenzie, D.R, 1983. The geometrical evolution of normal fault systems. J. Struct. Geol., 5: 471-482. Kelly, S.R.A., Whitham, A.G., Koraini, A.M. and Price, S.R, 1998. Lithostratigraphy of the Cretaceous (Barremian-Santonian) Hold with Hope Group, NE Greenland. J. Geol. Soc., London, 155: 993-1008. Kittilsen, J.E., Olsen, R.R., Marten, R.F., Hansen, E.K. and Hollingsworth, R.R., 1999. The first deepwater well in Norway and its implications for the Upper Cretaceous play, Vcring basin. In: A.J. Fleet and S.A.R. Boldy (Editors), Petroleum Geology of Northwest Europe: Proceedings of the 5th Conference. Geological Society, London, pp. 275-280. Koch, L. and Haller, J., 1971. Geological map of East Greenland 72~176 Lat. (1 : 250,000). Medd. Gr~nl. 183, 13 map sheets. Larsen, M., Nedkvitne, T. and Olaussen, S., 1999. A coarse-grained clastic wedge (Barremian-Aptian) in North-East Greenland sedimentology, sequence stratigraphy and regional implications. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway Palaeozoic to Recent. Extended Abstracts, Grieghallen 3-5 May, 1999, Norwegian Petroleum Society, Stavanger, pp. 179-182. Maync, W., 1947. Stratigraphie der Jurabildungen Ostgr6nlands zwischen Hochstetterbugten (75~ und dem Kejser Franz Joseph Fjord (73~ Medd. Gr~anl. 132(2), 232 pp. Maync, W., 1949. The Cretaceous beds between Kuhn Island and Cape Franklin (Gauss Peninsula), northern East Greenland. Medd. Gr0nl. 133(3), 291 pp. Mohrig, D., Elverhr A. and Parker, G., 1999. Experiments on the relative mobility of muddy subaqueous and subaerial debris flow, and their capacity to remobilize antecedent flows. Mar. Geol., 154: 117-129. Morton, R.C. and Grant, S., 1998. Cretaceous depositional systems in the Norwegian Sea: heavy mineral constraints. Am. Assoc. Pet. Geol. Bull., 82: 274-290. Nemec, W., 1990. Aspects of sediment movement on steep delta slopes. Int. Assoc. Sedimentol. Spec. PUN., 10: 29-73. N~ahr-Hansen, H., 1993. Dinoflagellate cyst stratigraphy of the Barremian to Albian, Lower Cretaceous, East Greenland. GrOnl. Geol. Unders. Bull., 166: 1-171. Pickering, K.T., Hiscott, R.N. and Hein, F.J., 1989. Deep-Marine Environments: Clastic Sedimentation and Tectonics. Unwin Hyman, London, 416 pp. Price, S.R and Whitham, A.G., 1997. Exhumed hydrocarbon traps in East Greenland: analogs for the Lower-Middle Jurassic play of Northwest Europe. Am. Assoc. Pet. Geol. Bull., 81:196-221. Price, S.E, Brodie, J., Whitham, A.G. and Kent, R., 1997. Mid-Tertiary rifting and magmatism in the Traill el region, East Greenland. J. Geol. Soc., London, 154:419-434. Shanmugam, G., Lehtonen, L.R., Straume, T., Syvertsen, S.E., Hodgkinson, R.J. and Skibeli, M., 1994. Slump and debris-flow dominated upper slope facies in the Cretaceous of the Norwegian and northern North Sea (61-67~ implications for sand distribution. Am. Assoc. Pet. Geol. Bull., 78: 910-937. Stemmerik, L., Christiansen, F.G., Piasecki, S., Jordt, B., Mar-

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cussen, C. and N~ahr-Hansen, H., 1992. Depositional history and petroleum geology of the Carboniferous to Cretaceous sediments in the northern part of East Greenland. In: T.O. Vorren, E. Bergsager, ~O.A. Dahl-Stamnes, E. Holter, B. Johansen, E. Lie and T.B. Lurid (Editors), Arctic Geology and Petroleum Potential. Norwegian Petroleum Society, Special Publication 2. Elsevier, Amsterdam, pp. 67-87. Stemmerik, L., Dam, G., Noe-Nygaard, N., Piasecki, S. and Surlyk, F., 1998. Sequence stratigraphy of source and reservoir rocks in the Upper Permian and Jurassic of Jameson Land, East Greenland. Geol. Greenl. Surv. Bull., 180: 43-54. Surlyk, F., 1977. Mesozoic faulting in East Greenland. In: R.T.C. Frost and A.J. Dikkers (Editors), Fault Tectonics in NW Europe. Geol. Mijnbouw, 56:311-327. Surlyk, F., 1978. Submarine fan sedimentation along fault scarps on tilted fault blocks (Jurassic-Cretaceous boundary, East Greenland). Gr~anl. Geol. Unders. Bull., 128: 1-108. Surlyk, F., 1984. Fan-delta to submarine fan conglomerates of the Volgian-Valanginian Wollaston Forland Group, East Greenland. In: E.H. Koster and R.J. Steel (Editors), Sedimentology of Gravels and Conglomerates. Mere. Can. Soc. Pet. Geol., 10: 359-382. Surlyk, F., 1989. Mid-Mesozoic syn-rift turbidite systems: controls and predictions. In: J.D. Collinson (Editor), Correlation in Hydrocarbon Exploration. Norwegian Petroleum Society, Graham and Trotman, London, pp. 231-241. Surlyk, F., 1990. Timing, style and sedimentary evolution of Late Palaeozoic-Mesozoic extensional basins of East Greenland. In: R.ER Hardman and J. Brooks (Editors), Tectonic Events Responsible for Britain's Oil and Gas Reserves. Geol. Soc., Spec. Publ., 55: 107-125. Surlyk, F., 1991. Sequence stratigraphy of the Jurassic-lowermost Cretaceous of East Greenland. Am. Assoc. Pet. Geol. Bull., 75: 1468-1488. Surlyk, F., 2001. The Jurassic of East Greenland: a sedimentary record of thermal subsidence, rift initiation and climax. In: F. Surlyk and J.R. Ineson (Editors), The Jurassic of Denmark and Greenland. Geol. Denm. Surv. Bull. (in press). Surlyk, F. and Clemmensen, L.B., 1983. Rift propagation and eustacy as controlling factors during Jurassic inshore and shelf sedimentation in northern East Greenland. Sediment. Geol., 34: 119-143. Surlyk, F. and Noe-Nygaard, N., 2000. Jurassic Sequence Stratigraphy of East Greenland. GeoRes. Forum, 6: 357-366. Surlyk, F., Clemmensen, L.B. and Larsen, H.C., 1981. Post-Paleozoic evolution of the East Greenland continental margin. In: J.W. Kerr and A.J. Ferguson (Editors), Geology of the North Atlantic Borderlands. Mem. Can. Soc. Pet. Geol., 7: 611-645. Surlyk, F., Dam, G., Engkilde, M., Hansen, C.F., Larsen, M., NoeNygaard, N., Piasecki, S., Therkelsen, J. and Vosgerau, H., 2001. Jurassic stratigraphy of East Greenland. In: F. Surlyk and J.R. Ineson (Editors), The Jurassic of Denmark and Greenland. Geol. Denm. Surv. Bull. (in press). Vischer, A., 1943. Die postdevonische Tektonik von Ostgr6nlands zwischen 74~ und 75~ Br., Kuhn O, Wollaston Forland, Clavering 0 und angrenzende Gebiete. Medd. Gr~anl., 133(1): 1-195. Vosgerau, H., Alsen, R, Therkelsen, J., Stemmerik, L. and Surlyk, F., 2001. Early rift sedimentation on an antithetic Jurassic fault block, Traill 121,East Greenland. Geol. Greenl. Surv. Bull. (in press). Whitham, A.G., Price, S.R, Koraini, A.M. and Kelly, S.R.A., t999. Cretaceous (post-Valanginian) sedimentation and rift events in NE Greenland (71-77~ In: A.J. Fleet and S.A.R. Boldy (Editors), Petroleum Geology of Northwest Europe: Proceedings of the 5th Conference. Geological Society, London, pp. 325-336. Ziegler, R, 1988. Evolution of the Arctic-North Atlantic and the Western Tethys. Am. Assoc. Pet. Geol., Mem., 43: 1-198.

Geological Institute, University of Copenhagen, Oster Voldgade 10, DK-1350 Copenhagen K, Denmark Geological Institute, University of Copenhagen, Oster Voldgade 10, DK-1350 Copenhagen K, Denmark

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Cretaceous-Tertiary palaeo-bathymetry in the northern North Sea; integration of palaeo-water depth estimates obtained by structural restoration and micropalaeontological analysis R. Kyrkjebo, T. Kjennerud, G.K. Gillmore, J.I. Faleide and R.H. Gabrielsen

Temporal and spatial variations in palaeo-water depth are crucial parameters in basin analysis since changes in palaeo-bathymetry detail the amount of sediment underfill during basin evolution. By carefully integrating seismic-stratigraphic observations with palaeo-water depth estimates from structural restoration and micropalaeontological data, changes in accommodation space throughout the Cretaceous-Tertiary post-rift interval are documented on a regional scale in the northern North Sea. Since it is not possible to determine the palaeo-water depth exactly, we have focused on determining most likely water depth figures, and identifying the principal shallowing and deepening trends. The inferred trends from the investigated wells are generally in good agreement with each other on a regional scale, especially when the tectonic position within the basin is taken into account. The inferred general trends are: (1) general shallowing superimposed on several transgressive/regressive events during the Early Cretaceous; (2) deepening from the early Cenomanian to mid-Campanian; (3) shallowing from the mid-Campanian to latest Maastrichtian; (4) deepening in the Early to Late Paleocene; (5) shallowing from the Late Eocene to Late Miocene; (6) deepening from the Late Miocene to Early Pliocene; (7) shallowing during Pliocene time. The early Cenomanian to latest Maastrichtian and the late Eocene to Pliocene events correspond with changes in eustatic sea level, but the deepening/shallowing trends were probably amplified by tectono-thermal effects. The events in the Early Cretaceous, Early to Late Paleocene, and Late Miocene to Early Pliocene cannot be explained by the eustatic sea-level curve, and therefore need to be explained by purely tectono-thermal events.

Introduction

Palaeo-water depths defining the shape of the basin are crucial parameters in basin analysis since changes in palaeo-bathymetry detail the amount of sediment underfill in basins (Gradstein and B~ickstr6m, 1996). In basin modelling, palaeo-water depth is an important input parameter that controls the measured subsidence/uplift. However, palaeo-water depth is difficult to determine, and depends largely on the quality and quantity of data from boreholes and seismic data. Even in intensively investigated basins such as the North Sea Basin, minor attention has been paid to palaeo-bathymetry. Berggren and Gradstein (1981), Gradstein et al. (1994) and Gradstein and B~ickstr6m (1996) outlined Cenozoic palaeo-bathymetric trends using micropalaeontological data from several wells in the North Sea. Other workers (Jones, 1988; Sloan, 1995) focused on recognition of Palaeogene water depths. Bertram and Milton (1989), and Joy (1993) assessed palaeo-water depth by combin-

ing key horizons in wells formed at a "known" water depth with observed correlative sedimentary thicknesses. The present analysis of palaeo-water depth development of the Cretaceous-Tertiary of the northern North Sea Basin between 58~ and 62~ (Fig. 1) is rooted in micropalaeontological analyses, structural restoration ("back-stripping") and seismic-stratigraphic considerations. Finally, the results of these analyses have been tested against general geological information from the study area. The procedure in this study was as follows. (1) 2D reflection seismic data were interpreted and analysed by seismic-stratigraphic methods. The data were calibrated against cores, cuttings and wireline logs from more than 50 North Sea wells. Altogether, 16 seismic-stratigraphic units were defined (Jordt et al., 1995; Gabrielsen et al., 2001). (2) Structural restoration along four regional transects was performed to assess palaeorelief (Kjennerud et al., 2001). (3) Micropalaeontological data from 12 wells were

Sedimentary Environments Offshore Norway - Palaeozoic to Recent edited by O.J. Martinsen and T. Dreyer. NPF Special Publication 10, pp. 321-345, Published by Elsevier Science B.V., Amsterdam. 9 Norwegian Petroleum Society (NPF), 2001.

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R. K y r k j e b O e t al.

Fig. 1. The northern North Sea Basin comprises several prominent structural features, as displayed on this structural map of earliest Ryazanian level. The data base utilised in this study contained the four displayed transects and twelve key wells. ESB = East Shetland Basin; E S P -- East Shetland Platform; L T = Lomre Terrace; M g B = Magnus Basin; M r B - Marulk Basin; M F B = MSlOy Fault Blocks; SG - Sogn Graben; SB = Stord Basin; U t t = Utsira High; VG = Viking Graben;/]G = Asta Graben; U T - Uer Terrace; H P = Horda Platform.

re-analysed, and palaeo-water depths for each well were determined (Gillmore et al., 2001). (4) Results from the analyses, as described above

(1-3), were integrated for each well, and for relevant structural positions, to define deepening/shallowing palaeo-water depth trends.

Cretaceous-Tertiary palaeo-bathymetry in the northern North Sea

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(5) Finally, all available data were utilised to produce palaeo-bathymetric maps for the entire study area at different stages in basin evolution. In this paper we discuss the integration of the results from Kjennerud et al. (2001) and Gillmore et al. (2001), supported by information from the regional aspects of Cretaceous-Tertiary evolution (e.g. Jordt et al., 1995; Gabrielsen et al., 2001). Database and method

Conventional multi-channel reflection seismic surveys (NVGT88, NVGTI92, SG8043, NNST89, NNSTI91, VGST89, SBGS87, and CNST82) were interpreted. Sixteen regional post-rift seismic units were defined in the reflection seismic data by seismic stratigraphy and correlated to well data, calibrated to the chronostratigraphic time-scale of Gradstein and Ogg (1996) (Fig. 2), and mapped in the entire study area. Six of the seismic units are of Cretaceous age (K1 to K6: base Ryazanian-Maastrichtian). Subdivision of the Cainozoic (Base Tertiary-Present) sequence into ten units (CSS-1 to CSS-10) were adopted from Jordt et al. (1995). Datings are partly uncertain for the Miocene of the northern North Sea, and in the seismic interpretation it was not possible to differentiate between the Miocene seismic units CSS-5, CSS-6 and CSS-7 (Jordt et al., 1995). Therefore, palaeo-water depth estimates were not obtainable for each Miocene unit by structural restoration. Four regional crustal transects (Figs. 1 and 3) were constructed by combining conventional seismic lines with deep seismic data, gravimetric data and magnetic data. All transects were depth-converted using velocity information from wells, interval velocities from stacking velocities and velocities from seismic refraction data (ESP). Constrained by the transects, structural restoration was performed using the software LOCACE (Beicip Franlab, 1996). The palaeorelief was restored along the regional profiles using the following criteria: (1) depositional geometries; (2) sedimentological evidence of sub-aerial exposure and shallow water depth; (3) erosional unconformities; (4) tectonic movements. The method and discussions concerning the results from the structural restoration are described in Kjennerud et al. (2001). Since it is not possible to differentiate between continental onlap, coastal onlap and marine onlap without having palaeo-bathymetric information from sediment composition or fossil content (Bertram and Milton, 1989; Roberts et al., 1993), the most critical pitfall in this restoration method is the potentially incorrect definition of zero water depth. Therefore, the method provides estimates of palaeorelief along the transects at the base of each

Fig. 2. The seismic-stratigraphic framework comprises sixteen postrift seismic units calibrated to chronostratigraphy (Gradstein and Ogg, 1996).

seismic unit rather than absolute palaeo-water depth (Kjennerud et al., 2001). The structural restoration trends as recorded through time (hereafter denoted SRT), are displayed as a red curve for positions corresponding to the wells in Figs. 4-9. In most cases this

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Fig. 3. Four regional geoseismic sections (transect 1-4) were basis for the structural restoration. The transects were depth converted and displayed with focus on the post-rift interval (see stratigraphic correlation chart in Fig. 2).

basin relief can be regarded as a period of minimum water depth (Kjennerud et al., 2001). It is, however, particularly emphasised that sub-aerial topography cannot be determined by this method. In addition, structural restoration gives palaeo-water depth estimates only at the base/top of each seismic unit. Twelve wells from the Norwegian part of the northern North Sea were studied in great detail.

These wells occupy different structural and geographical settings within the basin and are linked to the transects (Figs. 1 and 3). Wells were transferred to the transects by projection along the actual structural element. Lateral morphological variations observed in the seismic reflection data away from the transects should therefore be taken into account when comparing results from

Cretaceous-Tertiary palaeo-bathymetry in the northern North Sea

these wells with the structural restoration. This is particularly relevant for wells 31/6-3 and 33/9-18. Wells 15/5-3 and 34/7-1 are not directly linked to the transects (Fig. 1), and have been used for regional calibration only. The well data were analysed focusing on lithological composition, dating and analysis of micropalaeontological data to assess palaeo-water depths. The lithological characteristics were also used in the decompaction of the transects (Kjennerud et al., 2001). The results are summarised in diagrams for each well (Figs. 4-10) also including a link to the interpreted seismic horizons. The micropalaeontological methodology utilised to assess palaeo-water depth is reported in Gillmore et al. (2001). As pointed out by Roberts et al. (1993), palaeobathymetric information cannot exclusively be documented from seismic stratigraphy. On the other hand, studies of palaeo-water depth based on micropalaeontology from single wells would only give a depth approximation in the well location. Therefore, regional understanding of palaeo-water depth depends upon seismic-stratigraphic calibration. Uncertainties associated with palaeo-water depth estimates based on micropalaeontological analysis tend to be large, and increase with increasing depths. In most published work (e.g. Jones, 1988; Gibson, 1988; Gradstein and B~ickstrOm, 1996) water depth has been related to basin position relative to the shoreline or to palaeoenvironment, and not directly to depth (for further discussion see Gillmore et al., 2001). The micropalaeontological analysis (Gillmore et al., 2001) gave a most likely range (hereafter denoted MLR) estimate of palaeo-water depth for each sample at the corresponding time, displayed as solid error bars in Figs. 4-9. The MLR is commonly in the order of 300 m (or more), and in some cases includes an additional maximum/minimum range as indicated by dotted error bars. In the next step, a most likely trend through time (hereafter denoted MLT) was determined (Gillmore et al., 2001) for each well (displayed as a blue curve in Figs. 4-9). In the present study, results from each well are summarised, and trends from different wells in the same general structural position compared. Palaeowater depth trends along each transect are also compiled in order to test consistence in the data sets when different structural positions are compared. In this procedure, the sixteen Cretaceous-Tertiary regional seismic-stratigraphic isopach maps were also utilised (Jordt et al., 1995; Gabrielsen et al., 2001). Finally, integrated palaeo-water depth curves, based on the two methods and other geological information from the study area will be discussed.

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Post-rift palaeo-bathymetric development In the following, establishment of the post-rift palaeo-water depth trends for different structural positions is emphasised. The sensitivity for shallowing/deepening trends is considered at best to be 100 m.

Tampen Spur Well 3 3 / 9 - 1 8 Well 33/9-18 is located in the hanging wall of the Statfjord fault block at the Tampen Spur (Figs. 1, 3 and 4a). Micropalaeontological data are available for the time interval late Turonian-Quaternary. Cretaceous. The earliest Ryazanian to early Hauterivian and the intra-Barremian to intra-Turonian intervals are interpreted as hiatuses caused by sub-aerial exposure. Only a thin (7 m) sequence of sediments is preserved, but not sampled, between the two possible erosional events. Hence, no information on the palaeo-water depth is available from this interval, other than possible indication of sub-aerial exposure related to the hiatuses. Due to lack of micropalaeontological evidence for the Early Cretaceous, the trend from the structural restoration was given priority in the final palaeowater depth curve (Fig. 4a). For the entire Cretaceous interval, palaeo-water depth estimates based on structural restoration suggest shallower palaeo-water depth than the minimum estimates extracted from micropalaeontological data (Fig. 4a). The differences in depth estimates of the MLT and SRT can most likely be explained by the possibility that the available accommodation space was greater in the well position than in the transect analysed (Kjennerud et al., 2001). Therefore, the Late Cretaceous trend relies entirely on the MLT, which follows the mean values of the MLR estimates (Gillmore et al., 2001). It is, however, realised that the amplitude of the deepening event in earliest Cenomanian to intra-Turonian times may have been overestimated for the well position when combining the SRT and the MLT. Also large uncertainties, in some cases MLR of 200-750 m, are associated with the palaeo-water depth estimates of the Late Cretaceous (Fig. 4a). At the Cretaceous/Tertiary boundary, zero water depth was obtained by structural restoration, while the interpretation of the micropalaeontological data suggests a palaeo-water depth of 300 m (Fig. 4a). Therefore, the results offered by the structural restoration for the Cretaceous/Tertiary boundary is taken as an indication that the basin-relief was low, but the data cannot be used to estimate the absolute palaeowater depth (Kjennerud et al., 2001). Hence, the MLT

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Cretaceous-Tertiary palaeo-bathymetry in the northern North Sea

result is favoured as palaeo-water depth estimate for the Cretaceous/Tertiary boundary.

Tertiary. The MLT for the Tertiary is supported by the SRT, when the general trends are compared (Fig. 4a). The deviations found for instance for the Eocene can be explained by better time resolution in the MLT than in the SRT. The MLR for the Early Eocene suggests a water depth of 200-600 m, and the MLT is estimated to be in the uppermost part of the MLR (Gillmore et al., 2001). Based on the present data it could not be decided with any certainty whether sub-aerial/submarine erosion or non-deposition caused the Middle Miocene (CSS-6) hiatus. A palaeo-water depth range of 0-200 m is suggested based on micropalaeontological data for the Early Miocene (Fig. 4a). Assuming that the Middle Miocene hiatus was caused by sub-aerial erosion, a shallowing of the basin from Early to Middle Miocene time is evident. The interpreted MLR for the Late Miocene is 100-500 m. It is suggested that the MLT is in the lowermost range of the estimated palaeo-water depth interval during the Late Miocene. In general, results from the two methods are in good agreement when the shallowing/deepening trends are compared (Fig. 4a).

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ing into account that sub-aerial exposure during parts of the Eocene-Pliocene interval is evident, several shallowing/deepening events seem to have occurred (Fig. 4b).

Summary Wells 34/7-1 and 33/9-18 located at the Tampen Spur area (Figs. 1 and 10b) both indicate a general shallowing from the Middle Eocene to Late Miocene, but especially the Neogene is associated with uncertainties. Due to lack of data, information of the pre-Eocene palaeo-water depth was not available from well 34/7-1 to be compared to the results from well 33/9-18 (Fig. 4). Therefore, a general palaeowater depth development for the Tampen Spur relies heavily on the results from well 33/9-18 (Figs. 4 and lOb).

Viking Graben Well 34/11-1 Well 34/11-1 is located on an intra-basinal high in the northern Viking Graben (Figs. 1, 3 and 5a). From this well micropalaeontological data are available for the interval from the Early Ryazanian to Late Oligocene.

Cretaceous. Most of the Early Cretaceous is Well 34/7-1 Well 34/7-1 is located at the crest of the Snorre fault block at the Tampen Spur (Figs. 1 and 4b). The well provided single palaeo-water depth estimates based on micropalaeontological data from the interval Late Eocene to Present. The well is not linked to any transect.

Tertiary. Three hiatuses (Late Eocene, late-EarlyLate Oligocene and late-Early-latest Miocene) are encountered in this interval, all probably related to sub-aerial exposure. Since it was not possible to state at which time sub-aerial exposure took place, a most likely trend was difficult to assess (Gillmore et al., 2001). The analysis suggests MLR of 300600 m in the Middle Eocene, 100-600 m in the Early Oligocene, 200-500 m in the Late Oligocene, 100-300 m in the earliest Miocene, 100-400 m in the earliest Pliocene, and 100-200 m in the Late Pliocene. This could possibly indicate a general shallowing from the Middle Eocene to Quaternary. Tak-

characterised by hiatuses or condensed intervals (Fig. 5a). Palaeo-water depth estimates from the preserved Lower Cretaceous strata gave MLR of 50-200 m for the Ryazanian, 100-500 m for the intra-Barremian and 200-500 m for the latest Albian. Gillmore et al. (2001) suggested that the MLT for the Early Cretaceous followed the lowermost part of the MLR. Thus, it is believed that general deepening occurred from the earliest Ryazanian to the early Campanian, and that the interpretation of micropalaeontological data yielded reasonable depth estimates for this interval. Palaeo-water depth estimates obtained by structural restoration of transect 1 suggest deeper basin conditions for the Ryazanian-latest Turonian than suggested by the MLT (Fig. 5a). Also the indicated MLR and additional maximum ranges point to possible deeper basin conditions than suggested by the MLT. Unfortunately, it was not possible to decide what caused the hiatuses, and it is therefore realised that the Early Cretaceous evolution could have been

Fig. 4. Summary of the interpreted palaeo-water depth from wells 33/9-18 (a) and 34/7-1 (b). The most likely trend (MLT) based on micropalaeontological data is displayed in blue, whereas the structural restoration trend (SRT) is in red. Horizontal black bars indicate the most likely range based on micropalaeontological data (MLR). Hiatuses are indicated in the lithology column, where also dominant lithologies - - shales/mudstones in green, sandstones in yellow and limestones in blue are included. The integrated palaeo-water depth curve is plotted in green.

Fig. 5. Summary of the interpreted palaeo-water depth from wells 34/11-1 (a) and 30/4-1 (b). For detailed explanation, see Fig. 4.

I t...,

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Fig. 7. Summary of the interpreted palaeo-water depth from wells 24/9-1 (a) and 15/5-3 (b). For detailed explanation, see Fig. 4.

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r Fig. 8. Summary of the interpreted palaeo-water depth from wells 35/11-3 (a) and 35/12-1 (b). For detailed explanation, see Fig. 4.

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Fig. 9. Summary of the interpreted palaeo-water depth from wells 31/6-3 (a) and 17/3-1 (b). For detailed explanation, see Fig. 4.

Cretaceous-Tertiary palaeo-bathymetry in the northern North Sea

more complex than that reflected by the suggested weak deepening trend. For the Cretaceous/Tertiary transition the micropalaeontological data indicate palaeo-water depth in the order of 300 m (MLT), whereas structural restoration offered zero water depth, which is taken to indicate low basin relief (Kjennerud et al., 2001). Apart from this, the trends obtained by structural restoration support the trends indicated by the MLT (Fig. 5a). Tertiary. A deepening occurred in the Paleocene, probably reaching a palaeo-water depth in excess of 600 m. This depth was preserved during the Eocene. It is noted, however, that shallowing is indicated by the MLT for the latest Paleocene (Fig. 5a). The SRT is considered to be too deep for the Eocene, since it suggests water depths greater than the maximum range obtained by micropalaeontological data. As indicated by the maximum range, it is also noted that palaeo-water depth could have been in excess of 500 m for the Eocene. A general shallowing from latest Eocene to Late Oligocene time is indicated by the integrated curve. No micropalaeontological data exist which support the near-zero water depth suggested by the structural restoration for the Miocene. However, the presence of sediment in unit CSS-5, which thins westwards, indicates that a certain palaeo-water depth existed in the Early Miocene. From the Late Miocene to Present, the results rely entirely on the structural restoration (Fig. 5a). Well 30/4-1 Well 30/4-1 is located on an intra-basinal high in the northern Viking Graben (Figs. 1, 3 and 5b). The interpretation of palaeo-water depth in this well is hampered by sparse micropalaeontological data and by the lack of recordings of Tertiary deep-water faunas in the well report (Gillmore et al., 2001).

Cretaceous. It is assumed that the hiatus indicated for the early Ryazanian (Fig. 5b) is related to extreme condensation, and it is most likely that deep basin conditions existed as suggested by the structural restoration. The few existing micropalaeontological data of the Cretaceous suggested palaeo-water depths of 150-200 m for the late Albian, and 300-400 m in the late Campanian (Fig. 5b). These estimates were given priority when combining the MLT with the SRT. Both methods indicate shallowing during the Maastrichtian. Palaeo-water depth of 150 m most likely occurred at the Cretaceous/Tertiary boundary, as indicated by the micropalaeontological data.

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Tertiary. Uncertainties in the MLR palaeo-water depth estimates in the Tertiary are in the order of 100-500 m for the Palaeogene. It is also noted that the MLT for the Palaeogene is in the upper range of the MLR. When comparing with the EoceneOligocene interval, deviations are observed in absolute palaeo-water depth estimates from the two methods. Gillmore et al. (2001) discussed problems related to recognition of deep-water faunas in the Eocene samples from well 30/4-1. They claimed that deeper basin conditions existed for the Eocene than suggested by the MLR. Better constraints exist for the Early Miocene estimates for which MLR of 0-100 m is suggested (Fig. 5b). A hiatus of erosive nature was tracked, covering the Middle to Late Miocene interval. However, it is assumed that deepening started in the Late Miocene and a palaeo-water depth of more than 400 m was reached in the earliest Pliocene, as suggested by the SRT. The MLT and SRT seem to be in agreement, but the low amount of micropalaeontological data makes it impossible to assess a MLT for the entire post-rift interval. Therefore, the integrated palaeo-water depth curve relies mainly on the SRT due to incompleteness of the MLT. Where estimates from micropalaeontological data exist, these are taken into account (Fig. 5b). Well 30 / 10-6 Well 30/10-6 is located on an intra-basinal high in the northern Viking Graben, close to 60~ (Figs. 1, 3 and 6a). Minor lateral variations in topography between the well and the transect position are observed in the seismic reflection data. In addition, the well includes the most complete record of post-rift strata in the Viking Graben (Fig. 6a). Micropalaeontological data coverage for the Early Cretaceous interval is, however, sparse.

Cretaceous. Ingeneral, good agreement exists between the MLT and SRT (Fig. 6a). However, some contradictions are noted. The SRT indicated shallowing followed by deepening for the Early Cretaceous. By contrast, general deepening is suggested by the MLT. Therefore, the following assumptions were made for the Early Cretaceous integrated palaeowater depth trend (Fig. 6a). (1) Palaeo-water depth between 200 and 300 m existed in the earliest Ryazanian, as suggested by the structural restoration. (2) The proposed most likely estimates from the interpreted micropalaeontological data are valid. (3) The hiatus during Aptian time was caused by sub-aerial erosion, consequently, zero palaeo-water depth was reached in the intra-Aptian.

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Fig. 10. Comparison of the post-rift palaeo-water depth trends (in green) as obtained from the twelve selected North Sea wells. The six wells located in the Viking Graben are displayed in (a), the others in (b). The Haq et al. (1987, 1988) curve is included (in red) for comparison in all wells.

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faO r r

336 From the intra-Cenomanian to middle Campanian the MLR was typically 200-400 m (Fig. 6a). A slight shallowing is seen in the MLT curve from the early Cenomanian to the early Santonian, followed by deepening until the late Campanian (400 m). In contrast, the SRT indicates rapid deepening during earliest Cenomanian (K-3) to a water depth close to 600 m, followed by slight shallowing until latest Campanian time (Base K-6). Large uncertainties are suggested for the MLR in the late Campanian (Fig. 6a), which indicate that water depths of up to 1000 m could have existed. It is also noted that the MLT followed the lowermost range of the MLR estimates for late Campanian and Maastrichtian times (Fig. 6a). Water depths of approximately 150 m existed in the latest Maastrichtian (Gillmore et al., 2001). Consequently, zero palaeowater depth, as suggested by structural restoration at the Cretaceous/Tertiary boundary, is rejected as an absolute palaeo-water depth estimate. The integrated curve relies mainly on the MLT for early Cenomanian to latest Maastrichtian times. However, a slightly deeper palaeo-water depth, as constrained by the SRT, is suggested for the late Campanian.

Tertiary. The MLT (Gillmore et al., 2001) was proposed to coincide with the mean value of most of the reported range estimates (Fig. 6a). The structural restoration trend supports the MLT for the Tertiary, but indicates shallower depth for the Eocene to Late Miocene (Kjennerud et al., 2001). Zero palaeo-water depth was obtained by structural restoration for the Middle Miocene, followed by deepening in the Late Miocene (Kjennerud et al., 2001). In contrast, general shallowing during the Miocene-Pliocene is suggested by the MLT. The integrated curve is identical to the MLT for the Early Paleocene to the Late Miocene. Deeper palaeo-water depth, as constrained by the SRT, is suggested for Late Miocene to Pliocene times based on the assumption that the preserved sediments were deposited in a certain accommodation space (Fig. 6a). Well 29/6-1 Well 29/6-1 is located on the transition between the Viking Graben and the Hild-Alwyn Alignment (Figs. 1, 3 and 6b). The well report contains micropalaeontological information for the mid-Campanian to Quaternary interval.

Cretaceous. A hiatus caused by sub-aerial erosion was tracked for the entire Early Cretaceous interval (Fig. 6b). From the SRT data it is assumed that zero palaeo-water depth occurred in the latest Albian.

R. KyrkjebO et al.

The MLR indicates a possible palaeo-water depth for the mid-Campanian and the early Maastrichtian to be in the order of 100-600 m. Less uncertainty, 200-500 m, was affiliated with the estimate for the latest Maastrichtian. From the mid-Campanian to the latest Maastrichtian the most likely palaeo-water depth was around 300 m (Gillmore et al., 2001). It is assumed that the shallowing as seen by SRT was related to a low basin relief at the Cretaceous/Tertiary boundary, rather than an indication of zero water depth (Kjennerud et al., 2001).

Tertiary. The MLT for the Tertiary is suggested to be deepening from late-Early Paleocene to Late Paleocene followed by general shallowing to the Early Pliocene (Fig. 6b). A water depth of 600 m in the mid-Paleocene is constrained by the MLR of 500-750 m for this interval. Also the MLT for the Early Miocene of 150 m, is uncertain (MLR of 100200 m). In contrast, the Eocene palaeo-water depth may have been 200-750 m. The MLT is suggested to have been in the uppermost MLR for the Early Eocene, while the Middle to Late Eocene MLT was suggested in the lower MLR (Gillmore et al., 2001). Minor contradictions exist when the MLT and SRT are compared (Fig. 6b). For the Eocene, shallowing of palaeo-water depth is indicated by the MLT, but not supported by the SRT. In this case, it is possible that larger sediment thickness seen on the transect compared with the true well position could be taken as evidence for deeper basin conditions (Kjennerud et al., 2001). The observed hiatuses for the middle-Late Miocene and for the latest Miocene to Early Pliocene were probably caused by sub-aerial erosion. Both methods indicate deeper palaeo-water depth :in the Early Pliocene than in the Late Miocene. The timing of the deepening event is unfortunately not constrained. By attributing the Late Miocene-Early Pliocene hiatus to sub-aerial erosion, the deepening must have taken place in the Early Pliocene. The post-rift palaeo-water depth trend, as seen from well 29/6-1, relies on the SRT for the earliest Ryazanian to mid-Campanian. For the rest of the interval the MLT is adopted in the integrated curve (Fig. 6b). Well 24/9-1 Well 24/9-1 is located in the southern Viking Graben (Figs. 1, 3 and 7a). Micropalaeontological data are available from the Paleocene to Late Oligocene.

Cretaceous. For well 24/9-1 the Cretaceous palaeo-water depth trends mostly rely on the SRT (Fig. 7a). The suggested zero palaeo-water depth at

Cretaceous-Tertiary palaeo-bathymetry in the northern North Sea

the Cretaceous/Tertiary boundary was, however, rejected because deep basin conditions existed in the earliest Paleocene (Gillmore et al., 2001), and because regional low relief not attributed to zero palaeowater depth has been suggested for other Viking Graben wells. Tertiary. For the Paleocene to the Early Oligocene, the MLT was adopted in the integrated curve (Fig. 7a). The MLR indicates that palaeo-water depth of 500-750 m existed in the Early Paleocene. Hence, a deepening is assumed to have started in the earliest Paleocene. The MLR of the middle Paleocene was of 200-600 m while the Late Paleocene MLR was estimated to be 200-500 m. Thus, the MLT is suggested to be in the uppermost MLR, and indicates slight shallowing from Early to Late Paleocene times. A palaeo-water depth in excess of 400 m (MLT) is proposed for the early-Middle Eocene, followed by a shallowing in the Late Eocene to Early Oligocene, as supported by the SRT (Fig. 7a). Ranges of 200750 m indicate the uncertainties associated with these estimates. From the well completion log a hiatus is indicated in the middle-Late Miocene, probably related to subaerial erosion. The deepening in the Late Miocene as suggested by the SRT post-dated the hiatus, or took place in the youngest time interval included in the hiatus (Fig. 7a).

Well 15/5-3 Well 15/5-3 is located in the southern Viking Graben (Figs. 1 and 7b). The well report contains micropalaeontological data from the entire post-rift interval. For the Early Cretaceous and Late Miocene only sparse data exist. Quantification of hiatuses from this well was not performed, and the well is not linked to a structural restored transect. The MLR is typically in the order of 200-500 m (Fig. 7b). Exceptions are: (1) the MLR for the Eocene is 500-750 m, supplemented with possible maximum palaeo-water depth up to 1000 m for the Early Eocene samples; (2) the MLR for the Early Paleocene, Turonian and latest Albian samples are 200-750 m; (3) less uncertainty is affiliated with samples for the Campanian, which have a MLR of 500-600 m. The suggested post-rift palaeo-water depth development for the well 15/5-3 relies entirely on the MLT (Fig. 7b).

Summary In general, the integrated palaeo-water depth curves from the investigated Viking Graben wells are consistent (Fig. 10a). However, it is noted that for the Early Cretaceous discrepancies exist. The in-

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tegrated palaeo-water depth curves for wells 24/9-1 (Fig. 7), 29/6-1 (Fig. 9) and 30/4-1 (Fig. 7), suggest shallowing from early Ryazanian to latest Albian times. This is supported by the general results from the structural restoration of the four transects, which indicate shallowing of the palaeo-water depth in the Viking Graben during the Early Cretaceous (Kjennerud et al., 2001). It is noted that for the deepest part of the basin (the interior graben axis) at the earliest Ryazanian level, the shallowing was considerable. In contrast, the integrated palaeo-water depth curves of wells 34/11-1 (Fig. 6), 30/10-6 (Fig. 8) and 15/5-3 (Fig. 7), which rely mainly on the MLT, are characterised by several shallowing/deepening events, but generally suggest slight deepening from early Ryazanian to latest Albian time. The differences in palaeo-water depth trends for the Early Cretaceous can be partly explained by local variations in sediment infill history, perhaps combined with complex subsidence pattern (Gabrielsen et al., 2001). From the earliest Cenomanian deepening took place as seen in the entire Viking Graben area (Fig. 10a). It is noted, however, that the deepest basin conditions occurred at a different time. For the wells that contained good micropalaeontological data, deepest basin conditions are suggested for the mid-late Campanian, corresponding to the seismic unit K-5. Thus, the structural restoration method could not detect the deepest palaeo-water depth (Kjennerud et al., 2001). From early-late Campanian to earliest Paleocene shallowing took place, followed by rapid deepening in the Early to Late Paleocene. A palaeo-water depth in excess of 500 m commonly exist in the Viking Graben during parts of Paleocene and Eocene times (Fig. 10a). It is noted that shallowing started in the Late Paleocene in 24/9-1 (Figs. 1 and 10) and 29/6-1 (Figs. 1 and 11). This may be attributed to closeness to CSS-1 depocentres and most likely as a result of a greater sedimentation rate in the southern Viking Graben (Jordt et al., 1995). The other Viking Graben wells are located in a more distal position (Fig. 1), and did not experience shallowing before Late Eocene time (Fig. 10a). Overall shallowing followed from Late Eocene until Miocene time. Zero palaeo-water depth probably existed in parts of the basin in the Middle Miocene, as inferred from the observed hiatuses, but seismic reflection data of higher resolution combined with more detailed sedimentological and high-resolution biostratigraphic investigations is needed to fully understand the water depth development during the Miocene. Our data suggest that deepening occurred in the Late Miocene to Early Pliocene, followed by shallower basin conditions during Late Pliocene time (Fig. 10a).

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Fig. 11. (a-d) The post-rift palaeo-water depth development as seen along the four transects.

Cretaceous-Tertiary palaeo-bathymetry in the northern North Sea

Fig. 11 (continued).

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340 The Lomre and Uer Terraces

Well 3 5 / 1 1 - 3 Well 35/11-3 is located on the crest of a major fault block at the Lomre Terrace (Figs. 1, 3 and 8a). Sparse micropalaeontological data exist for the interval earliest Ryazanian to the latest Coniacian, and no micropalaeontological data are available from the Miocene to Present interval. Cretaceous. The seismic units K-1 and K-2 were impossible to differentiate on seismic data (Fig. 8a) at the well position. The well contained about 100 m of preserved sediments of Early Cretaceous age. In the structural restoration it is assumed that zero palaeo-water depth existed in this position, at the base of the seismic units K-l, K-2, K-3, K-4 and K-5, which could be supported by the recognition of hiatuses corresponding to Base K-2, Base K-3 and Base K-4. On the other hand, zero palaeo-water depth may not be realistic for the entire interval. Considerable palaeo-water depth (MLR of 100-500 m) may have existed in parts of the Early Cretaceous, as suggested by the micropalaeontological data. Therefore, the palaeo-water depth trend from the earliest Ryazanian to latest Coniacian strongly relies on: (1) the estimates based on micropalaeontological data; (2) assumptions that the hiatuses were caused by subaerial erosion, and (3) that the sub-aerial exposure took place in the upper time interval of the hiatuses. The deepening as obtained by the structural restoration for the intra-Cenomanian to the latest Turonian is, by the reasoning in (2) above, suggested to have taken place in the Santonian, as this estimate is constrained by the oldest sediment present above the unconformity (Fig. 8a). From early Santonian to Early Miocene times, the integrated curve relies on the MLT (Fig. 8a). The MLT is mainly supported by the SRT, it is assumed that the zero palaeo-water depth suggested by SRT at the Cretaceous/Tertiary boundary implies that the basin had no relief, rather than occurrence of zero palaeo-water depth (Kjennerud et al., 2001). Tertiary. In general, structural restoration and micropalaeontological analysis offer similar trends for the Tertiary. However, it is noted that the shallowing as indicated by the MLT for the Eocene is not supported by SRT, which indicates prevailing of deeper palaeo-water depth during the Eocene (Fig. 8a). A hiatus, possibly caused by sub-aerial erosion, was tracked for the middle-Late Miocene time interval (Fig. 8a). No other palaeo-water depth estimates based on micropalaeontological data are available for the Neogene interval. From the Early Miocene to

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Present, the integrated curve relies on the SRT. It is assumed that the indicated deepening in the Late Miocene post-dates the hiatus. The MLR of the estimates obtained from 35/11-3 indicate that the palaeo-water depth was commonly above 200 m (Fig. 8a). The MLR combined with the zero palaeo-water depth associated with the hiatuses indicate that fluctuations in palaeo-water depth could have been more detailed than displayed in the MLT.

Well 35/12-1 Well 35/12-1 is located at the Uer Terrace (Figs. 1, 3 and 8b). Micropalaeontological data are available for the interval middle Albian to latest Paleocene, and from the Pliocene to Quaternary. Cretaceous. In well 35/12-1, a hiatus caused by sub-aerial erosion is found for the earliest Ryazanian to the late Barremian (Fig. 8b). For the latest Albian hiatus it is not possible to conclude whether it was caused by erosion or non-deposition. The uncertainty in the estimates based on micropalaeontological data is in most cases reported to be of 200-500 m. Occasionally, better constraints seem to be evident, as seen for the Turonian to Coniacian MLR (Fig. 8b). For the early Ryazanian to intra-Aptian the SRT is adopted in the integrated curve, and it is assumed that the palaeo-water depth of around 300 m, as suggested by micropalaeontological data, for the early-late A1bian is reasonable. Hence a deepening occurred in late Aptian to middle Albian times. The shallowing as seen from the MLT for the late Albian is not supported by the SRT. The SRT displayed deeper palaeo-water depth for the Turonian to Coniacian interval. However, the palaeo-water depth for this interval is constrained by strong micropalaeontological evidence suggesting it to be shallower than 200 rn (Fig. 8b). Hence no deepening, as suggested by the SRT, took place during the Cenomanian to Turonian. On the other hand, the SRT had slightly shallower palaeo-water depth estimates than the MLT for the middle Campanian to latest Paleocene (Fig. 8b). The differences between the MLT and the SRT could be caused by differences in topography when the transect position is compared to the true well position (Fig. 1) (Kjennerud et al., 2001). Therefore, it is assumed that the MLT is the most reliable palaeo-water depth indicator for the Cretaceous interval (Fig. 8b). Tertiary. The MLT suggests deepening in the Early Paleocene to a palaeo-water depth in excess of 300 m, followed by a slight shallowing in the

341

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Late Paleocene. The palaeo-water depth was close to 200 m in the latest Paleocene (Fig. 8b). A hiatus, caused by sub-aerial erosion, occurred in the earliest Eocene to Early Pliocene. Therefore, the Eocene-Present trend relies on the SRT (Fig. 8b).

Summary The palaeo-water depth trends in wells 35/11-3 and 35/12-1 are in good agreement (Figs. 8 and 10b). Deeper basin conditions, associated with a N S-trending depression in the Aptian-Cenomanian are suggested for well 35/12-1. The differences in absolute palaeo-water depth for the Tertiary interval may be explained by the distance from the pre-existing shoreline.

Horda Platform/Stord Basin/Asta Graben Well 3 1 / 6 - 3 Well 31/6-3 is located in the northern part of the Stord Basin of the Horda Platform (Figs. 1, 3 and 9a). Micropalaeontological data exist for most of the preserved strata in the well. The most important observations made in 31/6-3 are the recognition of several hiatuses.

Despite the distance from the well to the transect the SRT is taken as valid. However, the absolute palaeo-water depth offered by the structural restoration is affiliated with uncertainties, and in such cases, as seen for the earliest Cretaceous and Paleocene (Fig. 9a), the available palaeo-water depth estimates from micropalaeontological data are given priority. Additionally, the integrated curve relies on the assumption that sub-aerial exposure took place at the end of each encountered hiatus. The hiatuses were as follows: (1) late Barremian-late Albian, caused by sub-aerial erosion; (2) late Aptian-late Albian, caused by sub-aerial erosion; (3) late Albian-early Cenomanian, caused by sub-aerial erosion (?); (4) latest Turonian-latest Campanian, non-deposition (?); (5) late Maastrichtian-earliest Paleocene, caused by subaerial erosion; (6) earliest Eocene-latest Pliocene, caused by sub-aerial erosion. It is also noted that the SRT does not support zero palaeo-water depth associated with the encountered Early Cretaceous hiatuses (Fig. 9a). For the Eocene to the Quaternary, the integrated palaeo-water depth curve is identical to the SRT (Fig. 9a), as no other data are available.

Well 17/3-1 Cretaceous. Micropalaeontological data indicate that the palaeo-water depth for the Valanginian to early Barremian was most likely less than 200 m (Fig. 9a). In contrast, the SRT offered deeper palaeo-water depth in the Early Cretaceous. It is likely that the MLT is reasonable since transect 2 is located in a position that in the Early Cretaceous experienced deeper water depth. However, the MLR of 100-500 m for the early Barremian shows that deeper basin conditions could have existed, as suggested by the SRT. For the late Albian the MLR of 200-500 m is suggested, and the most likely palaeo-water depth was probably around 300 m (Fig. 9a). The palaeo-water depth estimates of the Cenomanian to latest Turonian are well constrained and have an uncertainty range of 100 m displayed by the MLR. Slight shallowing is suggested by the MLT, from 300 m in the early Cenomanian to below 200 m in the latest Turonian (Fig. 9a). One estimate (MLR 100-300 m) was done for the Maastrichtian, which suggests a palaeo-water depth of 150 m.

Tertiary. Palaeo-water depth estimates based on micropalaeontological data are available for the Paleocene and the Quaternary intervals (Fig. 9a). The most likely palaeo-water depth in the Paleocene is suggested to be around 600 m, which is in the uppermost part of the MLR. For the Quaternary, the palaeo-water depth is determined to be around 300 m.

Well 17/3-1 is located in the Asta Graben (Figs. l, 3 and 9b). Micropalaeontological data were available for the Maastrichtian and the Early Cretaceous interval (Fig. 9b).

Cretaceous. Two hiatuses are recorded, one of late Valanginian to early Hauterivian age, the other of late Barremian age (Fig. 9b). It is not possible with any certainty to decide the nature of these hiatuses. Another hiatus, probably related to non-deposition, is found for the earliest Cenomanian to middle Turonian (Fig. 9b). The most likely palaeo-water depth of the earliest Cretaceous is suggested to have been around 300 m. The uncertainties related to the palaeo-water depth estimates (MLR) are commonly 200-500 m. The MLT indicates slight shallowing from the earliest Ryazanian to a palaeo-water depth of around 400 m in latest Albian (Fig. 9b). The difference in the estimated palaeo-water depth for the earliest Ryazanian, when the results from the two methods are compared, can be explained by varying the topography between the true well position and the corresponding position at transect 4 (Fig. 1). This may also be evident for the palaeo-water depth estimates for the Maastrichtian. The MLT indicates that slight shallowing took place during the Maastrichtian. The palaeo-water depth was probably close to 300 m at the Cretaceous/Tertiary boundary (Fig. 9b).

342 The MLT is adapted for the Early Cretaceous and the Maastrichtian. For the rest of the interval, the results rely on the SRT (Fig. 9b).

Summary Some striking differences are noted when the palaeo-water depth trends for 31/6-3 and 17/3-1 (Figs. 9 and 10b) are compared. In 31/6-3, several shallowing/deepening events for the Early Cretaceous are suggested by indications of repeated sub-aerial exposures. Hiatuses are also tracked for the Early Cretaceous in 17/3-1 (Fig. 9b), but in this well it is not possible with any certainty to decide whether the hiatuses were caused by sub-aerial exposure or not. Micropalaeontological data suggest deepening to be the most likely scenario. In contrast, the result of the structural restoration of transect 2 suggests a shallowing.

Discussion It should be emphasised that the palaeo-water depth as determined in the present study ideally equals relative sea level minus accumulated sediment (Jervey, 1988; Posamentier et al., 1988; Posamentier and Vail, 1988). A seismic-stratigraphic framework provides control on the accumulated sediment thickness, but it has not been possible to differentiate between the tectonic and eustatic influence on the available accommodation space with any certainty. It is generally accepted that eustatic sea-level changes influence deposition. Such influences are important in the study area and indeed have been applied in the analysis of the North Sea (Vail and Todd, 1981). However, the eustasy-driven seismic sequence-stratigraphic model of Vail et al. (1977), and attempts to build global eustatic curves (e.g. Haq et al., 1987, 1988), have been debated for both general (for an overview see Nystuen, 1998) and local (North Sea) reasons (e.g. Cloetingh, 1991). Despite uncertainties, the Haq et al. (1987, 1988) eustatic curve has been extensively applied in basin modelling and used as a standard reference curve for comparison of regional and local relative sealevel changes (Nystuen, 1998). We suggest that the Haq et al. (1987, 1988) long-term eustatic curve is appropriate to be compared to the recognised palaeowater depth trends in the present study. The Haq et al. (1987, 1988) long-term curve displays the best estimate of maximum eustatic sea level from the Triassic to Present (reference level). For the Cretaceous-Cainozoic interval it is noted that the Haq et al. (1987, 1988) curve varies between 50 and 250 m, and that the eustatic change had low amplitude compared to the suggested deepening/shallowing events as suggested in the present study. A fall in eustatic sea level occurred from the ear-

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liest Ryazanian to Valanginian followed by a rise, which lasted until the Turonian (Haq et al., 1987, 1988). High eustatic sea level was encountered for the Turonian to the mid-Campanian interval, followed by fall until the Late Paleocene. The suggested deepening/shallowing trends from our study of the Cretaceous interval are in agreement with the Haq et al. (1987, 1988) long-term curve. We therefore claim that the observed curve includes an eustatic signal that is amplified by tectono-thermal effects. It is noted, however, that the long-term curve is not in concert with the relative shallowing observed by us for the Early Cretaceous, implying that if this event is real it had a purely tectono-thermal origin. A minor rise in eustatic sea level took place in the latest Paleocene, followed by a general fall until the Present (see also Haq et al., 1987, 1988). The prominent deepening events, as indicated for the Early to Late Paleocene and for the Late Miocene to Early Pliocene from the present data, can therefore not be explained by the long-term eustatic curve. By contrast, the shallowing, as suggested in the present study, for the Eocene to Late Miocene, corresponds well with the fall in eustatic sea level for this interval, whereas the Early Oligocene shallowing has considerably higher amplitude than predicted by the Haq et al. (1987, 1988) long-term curve.

Summary and conclusions By careful integration of seismic-stratigraphic observations of palaeo-water depth estimates from structural restoration and micropalaeontological data, we have documented changes in accommodation space and sediment supply/distribution throughout the Cretaceous-Cainozoic post-rift interval on a regional scale. Because it is not possible to give exact figures of the palaeo-water depth, we have focused on the determination of most likely depth intervals, combined with the identification of shallowing and deepening trends. The inferred trends from the investigated wells were generally in good agreement on a regional scale, and also when the tectonic position within the basin was taken into account. Palaeo-water depth development as seen along the four regional transects in Fig. 11, displays the following general trends. (1) General shallowing during the Early Cretaceous. (2) Deepening from the early Cenomanian to the mid-Campanian. (3) Shallowing from the mid-Campanian to the latest Maastrichtian. (4) Deepening in the early-Late Paleocene. (5) Shallowing from the Late Eocene to Late Miocene.

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Fig. 12. Northern North Sea palaeo-water depth maps for the late Ryazanian (top left), the earliest Maastrichtian (top right), the earliest Eocene (botton left) and the Early Oligocene (bottom right).

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(6) Deepening from the Late Miocene to Early Pliocene. (7) Shallowing during the Pliocene. These effects are, however, modified according to tectonic position within the basin. The Late Jurassic-Early Cretaceous rifting phase exerted a profound influence on post-rift basin morphology and depositional pattern in the northern North Sea (see also Gabrielsen et al., 2000). As a result, the Early Cretaceous northern North Sea Basin morphology encompassed a complex mosaic of several islands, shoals and bays (Fig. 12, top left). The Viking Graben, Sogn Graben and the Stord Basin probably experienced palaeo-water depths of more than 500 m, and acted as depocentres in the Early Cretaceous. Due to sediment infill, the submarine relief was reduced and a general shallowing of the palaeo-water depth is seen. Still, the presented palaeo-water depth estimates indicate that several regressive/transgressive events took place (Figs. 4-9). During the early Cenomanian to mid-Campanian the inherited rift topography gradually became filled in and levelled out (Fig. 11), at the same time as deepening occurred. The pre-existing source areas of the local islands and shoals and the shelf area to the east and west became drowned and the basin configuration changed to a wide uniform basin with a smoother bathymetry (Fig. 12, top fight). Still, the deepest basin condition existed along the Viking Graben axis and its northwards elongation. From mid-Campanian to earliest Paleocene time, shallowing took place (Fig. 11). The structural restoration suggests that the basin relief was reduced to zero at the Cretaceous/Tertiary transition (Kjennerud et al., 2001), but a palaeo-water depth of around 200 m most probably existed as indicated by micropalaeontological data. All wells investigated point to deepening during the Early to Late Paleocene. The basin axis shifted eastward relative to the interior Viking Graben axis (Fig. 12, bottom left). To the east and west, on the Norwegian Mainland and the East Shetland Platform, basin flank uplift took place, and prograding wedges that built into the basin dominated the depositional style. The depocentres at the transitional shelf provided sediment supply for mass flows and submarine fans that entered the deepest part of the basin during the Paleocene/Eocene (McGovney and Radovich, 1985). From the latest Eocene the basin became shallower and narrower (Fig. 11), and in the Early Oligocene a shallow threshold separated lows of deeper palaeowater depth to the south and north (Fig. 12, bottom fight). To the east, major parts of the Horda Platform and the Stord Basin were sub-aerially exposed.

Sub-aerial exposure commonly occurred in the Oligocene/Miocene. However, the available information from the wells does not point to these areas being exposed simultaneously. The results of the micropalaeontological interpretation and the structural restoration both suggest relative deep basin condition in the Early Pliocene (Fig. 11). Consequently, deepening must have taken place, probably in the Late Miocene to the Early Pliocene. In conclusion, therefore, it seems reasonable that some of the deepening/shallowing events cannot be explained by the eustatic sea-level curve, and therefore need to be accounted for by tectono-thermal events. However, further investigation and subsidence analysis is needed to test the validity of, and mechanisms behind, the post-rift tectonic events in the northern North Sea.

Acknowledgements The paper is publication no. 1 in a series from the project "Tectonic Impact on Sedimentary Processes in the Post-Rift Phase; Improved Models". The project was supported by the Research Council of Norway through grant no. 32842/211. The authors would like to thank the companies participating in the project (Amoco Norge Norway Oil Company, den Norske Stats Oljeselskap (Statoil), Mobil Exploration Norway Inc., Norsk Agip A/S, Norsk Hydro ASA, Phillips Petroleum Company Norway, Saga Petroleum ASA), and for permission to publish the results.We would also like to thank the project members for inspiring discussions, and particularly head of the Project Managing Board, Dr. Gunn Mangerud of Norsk Hydro ASA for strong support in the last stage of the project work. Specials thank to Felix M. Gradstein for discussions and constructive comments concerning the palaeo-water depth, and to Tone Gjelsvik (UiB) and Befit Fossum (IKU) for technical assistance. Reviews of K. SOgaard and G. Mangerud are greatly appreciated.

References Beicip Franlab, 1996. LOCACE 2.2 User's guide. Berggren, W.A. and Gradstein, EM., 1981. Agglutinated benthonic foraminiferal assemblages in the Palaeogene of the Central North Sea; their biostratigraphic and depositional environmental significance. In: L.V. Illing and G.D. Hobson (Editors), Petroleum Geology of the Continental Shelf of North-West Europe: Proceedings of the Second Conference. Institute of Petroleum, London, pp. 282-288. Bertram, G.T. and Milton, N.J., 1989. Reconstructing basin evolution from sedimentary thickness; the importance of palaeobathymetric control, with reference to the North Sea. Basin Res., 1: 247-257. Cloetingh, S., 1991. Tectonics and sea-level changes: a controversy? In: D.W. Mtiller, J.A. McKenzie and H. Weissert (Editors), Controversies in Modern Geology - - Evolution of Geological The-

Cretaceous-Tertiary palaeo-bathymetry in the northern North Sea ories in Sedimentology, Earth History and Tectonics. Academic Press, London, pp. 249-277. Gabrielsen, R.H., Kyrkjebo, R., Faleide, J.I., Fjeldskaar, W. and Kjennerud, T., 2001. The Cretaceous post-rift development in the northern North Sea. Pet. Geosci., in press. Gibson, T.G., 1988. Assemblages characteristics of modern benthic foraminifera and application to environmental interpretation of Cenozoic deposits of Eastern North America. Rev. Paleobiol., 62: 255-265. Gillmore, G.K., Kjennerud, T. and Kyrkjebr R., 2001. The reconstruction and analysis of palaeowater depths: a new approach and test of micropalaeontological approaches in the post-rift (Cretaceous to Quaternary) interval of the northern North Sea. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 365-381 (this volume). Gradstein, F.M. and B~ickstr6m, S.A., 1996. Cainozoic biostratigraphy and palaeobathymetry, northern North Sea and Haltenbanken. Nor. Geol. Tidsskr., 76: 3-32. Gradstein, F.M. and Ogg, J., 1996. A Phanerozoic time scale. International Union of Geological Science (IUGS), Ottawa, ON, Episodes 19(1-2), pp. 3-6. Gradstein, F.M., Kaminski, M.A., Berggren, W.A., Kristiansen, I.L. and D'Iorio, M., 1994. Cainozoic Biostratigraphy of the North Sea and Labrador Sea. Micropaleontology, 40 (Suppl., 1994), 152 pp. Haq, B.U., Hardenbol, J., Vail, RR., 1987. Chronology of fluctuating sea-levels since the Triassic. Science, 235:1156-1167. Haq, B.U., Hardenbol, J. and Vail, RR., 1988. Mesozoic and Cenozoic chronostratigraphy and cycles of sea-level changes. In: C.K. Wilgus, B.S. Hastings, C.G.St.C. Kendall, H.W. Posamentier, C.A. Ross and J.C. Van Wagoner (Editors), Sea-Level Changes: An Integrated Approach. Soc. Econ. Paleontol. Mineral., Spec. Publ., 42: 71-108. Jervey, M.T., 1988. Quantitative geological modelling of siliciclastic rock sequences and their seismic expression. In: C.K. Wilgus, B.S. Hastings, C.G.St.C. Kendall, H.W. Posamentier, C.A. Ross and J.C. Van Wagoner (Editors), Sea-Level Changes: An Integrated Approach. Soc. Econ. Paleontol. Mineral., Spec. PUN., 42: 47-69. Jones, G.D., 1988. A paleoecological model of Late Paleocene "flysch-type" agglutinated foraminifera using the paleoslope transect approach, Viking Graben, North Sea. In: Second Workshop on Agglutinated Foraminifera, Vienna Proceedings. Abh. Geol. Bundesanst., 41: 155-227. Jordt, H., Faleide, J.I., Bjorlykke, K. and Ibrahim, M.T., 1995. Cenozoic sequence stratigraphy of the central and northern North Sea basin: tectonic development, sediment distribution and provenance areas. Mar. Pet. Geol., 12: 845-880. Joy, A.M., 1993. Comments on the pattern of post-rift subsidence in the Central and Northern North Sea Basin. In: G.D. Williams

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and A. Dobbs (Editors), Tectonics and Seismic Sequence Stratigraphy. Geol. Soc. London, Spec. Publ., 71: 123-140. Kjennerud, T., Faleide, J.I., Gabrielsen, R.H., Gillmore, G.K., Kyrkjeb0, R., Lippard, S.J. and L0seth, H., 2001. Structural Restoration of Cretaceous-Cenozoic (post-rift) Palaeobathymetry in the northern North Sea. In: O.J. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway Palaeozoic to Recent. Norwegian Petroleum Society (NPF), Special Publication 10. Elsevier, Amsterdam, pp. 347-364 (this volume). McGovney, J.E. and Radovich, B.J., 1985. Seismic stratigraphy and facies of the Frigg fan complex. In: O.R. Berg and D.G. Woolverton (Editors), Seismic Stratigraphy, II. An Integrated Approach to Hydrocarbon Exploration. Am. Assoc. Pet. Geol., Mem., 39: 139-154. Nystuen, J.E, 1998. History and development of sequence stratigraphy. In: EM. Gradstein, K.O. Sandvik and N.J. Milton (Editors), Sequence Stratigraphy Concepts and Applications. Norwegian Petroleum Society (NPF), Stavanger, Special Publication 8, pp. 31-116. Posamentier, H.W. and Vail, RR., 1988. Eustatic control on clastic deposition II sequence and system tract models. In: C.K. Wilgus, B.S. Hastings, C.G.St.C. Kendall, H.W. Posamentier, C.A. Ross and J.C. Van Wagoner (Editors), Sea-Level Changes: An Integrated Approach. Soc. Econ. Paleontol. Mineral., Spec. Publ., 42: 125-154. Posamentier, H.W., Jervey, M.T. and Vail, RR., 1988. Eustatic control on clastic deposition I sequence and system tract models. In: C.K. Wilgus, B.S. Hastings, C.G.St.C. Kendall, H.W. Posamentier, C.A. Ross and J.C. Van Wagoner (Editors), Sea-Level Changes: An Integrated Approach. Soc. Econ. Paleontol. Mineral., Spec. Publ., 42: 109-124. Roberts, A.M., Yielding, G. and Badley, M.E., 1993. Tectonics and bathymetric controls on stratigraphic sequences within evolving half-graben. In: G.D. Williams and A. Dobbs (Editors), Tectonics and Seismic Sequence Stratigraphy. Geol. Soc. London, Spec. PUN., 71: 87-121. Sloan, B.J., 1995. Eocene Sequence Stratigraphy of the North Sea Basin. Dissertation for the Degree of Doctor Philosophy, University of Texas at Austin, 214 pp. Vail, RR. and Todd, R.G., 1981. Northern North Sea Jurassic unconformities, chronostratigraphy and sea-level changes from seismic stratigraphy. In: L.V. Illing and G.D. Hobson (Editors), Petroleum Geology of the Continental Shelf of North-West Europe: Proceedings of the Second Conference. Institute of Petroleum, London, pp. 216-235. Vail, RR., Mitchum, R.M., Jr., Todd, R.G., Widmier, J.M., Thompson, S., III, Dangree, J.B., Bubb, J.N. and Hatlelid, W.G., 1977. Seismic stratigraphy and global changes of sea-level. In: C. Payton (Editor), Seismic Stratigraphy Applications to Hydrocarbon Exploration. Am. Assoc. Pet. Geol., Mere., 26: 49-212.

Geological Institute, University of Bergen, N-5007 Bergen, Norway; E-mail: [email protected] SINTEF Petroleum Research, N-7034 Trondheim, Norway University College Northampton, School of Environmental Science, Northampton NN2 7AL, UK Department of Geology, University of Oslo, N-0316 Oslo, Norway Geological Institute, University of Bergen, N-5007 Bergen, Norway

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Structural restoration of Cretaceous-Cenozoic (post-rift) palaeobathymetry in the northern North Sea T. Kjennerud, J.I. Faleide, R.H. Gabrielsen, G.K. Gillmore, R. Kyrkjebe, S.J. Lippard and H. Leseth

Palaeobathymetric variations and palaeobasin shape are important for recognition of tectonic phases and distribution of sediment. We present an approach for restoring palaeobathymetry from interpreted, depth-converted seismic sections. The method involves section balancing/restoration techniques for extensional regimes and seismic sequence stratigraphy. The method is based on regional profiles that cover the entire basin using fixed water-depth points such as coal layers, shallow marine sand and subaerial unconformities, which are used to calibrate areas along the profile. Geometric information in the under- and overlying seismic sequences is used to shape the palaeobasin. The palaeowater-depths and basin gradients of the Cretaceous and Tertiary post-rift interval in the northern North Sea have been restored. The results show a highly segmented basin with high gradients and locally large water-depths in the earliest Cretaceous. A shallowing and a lowering of the basin gradients occurred in the early Cretaceous. A deepening and widening of the basin was initiated in the middle Cretaceous. During the late Cretaceous the water-depth and the basin gradients decreased. A deepening occurred in the Palaeogene resulting in high gradients on the flanks of the basin. The basin became shallower and the relief became smaller in the early Neogene. The basin deepened in the late Neogene and then became shallower to the present day. The palaeobathymetric trends correlate well with results obtained from micropalaeontological analysis of palaeowater-depth in the same area.

Introduction

A good understanding of palaeobathymetric variations is important for the recognition of tectonic phases in sedimentary basins (Bertram and Milton, 1989). Palaeobathymetry is also important for interpreting the distribution of deep marine sands and source rocks within a geological time frame. Palaeobathymetry is difficult to reconstruct in deep marine settings, because palaeontological and sedimentological parameters cannot distinguish between significant water-depths (e.g. 500 m or 1000 m). In most rift basins the sedimentological and palaeontological approaches are based on the analysis of well data, which commonly originate from flank positions or highs within the basin. Hence, information on the palaeobathymetry of the deeper parts of the basin, which is very important in improving the understanding of the tectonic impact during the post-rift stage, may be scarce. The available data are therefore commonly not sufficient to accurately determine the palaeobasin shape. As part of a larger study (the project Tectonic Impact on Sedimentary Processes in the Post-Rift Phase Improved Models), information about palaeobathymetry along four key transects for the post-rift stage in the northern North Sea was required. In

previous work, palaeobathymetry has been obtained mainly by the analysis of wells. Micropalaeontological work, with reference to the North Sea, has focused on Cenozoic palaeobathymetry (Berggren and Gradstein, 1981; Gradstein et al., 1994; Gradstein and B~ickstr6m, 1996). Jones (1988) and Sloan (1995) focused on Palaeogene palaeobathymetry, while Gradstein et al. (1999) presented Cretaceous water-depths. In addition Bertram and Milton (1989) and Joy (1993) estimated palaeobathymetry by key horizons in wells formed at known water-depths combined with observed correlative sedimentary thickness. The aim of this paper is to describe a structural restoration method, which restores palaeobathymetry and basin gradients, based on geometrical constraints (seismic stratigraphical relationships) and points of zero water-depth, and to present the results for four restored transects. The restorations have been performed on the Cretaceous and Tertiary post-rift interval in the northern North Sea. The present paper is part of a workflow, where first a palaeobathymetric data set was constructed along four key transects (this paper). Micropalaeontological analysis was then performed for 12 wells with respect to palaeobathymetry (Gillmore et al., 2001). An integration of the results was then performed by Kyrkjeb~ et al. (2001) in the well positions. Finally, the palaeobathymetric data

Sedimentary Environments Offshore Norway - Palaeozoic to Recent edited by O.J. Martinsen and T. Dreyer. NPF Special Publication 10,

pp. 347-364, Published by Elsevier Science B.V., Amsterdam. 9 Norwegian Petroleum Society (NPF), 2001.

348 were integrated with the regional seismic database and isopach maps to construct palaeobathymetric maps, two of which are presented in the present paper. Methods

The basis for the restoration of profiles in extensional regimes was established in the 1980s (e.g. Verral, 1981; Gibbs, 1983, 1984). The methods applied in the early studies were based on the assumption that vertical simple shear dominates the deformation in the hangingwall of listric faults, and these papers were to some degree focused on the prediction of fault shape based on geometries in the hangingwall. White et al. (1986) modified this method in order to accommodate inclined shear. Due to lack of data on palaeowater-depth, several early workers were forced to assume the basin had negligible relief, although the authors knew that this was not entirely correct (see discussion in Gabrielsen et al., 1986). In previous work, palaeobathymetry has been restored on geological profiles using the flexural backstripping approach of Kusznir et al. (1991). The flexural backstripping method restores the basin geometry in a stepwise manner, taking the thermal subsidence, flexural unloading and the decompaction into account. Thermal subsidence is obtained from the model of McKenzie (1978). This implies that the ~3 factor must be known for each time step. Hence these three factors constrain the palaeobathymetry. Roberts et al. (1998) particularly emphasise the use of known palaeobathymetric constraints in the flexural backstripping method. In practice this means areas of zero or near zero water-depth, which is also emphasised in the present work. Based on simple shear as the deformation mechanism and geometrical relationships and observations relevant to relief, a method has been developed in the present study, using the section restoration software LOCACE (Beicip Franlab, 1999). The target was to generate a separate palaeobathymetric data set, other than that from the micropalaeontology (Gillmore et al., 2001). Two separate palaeobathymetric data sets allow for quality control on this parameter. The method described below will provide palaeobathymetry along selected profiles, which enables us to get estimates between wells. Vertical simple shear is considered to be a good approximation of the hangingwall mode of deformation, since the study is carried out in the post-rift interval, which is dominated by vertical displacements. This is consistent with Airy isostasy. Hence, flexure is not considered in the restorations at present. The idea behind the present restoration method is that it is possible to quantify palaeobathymetry/topo-

T. Kjennerud et al. graphy by using geometrical information extracted from seismic sequences. This includes geometrical factors below and above the surface at which the restoration is performed. This information can then be used to restore the basin at the actual stratigraphic level. In addition, inferred areas of zero or near-zero water-depth must be incorporated in the shaping of the basin. Palaeobathymetric interpretation adapted from micropalaeontology is left out in order to keep these two approaches separate. This allows us to compare the two methods and their respective results (Kyrkjebo et al., 2001). The restorations in the present work were performed starting with the present-day geometry on depth-converted interpreted seismic sections. For each restoration, the restored section was decompacted, using the decompaction curves of Sclater and Christie (1980), which are regarded as adequate, since their work was based on North Sea data. It is important that the restoration is performed on profiles covering the entire basin. Every point across the profiles where the palaeobathymetry is well constrained (e.g. zero or near-zero water-depth) should be utilised for calibration of the profile. Zero or near-zero water-depth points include coal horizons, shallow marine sands and erosional hiatuses. Two general approaches, which may be combined, have been recognised for the post-rift stage of the northern North Sea:

Deep marine infill Sedimentary units that progressively onlap the basin flanks and show maximum thickness in the basin centre, contribute to the infilling of relief created at an earlier stage. This is characteristic for most of the Cretaceous sequence in the northern North Sea, where these sediments are mainly fine-grained mudstones and siltstones, deposited as pelagics and muddy turbidites. Thus, the maximum sediment thickness will also record the maximum bathymetry at a given time stage. This can be seen in many recent examples away from large sources of sediment input (e.g. the present-day sedimentation in Lake Malawi: Scholz and Finney, 1994). In order to restore the relief in this type of setting, the thickness variations of the infilling unit are used to create a relative relief (Fig. 1). This is done by flattening the top of the unit, or if known, restoring it according to the actual relief at that stage. The uppermost unit is then decompacted, and the older units are restored accordingly. By the use of known calibration points, as determined by biostratigraphical, sedimentological and/or seismic sequence stratigraphical methods, the palaeobasin relief is constructed. It is of importance

Structural restoration of Cretaceous-Cenozoic (post-rift) palaeobathymetry in the northern North Sea

349

Fig. 1. Restoration of deep marine infill sequences.

to determine as many calibration points with known water-depth as possible. In this study the most common fixed points are those that fulfill the criterion of zero water-depth. It is crucial then to adjust only the rotating part of the slopes. Finally, the section is decompacted. It can be difficult to distinguish between continental onlap, coastal onlap and pelagic onlap. This can, however, be further constrained by firmly determined fixed water-depth points. If these are not constrained near the onlap, the restored section will represent relief rather than absolute palaeobathymetry.

Prograding sequences Prograding sequences are commonly identified in reflection seismic data as sets of prograding clinoforms in the proximal part of the system, and parallel

Fig. 2. Restoration of prograding sequences.

reflections in the more distal part of the system (Vail et al., 1977). The approach for restoring the palaeobathymetry of prograding sequences is based on the geometry of the clinoforms, and the thickness distribution within the prograding unit (Fig. 2). The section is decompacted and the topsets are put at sea level. The palaeobathymetry is determined by calculating the inverse thickness of the prograding unit, where the thick prograding part is assumed to represent the shallow part of the section, and the thinner part, characterised by parallel reflections, is interpreted to represent the deep marine part of the section. The older units are restored accordingly and

350 decompacted. The reconstruction provides the shape of the basin in the part where the clinoforms occur. Ideally, the geometry of the infilling sediments should be used to quantify the relief in the distal part. The main uncertainty for this restoration is associated with the accuracy in determining lithological units present, since this will influence the decompaction. For each time step, faults and eroded units are restored. Finally results from micropalaeontology can be integrated with the structural palaeobathymetric restoration. The micropalaeontological estimates will have to be treated as fixed points. The restoration then incorporates the results from micropalaeontology, while still representing the relief described by the seismic sequences. This integration should be performed as a last step, so that it does not influence the pure geometrical approach in the present study. The results presented on sections in the present paper are purely based on the structural restoration. The integration was performed in the construction of the palaeobathymetric maps. KyrkjebO et al. (2001) present the integration of results from the structural restoration with micropalaeontology performed on key wells.

Map construction Based on the same principles as those stated above, palaeobathymetric maps may be constructed by using information from seismic isopach maps for each unit. Ideally the seismic isopach maps should be decompacted. Firstly, the palaeobathymetries from the restored transects are posted onto the map. In the present study, we distinguish between subaerial exposure, 0-200 m, 200-500 m and more than 500 m water-depth on the maps. For deep marine infill units, the overlying seismic isopach maps are used when drawing the contours. The structural map was used to guide the contours for the lowermost stages. For prograding units, the underlying prograding isopach is used in the map construction. In addition, palaeobathymetric interpretations from micropalaeontology are posted on the map and integrated. The northern North .Sea

Following the end of the Caledonian orogeny in late Silurian to early Devonian times, the northern North Sea (Fig. 3) has experienced two major phases of extension in the Permian-early Triassic and Bajocian/Bathonian to late Ryazanian (Ziegler, 1990; F~erseth, 1996). Prior to the latest rifting episode, the depositional environment was deltaic during the deposition of the Brent Group (e.g. Johannesen et al., 1995). The underlying rift structures played an important role in defining the various sub-basins in

T. Kjennerud et al. the early phase of the post-rift interval (N0ttvedt et al., 1995) and the syn-rift water-depth from the Oxfordian and onwards was significant in the rift axis (Rattey and Hayward, 1993). The Cretaceous post-rift succession in the northern North Sea has been given relatively little focus in the literature. This is probably because it is regarded as having low prospectivity. Work on the Cretaceous includes Skibeli et al. (1995), on the Lower Cretaceous sequence stratigraphy and depositional environments, and Gabrielsen et al. (2001), which describes the post-rift basin configuration. Maximum subsidence occurred along the rift axis and the strata are generally progressively onlapping the basin flanks (NCttvedt et al., 1995). The sediments deposited were mainly clays and fine-grained silts, with some deep marine sands deposited in the Agat region in the northern part of the area and limestone around basin margins (Isaksen and Tonstad, 1989). This is probably because the basin was very deep so that the sediment input to this central part was greatly reduced (NOttvedt et al., 1995). In the present study, the Cretaceous has been divided into six seismic units (Fig. 4). Two of the units cover the Lower Cretaceous and four cover the Upper Cretaceous. The Cretaceous sediments have their maximum thicknesses along the rift axis, and thus also recorded the maximum relief (Fig. 5). Gabrielsen et al. (2000) divided the Cretaceous post-rift stage in the northern North Sea into the incipient, middle and late post-rift stages. The incipient post-rift stage covers units K1 and K2 (Lower Cretaceous) and was characterised by local subsidence. The major structural features were inherited from the syn-rift basin. The crests of the major fault blocks were either islands or shoals. The middle postrift stage includes units K3 and K4 (CenomanianTuronian), and was characterised by the internal basin relief becoming gradually filled by sediments. The mature post-rift stage, which consists of units K5 and K6 (Coniacian-Maastrichtian), displayed the most significant reorganisation of the basin configuration. This stage was characterised by a widening of the post-rift basin, complete drowning of the fault block crests and overstepping of the interior rift margins. The sequence stratigraphy of the Cenozoic has been given a lot more attention than the Cretaceous (e.g. Rundberg, 1989; Galloway et al., 1993; Jordt et al., 1995). Jordt et al. (1995) divided the Cenozoic into ten seismic sequences (CSS-1 to CSS-10) (Fig. 4). This scheme has been adapted in the present study. There is a distinct break in sedimentation and subsidence pattern at the Cretaceous-Tertiary boundary and sediments of Danian age are apparently missing from the northern North Sea (Jordt et al., 1995). In the Palaeocene, a marked change occurred

S t r u c t u r a l r e s t o r a t i o n o f C r e t a c e o u s - C e n o z o i c (post-rift) p a l a e o b a t h y m e t r y in the n o r t h e r n N o r t h S e a

351

Fig. 3. Structural map of the northern North Sea Basin at late Ryazanian level. The database utilised in this study contained the four displayed transects and twelve key wells. In addition, regional seismic data coverage was available together with a large number of well logs. ESB -- East Shetland Basin; ESP = East Shetland Platform; L T - Lomre Terrace; MgB - Magnus Basin; MrB - Marulk Basin; M T = Mfilcy Terraces; SB = Stord Basin; U H - Utsira High; VG = Viking Graben; WG = Witch Ground Graben; AG = Asta Graben.

in the study area, as coarse clastic sediment entered the basin from the west (Milton et al., 1991) as a response to uplift of the East Shetland Platform. During the late Palaeocene-early Eocene (CSS-1 to CSS-2), coarse clastics were sourced from the East Shetland Platform, with some minor supply from southern Norway (Jordt et al., 1995). The sediment entering the basin from the east was mainly clay (Isaksen and Tonstad, 1989). In the central part of the basin, parallel seismic facies can be seen in the reflection seismic data, indicative of a low-energy environment. The early Oligocene (CSS-3) sequence indicates uplift and rejuvenation of sources along both margins of the basin. The late Oligocene sequence (CSS-4) was dominated by eastward progradation from the western margin. The Miocene is characterised by a widespread hiatus in the northern North

Sea, indicated by biostratigraphic data (Gradstein and B~ickstrOm, 1996; Martinsen et al., 1999). According to Jordt et al. (1995), Lower Miocene sediments (CSS-5) are thin or absent and may represent shallow marine deposits with some renewed outbuilding from the west and the Middle Miocene (CSS-6) is missing or condensed. CSS-7 represents the Upper Miocene and is mainly deposited in the southern part. The Pliocene sequence (CSS-8) represents major prograding units associated with the uplift of southern Norway. The top of CSS-8 is marked by a major unconformity, which cuts across older units on the eastern flank of the basin. CSS-9 represents further outbuilding and infilling of relief, whereas CSS-10 represents late Quaternary glacial deposits deposited across the base Quaternary unconformity.

T. Kjennerud et al.

352

Palaeobathymetry from the structural restoration Palaeorelief has been restored on four regional crustal transects (Fig. 3). These transects have been constructed from a combination of conventional seismic lines, deep seismic data, gravimetric and magnetic data. The transects were depth-converted using velocity information from wells, seismic stacking velocities and velocities from seismic refraction data (ESP).

The Cretaceous development The Cretaceous has been divided into six seismic units in the study area, all of which are present in the Viking Graben area on Transects 1-3, while only five where defined on Transect 4. Fewer units were detected in the platform areas. The deep marine infill scenario has been used in all the Cretaceous restorations, with some minor exceptions for Transect 4, where Volgian progradations were utilised to determine late Ryazanian palaeobathymetry in the Stord Basin (Fig. 6). The validity of the Cretaceous restorations relies on correct identification of the zero water-depth points used in the restoration (Fig. 7). In the inverse restoration, the relief was assumed to have been smooth at the Cretaceous-Tertiary boundary for all the four transects. This is based on the following observations. (1) The seismic reflections of unit K6 are parallel to those of CSS-1 in the central parts of the basin. (2) The K6 isopach (Fig. 8) indicates that all the rift relief was filled in. (3) The micropalaeontological estimates of palaeobathymetry are relatively similar in all of the study area (Gillmore et al., 2001), typically varying Fig. 4. Seismic units used in the present study.

Fig. 5. Cretaceous units K1-K4 in the northern Viking Graben. The vertical exaggeration is 20 times.

Structural restoration of Cretaceous-Cenozoic (post-rift) palaeobathymetry in the northern North Sea

353

Fig. 6. Volgian progradation in the Stord Basin. Section is squashed. The vertical exaggeration is 25 times.

7~ ..~~

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Seismic section through the Kvitebjern structure. The Kvitebjem fault ~3~ block was utilised as a ~,,~,,~::; 2::,':z! zero water depth point ~~," in the late Ryazanian ~ ~ . : restoration, although it :~\~,:.~.,. ?." may have been just a ......"~ ~' '~' shallow area (see Kyrk~i;~~ jebo et el., this volume).

.

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.'~. 9 ;,~..~,~.~: ~, ,~.,,:',,..~....,,I T " - - ~ : " :'~" " ~ ' ~ = - ' ~ , ' ~ ; ~

~~%",'.,;.~.~7.~..,;* sediments in the near' E ;~<,,,, ..... ~..' .i"~; i~'. .,........... !.i-"-' ,,. ~" ~ ~ < , . ~ . ~ ~:'; by 29/6-1 well (down o ~ ~ ; ; ! ~ ~ ~ : flank) is of Cenoman,,'=,;....,. ......... ....... ;:~< ian age(see Kyrkjebe '~""~" "i"~"~';'' ~..... ~ et el., this volume). It ~,.,.,..,,,,,. 9 : . ,~.!: j ~~,~. ~,,~,. , .:. . .~. . . "" has been inferred that the crest of the struc~ ~ ture drowned in early ~ ) 1<5 time. " '

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Seismic section through the Lomre Terrace (35/11-3). The 35/11-3 was exposed or shallow up until the Coniacian (see Kyrkjebe et el., this volume).

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Seismic section through the Hild- AIwyn alignment. The

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Seismic section through the eastern flank of the Utsira High, near transect 4. Latest Cretaceous rocks lie unconformably on Jurassic and older rocks on the Utsira High. It has been inferred that the Utsira High drowned in K5 time.

;~,..~",',,~:'k['~~:~ ,.~.,:,;., ; .,,, ,,.. ,~,;~ ,, ,,+~. - , ; , ~ ! ; , . , , ,:,..,~ ry",:'" " ':*~.... '"":"~;""~"~ ~ 5 km

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Seismic section through the area between the Viking Graben and the Horde Platform on transect 2. According to Skibeli et el. (1995) this area was possibly exposed or shallow until middle Cretaceous times. It has been inferred that this area drowned in the latest Cenomanian.

5 km

Fig. 7. Documentation of zero water-depth for various structural highs. The numbers on each section refer to their location in Fig. 3.

T. Kjennerud et al.

354

Fig. 8. Seismic isopach maps of the Cretaceous units K1 and K6.

between 150 m and 300 m. Even if there was a certain relief in the basin in the latest Cretaceous, it must have been modest having a wavelength of the order of 250 km. Areas of zero water depth and shoals have been described in Fig. 7. In addition it has been inferred that the Gullfaks block drowned in K5 time. This is based on the fact that latest Cretaceous sediments are overlying Bathonian sediments. The 31/2-8 well is situated on the western crest of the Horda Platform and contains 3 m of Cretaceous sediment. This area probably became submerged during K5 time. The East Shetland Platform was a regional high in the Mesozoic and Tertiary (Platt and Cartwright, 1998). In many parts, it probably did not drown before K5/K6 time. According to Cherry (1993), Maastrichtian sediments lie on top of Late Jurassic sediments on the edge of the East Shetland Platform in the Brae area. In addition it is likely that the fault block crests in the East Shetland Basin and on the Horda Platform were exposed or shallow in the early Cretaceous. The incipient post-rift stage is represented by the restoration of late Ryazanian (Fig. 9a, Fig. 10a), early Aptian (Fig. 9b, Fig. 10b) and earliest Cenomanian (Fig. 9c, Fig. 10c). The late Ryazanian water depth along the Viking Graben axis was 900 m in the north and 400 m in the south. In the western sub-platform area, Kvitebj~rn and Gullfaks fault blocks, the tilted

blocks in the East Shetland Basin and the HildAlwyn Alignment acted as shoals or were exposed. The sub-basins in between generally had water-depths less than 200 m, with the exception of the Magnus Basin, which was deeper than 500 m. Further west, the East Shetland Platform was exposed. The situation in the eastern sub-platform area was similar, where the crest of the Horda Platform, the Uer and Lomre Terraces, and the crest of the Huldra fault block and the Utsira High, were either exposed or shoals. In addition, the fault block crests on the Horda Platform were shallow or exposed areas. The sub-basins dominantly had water-depths of less than 200 m. The maximum water-depth in the Stord Basin was 600 m. In summary, the northern North Sea was a highly segmented basin in the late Ryazanian, and the rift axis was deeper in the north than in the south. The southern part of the Viking Graben was probably shallower because of higher clastic input in this area during the Late Jurassic. A palaeobathymetric map of the late Ryazanian is presented in Fig. 11. The K1 isopach (Fig. 8), which was used to construct the palaeobathymetric map, shows the uneven distribution of the K1 sediments, which reflects the rift segmentation. The palaeobathymetric map shows a basin with many islands and deep areas. Three distinctive sub-basins existed in the Viking Graben area,

Structural restoration of Cretaceous-Cenozoic (post-rift) palaeobathymetry in the northern North Sea

355

Fig. 9. Restoration of the Cretaceous development on Transects 1 and 3.

each separated by a transfer zone. The Stord Basin was also deep. The early Aptian represented a beginning of the infill of the relief created in the late Jurassic rifting.

The fixed water-depth points remained constant on all of the transects except for the Kvitebjorn fault block on Transect 1, which had been drowned and was covered in sediment. In all the restorations, the basin was

T. Kjennerud et al.

356

Fig. 10. Restoration of the Cretaceous development on Transects 2 and 4.

markedly shallower in the early Aptian. From north to south, the water-depth along the basin axis varied between 400 m and 150 m. The platforms and terraces remained shallow. From the restorations of the early

Aptian, it is clear that the early post-rift subsidence was governed by the segmentation of the rift. The earliest Cenomanian was similar to the early Aptian. The subsidence was still mainly focused

Structural restoration of Cretaceous-Cenozoic (post-rift) palaeobathymetry in the northern North Sea

357

Fig. 11. Palaeobathymetricmap for the late Ryazanian.

within each individual sub-basin, and the waterdepths stayed relatively constant. The south to north deepening is evident also on this level. Kyrkjebo (1999) quantified the vertical movements of the post-rift phase. According to this work the shallowing in the incipient post-rift stage has been attributed to tectonic uplift, and was probably caused by renewed activity within the North Atlantic rift system. The basin gradients during the incipient post-rift stage were generally high, especially in the areas bordering the Viking Graben. As sediment was filling into the rift topography, the basin gradients gradually decreased. The middle post-rift stage is represented by the late Cenomanian (Fig. 9d, Fig. 10d) and the earliest Coniacian (Fig. 9e, Fig. 10e) restorations. The late Cenomanian represented a change in the basin development, as the northern part of the basin started to subside with a longer wavelength. At the same

time the basin became markedly deeper. The palaeobathymetry along the basin axis was 600 m in the north, 700 m on Transect 2 and 450 m in the south. The situation in the western sub-platforms remained mostly unchanged, while the eastern sub-platforms in the northern part of the basin started to subside. This resulted in the drowning of the Huldra fault block. In the northern part of the study area, the reflectors of the infilling K4 unit show a divergent pattern with thickening towards the Tampen Spur area (see reflection pattern in Fig. 5). The increase in bathymetry is best constrained for the northern two lines, as these have been tied together in the restoration. The Viking Graben restoration on Transect 4 was based on the geometry of an outbuilding sedimentary wedge from the Utsira High. The basin gradients decreased on the eastern flank of the Viking Graben on the northern two transects, while an increase occurred in the southern part of the study area. This increase

358

T. Kjennerud et al.

occurred because the subsidence was more focused in the southern Viking Graben. In the earliest Coniacian, the water-depth along the basin axis was 600 m in the northern part of the basin, 700 m in the middle (on Transect 2) and 200 m in the south. The western sub-platforms remained unchanged, while the eastern sub-platform areas were being caught up in the basin-wide subsidence. In the east, the Lomre Terrace and the area between the Viking Graben and the Horda Platform on Transect 2 were drowned. Kyrkjebr (1999) attributes the increased waterdepths in the middle post-rift phase to thermal cooling and contraction of the lithosphere. The mature post-rift stage is represented by the earliest Maastrichtian (Fig. 9f, Fig. 10f) and earliest Palaeocene (Fig. 9g, Fig. 10g) restorations. The rift topography was filled in during the deposition of the K5 unit in the northern part of the basin.

There was still a certain control on morphology by the underlying rift structures in the southern part of the study area, but otherwise the Viking Graben and surrounding areas were subsiding as a unified basin in the earliest Maastrichtian. No areas were exposed in the earliest Maastrichtian. A significant change is that the basin axis was shifted westwards, in the northern part of the basin, compared with the previous stages. The water-depth in the basin axis was 400 m in the north and 200 m in the south. Fig. 12 shows a palaeobathymetric map for the earliest Maastrichtian. The infilling K6 isopach (Fig. 8) is clearly part of a different type of basin development than the K1 isopach. The palaeobathymetric map shows a lower relief and a widening of the basin towards the north. The gradients were markedly lower than the earlier stages. In the earliest Palaeocene restoration (Fig. 9g, Fig. 10g), the relief has been inferred to have been smoothened as stated above. All the transects were

Fig. 12. Palaeobathymetricmap for the earliest Maastrichtian.

Structural restoration of Cretaceous-Cenozoic (post-rift) palaeobathymetry in the northern North Sea restored flat for this time stage. This does not, however, imply that the basin was exposed at sea level as indicated in the restorations, as it is not possible to determine water-depth using the present method when there is no or very little relief. According to Kyrkjebo (1999), tectonic subsidence ceased in the latest Cretaceous and thermal equilibrium was reached.

Cenozoic development The Cenozoic development was markedly different to that of the Cretaceous. While the Cretaceous can be characterised by gradual infill of already existing rift relief, the Cenozoic is characterised by episodic uplift and sediment outbuilding from eastern and western source areas. As with many of the Cenozoic restorations, the earliest Eocene (Fig. 13h, Fig. 14h) was restored by using the prograding sequences approach as described above. The topsets of the prograding clinoforms in the western part of the study area have been used as zero water-depth points because of the presence of coal and shallow marine sands (e.g. Rochow, 1981; Ziegler, 1990). Transect 1 does not cross any of the main Palaeocene depocentres, which represents progradation in the Palaeocene. Transect 3 intersects both the eastern and western depocentres, and had an important control on the palaeobathymetric estimates on Transect 1. Transects 2 and 4 both pass through important western depocentres. Compared with the late Cretaceous, the basin axis had shifted slightly eastwards, due to high sediment input from the west. Along the basin axis, the water-depth was 700 m in the north and 450 m in the south. The basin gradients were characterised by being high near the palaeocoast, but otherwise gentle. Kyrkjeb~a (1999) thought the increase in water depth during the Palaeocene to be connected to acceleration in subsidence for the entire basin, caused by the arrival of the Iceland plume. Outbuilding from the west continued during the Eocene (Jordt et al., 1995), resulting in a further eastward shift of the basin axis and a narrower basin. Little sediment entered the basin from the east during the Eocene. The earliest Oligocene restoration (Fig. 13i, Fig. 14i) shows the active progradation of the western part of the basin. In the eastern part, the Eocene sediments are draping Palaeocene progradational units. In the restorations it is clearly seen that the deepest part of the basin has been shifted eastwards. The water-depths were 800 m in the north and 450 m in the south along the basin axis. The basin gradients remained high near the coast. Outbuilding from the west continued in the early Oligocene (Fig. 13j, Fig. 14j), but at this time, much of the accommodation space was filled. The maxi-

359

mum water-depth was 300 m in the north and 200 m in the south. The gradients were low in the early Oligocene, compared with the previous stage. This shallowing in water-depth may be explained by an uplift event, which affected the northern North Sea and south Norway (Kyrkjeb~a, 1999). The uplift of the basin continued into the Miocene. The sediment input from the west continued in the late Oligocene. The upper boundary of the CSS-4 sequence was modified by clay diapirism (Jordt et al., 1995). Neither the geometry of the CSS-4 sequence, nor the infilling Miocene CSS-5 sequence, is indicative of large relief. Because of a lack in both seismic and stratigraphic resolution in the Neogene, the top CSS-4 restoration (Fig. 13k, Fig. 14k) is thought to represent most of the Miocene. In most wells, the Miocene is dominated by a large hiatus (e.g. Gradstein and B~ickstr6m, 1996; Martinsen et al., 1999), which may be due to very few samples. It is an open question whether the basin stayed relatively deep during the Miocene and received very little sediment, or was at or near sea level. Martinsen et al. (1999) interpreted the Miocene hiatus to be due to non-deposition. In the present study, the basin was restored at sea level, although a minor relief has been indicated in the restorations of Transects 3 and 4. The earliest Pliocene restorations (Fig. 131, Fig. 141) illustrate the situation just prior to the Pliocene CSS-8 outbuilding from the east. This was associated with uplift of southern Norway and resuited in tilt and extensive erosion in the eastern part of the sections (Jordt et al., 1995). The restorations illustrate the space needed to accommodate the prograding CSS-8 sequence. The restoration was based on decompacting the clinoforms of the infilling CSS-8. This resulted in the following maximum water depths from north to south; 500 m on Transect 1, 650 m on Transect 3, 500 m on Transect 2 and 600 m on Transect 4. The deepening was associated with an increase in basin gradients. The deepening in the Pliocene was caused by subsidence of the basin (Kyrkjeb~a et al., 2000). Fig. 13m and Fig. 14m show a restoration of the relative relief in the earliest Quaternary. The deep created in the earliest Pliocene was now filled in, leading to shallower water-depths. The erosive relief of the Norwegian Trench can be seen in the eastern part of the section. Fig. 15 summarises the palaeobathymetric development in the deepest point for each transect. Discussion and conclusions In the restoration method used in the present study, the main uncertainties are associated with depth con-

360

T. K j e n n e r u d

e t al.

Fig. 13. Restoration of the Tertiary development on Transects 1 and 3. E S P = East Shetland Platform; M B = Magnus Basin; E S B = East Shetland Basin; G = Gullfax; K = KvitebjCrn; l i p - Horda Platform; H A = Hild-Alwyn Alignment; H - Huldra; L T = Lomre Terrace; U T = Uer Terrace. v e r s i o n a n d c o m p a c t i o n . T h e choice of velocity m o d e l

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361

Structural restoration of Cretaceous-Cenozoic (post-rift) palaeobathymetry in the northern North Sea

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played as the dotted line in the palaeobathymetric curves for the Cretaceous, where the deep marine infill approach described above has been used in the restora-

T. Kjennerud et al.

362

Fig. 15. Palaeobathymetric pseudowells for Transects 1-4.

tion (Fig. 15). It can be seen from this figure, that the "undecompacted" estimates follow the trends as those models obtained when decompaction was taken into account for the infilling unit. The most crucial constraints in the present work are areas that represent zero water-depth or shoals. Particular care has been put into quality controlling these. The restorations presented here may be further improved in the future when additional shoals/zero water-depth points are added, particularly for the earliest Cretaceous. The construction of the palaeobathymetric maps worked both as a visualisation of the basin at a given time and also as a quality control. The map exercise had crucial importance as a quality control because the maps were an integration of the restored transects, the seismic isopachs and the results from micropalaeontology. The basin shapes suggested by the 2D reconstructions correlate with the relative basin morphology indicated in the seismic isopach maps. This proves the power of using actual depositional geometries in palaeobathymetric reconstructions. The integration study performed by Kyrkjebr et al. (2001) between the present study and the palaeobathymetry obtained from micropalaeontology by Gillmore et al. (2001) proved that the two data sets

generally correlated both in shallowing-deepening trends and also in many cases on absolute values. This works as a quality control on both approaches and also justifies both. The post-rift basin gradients have been quantified in the present work. This is valuable additional information to what is possible with the micropalaeontological approach, which is based on single wells (e.g. Gillmore et al., 2001). A schematic evolution of the Cretaceous-Cenozoic post-rift relief is presented in Fig. 16. The incipient post-rift stage was characterised

Fig. 16. Schematic development of the basin relief during the Cretaceous and Cenozoic in the northern North Sea.

Structural restoration of Cretaceous-Cenozoic (post-rift) palaeobathymetry in the northern North Sea

by diversely subsiding sub-basins and high gradients. The middle post-rift stage was characterised by lower gradients and a gradual infill of the basin relief. The mature post-rift stage was characterised by a widening of the basin and a low relief. The early and middle Palaeogene was characterised by high gradients near the basin margins and low relief in the basin axis. The late Palaeogene to middle Neogene was characterised by low relief, while the gradients increased in the late Neogene, and decreased up until the present day. In summary the following general palaeowaterdepth trends are observed in the reconstructions. (1) A segmented locally, deep basin in the late Ryazanian. (2) In the early Cretaceous the northern North Sea became shallower. (3) The basin started to become both wider and deeper in the Cenomanian. This trend continued into the K5 (Coniacian-Campanian) unit. (4) Shallowing occurred in the latest Cretaceous. (5) The inherited Jurassic rift relief was eliminated by the latest Cretaceous. (6) The basin deepened in the Palaeocene and stayed deep during the Eocene. (7) A shallowing occurred in the Oligocene and continued into the Miocene. (8) The basin became deeper in the Pliocene. (9) The basin has become shallower up to the present day.

Acknowledgements The authors would like to thank the following for financing the project Tectonic Impact on Sedimentary Processes in the Post Rift Phase: The Research Council of Norway, Agip, Amoco, Mobil, Norsk Hydro, Phillips Petroleum, Saga Petroleum and Statoil. Stein Fanavoll, Felix Gradstein, Morten Smelror and Joar Sa~ttem are acknowledged for their interesting discussions on palaeobathymetry. Befit Fossum is acknowledged for drafting the figures. The manuscript has benefited from reviews by S. Corfield and R. F~erseth.

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the 4th Conference. Geological Society, London, pp. 215-249. Roberts, A.M., Kuznir, N.J., Yielding, G. and Styles, E, 1998. 2D flexural backstripping of extensional basins: the need for a sideways glance. Pet. Geosci., 4: 327-338. Rochow, K.A., 1981. Seismic stratigraphy of the North Sea "Paleocene" deposits. In: L.V. Illing and G.D. Hobson (Editors), Petroleum Geology of the Continental Shelf of North-West Europe. Institute of Petroleum, London, pp. 255-266. Rundberg, Y., 1989. Tertiary History and Basin Evolution of the Norwegian North Sea between 60~176 - - An Integrated Approach. Ph.D. Thesis, Norwegian Institute of Technology, Trondheim, 292 PP. Scholz, C.A. and Finney, B.E, 1994. Late Quaternary sequence stratigraphy of Lake Malawi (Nyasa), Africa. Sedimentology, 41: 163-179. Sclater, J.G. and Christie, EA.F., 1980. Continental stretching; an explanation of the post-mid-Cretaceous subsidence of the central North Sea basin. J. Geophys. Res., 85: 3711-3739. Skibeli, M., Barnes, K., Straume, T., Syvertsen, S.E. and Shanmugam, G., 1995. A sequence stratigraphic study of Lower Cretaceous deposits in the northernmost North Sea. In: R.J. Steel, V.L. Felt, E.E Johannesen and C. Mathieu (Editors), Sequence Stratigraphy of the Northwest European Margin. Norwegian Petroleum Society (NPF), Special Publication 5. Elsevier, Amsterdam, pp. 361-372. Sloan, B.J., 1995. Eocene Sequence Stratigraphy of the North Sea Basin. Ph.D. Dissertation, University of Texas, Austin, 214 pp. Vail, ER., Mitchum, R.M., Todd, R.G., Widmier, J.M., Thompson, S., Sangree, J.B., Bubb, J.N. and Hatlelid, W.G., 1977. Seismic stratigraphy and global changes of sea level. In: C.V. Payton (Editor), Seismic Stratigraphy - - Applications to Hydrocarbon Exploration. Am. Assoc. Pet. Geol., Mem., 26: 49-212. Verral, E, 1981. Structural interpretation with application to North Sea problems. Joint Association for Petroleum Exploration Courses, Course Notes, 3. White, N.J., Jackson, J.A. and McKenzie, D.E, 1986. The relationship between the geometry of normal faults and that of sedimentary layers in their hangingwalls. J. Struct. Geol., 8: 897-909. Ziegler, EA., 1990. Geological Atlas of Western and Central Europe 1990. Shell Internationale Petroleum Maatschappij B.V., The Geological Society, London, 239 pp.

SINTEF Petroleum Research, N-7465 Trondheim, Norway Department of Geology, University of Oslo, N-0316 Oslo, Norway Geological Institute, University of Bergen, N-5007 Bergen, Norway University College Northampton, School of Environmental Science, Northampton NN2 7AL, UK Geological Institute, University of Bergen, N-5007 Bergen, Norway Department of Geology and Mineral Resource Engineering, NTNU, N-7465 Trondheim, Norway Statoil Research Centre, N-7005 Trondheim, Norway

365

The reconstruction and analysis of palaeowater depths: a new approach and test of micropalaeontological approaches in the post-rift (Cretaceous to Quaternary) interval of the northern North Sea Gavin K. Gillmore, Tomas Kjennerud and Rune Kyrkjebo

This paper explores and tests a new eclectic approach to micropalaeontological data which are used to reconstruct palaeowater depths in the Norwegian North Sea in the period from the Cretaceous to Quaternary. The new ideas and evidence presented here promote a greater understanding of palaeobathymetry and basin evolution. This account focuses in particular upon Cretaceous and Cenozoic sediments in this region although the new approach has general applicability. Such information is important because palaeobathymetric variations and palaeobasin shape are essential for the recognition of tectonic phases, and understanding the distribution of sediments and source rocks. Palaeobathymetric curves with upper and lower depth limits derived in part from statistical analyses of micropalaeontological data were constructed to present the most likely depth variations through time in the region. The palaeobathymetric curves obtained from micropalaeontological investigations are compared with the palaeobathymetric variations that were suggested by independent structural restorations for the region. The outcomes of these two approaches are shown to be essentially similar which implies the soundness of this new approach to micropalaeontological data. Both the structural restoration and micropalaeontological results point to the following: a shallowing in the palaeowater depths in the Early Cretaceous; a deepening in the middle Cretaceous; a shallowing in the Late Cretaceous; a deepening in the Palaeogene; a shallowing in the early Neogene, and a deepening in the late Neogene (Pliocene). The Quaternary interval in wells in this study illustrates the complexity of the palaeoenvironmental record over this period of time, with some wells clearly showing subsidence while others suggesting relative uplift. This complexity is the result of the impact of ice movement and multiple glaciations and subsequent isostatic re-adjustments.

Introduction

In past studies, many methods have been used to examine palaeoenvironments and palaeobathymetry, but this has rarely been done in a systematic manner. A number of authors discuss palaeowater depths with little real basis for the water depths assigned to particular genera and species (for example, Bowen, 1954). In particular, Van der Zwaan et al. (1999) point out that although benthic foraminifera have often been used to estimate palaeowater depths, there is a lack of any substantial theoretical or observational basis for this. It is hoped that this study goes some way towards addressing this imbalance. One of the main objectives of the present and complementary studies (Kjennerud et al., 2001) was to map and describe the basin configuration and hence the palaeobathymetry of the northern North Sea from the post-rift differential subsidence of the JurassicCretaceous basin of the northern and southern Viking Graben through to the Quaternary. Other objectives

included enhancing understanding of the processes which contributed to the post-rift basin configuration, and how these time-dependent processes influenced the basin development. An understanding of the basin infilling processes, and thereby distribution of sand in the system, was an important element of this work. Information was required on the palaeobathymetry and palaeobasin morphology of the northern North Sea along four key transects for the post-rift stage (Fig. 1). The present study, by integrating palaeowater depth estimates from both micropalaeontological data and structural restoration (Kjennerud et al., 2001) has highlighted changes in sediment accommodation space and sediment supply and their distributions throughout the Cretaceous to Quaternary on a regional scale (Kyrkjeb~ et al., 2001). Gillmore et al. (1998; 1999) demonstrated that computer modelling of the shape of sedimentary basins at any particular moment in time, together with the application of data on palaeowater depths can facilitate appropriate reconstructions of the long-term regional history of the

Sedimentary Environments Offshore N o r w a y - Palaeozoic to Recent edited by O.J. Martinsen and T. Dreyer. NPF Special Publication 10, pp. 365-381, Published by Elsevier Science B.V., Amsterdam. 9 Norwegian Petroleum Society (NPF), 2001.

G.K. Gillmore et al.

366

Fig. 1. Base Cretaceous structural map of the study area showing transects 1-4.

basin. Reliable modelling of palaeobathymetry and palaeobasin morphology also needs precise information on the nature and distribution of hiati within these Cretaceous to Quaternary sequences. A descrip-

tion of these is given elsewhere in this volume by Kyrkjebr et al.

The reconstruction and analysis of palaeowater depths: a new approach and test of micropalaeontological approaches

Previous work

Estimations of palaeowater depths are fundamental to most palaeoenvironmental analysis but are probably the hardest parameter to measure, not least because depth per se is not an environmental variable (Gibson, 1988). Depth has been aptly described by Gibson (1988) as "a location in three dimensional space in which numerous environmental variables.., are limiting factors". In the case of palaeowater depth estimates based upon the study of foraminifera, for example, their original depth distributions are likely to have been a function of several interacting factors: temperature, salinity, hydraulic energy, oxygen levels, nutrients and the nature of the substrate (Sliter and Baker, 1972; Gibson, 1988; Bosence and Allison, 1995; Spencer, 1996; Van der Zwaan et al., 1999). As Gooday (pers. commun., 1998) stated, an understanding of the ecology of foraminifera must be crucial to their use in reconstructing ancient environments including water depths. According to West (1976), the grouping of foraminifera using test structure and habit, arenaceous-calcareous and planktonic-benthic ratios, can be very successful in delimiting depth zones. Phleger (1960) regarded water depth as the most important environmental factor controlling the occurrence of foraminifera. He noted that there is a marked boundary shown by marine microfaunas at 70-125 m in mid-latitudes worldwide. This depth marks the lower limit of the seasonal temperature layer. Other microfaunal depth "boundaries" noted were found at 20 m, 1000 m and possibly 2000 m. Phleger (1960) suggested that the boundary at 2000 m may be due to hydrostatic pressure. Loeblich and Tappan (1964) noted that some particular genera have limited depth ranges and some individual species may exhibit an even more distinctive depth zonation. Said (1950), however, suggested that in the Red Sea, the faunal depth zonation observed was due to the organic content of the water and nature of the substrate. Benthonic species were observed as abundant near coastlines, regardless of water depth. Van der Zwaan et al. (1999) have also emphasized the importance of organic flux in deciding microfaunal abundance. Another inherent advantage of a micropalaeontological approach to the problem of determining palaeowater depths is the recognition that relative changes in the characteristics of microfaunal assemblage can give the necessary insights into palaeowater depths, and in particular, help to identify transgressive and regressive cycles, including the point of maximum water depth and the point of a decrease in water depth which suggests the beginning of a regressive cycle (Gibson, 1988). The palaeobathymetry of the post-rift interval from the Cretaceous to the Quaternary in the northern North

367

Sea has been given little attention in the literature, possibly because it is regarded as having low prospectivity. Skibeli et al. (1995) have worked on the Early Cretaceous sequence stratigraphy and depositional environments, while Hancock (1984) has reviewed the Cretaceous for the entire North Sea region. It is clear that underlying rift structures played an important role in defining the various sub-basins in the early phase of the post-rift interval. Maximum subsidence occurred along the rift axis and the strata generally show a backstepping character (NOttvedt et al., 1995). The sediments deposited in the Cretaceous were mainly clays and fine silts, with some deep marine sands deposited in the Agat area and limestone around the basin margins (Isaksen and Tonstad, 1989). General interpretation is aided because the sequence stratigraphy of the Cenozoic has undergone more scrutiny than for the preceding Cretaceous (e.g. Rundberg, 1989; Galloway et al., 1993; Jordt et al., 1995).

Definitions of palaeowater depth Gradstein and B~ickstr6m (1996) approached the problem of water depth by defining five depositional environments in Tertiary sediments of the northern North Sea and the offshore mid-Norway area. These were: (1) non-marine, terrestrial with no foraminifera; (2) shallow neritic, less than 100 m deep (contains benthic foraminifera but with low generic and species diversity and rare planktonics); (3) deep neritic, 100200 m deep (foraminiferal assemblages of varying diversity, planktonic foraminifera locally common); (4) upper bathyal, i.e. upper slope, 200-500 m deep (high generic and species diversity, common to frequent calcareous benthonic s, assemblages may be dominated by coarse agglutinants in the Palaeogene); (5) middle bathyal, i.e. middle slope, 500-750/1000 m or deeper (high foraminiferal diversity, Neogene assemblages include calcareous benthics, Palaeogene assemblages may be dominated by coarse and fine agglutinated foraminifera). Jones and Charnock (1985) from evidence in the modem North Atlantic presented their own definition of these environments, based on Hedgpeth (1957). These definitions of marine environments are essentially similar, and as a result, the Gradstein and B~ickstr6m (1996) definitions have been adopted in this study. Methods

The structural approach A "structural method" which has been applied to palaeowater depth reconstruction in this study area

O3 TABLE 1 Table of micropalaeontological techniques used to assess palaeobathymetry Techniques

Authors

A measure of:

Limitations

P:B ratios.

Murray (1976).

Nearness to shore. Upper depth limits (Gibson, 1988).

Limitations due to absence of planktonics; problems with selective dissolution occurs (i.e. more fragile selectively dissolved versus robust forms).

P: B and NCA ratios.

King et al. (1989); King (1989)

Sea-level changes.

Identification of NCAs? Few limitations but dissolution problems enhance NCA% due to syn-sedimentary or post-diagenetic dissolution of CaCO3.

Tau values.

Gibson ( 1988); Murray ( 1991 ).

Depth.

Relative depths only.

Generic/spp. composition.

Peet (1974); Murray (1991); Gradstein and B~ickstr6m (1996).

Environments of deposition.

Reliable? Depends on modern analogues. Less reliable going back through time.

Modern analogues.

Sliter and Baker (1972).

Diversity/environments of deposition.

Good if you can show that they are relevant.

Morphogroups related to trophic structures.

Jones and Charnock (1985); Koutsoukos and Hart (1990).

Different microhabitats, dominated by different forms (shapes).

Difficult to assess value in terms of palaeobathymetry for some morphogroups.

Species diversity (several techniques, see below).

Murray (1991).

Species (richness) and heterogeneity (evenness).

Different techniques may provide different answers.

Abundance/direct counts.

Nagy et al. (1988); Laursen and Andersen (1997).

Species richness (Peet, 1974).

Time consuming. Requires equal sample sizes. Studies rare.

Fisher alpha index.

Murray (1991); Ujetz (1996).

Species richness (Murray, 1991).

Will be reasonable guide to environment (depth?). Tendency for alpha to increase with sample size (Murray, 1991). Requires a minimum of 100 specimens.

Equitability.

Jones (1988); Murray (1991).

The distribution of individuals between species.

More reliable than some other techniques?

Shannon-Wiener/Weaver formulations (H (S)).

Kaminski et al. (1988); Murray (1991).

Diversity. A compound of richness (S) and evenness.

Reliability? The "information function" is "valid" only for an infinite sample (Murray, 1991). The function depends on sample size.

Brillouin formula.

Pielou (1966, 1979).

Diversity, for finite samples. Quicker and easier than S or H (S).

Not an acceptable index of heterogeneity (Peet, 1974). Crude index of little interpretational value (Murray, 1991).

Jaccard's coefficient.

Sliter and Baker (1972).

Assemblages/associations.

Not a widely used technique.

/ - m o d e cluster analysis.

Olsson and Nyong (1984); Kovach (1989).

Associations/biofacies.

Will only provide associations. Needs to be used with other techniques. Includes such multivariate techniques as Constrained Clustering of Kovach (1989, 1995).

q'-mode varimax factor analysis.

Kaminski et al. (1988).

Idealised composite assemblages.

Reduces large amounts of information into a condensed format, possible to lose data.

Order proportions.

Murray (1991); Ujetz (1996).

Diversity/environments based on wall structure.

Best used in more shallow water environments (Murray, 1991).

Morphological/size variations.

Murray (1976); Jones (1988).

Size/variations with depth.

Needs a transect to be of real value.

~,,~~

The reconstruction and analysis of palaeowater depths: a new approach and test of micropalaeontological approaches

is described at length by Kjennerud et al. (2001). It serves as a basis for comparison with the ideas and outcomes presented in this paper which are based upon various micropalaeontological approaches. This structural procedure involves section balancing/restoration techniques (using LOCACE software, Beicip Franlab, 1996) for extensional regimes together with seismic sequence stratigraphy. This structural method is based upon analysis of regional seismic profiles that generally traverse the basin (see Fig. 1). Coal layers were used as fixed water-depth points, as were shallow marine sands and subaerial unconformities, which were used to define points along the profile. Geometric information in the under- and overlying seismic sequences were also used to reconstruct the shape of palaeobasins.

The micropalaeontological approach The primary concern of this paper, however, is to present an eclectic approach to using micropalaeontological evidence to deduce palaeowater depths in this case in the northern North Sea. The various advantages and disadvantages of each approach explored in this study are summarised in Table 1. Those selected for use are discussed more fully below. A number of these approaches share similar limitations. The further back in the geological record sediments are investigated, clearly the lower number of species will be recovered that have survived into modern faunas, where their oceanographic tolerances can be determined. For example, in the Palaeogene and Cretaceous the number of species that exist in modern assemblages is less than 5% (Gibson, 1988). Obviously over time even the generic composition of faunas may also change significantly (Gibson, 1988). (1) The P : B % (i.e. planktonic:benthonic ratio) of Murray (1976, 1991; and Fig. 2). P : B % should show a general increase in planktonic foraminifera from the inner shelf to deeper waters (Murray, 1976; Gibson, 1988). This index can be used as a measure of nearness to shore and has been used by Gibson (1988) to suggest the upper depth limits of selective foraminifera (Table 1). Van der Zwaan et al. (1999) suggest that P : B ratios are sensitive to sediment transport, and to oxygen deficiency. (2) P: B% versus NCA% (percentage of non-calcareous agglutinants, see Fig. 2) as proposed by King (1989) and King et al. (1989). According to King (1989) this difference measure can be used to estimate sea-level change over time. Perhaps one of the limitations of this approach is the identification of NCAs in well-reports. High NCA values reflect poor water circulation, poor oxygenation, as well as stagnation of water at the seafloor. Such envi-

369

ronments may indicate relatively low sea levels and regressive regional episodes (King et al., 1989). High percentages of planktonic foraminifera usually indicate periods of relatively high sea level (King et al., 1989). Such conditions allow open oceanic circulation which is associated with diverse and abundant calcareous benthonic faunas. High oxygen concentrations are possible at the seafloor in association with marine transgressive phases. King et al. (1989) used the contrast between these ratios to divide the Cretaceous seas in the North Sea region into open and restricted phases. (3) Tau values (Gibson, 1988) have been used in this study, a procedure also followed by Murray (1991). This faunal index was defined by Gibson (1988) as the number of benthic species multiplied by planktonic percentage. This index is effectively a measure of relative water depth, i.e. the higher values indicate deeper waters. However, problems do exist with this index where there is low planktonic percentage, possibly due to dissolution. Dissolution of calcareous tests may play an important role in influencing their relative abundance (Murray and Alve, 1999). (4) Data on the ancient and modern tolerances of selected taxa have been used wherever possible, especially in terms of knowledge of their upper depth limits. This information has been used to define palaeowater depths following the ideas of Van Morkhoven et al. (1986) and Gradstein and B~ickstr6m (1996). It is, however, important to recognise that there may be variations in upper depth limits in different environmental settings, for example, cold versus tropical environments (Murray, 1991). (5) Modern generic and assemblage analogues have been used to assess environments of deposition and hence palaeowater depth (Risdal, 1960; Sliter, 1972; Sliter and Baker, 1972; Pearce, 1980; Gradstein et al., 1983; Murray, 1991). This particular analogue approach requires the demonstration that the analogues selected are relevant, especially where there is a large difference in location or time between the fossil community and the possible modern analogue. Sliter and Baker (1972) examined the bathymetric significance of Cretaceous benthic foraminifera (genera) from the Pacific margin and listed them in order of depth significance based on comparisons with modern faunas. They concentrated on inner shelf to lower slope forms. These authors noted that recurrent fossil genera that had been delineated by numerical analysis, and generic predominance patterns did match those of modern genera. It was pointed out that there were strong similarities between Cretaceous seas and those of today in the Pacific margin region. According to Gibson (1988), Sliter's (1972) approach is valid because his work in California was carried

370

G.K. Gillmore et al.

Fig. 2. Concise graph correlating an eclectic micropalaeontological analysis of Cretaceous sediments, comparing trends of tau (Gibson, 1988), NCA% (non-calcareous agglutinant percentage of King, 1989), P'B values (planktonic'benthonic foraminiferal ratio of Murray, 1991) and species diversity to assess palaeoenvironment and transgressive and regressive cycles for well 30/10-6.

The reconstruction and analysis of palaeowater depths: a new approach and test of micropalaeontological approaches

mE DATING

Early Pliocene

INFERRED 'ALEOWATER DEPTH (Gilly o r e et a/.,;999) Ainimum Most like1 Maximu

css-L

I

Common and diverse calcareous benthonic assemblage.Bul,mina elongafa recorded (inner-mid shelf. van Morkhoven el al., 1986).Tau increases downhole Suggesting a deepening through this interval.

css-7

-

1996). Ammonia beccarrii

Cjbfcrdordespachyderma (upper slope, van Morkhoven el al., 1986) occurs together with rare agglutinated benthonic and planktonic foraminifera. Within this intervalBulimrna aculeata occurs (slope and deeper, van Morkhoven el al., 1986) plus bolivinids (outer shell and upper slope. Sliter, 1972, Koutsoukos and Hart, 1990).Tau values indicate deeper waters than in the overlying intcrval. IncreasingP B % suggests more open Oceanic conditions. although NCA% also increases slightly downhole.

css-7 CSS-E

MICROFAUNAL COMMENTS

Cibicidids (sublitoralzone 01 Atkinson. 1971, down to 14m water depth; most abundant in mid shelf. Bhatia, 1957). cassidulinids (shelf to slope, Murray, 1991) and elphidiids (shelf. Murray, 1991;above lOOm water depth. Gradstein and %&Irom, (inner shelf O-Wm, Murray. 1991) and Pararofalia serrafa (Pararofalia, inner shelf. Murray, 1991) occur.

I

pp. 371-374

100

100-300 300

200

200-500

500

500-750 750

500

500-750 ?1OOC

Deeper intervalthan above suggested by presence by uvigerinids (water depths deeper than 1 Wrn , Murray, 1991).Calcareous benthonics indicate outer shelf to slope depths (e.g. PuNenia,Murray, 1991).Rare planktonics suggest oxygenated surface waters with oceanic circuiation.Increased tau values downhole suggesting deeper water conditions. Rare elphiiids and asterigerinids suggests shelfal influence. in the middle Miocene Melonis affine noted (upper slope, 200-500m. Gradstein and Backstrom. 1996; 100-18Om Murray, 1979) Also noted Ebrenbergma (outer shell to slope, Murray, 1991).Agglutinated foraminifera are rare sporadic Occurrenceof Martinotiel/acommunis (slope, Gradstein and Backstrom. 1996).

-

CSS-f

css-4

css-3

Eocene

500

CSS-2

Eocene Early Eocene

Early Eocens

Paleocene

The First Oownhole Occurrence (FDO's) of cyclamminlds. RecurvoJdes(submorphogroup AG -82 in slope environments. Koutsoukos and Hart, 1990). rhabdamminids (submorphogroup A, Jones and Charnock 1985) and Cnbrostomordes suggests upper slope (or deeper) deposition Increase in radiolaria downhde suggests more open oceanic surface water circulation or upwelling conditions IncreasingNCA4hdownhole

The FDO of Ammomarginulina auberfae suggests deeper water conditions (upper depth liml2500m North Atlantic. Charnock andJones, 1990).Labrospira scifula noted (upper depth limit 600m N. Atlantlc, Charnockand Jones, 1990 genus mostly middle slope, 700-1000m. Gradstein and BBckstr6m. 1996).plus rhabdamminids and small cystamminids (middle slope, Gradstein and Backstrom. 1996).Decrease in NCA% downhole, towards the base of this interval. Salcareous benthonk foraminifera Occur in this intervaltypical of outer shelf to upper slope environments (e.g. chilostomelllds. Murray, 1991).plus abundant planktonic foraminlfera (e.g.Subbotina lrnaperfa, shallow-intermediate depths. Hart, 1980). %her forms include Cibicidoidespraemundulus(primarily a lower slope and abyssal taxon, Tjalsma and Lohmann. 1983:upper depth limit around 500m. Berggren and Aubert. 1976).Increased tau values downhole suggest deeper water conditions. NCA values decreased while D:B% rose (bener-oxygenated waters). In the upper part of this interval diatoms (pyritised) occur (e.g. FenestreNa antqua and Coscinodiscus morsianus) plus a poor agglutinated assemblage consisting of rhabdammininds (upper slope or deeper, Jones and Charncock, 1985).The diatoms suggest the presence of upwelling conditions (Bonde, 1982:Mtiehner. 1996).while the poor agglutinated assemblage suggests poorly oxygenated bottom waters.

:ss-8 css-1

Paleocene indicates upper slope or deeper depths (Jonesand Charnock, 1985;Gradstein and Backstrom, 1996 The FDO's of Paleocene spiroplectamminids with increased numbers of rhabdamminids. ammodisclds, cystamminids, rzehakinlds, Cribrosfomoides and Reficulophr8@??ium Speijer el al , 1997) in this interval. ~

500

500

I

The presence of heterohellcids (predominant) with globigerinellids suggests shelfal deposition (Sliter, 1972).Calcareous benthonics include osangularis, gavelinellids (upper slope or deeper, Sllter, 1972).Rhizamminidssuggest slope (or deeper) conditions (Jones and Charnock, 1985;rhabdamminidssuggest outer shelf to upper slope, King el al., 1989) as does Recurvoides welten' (Charnock and Jones, 1990). Calcareous benthonks Include osangularids and gavelinelllds (generally upper slope or deeper, Slner, 1972). "rominent Pseudotextularia elegans elegans noted (sheifal. Sliter. 1972). Globotruncanidsoccur (deeper water slope, Sliter, 1972;Hart, 1980).Calcareous benthonics include lenticulinids(outer shelf to slope, Murray, 1991) and quadrimorphinids (shelf and Jpper-middle slope morphotypes, Koutsoukos and Hart, 1990).Deeper waters than the intervalabove suggested by common Rbizammrna spp and Recurvoides Walter; (Charnock and Jones, 1990). ;he presence of Archeoglobigerina suggests possibly shallower conditions at the base of this section in the late Campanian (shallow-intermediate depths. Hart, 1980).Small cystamminids occur towards the base of this interval (in the early Maastrichtian) suggesting leeper water conditions (slope, typically middle and deeper, Gradstein and Backstrom. 1996) Deeper waters at the top of this interval are suggested by forms such as Reussella szajnochae noted by Widmark and Speijer (1997) in their 'deep bathyal assemblage'. Noted by Butsoukos and Hart (1990) in middle-outershelf to upper-middle slope depths. Also noted was Globofruncanella havanensis (slope, Sllter, 1972).

(earliest?) L.Maastricht.

Slobotruncanids common (slope. Sliter, 1972) Rzehakinldsoccur (middle slope typically. Gradstein and Backstrom. 1996) together with calcareous benthonics e.g. Praeblimha (upper slope or deeper, Slter, 1972). The latter genus is more 'ecologically significant' in middle slope depths (Slter, 1972).

Campanian

Santonian

In the late Turonian both Dicdrinella (deep marine. Hart. 1980) and WhiteineNa (intermediate and shallow-intermediate. Hart. 1980) occur. In the Coniacian the presence of qwdrimorphinids (shelf and upper to middle slope morphotype. Koutsoukos and Hart, 1990), globotruncanids (slope. Sliter. 1972) and Rhizammind8afbysiphonspp. (slope and deeper, Jones and Charncck, 1985).Relativdy rare heterohelicids occur suggesting more shelfal conditions (Sliter, 1972).Stensbeinids in the early Santonian suggest outer shelf to upper slope depths (Gradstein and Backstrom, 1996).Higher tau values suggest deeper waters. Late Santonian faunas include radiolaria suggesting open oceanic conditions (Dyer and Copestake. 1989). Decreased tau values suggest shallower conditions at the top of this interval, while towards the base small cystamminids suggest slope condtions (Gradstein and Backstrom, 1996).Heteroheliclds were also noted suggesting shelfal deposition, akhough they m u r in reduced numbers on the slope (Sliter, 1972).In the early part of the Campanian spiroplectammininds (Upper slope. S l W 1972) occur together with hedbergellidsand calcareous benthics. Saccamminids were noted (submorphogroup El, Jones and Charnock. 1985) suggesting upper slope or deeper depths Dorothia crassa was noted (Dorofbiaoccurs in outer shelf and deeper waters, SlRer. 1972).Gioborolalidsnoted (outer shelf and upper slope, Koutsoukos and Hart. 1990).together with trltaxids (outer shelf to slope, Speijer el al., 1997).

Early

Santonian

Coniacian

r

Uirofaunas dominated by planktonic foraminifera.Calcareous benthonicsinclude Stensioerna beccarriiforrnis(slope and abyssal, van Morkhoven el al., 1986).Agglutinated species include Mafanzia varians (= Rernesella varians). a calcareous agglutinant (water depths typically j0-200m).The calcareous benthonics present suggest shallower conditions than above, but tau values incraase downhole suggesting deeper conditions.

Turonian

The presence of common globotruncanids suggests deeper water deposition (Slier, 1972). Other planktonic foraminifera includes Dicarinellahagni, typical of deep marine environments (Hart. 1980).Common hedbergdlids suggests open shell deposition (Butt, 1979).Benthonk oraminifera include Uwgennamrnmajankor (upper sbpe. Kuhnt el al , 1989) and consistentiy common Rbizamminaspp. at the base of this 'interval' (upper slope or deeper, Jones and Charnock. 1985).Tau values suggest a deepening at the base of this interval followed by a ;hallowing and a further deepening downhole. This is supported by the NCA values.

Middle to Early Turonian

'100

100-500 500

Cenomanian

r== Cenomanian? t o u p. p . e r Albiar

200

200-500

500

I

7500

Early Barremian

7'0°

'100500?

Faunas inclvde hedbergdlids, globigerinellids, Whiteinella, with subordinate numbers of agglutinated and calcareous benthonic foraminifera. The piesence of hedbergelli dominated assemblages suggests deposition on an open shelf (Butt. 1979).Gavellinellids occur in deep neritic (outer shell) depths according to Koutsoukos and Hart (1990),and are an important part of an upper slope assemblage in the late Cretaceous (Sliter, 1972).Tau values decrease, suggesting shallower condtions. Tau values suggest deeper water conditions. particularly towards the base of this interval.Tftaxids noted (outer shell to slope, Speijer el al., 1997) plus ammodiscids (slope, Koutsoukos and Hart, 1990). saccamminids (submorphogroup 61, Jones and Charnock, 1985). upper slope to abyssal, 200->2250m).Calcareous benthonics occur, e.g.Quadrimorphina albertensis. Quao'rimorphina is typically noted in shelf and upper-middle slope depths (1.e. subrnorphogroup CH.43.1, Koutsoukosand Hart, 1990) in the Cretaceous. Stensioeinids were nlso noted (outer shen to upper slope in theTertiary. Gradstein and Backstrom, 1996).Cystammlnids were noted at the base of K3K2 (more common in middle slope depths, Gradstein and Backstrom, 1996) Low tau values suggest shallow conditions.The presence of Uvigerinammina rnoesiana suggests slope depths (Kuhnt el al., 1989).Other typical forms of this slope assemblage are Glomospira cbaroides. G. gaultina, Recurvoides ,Gavelinella 'barremianaformis' [= G. barremiana?),Lentrcu//na kug/er/ and L. muensferi.Gavelinellids noted by Sliter (1972) as an important part of an upper slope assemblage in the late Cretaceous. G. barremma belongs lo submorphogroupCH.-A5 of Koutsoukos and Hart (1990).middle outer neritic (she0 and upper slope depths. Lenticulina is an outer shell to upper slope genus (Murray, 1991).

Fig. 3. Composite generalised well analysis relating ages, seismic units and inferred palaeowater depths to selected microfaunal comments.

I

The reconstruction and analysis of palaeowater depths: a new approach and test of micropalaeontological approaches

out in an area where there was good stratigraphic and palaeogeographic control. (6) The concept of morphogroups or morphofunctional analysis (Gibson, 1988) has been recognised and employed by Jones and Chamock (1985), Corliss and Chen (1988) and Koutsoukos and Hart (1990). This approach facilitates an examination of different microhabitats, because they may be dominated by different shapes, or forms of tests. Koutsoukos and Hart (1990) pointed out that one of the functions of the foraminiferal test is to favour a particular mode of life (e.g. dwelling habits, feeding strategy) in a particular substrate niche. Jones and Charnock (1985) presented an ecological model that related four morphogroups based on feeding strategies and inferred life positions to water depth. Corliss and Chen (1988) proposed that distinct morphological trends in foraminifera and microhabitat patterns both exist and vary with depth in the modem Norwegian Sea. According to Geroch and Kaminski (1992) this approach has become an increasingly popular method of palaeoenvironmental analysis, but care must be taken in inferring feeding strategy groups and water depths. (7) Measures of foraminiferal species diversity have been frequently used in the past (Table 1). The Fisher alpha index (Murray, 1991; Ujetz, 1996) is one widely used measure of species diversity or richness. This index is a reasonable guide to the environment and possibly depth (Murray, 1991). There is a tendency, however, for alpha to increase with sample size for statistical reasons a feature that can complicate interpretation. (8) Multivariate analysis of taxon abundance data has been successfully used in many ecological and palaeoecological studies (Kovach, 1989) and is employed as one strand in this new eclectic study. A number of different multivariate methods have been used here, but they have different mathematical bases and assumptions that must be kept in mind. The aim of such studies is to detect and summarise any underlying trends or patterns. In Kovach's (1989) study, clustering of the Spearman's Correlation Coefficient using UPGMA (unweighted pair group averaging methods) gave distinct clusters of taxa and samples that could be easily interpreted. In the present study, the statistical procedure adopted was the "Constrained Clustering" provided by the MVSP software package (Kovach, 1995). This is the only available software package designed especially for geological applications which contains the Constrained Cluster Analysis technique. This method is used for the classification of sequential data (Vincent, 1995). Clustering proceeds as normal, but the objects to be fused are constrained by having to be adjacent in the data matrix. The calculated dendogram thus has the ob-

375

jects (samples) in the same order as the input matrix (Kovach, 1993). In this study, this feature has enabled sections to be examined whilst the samples are kept in the correct stratigraphic order. It is from the dendogram that the palaeoenvironmental interpretations used here have been derived.

New procedures devised In this study faunal associations were used to define maximum and minimum water depths. Within these faunal associations we looked at key species and suggested a best-fit palaeowater depth curve. No algorithm was used to define the latter. The faunal associations were recognised using a series of techniques previously outlined. Key species recognised were those for which reliable information already existed. By reliable information the authors mean: (1) modem analogues/relatives; and/or (2) published information on fossil assemblages established with reliable models concerning water depths. Data sources

Data for this particular project have been gathered from consultancy well reports for twelve Norwegian North Sea wells placed along or near the key seismic lines: 15/5-3, 17/3-1, 24/9-1, 29/6-1, 30/4-1, 30/10-6, 31/6-3, 33/9-18, 34/7-1, 34/11-1, 35/11-3 and 35/12-1 (Fig. 1). Palaeobathymetric models have been erected for Cenozoic and Cretaceous times in these wells. The palaeowater depth curves here presented are related to the sequences of Skibeli et al. (1995) and Bjerke (1998) for K1 to K6 (Cretaceous sediments), and Jordt et al. (1995) CSS 1-10 (Cenozoic sediments). Fig. 3 shows a composite generalised well analysis, relating ages, seismic units and inferred palaeowater depths to some selected microfaunal comments used in our account. Results

The present study explicitly attempts to improve the coherence and consistency of results likely to be obtained in all such palaeobathymetric and palaeoenvironmental studies, a goal suggested by Peet (1974). All the methods described in the previous section of this paper have been used in this study to assess the palaeoenvironment and possible transgressive and regressive cycles that have taken place in the northern North Sea from the Cretaceous to the Quaternary. The results of this new integrated and eclectic analysis are displayed on Fig. 2 for an example well (30/10-6). This plots and compares the trends of tau, NCA, P:B values and species diversity in a form of

376 a concise graph. In brief, this research demonstrates that it is possible to reconstruct water depths using consultancy well data of variable quality. The nature of ancient water depths in the Norwegian North Sea are explored in more detail below.

Structural and micropalaeontological palaeowater depth comparisons In the following we compare suggested water depths derived from independent structural approaches to the new integrated approach to micropalaeontological data. Palaeowater depth analysis are determined for twelve wells and selective depth curves are presented as representative examples (Fig. 4).

Cretaceous palaeowater depths A shallowing is suggested in the Early Cretaceous by the structural reconstruction (Fig. 4). This shallowing, however was only observed in micropalaeontological data from well 30/4-1, and implied in wells 34/11-1, 35/11-3, 35/12-1 (Fig. 4). Often Early Cretaceous sediments are missing in wells in this study; an unfortunate fact that makes the comparison of evidence impossible in many wells. The Early Cretaceous stratigraphic breaks noted in well 35/11-3 for example (Fig. 4) are suggested by Kyrkjebr et al. (2001) to be a result of subaerial erosion and exposure. A general deepening was noted in the mid-Cretaceous, on the basis of the structural restoration (Fig. 4). This is supported by the present analysis of micropalaeontological data from cores, side-wall cores and ditch cuttings for most wells (15/5-3, 17/3-1, 30/4-1 and 34/11-1, Fig. 4). In well 34/11-1, lower tau values at the base of K3 (Cenomanian) than towards the top of K3 and base of K4 (Cenomanian to Turonian) suggests a deepening up-hole. This is also supported by the microfaunal character. In well 31/6-3, this analysis suggests a shallowing in K3/4 times; but this interesting new event also seems to be fully supported by the structural restoration. In well 30/10-6 microfaunal analysis suggests a slight shallowing in K3/4 and into the base of K5/6 (Fig. 4, Coniacian to Maastrichtian), whilst in well 35/12-1 the new integrated approach shows a brief shallowing in K3/4 followed by a deepening in K5/6. There seems to be a slight delay before deepening events suggested by the structural restoration became noticeable by micropalaeontological evidence. This delay may have been due to the influence of local tectonic activity. The wells concerned may have been on structural highs, so that they affected the crests of these highs rather late (Fig. 1). In general, in the Late Cretaceous, there is a widespread shallowing suggested by the microfaunal evidence from all of

G.K. Gillmore et al.

the wells in the study area. The shallowest part within K6 (Maastrichtian) appears to be just below the K6/CSS1 (Maastrichtian/Paleocene) boundary. Above this point, there is a deepening into CSS1 suggested by structural restoration.

Cenozoic to Quaternary palaeowater depths The trend towards basin deepening suggested by the structural restoration is supported by the new analysis of micropalaeontological evidence. The deepest point was reached towards the top of CSS 1 (Fig. 4). This deepening continued into the lower part of CSS2 (Eocene) according to the structural approach. The analysis suggests a very brief shallowing at the Late Paleocene/Early Eocene boundary (Figs. 3 and 4). This inference is based on the presence of the diatom-rich Balder/Sele Formations with rare agglutinated foraminifera that was followed by the occurrence of a rich and diverse red-stained planktonic foraminiferal assemblage. The estimated tau values in wells 29/6-1, 33/9-18 and 34/11-1 (Fig. 4) suggest deeper water conditions within this planktonic-rich horizon, although a number of the species present suggest a shallowing (that is, the presence of calcareous benthonic forms typically noted in outer shelf to upper slope depths). Either way, there was a distinct faunal change at this time, perhaps as a result of changing oceanic conditions (more oxygenated bottom waters and better water circulation than in the underlying Balder/Sele Formations). The analysis here suggests a deepening within CSS2 (Eocene), with the presence of a rich and diverse agglutinated foraminiferal fauna in well 34/11-1 followed by a distinct shallowing in the Oligocene (CSS3-4). This is evident in most wells (Fig. 4). In well 34/11-1 there was a decrease in P : B % and tau in the Early Oligocene (CSS3), while in the Late Oligocene the abundance of sponge spicules together with a fairly rich and diverse calcareous benthonic fauna suggests shelfal to upper slope conditions. In several wells, this shallowing occurs mainly at the top of CSS4 (Fig. 4). The structural restoration suggests a distinct shallowing beginning at the base of the Oligocene, a conclusion supported by the present analysis of micropalaeontological data in some of the wells in this study. In the seismic interpretation (for example, for well 33/9-18) it was usually not possible to differentiate between the Miocene seismic units CSS5, CSS6 and CSS7 in the northern North Sea (Kyrkjebr et al., 2001). It could also not be decided with any certainty in this study whether subaerial/submarine erosion or non-deposition caused the Middle Miocene (CSS6) hiatus (KyrkjebO et al., 2001) in many wells, although Kyrkjebr et al. (2001) suggest that the missing middle-Late Miocene and latest Miocene to Early

e5

~..~~

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~176

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Fig. 4. Comparisons of palaeobathymetry from structural restoration and micropalaeontology. The inferred palaeowater depth curves shown have been constructed in the following ways. The solid error bars on the palaeowater depth curves represent our estimate of "most likely" range of palaeowater depths, while the dotted error bars indicate our estimate of the "possible" range of palaeowater depths. The blue palaeowater depth line defines our view of the "most likely" curve suggested by micropalaeontology. The red lines indicate the reconstruction of palaeobathymetry based upon consideration of the geological structure over time. The palaeowater depths for each environment (shelf, slope, etc.) are based on Gradstein and B~ickstr6m (1996). For information, these charts also show the sequences of Skibeli et al. (1995), Jordt et al. (1995) and Bjerke (1998), together with hiati (or stratigraphic breaks) and sedimentation rates.

e5

"-4 "-,1

378

Pliocene section of well 29/6-1 is a result of subaerial erosion (Fig. 4). Jordt et al. (1995) documented a large hiatus in the Miocene in the Norwegian North Sea. Sediments of this age are usually missing in the study area (Fig. 4). Martinsen et al. (1999) note that this Miocene break is characterised by significant erosion of the underlying Oligocene in the MOre Basin. This break varies time-wise depending on the extent of the Utsira Formation in the Central North Sea Basin axis (Martinsen et al., 1999). Martinsen et al. (1999) also suggest that the lower part of the Miocene break is an erosional vacuity with the upper part being a hiatus of 12 m.y. duration. Eidvin et al. (1999) indicate that this hiatus is related to tectonic events in the Middle Miocene. Both our micropalaeontological investigations and structural reconstruction suggest that the Norwegian North Sea Basin in the study area deepened again in the Pliocene (Figs. 3 and 4). Many wells show a deepening in the Pliocene sequence CSS8, a shallowing into the Quaternary CSS9, and a deepening into CSS10 (wells 17/3-1, 30/4-1, 30/10-6, 31/6-3, 33/9-18, 34/11-1, 35/11-3, Fig. 4). The deepening in sequence CSS8 is fully supported by the micropalaeontological evidence (for example, the presence of Oridorsalis u m b o n a t u s in well 29/6-1 suggesting slope depths or deeper; Mackensen et al., 1985; Murray, 1991; higher tau values in well 30/10-6 within the Early Pliocene than in the Late Pliocene) but the deepening in CSS 10 is not clear from this approach. This is apparent in wells 29/6-1 (Fig. 4) and 35/12-1 where micropalaeontology suggests a shallowing from CSS8 to CSS10. A few wells in the study area show an apparent shallowing within the Quaternary (wells 15/5-3, 24/9-1 and 34/7-1), but two of these lack structural results. The structural reconstruction in well 35/12-1 contradicts the general trend by suggesting an overall deepening throughout CSS8 to CSS10, without a shallowing within CSS9. This illustrates the complexity of the Quaternary palaeoenvironmental record. Discussion

A number of key points emerge as a result of these investigations. (1) The two independent methods (structural and micropalaeontological) for deriving past water depths are in substantial agreement. There is, for example, a distinct and rapid deepening in CSS3 into CSS2 shown by the microfaunas and structural restoration in all wells in this study. This event is usually associated with relatively high sedimentation rates (Fig. 4). Nevertheless, some contradictions occur. These centre mainly around the position of wells for micropalaeontological analysis compared to the restored transects.

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For example, in well 30/10-6 the structural restoration suggests a deepening in the Cenomanian (K34), while the micropalaeontology shows relatively little change, indeed a shallowing may have occurred locally at the base of K5. The Cenomanian deepening suggested by the structural analysis can be seen in other wells in the study area and is confirmed by micropalaeontology from those wells (e.g. well 34/11-1; Fig. 4). According to the micropalaeontological reconstruction there is a deepening that occurs in well 30/10-6 in the Campanian (Figs. 3 and 4), followed by a shallowing that is indicated by both micropalaeontology and structural work. Well 30/10-6 was located on the flanks of the Cretaceous deep-water basin (Kjennerud et al., 1999, 2001), so palaeowater depths calculated from micropalaeontology will reflect depths at that particular location, i.e. the basin flank. Other contradictions centre around variable data quality available for microfaunal analysis. In well 30/4-1 (Fig. 1) in CSS2 times palaeowater depths suggested by structural analysis were around 800 m; while microfaunal evidence suggests a possible 500 m water depth, the bulk of the microfaunal evidence indicates depths of around 300-400 m. When the contractors report was originally produced for well 30/4-1 in 1979 many deep-water indicator species were not recognised and listed, hence the suggested microfaunal depths are probably too shallow. Kjennerud et al. (1999) show that in the earliest Eocene well 30/4-1 would have been near the centre of the basin in water depths deeper than 500 m. In well 34/11-1 in CSS2 (Figs. 1 and 4) water depths have been suggested by the micropalaeontological approach of around 500 m with deepest depths around 750-800 m in mid-CSS2. The structural analysis in this well over this interval suggests water depths around 800 m throughout. This discrepancy between micropalaeontological and structural results for part of CSS2 in this well is possibly a result of facies control on the occurrence of deep-water agglutinated benthonic foraminifera, such as small cystamminids (Gradstein and B~ickstr/Sm, 1996). At the top of CSS2 sand is occasionally present in this well, with light greenish claystones predominating mid-CSS2 and below. Indeed, such cystamminids occur inconsistently at the top of CSS2, and do not occur consistently until mid-CSS2, Middle Eocene. (2) The new integrated/eclectic method has recognised features and trends not previously signalled by the structural method, but which are likely to be real and not statistical anomalies or inventions. (3) The present authors conclude that their new approach is robust and that it has been tested successfully against independent data derived from a long period of geological time.

The reconstruction and analysis of palaeowater depths: a new approach and test of micropalaeontological approaches

(4) The approach has revealed coherence between the behaviour of the various measures and parameters calculated from existing micropalaeontological data. Therefore this new approach has appropriate internal consistency. As a result of this research, we can now derive with greater confidence palaeowater depth trends based upon interpretations of micropalaeontological data and explore with greater confidence the nature and significance of past variations in sea level in the northern North Sea.

Conclusions This paper has successfully established estimates of palaeowater depth in the Norwegian North Sea through the use of modem foraminiferal species analogues and by calculating a series of surrogate statistical measures derived from existing information on identification and frequencies. These derivative measures have been compared and contrasted to produce a best micropalaeontologically based estimate of palaeowater depth changes from the Cretaceous to Quaternary, although the complexity of change in Quaternary palaeogeography and water depths in this area are shown to be far from fully realised by this analysis. This study has established which micropalaeontological approaches that in combination appear to give the most consistent results. A number of microfaunal methods to assess palaeoenvironment, and hence palaeowater depth, should be used, because according to Peet (1974) there is no sound basis for comparing the richness of a series of communities through using only a single index. A verification of this micropalaeontological method estimating palaeowater depth was obtained by comparing the outcomes of this new eclectic approach with palaeodepth estimates derived independently from a "structural method". The latter method involved section balancing/restoration techniques designed for extensional regimes in combination with interpretations of seismic sequence stratigraphy. This comparison revealed an essential similarity between the outcomes of the micropalaeontological and structural approaches; an outcome that encourages confidence in this new approach. The present study has emphasized the correctness of the comments by Gradstein and B~ickstr6m (1996) who pointed out the importance of palaeobathymetry for establishing the depositional history, subsidence and burial analysis of a region. An overestimation of palaeowater depth may lead to inadequate estimates of uplifts and an underestimation or reduction in subsidence rates. Ignorance of palaeobathymetric trends may therefore lead to an underestima-

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tion of basin subsidence and incorrectly assigned palaeoslopes (Gradstein and B~ckstr6m, 1996). Palaeobathymetry has also been restored in this study on regional transects, using this new combination of methods for structural restoration and microfaunal analysis. This has enabled restoration of the basin profile in the post-rift interval in the Norwegian North Sea region. A series of shallowing and deepening trends have been recognised in the study area. A shallowing was noted in the Early Cretaceous (based primarily on structural restoration), followed by a deepening in the mid-Cretaceous and a shallowing in the Late Cretaceous. The basin became deeper in the Paleocene, followed by a shallowing in the Oligocene to Miocene. The basin became deeper again in the Pliocene. The complexity of palaeowater depth change in the Quaternary interval is well illustrated here in a manner which has not previously been achieved. This complexity is a result of the interactions between various factors: eustatic sea-level changes, multiple glaciotectonic events, erosion and a lack of precise dating. The Norwegian North Sea Quaternary palaeogeography cannot be understood without a sound understanding of previous water depths and seafloor topography and vice versa. This requires further investigation. Finally, it is evident that a broader understanding of fluctuations in palaeowater depth can help to reveal tectonic events. Evidently significant changes in water depth may be associated with increased sedimentation rates. This is illustrated in wells 29/6-1 and 35/11-3 (Fig. 4) where a significant shallowing in the Oligocene sequence CSS3 is associated the significantly greater sedimentation rates. Tectonic events may have led to uplift and hence increased erosion of the source area.

Acknowledgements Thanks are due to SINTEF Petroleum Research, the Research Council of Norway, University College Northampton, U.K., NTNU, Universities of Bergen and Oslo, and clients (Agip, Amoco, Mobil, Norsk Hydro, Phillips, Saga, and Statoil) of the Tecsed (Tectonic Impact on Sedimentary Processes in the PostRift Phase - Improved Models) project for financing this research and permission to publish.

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fera. John Hopkins Press, Baltimore, MD, 297 pp. Pielou, E.C., 1966. Shannon's formula as a measure of species diversity: its use and misuse. Am. Nat., 100: 463-465. Pielou, E.C., 1979. A quick method of determining the diversity of foraminiferal assemblages. J. Paleontol., 15:1237-1242. Risdal, D., 1960. Foraminiferfaunaenes relasjon til dybdeforholdene i Oslofjorden, med en diskusjon av de senkvartaere foraminifersorter. (The bathymetrical relation of Recent Foraminiferal Faunas in the Oslo Fjord, with a discussion of the Foraminiferal Zones from Late Quaternary time). Nor. Geol. Unders., 226, 142 pp. Rundberg, Y., 1989. Tertiary History and Basin Evolution of the Norwegian North Sea Between 60 ~ and 62~ An Integrated approach. Unpublished Ph.D. Thesis, University of Trondheim, 292 pp. Said, R., 1950. The distribution of Foraminifera in the northern Red Sea. Cushman Found. Foraminiferal Res., Contrib., 1: 9-29. Skibeli, M., Barnes, K., Straume, T. Syvertsen, S.E. and Shanmugam, G., 1995. A sequence stratigraphic study of Lower Cretaceous deposits in the northernmost North Sea. In: R.J. Steel, V.L. Felt, E.R Johannesen and C. Mathieu (Editors), Sequence Stratigraphy on the Northwest European Margin. Sequence Stratigraphy of the Northwest European Margin. Norwegian Petroleum Society (NPF), Special Publication 5. Elsevier, Amsterdam, pp. 389-400. Sliter, W.V., 1972. Cretaceous foraminifers depth habitats and their origin. Nature, 239:514-515. Sliter, W.V. and Baker, R.A., 1972. Cretaceous bathymetric distribution of benthic foraminifers. J. Foraminiferal Res., 2: 167-183. Speijer, R.R, Schmitz, B. and van der Zwaan, G.J., 1997. Benthic foraminiferal extinction and repopulation in response to latest Paleocene Tethyan anoxia. Geology, 25: 683-686. Spencer, R.S., 1996. A model for improving the precision of paleobathymetric estimates: an example from the northwest Gulf of Mexico. Mar. Micropaleontol., 28: 263-282. Tjalsma, R.C. and Lohmann, G.R, 1983. Paleocene-Eocene bathyal and abyssal benthic foraminifera from the Atlantic Ocean. Micropaleontol. Spec. Publ., 4, 90 pp. Ujetz, B., 1996. Micropaleontology of Paleogene deep water sediments, Haute-Savoie, France. Universit6 de Genbve Publication du D6partment de Gdologie et Paldontologie, 22, 149 pp. Van der Zwaan, G.J., Duijnstee, I.A.R, den Dulk, M., Ernst, S.R., Jannink, N.T. and Kouwenhoven, T.J., 1999. Benthic foraminifers: proxies or problems? A review of paleoecological concepts. EarthSci. Rev., 46: 213-236. Van Morkhoven, F.P.C.M., Berggren, W.A. and Edwards, A.S., 1986. Cenozoic cosmopolitan deep-water benthic Foraminifera. Bull. Cent. Rech. Explor.-Prod. Elf-Aquitaine, Mem., 11,479 pp. Vincent, A.J., 1995. Palynofacies Analysis of Middle Jurassic Sediments from the Inner Hebrides. Unpublished PhD thesis, University of Newcastle upon Tyne, 474 pp. West, R.R., 1976. Trophic classification of benthic communities. In: R.W. Scott and R.R. West (Editors), Structure and Classification of Paleocommunities. Dowden, Hutchinson and Ross, Stroudsburg, PA, pp. 29-66. Widmark, J.G.V. and Speijer, R.R, 1997. Benthic foraminiferal faunas and trophic regimes at the terminal Cretaceous Tethyan seafloor. Palaios, 12: 354-371.

UniversiO' College Northampton, School of Environmental Science, Northampton NN2 7AL, UK SINTEF Petroleum Research, N-7034 Trondheim, Norway Geological Institute, UniversiO' of Bergen, N-5007 Bergen, Norway

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383

Outcrop-based classification of thick-bedded, deep-marine sandstones M.O. Badescu

The outcrop-based classification of thick-bedded deep-marine (TBDM) sandstones from sand-rich and mud/sand-rich deep-water systems is built upon the evaluation of facies associations of the TBDM sandstones, their geometry, lateral extent, and the variation of the sand-to-gross ratio (S/G) of the complexes containing TBDM sandstones. The study evaluates individual basin-fill complexes related to an equivalent stage of sedimentation. Three main generic and geometric types of TBDM sandstones were identified. Type A represents TBDM sandstones with a very good correlation (along and across the palaeocurrent direction). The S/G ratio of complexes comprising this type is low, but uniform (0.5-0.4). Type A TBDM sandstones were deposited in unconfined basins, or basins with little obstruction to flow. Type B represents thick sandstones with a good correlation along and poor correlation across the palaeocurrent direction. Complexes comprising type B TBDM sandstones present a highly variable S/G ratio (0.9-0.3). They represent amalgamated, channelised thick sandstones, deposited either in open or in constricted basins, with the poor lateral correlation being due to the elongated geometry of the channels. Type C represents sandstones with generally poor correlation. Complexes comprising type C TBDM sandstones locally present very high S/G ratios (0.9). Type C TBDM sandstones were deposited in topographically confined basins. The classification can be applied to the North Sea's Paleocene and Eocene reservoirs comprising TBDM sandstones. The Paleocene reservoirs are of large size and very sand-rich, with stacked pay zones. The Eocene reservoirs are more localised. Basin-floor topography defines their geometry and lateral extent. Type A and B TBDM sandstones are mostly recognised in the Paleocene reservoirs and type C TBDM sandstones are commonly recognised in the Eocene reservoirs. The study reveals that the TBDM sandstones present different spatial signatures that influence the reservoir behaviour in the early stages of production.

Introduction

Thick-bedded deep-marine sandstones (TBDM) from sand-rich and mud/sand-rich deep-marine depositional systems have been the subject of ongoing research for many years. The processes of transport and deposition of the TBDM sandstones, and the factors controlling their architecture and their geometry are still relatively poorly understood. The architecture and geometry of the TBDM sandstones have significant implications for reservoir quality (porositypermeability variations); hence an improved understanding of the nature of these deposits is critical for the accurate evaluation of reservoirs containing TBDM sandstones. In the last five years, there has been considerable debate concerning the nature and the mechanisms of deposition of the TBDM sandstones (Shanmugam and Moiola, 1994, 1995; Slatt et al., 1997; Lowe, 1997; Bouma et al., 1997; Shanmugam and Moiola, 1997; Shanmugam, 1999). The debate focused on core studies from the North Sea and on outcrop studies. The main concern is whether the TBDM sandstones are deposited by high-density turbidity

currents or by sandy debris flows. In the high-density turbidity current scenario TBDM sandstones are deposited by turbulent flow and are considered to be continuous, homogeneous and with little variation in the porosity-permeability characteristics. If the sandy debris flow scenario is accepted, the TBDM sandstones are deposited by freezing of laminar flow and are expected to be discontinuous, heterogeneous and with unpredictable geometry and porositypermeability characteristics. In previous work the influence of the basin-floor topography on the geometry and architecture of the TBDM sandstones was often neglected. The present study introduces important recognition criteria and common elements of TBDM sandstones based on an overview of TBDM sandstones in outcrops. The objective is to establish characteristics for thick-bedded sandstone successions in order to differentiate types of TBDM sandstones that could be recognised in the subsurface. This overview suggests that the basin-floor topography plays an important role in the geometry, shape and architecture of the TBDM sandstones. The idea proposed here is based on the evaluation of facies associations of the TBDM

Sedimentary Environments Offshore Norway - Palaeozoic to Recent edited by O.J. Martinsen and T. Dreyer. NPF Special Publication 10, pp. 383-405, Published by Elsevier Science B.V., Amsterdam. 9 Norwegian Petroleum Society (NPF), 2001.

3t34 sandstones, their geometry, lateral extent and the variation of the sand-to-gross ratio (S/G). This study evaluates individual basin-fill complexes related to an equivalent stage of sedimentation. The cases represent a variety of basin shapes and dimensions. The thickness of the complexes that were evaluated ranges from 40 to 215 m and their extent from 1 to more than 100 km. All the information used in this text was gathered from open literature, with the exception of the "Mames B leues" study, which represents own field results. Eighteen case studies were initially evaluated. From this data set only ten cases remained, after careful consideration of the sandstone beds' geometries, and after evaluating the nomenclature used so that those contributions could be compared without bias. This paper comments on more than 50 articles on thick-bedded sandstones and associated facies.

Deep-water sandstone facies characteristics and depositional settings Historically, prominent papers on turbidites and other deep-water deposits have been made by Kuenen (1957), Bouma (1962), Mutti and Ricci Lucchi (1972), Walker (1978), Normark (1970) and Shanmugam and Moiola (1988). TBDM sandstones have received special interest over many years. However, only in the past few years their origin and mechanics of transport and deposition have become the subject of serious debate. Much of our knowledge about turbidite facies and mechanisms of deposition of deep-marine sandstones is attributed to Bouma (1962), whose work was primarily related to the Annot sandstones, SE France, where his now classical Bouma sequence was described. Mutti and Ricci Lucchi (1972) developed a comprehensive fan model based on Apenninic outcrops. TBDM sandstones (facies B of Mutti and Ricci Lucchi, 1972) occur from the slope to the mid-fan deposits. Normark (1970) proposed a suprafan model based on recent turbidites of Californian borderland basins, and Walker (1978) incorporated the suprafan concept within the Mutti and Ricci Lucchi models (1972), resulting in a hybrid model, based on both ancient and recent deep-marine sediments. These models are shown in Fig. 1. Later, Mutti (1979) underlined the importance of the grain size and the volume of sediment that influence the facies characteristics and the architecture of the deep-water deposits and developed a new model of deep-marine clastics: efficient vs. inefficient systems. However, this model was widely criticised by Shanmugam and Moiola (1988), who also proposed a new classification scheme based on tectonic setting.

M. O. Badescu

Regarding mechanisms of deposition, Kuenen and Migliorini (1950) carried out the first experiments. Kuenen also proposed the term "turbidite" for the deposits of a turbidity current (Kuenen, 1957). Further experimental work was carried out by Middleton and Hampton (1973) who recognised four main endmember flow types based on the grain support mechanism (gravity, fluidised, liquefied and grain flows). Lowe (1982) suggested a classification and nomenclature based on rheology of the flow and the particle support mechanism. More recently, Shanmugam and Moiola (1995) claimed that TBDM sandstones could be deposited by sandy debris flows. Their novel, still debatable interpretation was based on a case study of the Jackfork Group and on cores from the North Sea, but not on experiments. Their interpretation was amply criticised by Slatt et al. (1997), Lowe (1997), and Bouma et al. (1997). Another interpretation was suggested by the experiments of Kneller (1995) who proposed that deposition of massive sands occurs due to non-uniformity in prolonged, quasi-steady, highdensity turbidity currents. His work on experiments took into account the flow velocity and its spatial and temporal variations. Shanmugam (1999) challenged this idea. Parallel with this work on outcrop and experimental studies, seismic and sequence stratigraphic concepts were developed (Vail et al., 1977; Mutti, 1985; Posamentier and Vail, 1988). New emphasis was put on temporal relationships between source area and basin fills. Mutti (1985) developed a model for the evolution of fan systems during rises and falls of relative sea level that incorporates the features of high- and low-efficiency fans into a single scheme (Types I, II and III). His model was developed for sand-rich delta-fed systems in tectonically active areas. TBDM sandstones can occur within systems I or II of these models (Fig. 2). More recently, Reading and Richards (1994) have introduced a novel classification scheme of turbidite systems in deep-water margins. The focus was on the grain size and the feeder system. It is not clear how one can distinguish between sand-rich and mud/sand-rich systems and the evolution of the systems in time was not taken into account. TBDM sandstones can belong to the sand-rich fans, mud/sand-rich fans, sand-rich ramps, mud/sand-rich ramps or sand-rich aprons (Fig. 3). These studies have led to the present state of the art, which still contains controversies in many respects. Although our knowledge has much increased, TBDM sandstones are still not satisfactorily understood. This problem is related to the fact that the origin of massive TBDM sandstones is due to a variety of processes involving both steady and unsteady

Outcrop-based classification of thick-bedded, deep-marine sandstones

385

Fig. 1. Historical overview of models for deep-water clastic systems. (A) Mutti and Ricci Lucchi (1972) - single fan model. (B) Normark (1970) point source, fan model observed in recent sediments. (C) Walker (1978) - point source, multiple fan model observed both in ancient and modern sediments. Reprinted by permission of the American Association of Petroleum Geologists, Tulsa, OK, whose permission is required for further use. -

Fig. 2. Application of sequence stratigraphy principles in turbidite systems, Mutti (1985). Reprinted with kind permission from Kluwer Academic Publishers, Dordrecht.

386

M. O. Badescu

Fig. 3. Overview of models for deep-water clastic systems. From Reading and Richards (1994). (A) Mud/sand-rich fan. (B) Sand-rich fan. (C) Mud/sand-rich ramp. (D) Sand-rich ramp. (E) Mud/sand-rich slope apron. (F) Sand-rich slope apron. Reprinted by permission of the American Association of Petroleum Geologists, Tulsa, OK, whose permission is required for further use.

flows. There are currently three scenarios for deposition of TBDM sandstones (Stow et al., 1996): (1) deposition from high-concentration turbidity currents, where grain concentration at the depositional surface is sufficiently high to cause hindered settling; some sands, containing scattered outsize clasts, may originate from this process; (2) deposition from freezing of sandy debris flows; (3) sediment collapse fallout. Another fact that complicates the problem of TBDM sandstones is that they can be deposited in a va-

riety of tectonic settings. Relative sea level, the source area and the basin-floor topography are interacting factors that have a major influence in the deposition of TBDM sandstones. TBDM sandstones are generally deposited during a relative fall in sea level, and the source of the TBDM sandstones must be very sandrich (Stow et al., 1985; Picketing et al., 1989; Stow et al., 1996). The models previously considered are too general and basically take into account only three (major) fac-

Outcrop-based classification of thick-bedded, deep-marine sandstones tors: tectonics, source area and relative sea-level variations. These models do not provide sufficient data for an accurate appraisal of an oil reservoir. More factors should be taken into consideration, if working on a smaller scale (up to 5-10 km). The local factors (basin-floor topography, differential compaction, local tectonics), at this scale will have a critical influence on the geometry and the architecture of these deposits. Information about reservoir aspects was mainly restricted to oil company reports for many years. Lately, these aspects were discussed in the literature (Chapin et al., 1994; Cossey, 1994), but the production behaviour received little attention. This paper describes an approach for predicting reservoir properties around exploration and appraisal wells.

Evaluation of outcrop descriptions The objectives of recognising facies-types characteristics for TBDM sandstones requires outcrop descriptions that give reliable vertical profiles combined with a sufficient spatial distribution of data. In

387

addition, the basin setting, provenance and synsedimentary geological events have to be fairly well known. The selection criteria (see Fig. 4 for details) used in this study are as follows: (1) occurrence of sandstone beds thicker than 1 m; (2) information on the architecture and geometry of the TBDM sandstones; (3) at least two well-described time-equivalent detailed sedimentary logs. From an original collection of twenty-five outcrop studies, initially a batch of eighteen appeared promising (Fig. 5). However, a careful analysis left only ten cases. The reason for omitting the other eight cases were: (1) poorly logged sections unsuitable for quantification; (2) severe post-sedimentary deformation prohibiting reliable correlations; (3) existing controversy on interpretation. For examples see Table 1. Several of these cases would appear interesting for more quantitative investigation, such as Tabernas (Spain). All cases provide important information regarding facies and facies associations, the mechanisms of deposition and their environment. However, the small

LITERATURE

I

4,

q, All sandstones beds thinner than 1 m

Sandstone beds thicker than l m Go on Information about environment (channels, interchannels, lobes) exists Go on

~/

,4, No information about bed continuity

Information about bed continuity

Go on

Well described sedimentary logs (SL)

At least 2 time-equivalent SL exist

Reject: Butano, Numidian Flysch Jackford (too controversial)

q,

~/

Go on

Reject: Capistrano, Oquirhh, Torlesse Venado

No information about environment

Poorly described SL

Reject: Tarcau, Tabernas

q, There are not even 2 time-equivalent SL

Annot Beardmore-Geraldton Cengio (PTB) Guipuzcoa-Jaizkibel Marnes Bleues (VB) Marnoso Arenacea Rock Shale Grit Tourelle Tyee Fig. 4. Criteria used for the selection of the case studies.

Reject: Matilija

M. O. Badescu

388 TABLE 1 Overview of outcrops Formation name

Age

Basin setting

Annot (Les Alpes Maritimes, France)

Late Oligocene-Early Eocene

Foreland

100/35

P

550-650/< 10

HDTC

Beardmore-Geraldton (Ontario, Canada)

Archaean

Fore- arc

80/30

M

<70/7-8

GF, TC

Butano (La Honda, California)

Oligocene

Strike-slip

160/120

P

1500-3000/?

SGF

Cengio (Tertiary Piedmont Basin)

Late Eocene-Early Miocene

Post-alpine deformation

60/20

P

170/8-4

TC

Guipuzcoa-Jaizkibel (Pyrenean Basin, Spain)

Early Oligocene

Compressional

50/10

D

800/< 10

Intermediate TC-HDTC

Jackfork (Oachita Basin Arkansas and Oklahoma)

Pennsylvanian

Back-arc

550/450

Les Marnes Bleues (Vocontian Basin, France)

Aptian-Albian

Pull-apart

50/20

M

800/0.5-6

HDTC

Marnoso-Arenacea (N. Apennines, Italy)

Miocene

Foreland

120/20

M

800/0.6-4

SGF

Matilija (Santa Ynez, California)

Oligocene

Fore-arc

~100/~50

M

400/0.1-2

SGF

Numidian Flysch (Riff-Tellian Atlas and S. Apennines, Italy)

Oligocene-Early Miocene

Foreland

100s/? 100

M?

Oquirrh (Utah)

Late Pennsylvanian

9

300/100

9

100/1-2

TC, GF

Rock Sandstone (St. Lucia, California)

Oligocene

Continental borderland

60/40

P

122-610/0.5-7.5 (1.1)

SGF, DF, FSF

Shale Grit (Carboniferous Central Pennin Basin)

Late Carboniferous

9

20/20

M

300/<3.5

HDTC

Tabernas (Betic Cordilleras, Spain)

SerravallianMessinian

Transform margin

28/12

P/L

Tarcau (East Carpathian, Rumania)

Paleocene-Eocene

Foreland

120/20

D

Torlesse (Terranes, New Zealand)

Jurassic-Cretaceous

Subduction complex

200/<5

M

Tourelle (Quebec, Appalachians)

Early Ordovician

Foredeep trough

60/10

M

Tyee (Oregon)

Oligocene

Fore-arc

250/120

M

Size, 1/w (km)

Source a

P vs. M

Thickness Fm./ individual beds (m)

9000-13,500/2-3

1000-3000/1-2

1000/1-6 400-800/< 10 9

Supply mechanism b

SDF vs. HDTC

TC redeposited as "Fluxoturbidites" and slumps

HDTC "Fluxoturbidites" SGF

500-1000/0.1-10 2000/<5

HDTC HDTC

a D = double; L = line; M = multiple; P = point. b DF = debris flows; FSF = fluidised sediment flows; GF = grain flows; HDTC = high-density turbidity currents; SDF = sandy debris flows; SGF = sandy grain flows; TC = turbidity currents.

Outcrop-based classification of thick-bedded, deep-marine sandstones

389

Comments f

References

8/1-1.5 (C)

SL incomplete

Stanley and Bouma (1962), Stanley (1975), Bouma and Coleman (1985), Ravenne and Vially (1987), Sinclair (1994)

L, C (MF)

2 - 3 / 1 - 2 (E)

Metasedimentary deformation; expect changes in thickness due to metamorphism

Barrett and Fralick (1989)

0.7-0.9

C (MF) L

No data

Too general SL, beyond quantification purposes

Nilsen (1985)

S, 0.4

L, C

2.8-4.4/ 0.4-2.4 (C)

Mainly composite SL; do not expect large outcropping

Gelati and Gnaccolini (1980), Cazzola et al. (1981, 1985), Cazzola and Rigazio (1983), Cazzola and Sgavetti (1984)

V, --~0.85-0.6

L, C (MF)

0.5-3/0.4-2 (E)

Expect tectonic to have a big signature on thickness and body communication

Van Vliet (1978, 1982), Rosell et al. (1985)

?V 0.2

Different scenarios. Last interpretation: BFF vs. SDF

1 vs. 100s (E)

Very controversial

Shanmugam and Moiola (1994, 1995, 1997), Slatt et al. (1994a,b, 1997), Lowe (1997), Bouma et al. (1997)

V, 0.5-0.9

C, L

1 - 1 0 / < 9 (C)

2D exposure; emphasis too much on sequence stratigraphy and not on quantification of sand-body geometry

Fries et al. (1984), Rubino (1988), Rubino (1989), Badescu et al. (1998)

S, 0.4

C, L? (OF)

Complete data set

Ricci Lucchi (1975, 1978, 1985), Ricci Lucchi and Valmori (1980)

V, <0.8

C, L? (OF)

I(E)

SL too general; bad reservoir analog, due to early diagenesis

Stauffer (1967), Link and Welton (1982)

V, ? 0.5

C, dykes, sills

1 (E)

Poor information

Wezel (1970), Beaudoin et al. (1987), Parize et al. (1987)

9

9

--~0.1 (E)

Poor information

Jordan ( 1981 ) Link and Nilsen (1980)

S/G c

Facies architecture d

V, 0.8-0.46

C (IF, MF)

V, ,-,0.9

Sand-body communication e (km)

15-120/<12 (C)

V, 0.9-0.33

?C (MF)

4/2 (E)

Good exposures, but only descriptive information on sand-body continuity

V, <0.6

C

1-3085/ 0.7-1 (C)

Moderate exposure

Walker (1966a,b)

V, <0.9

C

1-5/0.2-0.5 (C)

Too general SL, beyond quantification purposes

Kleverlaan (1989), Cronin (1995)

9

C, L

2-3? (E)

Only information on biostratigraphy, SL too general

Contescu (1966), Sandulescu (1989)

V, 0.9

'~

0.2 (E)

Too severe post-depositional tectonics; no correlation

MacKinnon and Howell (1985)

V, 0.5-0.77

C

?/1-2 (C)

No quantification of sand-body extent

Hiscott and Middleton (1979, 1980), Hiscott (1980)

S, 0.8-0.9

Ramp without C

> 2 / > 2 (E)

Descriptive, no numerical indication of sand-body extent

Chan and Dott (1983), Heller and Dickinson (1985)

c S/G = sand/gross; S = stable; V = variable. d BFF = basin floor fan; C = channels; IF = inner fan; L = lobes; MF = mid-tan; OF = outer fan. e c = calculated; E = estimated. f SL = sedimentary logs.

M. O. Badescu

390

Fig. 5. Global overview of TBDM sandstones in outcrops.

number of geological profiles and sedimentological logs makes the quantification of geometry and continuity of the beds difficult. With some exceptions, data on quantification of the geometry and lateral continuity of the thick beds are lacking. The best-described cases are Annot, Cengio, and Marnoso-Arenacea (Table 1). They provide a complete overview of the facies and facies associations, and the geometry and continuity of the thick beds. The other case studies have too much emphasis on the description of TBDM sandstones and their relation to the depositional environment and little emphasis on sandbody geometry and continuity of TBDM sandstones.

Overview of outcrop data of TBDM sandstones

Table 1 shows an overview of the outcrop studies used in text. Classification of outcrop data first requires a comparison of basin-wide parameters. Table 1 and Fig. 6 show that TBDM sandstones can be deposited in a variety of tectonic settings, from foreland basins (Marnoso) to fore-arc (Tyee) and strike-slip systems (Butano). The basins exhibit a variety of shapes and dimensions (see Fig. 6), which complicates even more the problem of quantification of the continuity of TBDM sandstone beds. The main control on the continuity of the beds is due to the shape and size of the basin and the direction of the palaeoflow. In most cases the basin-floor topography is irregular due to tectonics, diapirism, batholith intrusions, or differential compaction.

The outcrop sections pertain to similar stratigraphic packages representing a series of depositional cycles that include numerous thick sandstone beds. The sequence stratigraphic model explaining the process of sand accumulation is probably not the same for all outcrops and in fact the passive margin model described by the Exxon geologists (Vail et al., 1977; Posamentier and Vail, 1988) does not apply on these active tectonic settings. Sedimentary logs from outcrops are shown in Figs. 7 and 8. Fig. 6 locates the sedimentary logs within the basinal context. The following combinations are recognised. (1) Packages of fines (pelagites or hemipelagites), thin-bedded turbidites and TBDM sandstones (e.g., Marnoso-Arenacea and Cengio). Frequently, the TBDM sandstones are underlain by thin-bedded turbidites. (2) Packages of thick (non)graded sandstones overand underlain by thin turbidites or hemipelagites ("channelised packages") (e.g., Annot, Rock, and Shale Grit). (3) Thick blocky packages of non-graded sandstones with interbedded hemipelagites (e.g., Les Marnes Bleues). For each basin two sedimentary logs are shown. However, their spacing differs from 1 km in Les Marnes B leues to 20 km in the Tourelle example. The logs situated in the central, more or less axial parts of the basins are grouped in Fig. 7. The sedimentary logs in Fig. 8 show equivalent stratigraphic intervals but somewhat less close to the axial zones. In general, these logs are less sandy and emphasise the hetero-

Outcrop-based classification of thick-bedded, deep-marine sandstones

391

& ~

Marnoso-Arenacea

Ma 2

Mames Bleues NI a

I- ~ ~

Rock

Guipuzcoa-Jaizkibel TertiaryPiedmontan Basin (Cengio system) To 2

Shale Grit

Tourelle

Annot

Beardmore [] ~

Locationof the sedimentarysections discussed in text.

Q

~

Palaeocurrent direction

~

Exposures Outline of the basin

Scale 1/20,000 0

20

40

60Km.

Fig. 6. Schematic configuration of the outcrop areas of the basin and location of the measured sections discussed in text. For location, see Fig. 5.

geneous nature of the basin fills. Only in the case of the large almost unconfined Mamoso-Arenacea Basin there is evidence (Ricci Lucchi and Valmori, 1980) of continuity of TBDM sandstones over more than 10 kin.

Sedimentology of the TBDM sandstones The visible differences between the outcrop sedimentary logs (Figs. 7 and 8) can be associated with the occurrence of different types of TBDM sandstones on the basis of detailed outcrop descriptions. It is possible to group the various sedimentological

characteristics (Tables 1 and 2) of TBDM sandstones into four sandy facies associations, described below.

Non-channelised, thick-bedded sandstones The following features are characteristics for this type of TBDM sandstones. (1) They are non-channelised. (2) The bed thickness varies from 1 to 10 m. Most of the beds, thicker than 3 m, are amalgamated and pebbly at the amalgamation contact, with flat and continuous erosional surfaces. Their lateral extent varies from 1.5 to tens of kilometres.

GO r ['O

Fig. 7. Measured sections from outcrops in the axial zones of the basins shown in Fig. 6. MA1, C1, S1, A1, R1, G1, Tol, B1, Tyl, Jbl and MB1 represent the sedimentary sections as they are shown in Fig. 6.

Outcrop-based classification of thick-bedded, deep-marine sandstones

393

TABLE 2 Quantification of data Basin

Basin dimensions, length/width/thickness (km/km/m)

Name of the bed

End-member type

Annot (Les Alpes Maritimes)

100/35/600

level 4 level 6 level 8 level 10 level 12 level 14 level 16 level 18 level 20

B, B, B, B, B, B, B, B, B,

spoon-like spoon-like spoon-like spoon-like spoon-like spoon-like spoon-like spoon-like spoon-like

Continuity parallel a (km) 8(c) 13(c) 18(c) 21 (c) 24(c) 28 (c) 33 (c) 38 (c) 43 (c)

Beardmore (Ontario)

80/30/80

LA4 section 5 LA4 section 12 LA4 section 14

B, spoon-like B, spoon-like B, spoon-like

3 (e) 2-3 (e) 3 (e)

Cengio (Tertiary Piedmont Basin)

60/20/170

Vala member Mioglia member (inf.) Cengio, L1 Cengio, L2 Cengio, L3 Cengio, L4 Cengio, L5 Cengio, L6 Cengio, L7 Cengio, L8

C, pillow-like A, sheets C, pillows C, pillows C, pillows C, pillows C, pillows C, pillows A, sheets A, sheets

3.2 (max.) 1.5-1.75 2.8 (c) 3.3 (c) 2.7 (c) 2.7 (c) 2.7 (c) 3.52 (c) 4.4 (c) 4.4 (c)

Guipuzcoa (Pyrenean)

50/10/50

Jaizkibel beds from megacycle 1 A, sheets Guipuzcoan midfan-channelised B, spoons sands, megacycle 2

Les Marnes Bleues (Vocontian Basin)

50/20/800

La megaturbidite (Rosans)

B, spoons

A1 A15 A28 A30 A50 A70 A78 A92 Contessa D21 COL1 COL2 D74 COL3 D98 D139 D161 COL5

A, A, A, A, A, A, A, A, A, A, A, A, A, A, A, A, A, A,

Marnoso-Arenacea (Northern Apennines)

120/20/800

sheets sheets sheets sheets sheets sheets sheets sheets sheets sheets sheets sheets sheets sheets sheets sheets sheets sheets

Rock (St. Lucia)

60/40/max. 600

section A

B, spoons

Shale Grit (Pennin Basin)

20/20/up to 300

channel at Deep-Fagney Clough channel 3 (Grindslow shale) channel at Gathering Hill

B, spoons B, spoons B, spoons

canyon system system 1

B, spoons

channels in FA 1 channels in FA2 channel 1 in FA3 channel 2 in FA3 channel 3 in FA3 channel 4 in FA3 channel 5 in FA3 channel 6 in FA3

B, B, B, B, B, B, B, B,

Tabernas (Betic Cordilleras) Tourelle (Quebec)

a c = calculated; e = estimated.

28/12/up to 1000

spoons spoons spoons spoons spoons spoons spoons spoons

>0.5

17 (max.) 2.75 (min.) 120 (c) 50 (c) 15 (c) 30 (c) 120 (c) 120 (c) 80 (c) 60 (c) 120 (c) 40 (c) 120 (c) 120 (c) 40 (c) 120 (c) 100 (c) 120 (c) 120 (c) 120 (c)

Continuity perpendicular a (km) 1.4 (min.) 1.12 (min. exposed)

l(e) l(e) 1-1.5 (e)

>0.4

>0.8 >1.2 >1.7 >2.4 >0.075 0.05 (e) 11 (max.) 1.2 (rain.) 16 (c) l0 (c) 16 (c) 7 (c) 16 (c) 16 (c) 14 (c) 10 (c) 16 (c) 16 (c) 16 (c) 8 (c) 16 (c) 16 (c) 8 (c) 16 (c) 8 (c) 16 (c) 2 (max.)

3.85 (min.) 1 (min.)

l(c) 5 (c) --~1-2 (e)

0.8 (max.) l(c) 0.74 (c) 1 (c) 0.2-0.5 0.06-0.17 0.06 (min.) 0.23 (min.) 0.23 (min.) 0.26 (min.) 0.26 (min.) 0.12 (min.) 50 (e)

(e) (e) (e) (e) (e)

394

M. O. Badescu

Fig. 8. Correlative measured sections from outcrops lateral of the axial zones of the basins shown in Fig. 6. MA2, C2, $2, A2, R2, G2, To2, B2, Ty2, Jb2 and MB2 represent the sedimentary sections as they are shown in Fig. 6.

(3) The bedding surfaces are generally sharp, even and parallel, although cut-and-fill features may be abundant in amalgamated packages. These features are regarded to be the product of sheet-flow erosion by a single turbidity current. They may influence the variation in vertical permeability of the beds. (4) They are underlain and overlain by basin-plain fines (muddier turbidites, low-density turbidity currents or hemipelagites) or thin-bedded, Bouma-type turbidites. Amalgamated beds are rare in this type of bed succession. (5) The grain size is commonly medium to fine sand, but variation from fine to coarse sand exists. (6) They form thickening-upward or asymmetrical sequences. (7) They are generally massive, but parallel lamination, subtle grading and dish structures occur. Outsized mudstone clasts may occur at the top of the beds. (8) Amalgamation is common. Detailed data on the nature of amalgamation surfaces are not available. However, such surfaces are considered to be important for the vertical permeability of the sand bodies. The facies characteristics of the non-channelised thick-bedded sandstones indicate that they were deposited in outer fan environments, by waning, sandloaded gravity flows. The poor development of scour marks and the lack of the structures caused by trac-

tion (i.e., flute casts, groove marks, etc.) suggest that these deposits were the results of a transition between liquefied flows and high-density turbidity currents (see for example Marnoso-Arenacea, Mutti, 1985). Examples for this type of facies are the lobe units from Cengio (Tertiary Piedmont Basin) and the non-channelised thick-bedded sandstones from Marnoso-Arenacea. Channelised thick-bedded sandstones

These thick-bedded sandstones present the following characteristics. (1) The beds are channelised, with little lateral extent (individual channel width from 0.4 km to 1.5 km). (2) The bed thickness ranges from 1 to 10 m, and the beds are commonly amalgamated or stacked, forming channel-fill successions. (3) Their base is commonly sharp or channelised while the top is flat and can be channelised. (4) Stacked thick-bedded sandstones are most common and stacked beds are underlain and overlain by thin-bedded turbidites or fines. (5) Grain size varies from fine to coarse sand; normal- or inverse-graded pebbles and cobbles are common in the channel axis. Rip-up clasts occur in the lower part of the beds.

Outcrop-based classification of thick-bedded, deep-marine sandstones (6) Sandstone beds are commonly arranged in fining-upward and thinning-upward sequences. (7) Sedimentary structures are not very common; graded bedding, parallel lamination, groove casts, flute casts, load casts, horizontal lamination, current ripples, dish structures, and trace fossils are occasionally observed. (8) Amalgamated contacts are common. Two or more separate beds can be traced laterally into a single bed of sandstone. The amalgamated junctions may form permeability baffles and are expressed in different ways: (a) sharp, bed-parallel, with small load casts; (b) a row of angular or rounded mud flakes; (c) a change in grain size; (d) no grain size trend discernible. This facies type is interpreted to be deposited by various types of sediment gravity flows, including turbidity currents (mainly), grain flows, and fluidised sediment flows (Walker, 1966a,b; Link and Nilsen, 1980; Hiscott, 1980; Kleverlaan, 1989). Examples are bed units from Annot, Beardmore-Geraldton, Guipuzcoa, Jaizkibel, Matilija, Rock, Shale Grit, Tabernas, and Tourelle.

Slurry sandstones This type was recognised only in the Tourelle Formation (Hiscott, 1980) and is characterised by the following features. (1) They occur together with the channelised sands. (2) The bed thickness varies from 0.5 m to 2 m. (3) The bases are flat and sharp and the tops are graded (into shale). (4) They are underlain and overlain by shales. (5) The grain size is fine to coarse. The sand is dispersed within a mud matrix. (6) No obvious arrangement in sequences is observed. (7) Sedimentary structures include convolute bedding and locally developed pseudonodules. Rip-up clasts up to 1 m long and shale chips with random orientation are frequent. (8) Amalgamation is common. Slurry sandstones are interpreted to be the product of slumping of mud and unconsolidated sand. The absence of scouting at the base of the beds, the absence of the internal structures caused by the traction, the abundance of matrix and large fragments and the random appearance of the shale chips give reason to interpret them as the result of submarine debris flows formed by the slumping of mud and unconsolidated sand on a submarine slope (Hiscott and Middleton, 1979).

395

Thick-bedded sandstones confined by basin-floor topography The following features are characteristic for this type of beds. (1) The sandstones are channelised or non-channelised, but their confined aspect is not due to the channelling but to the basin topography. (2) The bed thickness ranges from 1 to 3 m. (3) The beds have either undulating, erosive bed boundaries or flat, sharp, non-erosive bed boundaries. (4) They form stacked sand-on-sand successions. Intercalated fines are eroded away by subsequent turbidity currents. (5) The grain size varies from very fine to coarse sand. (6) Sandstone beds together form blocky log patterns, rather than fining- or coarsening-upward sequences. (7) Sedimentary structures include dish structures and parallel and ripple lamination, but the general appearance is massive. (8) Amalgamation is characterised by a sudden change in grain size. These sandstone beds were deposited by highdensity turbidity currents in topographically confined basins. Examples of this facies are some lobes or channelised units from the Tertiary Piedmont Basin, units from the Vocontian Basin (Les Marnes Bleues), and Jaizkibel successions from the Guipuzcoa Basin.

Sand-to-gross ratio (S/G) Complexes of TBDM sands and associated facies present different S/G ratios (Fig. 9). S/G ratio varies in three styles. (A) Low (0.4-0.5) but stable S/G ratio (e.g., Marnoso-Arenacea, Cengio). This is likely to be the result of the large extent of the individual beds. (B) Highly variable (0.3-0.9) S/G ratio across the palaeocurrent direction and stable S/G ratio along the palaeotransport direction (e.g., Annot, Rock, Tourelle, Beardmore-Geraldton). This is the result of the presence of channels running more or less in the same direction. (C) Highly variable (0.5-0.9) S/G ratio both across and along the palaeocurrent direction (e.g., Les Marnes Bleues). These are partly associated with the pillow-type sandstones related to confined basins.

Bed continuity and shapes of the bodies Table 2 presents quantification on the continuity of the sandstone beds derived from ten outcrop studies. In some cases (Annot and Shale Grit), the bed by

M. O. Badescu

396

~.4

)-0.3 Fig. 9. Complexes of TBDM sandstones showing little thickness variation on a 10 km scale.

bed quantification was impossible, so that packets of sandstone beds were quantified. Out of these ten outcrop examples, nine are the same as in Fig. 6. The Tyee case was deleted because of lack of correlation data both as bed by bed and as sandstone packets. On the other hand, Tabernas was added because some sandstone beds correlations are available for that site. The absolute dimensions (length and width) of the beds directly depend on the basin size and shape and the direction of the palaeoflow. The more exact shape and lateral continuity of the sand bodies are strongly influenced by the particular tectonic style. Reading and Richards (1994) emphasised the complexity of the factors that control the architecture and the final shape of the deep-marine sandstone beds. A major problem is that many factors (the morphology of the source area, climate, shelf system, shape and type, variation in clastic input, etc.) cannot be recognised directly in the sedimentary facies and the geometry and shape of the beds. A new approach could be to compare the volume of the sandstone beds to the volume of the basin and relate this ratio to the tectonic style, character of the sources of the sand and S/G ratio. The basin-floor topography has a significant control on the geometry and on the shape of the thickbedded sandstones. Ricci Lucchi et al. (1985b) emphasized the influence of the basin-floor topography on the geometry of the thick-bedded sands deposited in small, tectonically controlled basins (lower units of the Cengio fan). Firstly, the basin-floor topogra-

phy influences the turbidity currents flow trajectory and beds with various shapes can be deposited. Secondly, sediments can be trapped in topographically controlled depressions, forming pillow-like, laterally discontinuous beds. It was also recognised in the subsurface that basin-floor topography caused by tectonic half-graben geometry (Ravnas and Steel, 1998) and by salt diapirism (Weimer et al., 1998) significantly influences the final geometry of the sandstone beds. Three basic types of turbidite system geometries were recognised in the outcrops. (1) Sheet-like geometries are characteristic for basin plain deposits. The beds are continuous on a basin scale. Non-channelised, thick-bedded sands exhibit a tabular shape. The thickness variation is insignificant, both across and along the palaeocurrent direction. (2) Ellipsoidal geometries are characteristic for channelised thick-bedded sandstone beds. Good continuity is expected downcurrent and poor continuity across the palaeocurrent direction. The thickness variation is significant across the palaeocurrent direction. The thickness variation is much lower downcurrent (see Table 2 for details on quantifications). Particularly for this type, the dynamics of the distributary sediment flow form an important factor that controls the geometry and the continuity of the sandstone beds. (3) Pillow-like TBDM sandstones, which are the third geometrical type of thick-bedded sandstones, were identified in few cases (some lobes or channelised units in the Piedmontan Tertiary Basin study, Les

Outcrop-based classification of thick-bedded, deep-marine sandstones Marnes Bleues, and Guipuzcoa-Jaizkibel). Here the basin-floor topography plays an important role on the geometry and on the lateral extent of the thick-bedded sandstones. Pillow-like, thick-bedded sandstones are characteristic for TBDM sandstones deposited in confined basins. A significant thickness variation is observed both along and across the palaeocurrent.

Main generic and geometrical types of the TBDM sandstones found in the outcrop studies The above-described geometries and sedimentological features of the TBDM sandstones recognised in the outcrops can be grouped in three main generic and geometrical types (Fig. 10). Type A represents continuous, tabular thick sandstone beds. The minimum distance of continuity is dependent on the basin shape and size and it varies from a few kilometres to tens of kilometres (depending on the basin dimensions). Bed-thickness gradient is low. Type A thick-bedded deep-marine sandstones represent the sandy facies association of non-channelised, thick-bedded sandstones. They were deposited in unconfined basins, or basins with little obstruction to flow, which is the reason why they correlate over long distances. However, examples (Marnoso-Arenacea) show that successive beds tend to compensate the topography created by the deposition of a previous bed.

397

Type B represents thick ellipsoidal sandstone beds. The minimum distance of continuity is dependent on basin shapes and sizes and the palaeocurrent direction. Good continuity on the order of tens of kilometres is expected along the palaeocurrent direction, and poor continuity on the order of few kilometres (depending on the basin size) is expected across the palaeocurrent direction. Bed-thickness variation is low along the palaeocurrent, and high across the palaeocurrent. The type B thick-bedded deep-marine sandstones represent channelised thick sandstones, deposited either in open or in confined basins. The slurry sandstones occur together with the channelised thick-bedded sandstones. Their poor lateral correlation is due to the geometry of the channels. Type C represents thick pillow sandstone beds. The minimum distance of continuity is defined by the basin topography and the basin size and shape. Beds are continuous on a range from hundred metres to a few kilometres, both along and across the palaeocurrent direction. The bed-thickness gradient varies significantly. Type C beds represent thick-bedded sandstones that were deposited in topographically confined basins. This is the reason why they correlate over short distances. However, the dimensions of the depressions that cause confinement can also vary, depending on the causes that generated the highs and lows on the basin topography. Causes of confinement could be: (a) graben/half-graben faulting (e.g., Cengio, Cazzola

Type A TBDM sands non-channelized

-

- tabular

shape

- good continuity

Type C TBDM sandstones channelized or n o t

-

- pillow-like

Type B TBDM sandstones channelized

- poor

shape

continuity in all directions

-

- spoon -

shaped

good continuity parallel to t h e p a l e o c u r r e n t poor continuity l a t e r a l t h e p a l e o c u r r e n t Fig. 10. Main generic and geometrictypes of TBDM sandstonebeds.

398 TABLE 3 Variation of types of thick-bedded sandstones in outcrop

et al., 1981); (b) salt or mud diapirism (e.g., Jaizkibel, Van Vliet, 1982); (c) differential compaction (e.g., Mames Bleues, Fries et al., 1984); (d) synsedimentary folding; (e) batholith anticlines. It is important to underline that more than one type can occur within the same basin (see Table 3). Hybrid types that are gradational between these three main types are the rule rather than the exception. This is mainly because: (1) factors controlling the architecture and the geometry of the TBDM sandstones (tectonics, source area, sea-level variations, and pre-existing topography) are variable and interrelated; (2) their relative importance can change in time and space; (3) TBDM sandstones are deposited from a variety of processes that give specific signature to the deposits. The most common type, recognised in this study, is type B: ellipsoidal, channelised thick-bedded sands. For this type of sandstone it is critical to know the direction of the palaeocurrent and how the basinfloor topography influences deposition, in order to predict the geometry and architecture of these sandstone beds. Good-quality 3-D seismics is probably the best method to delineate the shape of the individual channel-fill sandstones.

Linking prospects to outcrop analogues Although the outcrop studies are the only source of reliable information on the interwell scale, it is very difficult to identify an exact outcrop model for a given reservoir field. At present, there is not yet a clear methodology describing how deep-marine outcrop analogues can be used for the quantification

M. O. Badescu

of reservoir properties. A major problem is that the data from outcrops and subsurface reservoirs are in different formats. Transfer functions are used to convert outcrop data to subsurface data formats. These include synthetic seismic profiles, pseudo-logs and vertical sedimentological diagnostics. At this time, this study addresses only sedimentological diagnostics that can be recognised in cores. The main concern for using outcrop analogues for dimension predictions is that the lateral continuity of TBDM sandstones depends on the tectonic setting, basin shape, and its dimensions. However, a comparison of outcrop data with a number of reservoirs reveals that similarities in shape and facies characteristics occur. The classification of TBDM sands can be applied to the Tertiary of the North Sea. Fig. 11 illustrates the Paleocene-Eocene North Sea reservoirs that will be considered.

The Tertiary of the North Sea - - comparison of the Paleocene and Eocene reservoirs (source: Parker, 1994; Bowman, 1998) During the Tertiary, the North Sea sedimentation was controlled by a complex interplay between tectonic activity, eustasy, and hinterland characteristics. The major control on the siliciclastic supply during the Tertiary was the rifting of the GreenlandEuropean Plate in the Early Paleocene, with rejuvenation of older Mesozoic hinterlands and basin margins (Bowman, 1998). The volume and grain size of the clastic detritus feeding the major submarine fans were not constant in time. During the Paleocene, the volume increased gradually to its peak (midThanetian). The tectonic uplift of the hinterland was the major factor controlling sedimentation during the Paleocene. However, the impact of tectonic activity was not uniform. Differential uplift caused the development of geographically and temporally separate depocentres (Morton et al., 1993). The Paleocene is very sand-rich, involving stacked pay zones, within the distal part of the main submarine fan complexes. Type A and B TBDM sandstones are likely to occur. The Eocene is characterised by reduced rates of clastic input along the newly developed passive margins. Relative sea-level changes were the primary control on deposition. The Eocene is largely mudprone and contains localised submarine fans. The sandstones are typically clean and well sorted, and to some degree different from the more clay-rich, delta-fed systems of the Paleocene. The changes are a consequence of reworking sands in shoreline and upper shelf settings near the submarine canyon heads.

Outcrop-based classification of thick-bedded, deep-marine sandstones

399

Fig. 11. Paleocene and Eocene discoveries of the North Sea (after Bowman, 1998). (A) Paleocene. (B) Eocene. Reprinted by permission of the Geological SocietyPublishing House.

The Tertiary submarine fans of the North Sea are divisible into two main categories, with different characteristics and production behaviour. (1) The Early Paleocene fans of the Central Graben are large-scale basin-fill complexes. In the main depocentres the sand bodies have sheet-like geometries (type A). Outside these areas, the architectural patterns are dominated by channel fills (type B) with overbank deposits. The complexes of TBDM sandstones present higher S/G ratios along the main axis of flow and lower S/G ratios perpendicular to the flow. The sands are more clay-rich, which influences the reservoir permeability. Typically for the Early Paleocene reservoirs, the TBDM sandstones are interbedded with very continuous shales. This has a significant influence on the vertical permeability, compartmentalising the reservoirs. Notably, the direction of the palaeoflow is critical to achieve a good appraisal of these types of reservoirs. (2) The Late Paleocene-Eocene fans tend to be localised and smaller in size. The basin topography is the factor that defines the sediment architecture and geometry. The internal architecture is characterised by pod-like (type C) or elongate bodies. The complexes of TBDM sandstones have steep, abrupt margins and comprise homogeneous masses of clean,

well-sorted sand, affected by slumping and liquefaction. This modifies the primary depositional geometry. In contrast with the Paleocene reservoirs, the Eocene reservoirs contain cleaner sands, with higher vertical permeability. Although the sandstone beds are very thick in some wells they are thinner elsewhere. The geometry of the TBDM sandstones is complex and therefore very difficult to predict. However, good-quality 3-D seismic can help with this problem. The TBDM sandstones bodies contain very few shale intercalations that could act as barrier for fluid flow. Table 4 shows an overview of the Tertiary reservoirs of the North Sea. Type A (sheet-like) and B (channels) sandstones occur interbedded with thick, laterally extensive shales that generally compartmentalise the reservoir. Reservoirs containing type A sandstones, interbedded with laterally continuous shales should receive special attention when planning water injection. Generally, these reservoirs present a very active aquifer and there is no need to inject water in the early stages of production. Examples such as Forties (Carman and Young, 1981; Wills and Peattie, 1990), Frigg (Mure, 1987a; Brewster, 1991), and Heimdal (Mure, 1987b; Grinde et al., 1994) show that water breakthrough

O O TABLE 4 Overview of the Tertiary reservoirs of the North Sea Aquifer e

Sand-body configuration f

References

33/2800

Not available

Topographically confined HDTC (type C)

Newton and Flanagan (1993), Harding et al. (1990)

Post-plateau EW (1971), 8 AW (1971-1976); 5 platforms with 48 PW (1976-1979), production since 1977

29/1500

Efficient, not anticipated, 3 WI (1984-1985)

Massive channels (type B) Levees and lobes (type A)

Mure (1987c), Brewster (1991)

Production (since 1988) EW (1974) + 2 AW (1984) 2 platforms

28.5/1000

Efficient, communication with Frigg

Channelised sediments Deposited by GF and TC (type B TBDM)

Mure (1987d)

Northeast Frigg (VG) Ypresian (T70)

Production (since 1983) EW (1974), 1 platform with 6 PW

28/1000

Efficient, communication with Frigg

GF and TC deposited in a channelised environment (type B TBDM)

Mure (1987c)

Odin (VG)

Ypresian (T70)

Production (since 1984) 2 EW (1973, 1975) Fixed platform with 11 PW

29.5/1000

Limited

Channelised sands ("fluxo-turbidites"); type B and HDTC non-channelised, type A

Nordgard Bolas (1987)

Gannet (CG)

Thanetian (T40) and Ypresian (T70)

Production, no information on field development

38/650-2800

Not available

T40: channelised (type B) T70: salt-confined, channels (type B TBDM) Lobes (extensive) (type A)

Armstrong et al. (1987)

Gryphon (VG)

early Ypresian (T50)

Development drilling EW (1987), 13 AW + ST

not available

Not available

Stacked massive sands,topographically confined (type C TBDM)

Newman et al. (1993), Timbrell (1993)

Guillemot D (CG)

Thanetian (T40 § T50)

Production EW (1969), EW (1988), 2 AW (1988)

not available

Not available

T40: sheet-like sand bodies (A) T50: shoestring channels (B)

Banner et al. (1992)

Balder (VG)

Thanetian (mainly)

Production

Active, no WI

Mounded sand bodies, with rapid pinch-outs (type C)

Hanslien (1987), Jenssen et al. (1993)

Forties (WGG, CG)

Thanetian (T40)

Mature, post-plateau 17EW + AW EW (1970), 4 AW (1971-1972) 4 fixed platforms (1975) + 5th platform (1985) WI (since 1976), GI

Very active, WI too early

HDTC deposits: braided channels, sheet-like geometry, topographically confined (types A and B TBDM)

Wills and Peattie (1990), Carman and Young (1981)

Reservoir a

Age b

Status/field development c

Alba (WGG)

Lutetian (T92-T98)

Early production (?1998) 16 AW + ST

Frigg (VG)

Ypresian (T70)

East Frigg (VG)

Ypresian (T70)

Porosity (%) / permeability (mD) d

33 (32-35) 1000-3000 27 (10-36) 700 (30-4000)

r~

t..., e5

|

TABLE 4 (continued) Reservoir

a

Nelson (WGG, CG)

Age b

Status/field development c

Thanetian (T40)

Early production (1994) 6 EW (196%1987) 6 AW, 1 platform jacket with 24 PW (including 7 WI)

Porosity (%)/ permeability (mD) d 20-25 100-300

Aquifer e

Sand-body configuration f

References

e5

Small, requires early WI

Channelised HDTC (type B)

Whyatt et al. (1991 ), Griffin et al. (1994)

t..., t..~.

Arbroath (CG)

Thanetian (T40)

Early production Satellite platform 6 PW + 4 WI

24 (3-30) 80 (1-2000)

Small

Stacked channelised sands with little lateral extent (type B)

Crawford et al. (1991)

Montrose (CG)

Thanetian (T40)

Post-plateau production 1 platform, 15 PW/6 WI

24 (3-30) 80 (1-2000)

Small

Stacked channelised sands with little lateral extent (type B)

Crawford et al. (1991 )

Cod (CG)

Thanetian (T40)

Production 3 EW + 9 DW (from which 7 are PW), 1 platform

17 (15-24)

Not available

Lenticular, channelised bodies (type B)

D'Heur (1987), Kessler et al. (1980)

Production (since 1986) DW (1975), 9 AW (1979-1983) floating production vessel 13 PW + 6 WI

25 (20-30) 20-3300

Small, early WI

Stacked lobes, sheet-like (type A) and channels (type B)

Tonkin and Fraser (1991 )

Balmoral (WGG)

early Thanetian (T30)

Cyrus (WGG, CG, VG)

early Thanetian (T30)

13 AW, 2 horizontal wells tied back from vessel

Heimdal (VG)

early Thanetian (T30)

Post-plateau production (1986) DW (1972), 2 AW (1975, 1981) 1 jacket platform with 10 PW

Sleipner East (VG)

early Thanetian (T20 + T30)

Production (1993, planned) DW (1981), 3 AW, 1 platform

Maureen (WGG)

Danian (T20)

Production (1983)" WI, GI

15 (15-100)

e~

t....,

~20 -~200 24-25 --~1 0 0 0 ~26 200-1500 Not available

Efficient

Sheet-like sandstones with lateral extent beyond the field (type A)

Mound et al. (1991)

Strong

Channelised sands (type B) and sheetlike sands in pressure communication

Mure (1987d), Grinde et al. (1994)

In pressure communication with Heimdal

Massive channels (type B)

Ostvedt (1987)

Small

3 thick, amalgamated lobes, with abrupt lateral pinch-out (type C)

Cutts (1991)

a CG = Central Graben; VG = Viking Graben; WGG - Witch Ground Graben. b T20-T98 = stratigraphic sequences proposed by Bowman (1998). c AW -- appraisal wells; WE = exploration wells; PW = production wells; ST = side tracks; WI - water injectors; GI = gas injectors. d Average/(min.-max.). e WI = water injection. f H D T M = high-density turbidity currents; GF = grain flows; TC = turbidity currents; TBDM = thick-bedded deep-marine sands. 4x O

M. O. Badescu

402

occurred due to too early water injection and pockets with significant volumes of remaining hydrocarbons still exist. Reservoirs containing only type B (channels) TBDM sandstones, such as Nelson (Whyatt et al., 1991; Griffin et al., 1994), Arbroath (Crawford et al., 1991), Montrose (Crawford et al., 1991) and Cod (Kessler et al., 1980; D' Heur, 1987), present a complex geometry, and reservoir correlation is problematical. Their aquifer support is generally reduced and early water injection, eventually combined'with gas lift, is required. Reservoirs containing type C (pillows) TBDM sandstones are very problematical for the prediction of their lateral extent. They are generally small, localised, but present excellent, homogeneous internal porosity-permeability characteristics (for example: Alba, Harding et al., 1990; Newton and Flanagan, 1993; Balder, Hanslien, 1987; Gryphon, Newman et al., 1993; Timbrell, 1993). Type C sandstones are generally very clean, well sorted, with vertical permeabilities higher than 1 D. Their lateral extent is difficult to predict. Data on aquifer properties are not available, but it is likely that the aquifer has little lateral extent (sand bodies are confined by the topography). When economically feasible, water injection and artificial lift are required in the early stages, for maximal production in the first years.

Conclusions TBDM sandstones can be described by subdividing the sandstone beds into three main types: sheets, channels and pillows. Although it is understood that in detail one can distinguish between a large number of sand-body shapes, it is practical to simplify this variety into three main types, because this relates the reservoir characteristics to the expected reservoir behaviour. For the appraisal of an oil reservoir comprising mainly thick-bedded deep-marine sands one should bear in mind that: (1) thick-bedded deep-marine sandstones represent different architectural types, with specific spatial signatures; (2) on a basin scale, more than one type of thick-bedded sandstones can occur; (3) to predict reservoir architecture, it is critical to know the palaeoflow direction; (4) on a small scale, the basin-floor topography is the factor that mostly influences facies architecture of TBDM sandstones; (5) the three types TBDM sandstone beds can be differentiated in cores form their few diagnostic features, such as structures, occurrence and types of

the mudclasts, bed boundaries and facies associations (shale types); (6) for future planning it is necessary to define the palaeoflow direction and the interplay between synsedimentary topography and deposition.

Acknowledgements The work presented in this paper is part of Ph.D. research carried out at the Delft University of Technology, Holland. Thanks are due to K.J. Weber and M.E. Donselaar for continuous support and for valuable comments on earlier versions of the manuscript. I.M. Voiculescu is acknowledged for drawing Fig. 10 and for her warm support. The comments of A.H. Bouma and J.R Nystuen improved the final manuscript and are greatly appreciated. Discussions with C. Puigdefabregas and T. Lien were very fruitful and are acknowledged.

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Delft University of Technology, Faculty of Applied Earth Sciences, P.O. Box 5028, 2600 GA Delft, The Netherlands

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Use of integrated 3D seismic technology and sedimentology core analysis to resolve the sedimentary architecture of the Paleocene s u c c e s s i o n of the North Sea M. Cecchi. C. Guargena, L. Hansen, D. Rhodes and A. Roberts

The integration of extensive 3D seismic data sets with core sedimentology analysis from the Siri and Heimdal areas indicates that the stratigraphic architecture of the Paleocene succession of the North Sea is highly variable and fractal in nature. Moreover, we propose that, at least in some parts of this area, debris flows, slides and strong current reworking were the prevailing depositional processes operating during the Paleocene, in addition to the more commonly cited turbiditic ones. This interpretation helps to explain the occurrence of massive sands associated with slumps and olistostromes. Future wildcat and development projects associated with Paleocene plays must consider this variability of reservoirs and seals.

Introduction

About 30 years have elapsed since hydrocarbons were first discovered in Paleocene sediments of the North Sea, these sediments proving to be billion barrel reservoirs. More recent discoveries in the 1990s (Siri, Bittern, Jotun, Grane) have revitalised the interest for this play and have indicated that the Paleocene of the North Sea still represents an underexplored target. This sluggishness in exploration and exploitation may also be related to the understanding of the geological model, as perhaps also indicated by older Paleocene discoveries that still await to be developed. Traditionally, 2D seismic, wireline logs and core sedimentology have been the classical data sets and tools used to interpret the sedimentary architecture of the Paleocene succession. North Sea Paleocene sands have usually been interpreted to be deposited by deep-water fan turbidity currents (Newman et al., 1993; Jenssen et al., 1993; Newton and Flanagan, 1993; Timbrell, 1993). An alternative interpretation of North Sea Paleocene examples (debris flows and slumps) was given by Shanmugam et al. (1994). The aim of this paper is to firstly show that the sedimentary architecture of the Paleocene of the North Sea is different from area to area and can not be forced into fixed models. Also, for the Paleocene of the North Sea, facies distribution can be described as essentially fractal in nature. Secondly, we present evidence that many Paleocene sands could have been

deposited and reworked by mechanisms other than the more conventional "turbiditic" ones. These two points have a profound impact on our understanding of facies distribution. In this paper we show how the compilation of attribute maps from regional 3D seismic data sets, combined with core observations, may provide a step forward in the interpretation of the stratigraphic and sedimentary architecture of the Paleocene succession of this area. This approach has been used by oil companies during the last years, but, to date, the number of recent published papers on the subject is quite limited. We believe that this approach is important for future exploration and exploitation projects dealing with this sedimentary sequence. Two case histories are presented, taken from a database of fifteen Enterprise North Sea fields and discoveries (Fig. 1). Our interpretation challenges the classical "turbidite deep basin fan" interpretation, which we believe to represent more of a mind-set than an objective interpretation. For reasons of confidentiality, most of the data are reported in a general way. Preliminary results from ongoing petrographic studies are also presented. Siri area

Linear and S-shape features from 3D seismic The isochore of the Lower Paleocene (Top Chalk to Top Lista Seismic Markers) show a striking o c -

Sedimentary Environments Offshore Norway - Palaeozoic to Recent edited by O.J. Martinsen and T. Dreyer. NPF Special Publication 10, pp. 407-419, Published by Elsevier Science B.V., Amsterdam. 9 Norwegian Petroleum Society (NPF), 2001.

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Fig. 1. Study area and Enterprise's database superimposed over a regional isopach of the Paleocene succession. Warm colours indicate a thicker section. The Siri Field in the Central Graben and one South Viking Graben example (Heimdal area) are covered in this paper. The asterisks refer to the location of the Norfolk and Flemish ridges cited in the text.

Fig. 2. The 3D isochore of the Lower Paleocene Lista Formation in the Siri area shows rectilinear, parallel and S-shape geometries. These correspond to narrow mounds on seismic sections (arrows), here interpreted to represent sedimentary ridges resulting from strong current reworking. A fan geometry is visible to the south west.

currence of linear and parallel to S-shape, N E SW-oriented features in most of this area (Fig. 2). Length of these bodies is 30-40 km. These features correspond to narrow but prominent mounds on seismic sections (ca. 40 ms TWT, i.e. ca. 50 m thick). Towards the southeast, the overall pattern changes to a fan-shape geometry. This seismic package is

contained within the so-called "Siri fairway", a depression within the underlying Cretaceous-Paleocene chalk. The origin of the fairway is still a matter of debate, and there is no apparent link to the underlying structural grain of the chalk. Interestingly, the fairway is aligned with the present day Skagerrak to the east.

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TABLE 1 Paleocene sedimentary facies of cores from the Siri area Facies

Sedimentary structures

Interpretation

Relative abundance

1. Fine-grained, massive sands

Sharp lower and upper contacts No/reverse grading Dish structures Planar fabric Rafted/floating clasts Parallel and trough cross-lamination

Sands transported by debris flow mechanisms and reworked by strong submarine (bottom?/tidal?) currents

>80% (up to 50% current reworked)

2. Brecciated sands and shales

Sharp lower and upper contacts Brecciated appearance Slumps Rafted/floating clasts

Sands and shales transported by debris flow mechanisms with higher percentage of fines; probably ignited by submarine slides; associated to slumps; could be classified as olistostromes AA

< 10%

3. Shales with rippled sandy intervals

Truncated ripples, laminated shales

Sands and shales reworked by submarine currents

< 10%

4. Graded sands

Sharp lower contact and gradual upper contact Fining upward, normal grading

Sands transported by turbidite flows and deposited by settling

<10%

5. Laminated shales

Parallel lamination

Pelagic deposits

< 10%

Massive sands and current reworking

Core observations (250 m of cores) indicate the presence of the following sedimentary facies (Table 1): (1) fine-grained, massive sands with occasional cross-bedding; (2) shales with abundant ripple sands; (3) brecciated shales and sands; (4) sands with graded bedding; (5) shales (Figs. 3-5). We interpret as turbidites those deposits characterised by the Bouma sequence (Bouma, 1962) and

conform to the original definition of Dott (1967) and Middleton and Hampton (1973), i.e. sands transported by a turbulent flow and deposited by settling. According to Shanmugam and Moiola (1995) and Shanmugam (1996a,b) the terms "sandy or muddy debris flow deposits" should be used instead of "high-density turbidites" (Lowe, 1982; Kneller, 1995; Kneller and Branney, 1995) to describe deposits with the sedimentary features as detailed in Tables 1 and 2 below (Facies 1 and Facies 2). Without entering in

Fig. 3. Paleocene facies in well cores from the Siri area, Central Graben. (A) massive, fine-grained sandstones with cross-lamination (red arrows) are interpreted in this paper as sandy debris flows (SDF), reworked by strong submarine currents. These sands alternate with brecciated shales and siltstones here interpreted as muddy debris flows (MDF). Note rafted and floating clasts (RC, FC) in the sands. (B) same facies as in (A), plus normal graded layers, interpreted as turbidites (TUR, yellow arrows indicate normal grading). Note the sharp upper and lower contacts of SDF and MDF (black arrow), and the presence of bed chunks (BC) within MDF. Vertical scale 1 m.

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Fig. 4. Paleocene facies in well cores from the Siri area, Central Graben. (A) High-angle truncated ripples (red arrows) are interpreted in this paper as the results of submarine currents, comparable to those described by Shanmugam et al. (1993) from the Pliocene-Pleistocene of the Gulf of Mexico (CRS: current reworked sediments). (B) Sandy debris flow (SDF) and muddy debris flow (MDF) facies as in Fig. 3. Note trough cross-lamination (black arrow), interpreted as submarine current reworking. Also note rafted clasts (RC) in SDF and MDF. Vertical scale 1 m.

TABLE 2 Paleocene sedimentary facies of cores from the Heimdal area Facies

Sedimentary structures

Interpretation

Relative abundance

1. Massive sands

Sharp lower and upper contacts No/reverse grading Dish structures Planar fabric Rafted/floating clasts

Sands transported by debris flow mechanisms and reworked by submarine currents

> 50-60%

2. Brecciated sands and shales

Sharp lower and upper contacts Brecciated appearance Slumps Rafted/floating clasts "sheared matrix" Boudins Sand injections

Sands and shales transported by debris flow mechanisms with higher percentage of fines; probably triggered by submarine slides; associated to slumps; could be classified as olistostromes AA

3. Graded sands

Sharp lower contact and gradual upper contact Fining upward, normal grading Truncated ripples

Sands transported by turbidite flows, deposited by settling and reworked by submarine currents

4. Shales with rippled sandy intervals

Truncated ripples Laminated shales

Sands and shales reworked by submarine currents

<5%

5. Laminated shales

Parallel lamination

Pelagic deposits

<5%

the celebrated ongoing dispute (Slatt et al., 1997; Shanmugam and Moiola, 1997), we would stress that our association offacies indicates depositional mechanisms characterised by plastic, laminar flow and final en-masse deposition, e.g. debris-flows. We realise that it is not possible to discriminate between the transport of a "high-density turbidite" and a "debris flow" on the basis of the sedimentary facies alone,

20-30%

5-10%

since these reflect what happened in the last moment before the mass flow "froze". However, if the massive sand facies were deposited from "high-density turbidites", they should exhibit characteristics of fluid flow in which sediments are supported by fluid turbulence and, again, deposited by settling. However, a very low percentage of the cored intervals in this study show these "turbidite sensu stricto" facies. A1-

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Fig. 5. Paleocene facies in well cores from the Siri area, Central Graben. Note rafted clasts (RC) and reverse grading within a sandy debris flow layer (SDF), with sharp upper and lower contacts. Also note the brecciated facies of the muddy debris flow layers (MDF), with abundant bed chunks. MDF facies may be interpreted as olistostromes. Vertical scale 1 m.

though post-depositional processes, e.g. de-watering and liquefaction, can frequently destroy sedimentary features, it is hard to imagine that these processes would operate selectively on the turbidite facies only. Muddy debris flow facies could also be classified as olistostromes sensu Abbate et al. (1981a,b). We rule out the possibility that our Facies 2 (Table 1) represents injected sands in the Siri area for the following reasons: (1) detailed well-log correlations that define sequential depositional patterns and (2) lack of seismic evidence. Injected sands are common in the Paleocene of the North Sea and will be described in the next example.

Petrographic results Preliminary results of petrographic analysis (inhouse data) indicate that almost all the sandstones are very fine- to fine-grained, with quartz as the main component, with minor feldspar, mica and glauconite (Fig. 6). Grain size distribution is unimodal. Clay minerals include chlorite, kaolinite(?) and smectite (Fig. 7). Other preliminary results from S m - N d analysis (inhouse data) indicate that the mudstones and sandstones show similar provenance ages and, consequently, may have been derived from the same source area.

Interpretation 3D seismic indicates that Lower Paleocene sediments in the Sift area are organised into rectilinear to S-shape, parallel ridges up to 40 to 50 m in height and 30-40 km in length (Fig. 2). In places, these ridges are so straight and parallel that it is inconceivable that they could have been formed by gravity processes alone. Sedimentological core analyses indicate that the bulk of these sediments consists of gravity flow deposits, modified by current reworking on a large scale (seismic) and evidence on a small scale (cores) as our Facies 3 and (partly) Facies 1 (Fig. 4). Our interpretation is (Fig. 8) that strong currents and/or storms, or earthquake events, caused the dumping of large amounts of sands into the Siri fairway. These sands were then reworked by strong offshore (tidal?) currents running more or less parallel to the fairway, to give the sand bodies their rectilinear to S-shape forms. The literature on current reworked deep marine sediments is not very extensive and these sediments are, in a sense, not well known. Our interpretation relies upon our observations and the facies associations proposed by Mutti, 1992 (his fig. 10) and Shanmugam et al., 1993 (their figs. 3, 4, 8, 9, 10, 11, 12).

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Fig. 6. SEM core photographs of cored sandstones from the Siri area, Central Graben. (A) Glauconite (GI) grain, with carbonate (?) cement and smectite clay mineral around. (B) Quartz crystal (Qz) and quartz grains with clay minerals (probably kaolinite) around. (C) Coccoliths within smectite (Sm). (D) Quartz crystals within kaolinite (Kl).

Fig. 7. X-ray diffractogram from one sandstone sample of the Siri cores, indicating the presence of smectite.

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Fig. 8. Schematic sketch of the depositional model for the Paleocene succession of the Siri area, Central Graben.

Interestingly, these sand bodies are comparable in size and geometry to offshore tidal bars reported for the modern Norfolk and Flemish Ridges (offshore southeast UK) by Stride, 1982 (in Johnson and Baldwin, 1986). These modern sand bodies are constructed by strong tidal current reworking, in water depths between 10 and 50 m. Similar offshore tidal bars are also commonplace along the eastern coast of the US. This suggests that similar current reworking processes can result in similar geometries in different settings.

Heimdal area, South Viking Graben

Fans, channels and bird foot geometries at the regional to sub-regional scale In the greater Heimdal area of the South Viking Graben, 3D seismic amplitude extraction at Top Heimdal Formation (Paleocene) shows linear to sinuous, channelled, "bird foot", and fan-shape geometries (Fig. 9). These configurations are perhaps more commonly found in continental and transitional environments, rather than in deep-water settings. However, we do know that Paleocene sediments in this area were deposited in deep water (Skibeli, 1987; Jenssen et al., 1993). Moreover, some of the channels are sometimes bigger than the nearby "fan". These seismic features illustrate that this three-dimensional

distribution of depositional environments (and facies) is essentially fractal in nature. That is, depositional pattern and geometries at the large, regional scale are mirrored and can also be recognised at the smaller, sub-regional scale. In addition, certain sedimentary configurations may be the result of similar flow regimes even in different depositional settings. Within this "fractal" context, the next Viking Graben case represents a good example of how important the integration between 3D seismic and core observations can be. 3D seismic observations indicate in fact the occurrence of mounded to fan-shape geometries in the upper Heimdal section of the South Viking Graben area (Fig. 10). Isochores of Layers 1 and 2 (upper part of Heimdal Formation) indicate offset geometries. Another interesting seismic feature is the observation of injected sands (also observed at the core scale) associated to re-mobilised sand bodies (Fig. 11).

Massive sands, olistostromes and current reworking Core observations from wells (450 m of cores) indicate the occurrence of the following sedimentary facies (Table 2): (1) massive sands; (2) brecciated sands and shales; (3) graded sands; (4) shales with rippled sands; (5) laminated shales (Figs. 12 and 13).

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Fig. 9. Seismic amplitude extraction at the top of the Paleocene Heimdal Formation in the South Viking Graben, highlights shapes/geometries that are arranged in a fractal pattern, i.e. what is observed at the small scale is also repeated at the larger scale. (A) Sinuous to meandering channel geometries (middle to upper part of the figure) occur nearby a fan geometry (lower part of the figure). Note occurrence of depositional geometries like in (B) ("bird foot" shape) and (C) (fan shape) that are also common in continental depositional environments. Scale in (B) and (C) is voluntarily omitted.

These facies are very similar to those described for the S iri area, but occur in different proportions.

Interpretation 3D seismic indicates that mounds and fan-shape geometries are common in the Heimdal area and this may lead to the conclusion that lobe-channel turbidites were deposited within a deep-water basin floor fan. However, only 5-10% of the sands can be classified as true turbidites. The bulk of the sands is massive, with sharp lower and upper contacts, and yielding floating and rafted clasts. All sediments (including turbidite facies) exhibit evidence of current reworking, similar to the Sift area. We interpret the upper Heimdal depositional system in the Heimdal area as a sand-rich system (Fig. 14), where sands were mobilised from a shelf-slope area (Shetland Platform) and deposited down dip mostly as coherent

flows. The mechanism facilitating the mass movement could be explained as aquaplaning. That is, the sediment mass is not in contact with the substrate, due to a water film which acts as a lubricant. Associated slope slides were responsible for the observed olistostromes (brecciated sands and shales) and slump facies. In addition, both core and seismic observations indicate the widespread occurrence of injected sands (see also Jenssen et al., 1993). These sands were probably re-mobilised as a response to the loading and consequent compaction caused by the further deposition of mass flows onto the existing basin floor topography, resulting in re-mobilised sands. Basin floor slopes, related to underlying structural elements, were also a major control on sediment distribution acting as baffles and causing stacking of sand bodies, and causing additional reworking (Fig. 14).

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Fig. 10. 3D-isochore mapping highlights mounded and fan geometries of the Paleocene Upper Heimdal Formation. Vertical lines in the cross-section are well ties.

Conclusions

The use of seismic attribute analysis from extensive 3D seismic data sets has increased during these last years and has greatly improved the interpretation of the lateral distribution and variability of depositional systems. In our examples, we have shown that seismic attribute analysis must be integrated with core sedimentology to really improve the identification, distribution and variability of the sedimentary architecture of the North Sea Paleocene succession. This sedimentary architecture is highly variable, fractal in nature and not forcible into fixed models. The implications for petroleum exploration and exploitation are that (1) any seismic interpretation should always reflect a continuous small- to largescale loop, in order to place observations and interpretations in their right context and scale, and (2) any

seismic interpretation should be integrated with core sedimentology analysis. Core sedimentology shows that several depositional mechanisms were active during the Paleocene. Among these, slumps, debris flows and current reworking are here proposed in addition to the classic turbiditic mechanisms. Consequently, different depositional models and processes must be considered when different areas of the North Sea are evaluated for petroleum exploration and exploitation of "Paleocene" plays and fields. The interpretation of gravity flow deposits in our two examples relies upon the association of the sedimentary features described in this paper. The core material provides evidence that traction processes (debris flow and current reworking) were at least as important as the conventionally cited gravity and suspension processes (i.e. turbidite) during

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Fig. 11. Paleocene injected sands (encircled) can be observed at the seismic scale in the Heimdal Field Area, South Viking Graben. The mounded feature to the left of the sand injection is here interpreted as a re-mobilised sand body.

Fig. 12. Paleocene facies of well cores from the distal settings of the Heimdal Formation, South Viking Graben. Massive sands with sharp upper and lower contacts (black arrows) and yielding bed chunks (BC), floating and rafted clasts (FC, RC), are here interpreted as sandy debris flow deposits (SDF). Note the brecciated nature of the finer lithologies, interpreted as muddy debris flow deposits (MDF). Also note the occurrence of associated slumps and boudins (SL, BD). The bi-directional red arrow in (B) indicates the general "sheared" appearance of the finer lithologies. These facies could also be interpreted as olistostromes, analogous to those described from the Northern Apennines of Italy (Abbate et al., 198 l a,b). Vertical scale 1 m. The apparent stratal dip is due to well deviation.

the final emplacement of Paleocene sands. In addition, we consider that the occurrence of slumps and olistostromes favours a debris flow rather than a turbidite interpretation. Interestingly, the presence of smectite and other clay minerals could be the factor that increases the cohesion of these sand bodies and allows the cohesive or semi-cohesive transport en masse with successive freezing (Middleton and Hampton, 1973; Shanmugam, 1996b). According to

Vorren et al. (1998) and Vorren and Laberg (2001), this en-masse sediment movement could have been facilitated by aquaplaning (as described above), probably the only process making long-distance transport of sandy debris flows possible. Interpretations similar to ours have already been given by O'Connor and Walker (1993) and Shanmugam et al. (1995) for other North Sea Paleocene examples.

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Fig. 13. Paleocene facies of well cores from distal settings of the Heimdal Formation, South Viking Graben. Note truncated ripples (red arrows) in normal graded sandy turbidite layers (TUR) as well as in finer lithologies. These truncated ripples are here interpreted to be caused by the action of strong submarine currents. Compare with similar facies from the Siri area (Fig. 4). SDF: sandy debris flow; CRS: current reworked sediments. The yellow arrow indicates an injected sand. Vertical scale 1 m.

Fig. 14. Schematic sketch of the depositional model for the Paleocene succession of the Greater Heimdal area, South Viking Graben.

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There is little doubt that the Paleocene sands of the North Sea were initially deposited by gravity flow processes. However, our examples suggest that secondary (post-depositional) processes significantly contributed to the final architecture of these important reservoirs.

Acknowledgements We thank Enterprise Oil Norge for giving permission to present and publish this paper. Thanks to H. Dennis (Enterprise Oil Norge, Ltd.) who commented on an earlier version of the manuscript and supported this effort. Thanks to A. Ramirez, R. Jansen and Mike Talbot of the University of Bergen, Norway, who supplied preliminary petrographic results from the analyses of Sift well sandstones and shared useful ideas while logging the Sift cores. Thanks to C. Puigdefabregas and to an anonymous reviewer whose comments greatly improved the manuscript. Finally, many thanks to T. Dreyer (Editor) for his constructive comments.The ideas and interpretations reported here represent the authors' viewpoints and interpretations and do not necessarily reflect those of Enterprise Oil or its partners.

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The Maastrichtian and Danian depositional setting, along the eastern margin of the Mere Basin (mid-Norwegian Shelf): implications for reservoir development of the Ormen Lange Field J.G. Gjelberg, T. Enoksen, P. Kjsernes, G. Mangerud, O.J. Martinsen, E. Roe and E. V&gnes

The reservoir interval of the Ormen Lange Field (M~re Basin) consists of two units, a lower heterolithic unit mainly of Maastrichtian age (Jorsalfare Formation) and an upper sandstone unit of Danian age (Egga member, Vfile Formation). The Cretaceous-Tertiary boundary is located in the uppermost part of the unit. The Egga member has long been known as a potentially prolific reservoir interval in the Sl~rebotn Subbasin area, where it is up to 150 m thick, consisting of thick, amalgamated medium- to coarse-grained sandstone turbidites. The base of the Egga member in the Sl~rebotn Subbasin and in other areas close to the eastern margin of the M~re Basin is a significant unconformity, with the Danian succession usually overlaying strata of Campanian age. This unconformity developed during Maastrichtian times with shelf erosion and sediment bypass. The bypassed coarse clastic sediments were redeposited as turbidites further out in the M~re Basin, and represent the lower, heterolithic reservoir unit in the Ormen Lange Field. In early Danian time uplift and rotation of the Fennoscandian provenance area in the east led to extensive erosion and redistribution of sandy sediments that prograded westwards into the M~re Basin and gave rise to amalgamated turbidites of the Egga member. These turbidites probably filled a shallow intraslope basin or depression in the Sl~reboth Subbasin and laterally bypassed into the deeper part of the M~re Basin, where relatively thick turbidites accumulated to form the upper reservoir interval of the Ormen Lange Field. During latest Cretaceous and Palaeogene times the location of the Ormen Lange Field was an area of relatively rapid subsidence, probably with the development of a temporary basin floor depression where deep water sediments were deposited. Inversion and development of the Ormen Lange structure took place in latest Eocene-Early Oligocene time. A considerable change in mineralogy and ichnofabric occurs at the Cretaceous-Tertiary boundary, indicating significant changes in the depositional environment.

Introduction The aim of this paper is to describe and interpret the sedimentological and stratigraphical development of the reservoir interval of the Ormen Lange Field, and to explain the development in a subregional and tectonostratigraphic setting. To date four wells have been drilled on the Ormen Lange structure. The first well was drilled by Norsk Hydro in the southern part of block 6305/5-1 in 1997. The second well was drilled in the B P license in block 6305/7 in 1998 and recovered a reservoir similar to that in well 6305/5-1. The third well (6305/1-1) was drilled by Norsk Hydro in block 6305/1 to gain information about the reservoir development towards the north. The fourth well (6305/8-1), drilled by Norsk Hydro in 2000, was an appraisal well to the eastern margin of the reservoir. In this paper we concentrate

on reservoir development in well 6305/5-1, which is representative for the development of the reservoir in the southern part of the structure. In addition, the northern development will be described with reference to well 6305/1-1.

Structural setting The Ormen Lange structure is one of a series of Cenozoic dome structures formed in the NorwegianGreenland Sea (Blystad et al., 1995). The onset of doming and creation of the Ormen Lange structure closely coincided with the onset of sea-floor spreading in the Norwegian-Greenland Sea (Dor6 and Lundin, 1996 and Dor6 et al., 1997), suggesting a change in the far-field plate-tectonic stress as a possible causal mechanism. Significant compression occurred during Oligocene-Miocene time, and ap-

Sedimentary Environments Offshore Norway - Palaeozoic to Recent edited by O.J. Martinsen and T. Dreyer. NPF Special Publication 10, pp. 421-440, Published by Elsevier Science B.V., Amsterdam. 9 Norwegian Petroleum Society (NPF), 2001.

J. G. Gjelberg et al.

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Fig. 1. Location map of the Ormen Lange Dome in the MOre Basin, with main Jurassic-Cretaceous structural elements outlined.

pears to have continued up to the present (Vgtgnes et al., 1998). It seems likely that the Ormen Lange Dome has been controlled by underlying structural elements of Late Jurassic and older age. These structural elements comprise a set of horsts, grabens and half grabens bordered by mainly SW-NE- and N S-trending faults (Fig. 1). A less prominent but significant NW-SE trend parallel to the Jan Mayen Lineament is also present (Fig. 1). This lineament may have exerted an important control on early Cenozoic sediment distribution. Distribution of Paleocene reservoir sand was probably also influenced by local depocentres related to reactivation of underlying Jurassic and older structural elements (Jongepier et al., 1996). This is expressed in pronounced local increase of thickness across structural elements (Fig. 2). One depocentre was located at the position of the present Ormen Lange Dome, whereas a second was located in the SlCrebotn Subbasin. These depocentres developed due to increased subsidence or previous underfilling, and it is suggested that in periods topographic depressions periodically developed on the sea floor that

influenced the flow direction of turbidity currents. It is tentatively suggested that the shape of the Paleocene depocentre at the Ormen Lange Field closely resembles the form of the present dome. This may be related to subtle reactivation of the underlying structural elements in an extensional regime during DanianYpresian time followed by contractional reactivation of the same elements from the Ypresian to the present (V~gnes et al., 1998). The local depocentre at the Ormen Lange Dome was located on the floor of the much wider MOre Basin, whereas the depocentre in the SlCrebotn Subbasin was on the eastern slope of the MOre Basin, bounded to the northwest by the Gossen High, thus forming an intra-slope basin (Fig. 1).

Maastrichtian and Paleocene stratigraphy The cored interval of well 6305/5-1 comprises reservoir sand both of Maastrichtian (Springar/Jorsalfare Formation) and Danian age (informally called Egga member in this paper). The Egga member comprises the lower part of the Vfile Formation (Danian

The Maastrichtian and Danian depositional setting, along the eastern margin of the MOre Basin

423

Fig. 2. Seismic section across the Ormen Lange Dome showing some of the key interpreted horizons. Note the thickening of the Paleocene interval (base Tertiary-top Balder) as opposed to the thinning of the overlying succession. Location of well 6503/5-1 is shown in Fig. 1.

age). In addition to the Vgtle Formation, the Paleocene of the MOre Basin also comprises the Lista Formation (Selandian) and the Sele Formation of latest Paleocene-earliest Eocene age (Fig. 3). In the MOre Basin it is generally difficult to distinguish between the Sele and Lista Formations, based on the original definition of Isaksen and Tonstad (1989). They differentiated between the two formations on the basis of a slightly higher tuff content and a higher TOC in the Sele Formation, compared to the Lista Formation below, and an abrupt decrease of interval transit time downwards across the formation boundary. It is, however, easier to apply the mid-Norwegian Shelf stratigraphy of Dalland et al. (1988) for this area, who defined only one formation for the entire Paleocene, the Tang Formation, which consists mainly of dark grey to brown claystone with minor sandstone and limestone. The main change in the Tertiary stratigraphy between the northern North Sea area and the TrOndelag Platform area occurs close to the Jan Mayen Lineament (Fig. 1), and it is therefore reasonable to use this as the geographical border for the application of the two stratigraphic schemes. One of the most characteristic stratigraphic features of the Maastrichtian and Paleocene succession

along the eastern margin of the MOre Basin and neighbouring areas is the base-Tertiary unconformity that is well developed and present in all wells along the margin (Fig. 3). Most of the Maastrichtian and the Lower Paleocene stratigraphy is absent in these wells. The origin and nature of this unconformity is not well understood; however, it is tentatively suggested that it may represent a subaerial erosional surface related to a significant relative sea-level fall during the late Maastrichtian and Early Paleocene, where both the basin margin and platform areas may have been exposed. This possible relative sea-level fall was probably related to regional epeirogenic uplift of the mainland (e.g. Riis, 1996). In the MOre Basin, on the other hand, there is a continuous stratigraphic development throughout the Upper Cretaceous-lower Tertiary, with the actual K-T boundary at about 2780 mRKB in well 6305/5-1 (Fig. 4).

Facies description Seven process-related facies have been identified within the cored succession (Facies A-G). Facies identified are based on classical work on turbidite facies such as Bouma (1962), Walker (1967), Mutti and Ricci Lucchi (1972), Walker and Mutti (1973), Carter

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The Maastrichtian and Danian depositional setting, along the eastern margin of the MOre Basin

Fig. 4. Core log of well 6305/5-1. The cored interval comprises the Maastrichtian and Danian succession.

425

426 (1975), Walker (1978), Lowe (1982) and Pickering et al. (1986, 1989) and Allen (1991).

Facies A - - sandy high-density turbidite This facies consists of massive and vaguely laminated green sandstone (up to 4 m thick) with abundant green mudstone clasts (up to 2 cm in diameter). Deformation structures due to water escape (e.g. dish structures and contorted bedding) are common in most of the sandstone beds. Normal grading may occur in the uppermost part of individual beds, and inverse grading is relatively common in the lowermost part. However, in most cases, the grain size distribution is remarkably uniform with medium-, and occasionally medium- to coarse-grained sandstone dominating. Minor vertical alternations between these grain sizes are evident within individual beds. Another typical vertical arrangement of this facies is a massive or graded lower part, with non-scoured base, passing upwards into a massive interval with green mudstone clasts. The mudstones are usually well rounded, consisting mainly of chlorite and glauconite. Thin intervals with ripple lamination and convolute lamination may occur on top of some of these massive sandstone beds. The tipple laminated intervals are, however, very thin (few cm), and commonly poorly developed. These ripple laminated upper sandstone intervals may separately be classified as a "classical" turbidite (see below), indicating the close relationship between high-density turbidites and "classical"/low-density turbidites. Coal fragments are common, especially in the upper part of the beds. Pyrite aggregates, mica and siderite are common as additional minerals. Trace fossils present include Planolites and Palaeophycus. Carbonate cement occurs within a few beds of Facies A, and may be restricted to thin intervals, both in the upper and lower part. Our understanding and use of the terms highdensity turbidites follow the well established use as defined by Middleton (1967), Walker (1978), Mutti (1979), Lowe (1979, 1982), Surlyk (1984) and Nemec and Steel (1984). There is a gradual transition between these deposits and "classical .... Bouma" turbidites, and the Ta subdivision of the "Bouma sequence" reflects generally the same depositional mechanisms as the high-density turbidites. The thick beds of this facies may reflect sustained, near-steady flow conditions with a gradually rising depositional flow boundary (cf. Kneller and Branney, 1995).

Facies B - - low-density ("classical") turbidite Most of the fine- to medium-grained sandstone beds with well developed normal grading, and with a

J. G. Gjelberg et al. partly developed Bouma sequence (e.g. grading from a massive lower part, to parallel lamination and occasional ripple lamination at the top), are interpreted as low-density "classical" turbidites. Some sandstone beds with strongly contorted lamination have also been assigned to this facies. This facies is relatively rare in the cored succession and comprises only a few beds.

Facies C - - highly bioturbated green sandy mudstone The facies is common in the upper sand-rich part of the cored interval, where it separates individual sandy turbidites. Typical for the facies is that it fines upwards from a sandy high-density turbidite (Facies A) at the base, through an interval of strongly bioturbated and soft sediment deformed, mixed sand and green mudstone and finally terminates in a green mudstone. B ioturbation mainly comprises Planolites, Palaeophycus and occasional Rhizocorallium trace fossils. This facies resembles Facies D (see below); however, it differs slightly due to its thinner and sandier development. It represents strongly deformed and bioturbated turbidite tops with a gradual upwards transition to hemipelagic sediments.

Facies D - - hemipelagic green, greenish-grey and grey mudstones This facies is very common in the lower part of the cored interval (mainly below the K - T boundary), where it is present as relatively thick beds (up to 3 m) of alternating green-grey and greenish-grey strongly bioturbated mudstones, with occasional centimetrethick lenses of glauconite-rich sand. The mudstones consist mainly of a mixture of calcareous clay and siltstone, with minor elements of sand. Glauconite grains, pyrite aggregates and mica occur at most levels. Small shell fragments are common throughout, but in relatively small amounts. At two thin intervals there is a significant increase of carbonate, probably primarily related to increased deposition of chalk. The intensity and diversity of bioturbation are very high and in most cases the sediments are almost completely homogenised with no primary sedimentary structures or lamination preserved. The most common trace fossils are Zoophycos, Planolites, Palaeophycos and Helminthopsis horizontalis. Facies D is interpreted to represent mainly finegrained sediments deposited from suspension in a relatively deep water column (probably far below storm wave base), and formed the background sedimentation for most of the Maastrichtian succession. The in-

The Maastrichtian and Danian depositional setting, along the eastern margin of the MOre Basin terbedded sandstone lenses probably represent thin turbidites. The considerable thickness and time involved in each of these beds indicate that background sedimentation was more or less uninterrupted for a long period of time (> 100,000 years) between each sand transporting turbidity current reached the area. Facies E -

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contorted and broken up into sandstone fragments, associated with sand injections. Another level with deformation and associated sand injections is present at 2770.5 m, where rotated and truncated, laminated sandstone blocks with irregular sand injections occur. This deformation seems to be associated with a minor normal fault.

hemipelagic black mudstones

Stratigraphic development This facies is common in the middle part of the cored succession where it occurs just above the K-T boundary, consisting of black and dark-grey mudstones with millimetre-scale thin laminae of silt and very fine-grained sandstone. It is up to 3 m thick, and occurs over an interval of 8 m, interbedded with three high-density turbidite sandstones. The degree and diversity of bioturbation is generally very low. Planolites is one of the few trace fossils in this facies, which represents a highly different depositional setting compared to the green mudstones below, and it is tentatively suggested that it may represent significant changes in basin physiography. This facies is also interpreted to represent finegrained sediments deposited mainly from suspension in deep water, and again, formed the background sedimentation during the Early Paleocene. The siltstone and very fine-grained sandstone lenses and laminae within this facies probably represent thin very lowdensity turbidites. The considerable thickness of this facies indicates that it was deposited over a long period of time without interruption by turbidity flows. Facies F ~ chalk

This facies consists of chalk and occurs only in the lower part of the cored interval (core 6), as a single bed, 13 cm thick, immediately on top of a highdensity turbidite (2798.7 m). The lower boundary is sharp whereas the upper is transitional. It is strongly bioturbated (degree 5) with both Planolites, Zoophycos and possibly Taenidium satanassi trace fossils. As mentioned above, there are a few other beds within the Maastrichtian succession that also have increased carbonate content due to chalk; however, these beds are strongly mixed with clastic mud. Facies G m deformed s a n d s t o n e / m u d s t o n e with sand injection

Some of the sandstone beds are significantly deformed and differ from the contorted bedding related to water escape structures within turbidites (e.g. Facies B, see above). One such a deformed interval is located at the K-T boundary (Fig. 8), where the top of a relatively thick sandstone bed is strongly

Two formations are represented in the cored interval in well 6305/5-1: the Jorsalfare Formation (or the Springar Formation according to mid-Norwegian Shelf stratigraphy of Dalland et al., 1988) of Maastrichtian age, and the Vfile Formation (or the corresponding Tang Formation) of Danian age. The V~le Formation is present in the upper part of the cored succession and comprises two distinct lithological units: a mudstone-dominated lower part and a sandstone-dominated upper part (Fig. 4). The Jorsalfare Formation (Maastrichtian)

The top of the Jorsalfare Formation is dated, based on the last occurrence (LO) of Palynodinium grallator, at 2779.30 m in well 6305/5-1 (Fig. 4), taken as evidence for a latest Maastrichtian age. P. grallator is recorded in all three wells drilled on the Ormen Lange structure, and its last occurrence corresponds approximately to the top of the deformed sand bed (Facies G). This occurrence also corresponds to a distinct change in the overall fossil assemblages. The assemblages in the Jorsalfare Formation are dominated by relatively rich and diverse marine microplankton and species like Isabelidinium majae, Areoligera "horrida ", Heterosphaeridium ? heteracanthum, Alterbidinium acutulum and Isabelidinium cooksoniae have their last occurrences within this formation. The Cretaceous pollen Aquillapollenites spp. is present occasionally. Microfaunal assemblages are poor, but agglutinated foraminifera, indicative of upper bathyal depths, are present. There is a considerable difference between the Maastrichtian succession in this well compared to the development in other wells in the northern North Sea region: the presence of significant amounts of sand. Most of the Maastrichtian succession has been cored, and it shows a poorly defined sandier upwards development, combined with a thickening upwards of individual sandstone beds (Fig. 4). The lithology of the lower part of the Maastrichtian succession consists mainly of high-density turbidites (Facies A), interbedded with thick beds of green and grey, highly bioturbated mudstones of Facies D (Fig. 5). There are also a few interbedded low-density turbidites

428

Fig. 5. Alternation between thick greenish mudstones and turbidites in the Maastrichtian succession of well 6305/5-1. Note the thin chalk bed at 2798.7 m (arrow) located on top of a classical turbidite. This is the only chalk bed in the cored succession of well 6305/5-1, but may be widely distributed in the area since it most likely correlates with a similar chalk bed present in well 6305/7-1 at the same stratigraphic level.

(Facies B). Carbonate cement is relatively common within the turbidites, and results in a significant reduction of pore volume. A bed of bioturbated chalk, 13 cm thick, (Facies F) is present in the upper middle part of the succession (Fig. 6). It is suggested that the chalk was deposited from suspension in periods favourable for production of coccoliths in the water column. Such sudden changes may occur due to short-lived climatic changes or changes in basin circulation patterns. Fig. 5 shows the typical development of the Maastrichtian succession characterised by alternating greenish, strongly bioturbated mudstones with turbidites. The upper part of the Maastrichtian succession shows more or less the same facies as the lower one, with alternating high-density turbidites and green and dark-grey mudstones (Fig. 4). The main difference is, however, that individual turbidites are thicker, commonly amalgamated or separated by thinner mudstone intervals between the turbidites (Fig. 4). The stratigraphic development of the Jorsalfare Formation indicates slow background sedimentation mainly from suspension fallout in a well oxygenated,

J. G. Gjelberg et al.

Fig. 6. The chalk bed (Facies F) in the middle part of the cored Maastrichtian succession in well 6305/5-1 (see Fig. 5), showing strong bioturbation with high diversity. open marine basin. This was interrupted by several pulses of turbidity current input that entered the basin floor. There seem to be a gradual or slightly episodic change upwards in the fine-grained background sediments in the uppermost part of the Maastrichtian succession from greenish-grey, bioturbated mudstone to dark-grey, less bioturbated mudstone. Turbidites of Facies A, contorted and broken up into sandstone fragments, associated with sand injections occur at the top of the Maastrichtian succession (Facies G, Figs. 7 and 8).

The Egga member, VMe Formation (Danian) The cored section of the Vgtle Formation contains a typical assemblage reported from Danian strata elsewhere, including a number of last occurrences (from oldest to youngest) as last common occurrence (LCO) of Cerodinium diebelii, last abundant occurrence (LAO) of Trithyrodinium fragile, LCO of Spongodinium delitiense, LO of S. delitiense, LAO and LO of Senonisphaera inornata and LO of Spiniferites

The Maastrichtian and Danian depositional setting, along the eastern margin of the MOre Basin

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Fig. 7. Cored succession across the K-T boundary. The boundary is located at 2779.3 m and occurs at the top of a deformed bed characterised by brecciation and sand injection. Note the dark-grey, virtually non-bioturbated mudstones immediately above the boundary. These mudstones represent fine-grained background sediments that differ considerably from the Maastrichtian greenish, strongly bioturbated background sediments.

"magnifica". The presence of Carpatella cornuta in one of the wells indicates an oldest Danian age and is the first observation of this species in the mid-Norwegian area. It appears in extremely low numbers in association with S. delitiense close to the FO of Xenicodinium lubricum. Similar assemblages are recorded in all three wells appearing just above the LO of Palynodinium grallator. Other first occurrences (FO) include FAO of T. fragile (at 2778.6 m in well 6305/5-1). The top of the Egga member falls between the two regional markers: the LCO of Palaeocystodinium bulliforme and the LO of Alisocysta reticulata (Fig. 4). A. reticulata is regarded to present the top Danian implying that the reservoir is of late Maastrichtian to Danian in age. The V~le Formation consists of two distinct lithological units: a lower mudstone unit, and an upper sandstone unit. The sandstone unit is here referred to as the Egga member. The lower mudstone unit There is a significant change in background sedimentation immediately above the K-T boundary.

Fig. 8. Strongly deformed turbidite at the K-T boundary (2779.3 m) in well 6305/5-1. The width of the core is approximately 10 cm. Sand injection outlined in black.

This change is especially well expressed in well 6305/5-1. Below this level fine-grained material of Facies D (hemipelagic green, greenish-grey and grey mudstones) dominates. Above this level, however, mainly fine-grained sediments of Facies E (hemipelagic black, virtually non-bioturbated mudstones) constitute the background sedimentation, all the way up to the sandstones of the overlying Egga member (Fig. 7). The fine-grained background sediments occur in association with coarse- to very coarse-grained, high-density turbidites at about the same spacing and thickness as below (Fig. 7), and is associated with a pronounced increase of acoustic velocity (downwards) across the K-T boundary. The spacing and thickness of the associated turbidites do not differ significantly from the distribution of simi-

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lar facies in the Maastrichtian succession below. The low degree and diversity in bioturbation of the finegrained facies, however, suggest a highly different depositional setting compared to the green mudstones below, and it is tentatively suggested that significant changes in basin physiography and/or climate took place.

The Egga (sandstone) member In the overlying sandstone-dominated Egga member (Danian), a different development of the finegrained background sediments occurs. Greenish, strongly bioturbated facies (Facies C and D) again dominate. This change is most significant across the base of the Egga member. However, a gradual transition to greenish bioturbated mudstone is also recorded slightly below the base. The Egga member is sanddominated, consisting mainly of high-density turbidites (Facies A) in alternation with thin, finingupwards sandy mudstones of Facies C and D (Fig. 4). Individual turbidites range in thickness from a few decimetres to more than 3 m, and may be amalgamated into several metres thick successions (up to 9 m), without significant fine-grained sediments between. The high-density turbidites show a surprisingly uniform grain size distribution, ranging mainly between medium- and coarse-grained sandstone. The vertical distribution of the turbidites of the Egga member shows no obvious vertical arrangements, such as well defined fining- or coarsening-upwards trends. There are, however, two intervals that may be defined as thinning-upwards successions, one of which is located in the lower part of the member (from 2764 m to the top of core 2 in well 6305/5-1) and the other one from the base of core 1 to the base of a thick turbidite (2731.25 m) on the top of the cored succession (Fig. 9). These trends are generally poorly defined and do not justify extensive conclusions. It is tentatively suggested that the fining- and thinning-upwards successions in this member may reflect the in-fill of poorly defined distal channels on a submarine fan system or simply lateral switching of lobes. Based on facies interpretation and distribution together with detailed seismic mapping the Egga member in the central area of the Ormen Lange Dome is interpreted to represent a basin floor fan complex. According to the classification scheme of Reading and Richards (1994), it may be classified as a sand-rich, point-sourced submarine fan system. The regional depositional setting of the Egga member will be discussed more in detail below. The second well drilled by Norsk Hydro on the Ormen Lange Dome (well 6305/1-1) is located in the northern part of the structure and shows a heterolithic

Fig. 9. The upper part of the Egga member in the central part of the Ormen Lange structure (well 6305/5-1), showing amalgamated high-density turbidites separated by fine greenish mudstones and siltstones of hemipelagic origin. The turbidites thin upwards towards the uppermost bed of the cored succession which is relative thick. This interval represents the interval of the Egga member where they are thinnest.

development mainly consisting of fine-grained sediments, such as strongly bioturbated siltstones and mudstones in alternation with thin turbiditic sandstones (Fig. 10). This development indicates that the submarine fan terminates northward and develops into fan fringe facies. Seismic interpretation of the Egga member across the Ormen Lange Dome shows a gradual thinning both northwards and westwards. Mineral composition and provenance area

The sandstones of the Egga member in well 6305/5-1 are all classified as subarkoses according to the classification scheme of Pettijohn et al. (1972), with close to 10% feldspar. Mineralogical and element compositions of the fine-grained background sediments in the well indicate that there is a significant change across the K-T boundary. The concentrations of both kaolinite, mica/illite and calcite decrease significantly across the boundary, whereas smectite and quartz increase (Fig. 11). Minerals such as K-feldspar and dolomite seem to be evenly dis-

The Maastrichtian and Danian depositional setting, along the eastern margin of the MOre Basin

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Fig. 10. Typical sedimentary facies of the Egga member in the northern part of the Ormen Lange Dome (well 6305/1-1). Strongly bioturbated (Zoophycus) siltstones and mudstones are the dominant lithologies (A) and represent probably a mixture of thin turbidites and hemipelagic mudstones. There are a few classical sandy turbidites present in the cored interval (B) consisting of fine-grained, relatively well sorted sandstone.

tributed across the K - T boundary. There is a gradual change from green, strongly bioturbated mudstones to less bioturbated dark-grey and black mudstones upwards towards the boundary. However, the major changes both in mineralogy, ichnofabric and chemistry occur at the boundary itself, and may be related to a catastrophic event such as a meteorite impact (see below). Eight samples from the cored reservoir interval of well 6305/5-1 have been analysed for samariumneodymium isotope (Sm-Nd) provenance ages, aiming to get some indications of the provenance area for the Maastrichtian and Paleocene reservoir sands of the Ormen Lange Field. The Sm-Nd technique relies on the natural radioactive decay of 147Sm to 143Nd (+ 4He) with a half-life time of 106 Ga. By determining the 143Nd/144Nd and J47Sm/144Nd isotope

ratios of mantle-derived crustal rock it is possible to calculate a model age for the sample that reflects the time elapsed since the rock first formed from the mantle (Faure, 1986; Dalland et al., 1995). The distribution of the 147Sm/144Nd ratio shows that all samples fall in the range 0.108-0.119, with a relatively even distribution. This indicates that there is no evidence for fractionation of Sm/Nd ratios between the Maastrichtian and upper Danian (Egga member) sandstone complexes of the reservoir. The Nd concentration shows little variation (less than 20 ppm in all samples). There are no significant differences between the two sandstone complexes, although there is a slight tendency of lower concentrations in the lower sandstone complex. The 143Nd/la4Nd ratios plotted against the Sm-Nd provenance age (Fig. 12B) show that all samples fall in the

432

J. G. Gjelberg et al.

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same range with a provenance age between 1650 and 1670 Ma, with one exception for the lowermost sample that shows an age of 1760 Ma. The provenance age plotted against depth is shown in Fig. 12A and

shows an even distribution with no major differences above and below the Cretaceous-Tertiary boundary. All the samples fall within the normal North Sea "background" range of 1400-1800 Ma (Dalland et

The Maastrichtian and Danian depositional setting, along the eastern margin of the MOre Basin

433

the source for the Cretaceous sandstones at the eastern margin of the MOre Basin was mainly the Western Gneiss region, with no influence from any other major provenance area such as East Greenland.

Regional depositional setting and tectonostratigraphic development

Fig. 13. Sm-Nd isochron plot of samples from the Ormen Lange reservoir, compared with data from modern source areas represented by Norwegian and British river systems (data from Mearns, 1988 and Dalland et al., 1995). The samples both from the Jorsalfare Formation and the Egga member (Vfile Formation) fall in the field of both Norwegian river sediments and British river sediments south and east of the Moine Thrust. The latter is, however, not a possible source for the Ormen Lange reservoir due to the geographical position.

al., 1995). Fig. 13 shows the isotope data from this study together with the range of isotopes from Norwegian rivers, together with British river sediments west and south/east of the Moine Thrust (Mearns, 1988; Dalland et al., 1995). All isotope data show more or less the same range as the Norwegian mainland river sediments but correspond also with the British river sediments south and east of the Moine Thrust. The latter is, however, not considered as a possible source area for the Ormen Lange reservoir sand due to the long transport distance and complex basin configuration. It is therefore concluded that the provenance area for the sands of the Ormen Lange reservoir was the Norwegian mainland domain that exhibits a provenance age of 1.6 Ga, and there are no differences between the Danian and the Maastrichtian sandstone complexes in terms of Sm-Nd isotope composition, indicating the same provenance area for the sands. This is also in agreement with the heavy-mineral composition that is dominated by titanium oxides and garnet. By comparing the different heavy-mineral suites from the Jorsalfare Formation and the Egga member sandstones it is not possible to identify significant differences in composition that can be related to different source areas. So far, there are no published provenance studies from the Palaeogene succession of the mid-Norwegian Shelf. However, based on heavy-mineral constraints Morton and Grant (1998) analysed several intervals within the Cretaceous succession, and concluded that

The Ormen Lange reservoir succession is interpreted to represent a submarine fan extending in time from the Maastrichtian through most of the Danian, and closely related to the depositional system of the widespread sandy Egga member (late Danian) present along the eastern margin of the MOre Basin. The development of the Egga member was strongly associated with the development of the extensive unconformity at the K-T boundary, probably related to uplift and rotation of the Fennoscandian shield.

The Egga member at the eastern margin of the More Basin The Egga member has been known as a prominent sandstone interval in the S10rebotn Subbasin area since 1989 when it first was penetrated in well 6205/3-1 (Fig. 14). During a period from 1989 to 1994 several wells were drilled along the eastern margin of the MOre Basin from the Selje High in the south to the Fr0ya High in the north, and the Egga member is present in all these wells, consisting of thick, amalgamated medium- to coarse-grained turbidite sandstones, with a maximum total thickness close to 150 m (Fig. 14). The base of the Egga member in the S10rebotn Subbasin area and in other areas close to the eastern margin of the MOre Basin is a significant unconformity with the Danian succession overlying strata of Campanian age. In the Halten Terrace and Tr0ndelag Platform the base-Tertiary unconformity is also well developed, and may represent an angular unconformity in the more proximal regions to the east. None of the wells drilled on the Halten Terrace and Tr0ndelag Platform areas have proven sand above the unconformity and differ therefore considerably from the development in the MOre Basin. The K-T unconformity was initiated during Maastrichtian time probably with shelf erosion and sediment bypass. The bypassed coarse clastic sediments were redeposited as turbidites further out in the MOre Basin, and represent the lower reservoir unit in the Ormen Lange Field, consisting of alternating turbidite sandstones and hemipelagic claystones. This development suggests that the generation of turbidity currents occurred occasionally with long periods of quiescence.

4x

Fig. 14. Stratigraphical correlation of the Egga member in wells along the eastern margin of the MOre Basin, showing thickness variation and internal geometries.

The Maastrichtian and Danian depositional setting, along the eastern margin of the MOre Basin

Change of background sedimentation at the Cretaceous-Tertiary boundary A significant change in sedimentation took place across the K - T boundary. This change is mainly seen in the fine-grained background sedimentation which changes from greenish-gray, strongly bioturbated claystones and mudstones of Maastrichtian age into dark-grey, low-degree bioturbated mudstones of early Danian age. The Maastrichtian mudstones exhibit a high diversity of trace fossils mainly of the Zoophycus ichnofacies together with Helminthoida and Anconichnus, whereas the Danian mudstones immediately above the boundary show a very low-diversity assemblage (isolated Planolites). The change in background sedimentation is also reflected by the mineral composition that shows a pronounced change across the boundary (see above). The change in the background sedimentation occurs, however, slightly below the boundary, portrayed by a change from greenish, bioturbated mudstones to dark-grey mudstones in the Maastrichtian succession, reflecting a change in basin physiography that seems to occur gradually or episodically towards the K - T boundary. However, the most significant change, both with respect to lithology and ichnofabric, occurs at the K - T boundary itself. The observation in well 6305/5-1 suggests a similar development as that deduced from mass extinction in planktic foraminifera from several other localities, suggesting that the extinction can be interpreted as a catastrophic event that centred at the K - T boundary, and was superimposed on a gradual mass extinction that began in the late Maastrichtian and continued into the early Danian (e.g. Haslett, 1994; Molina et al., 1996). Other studies, on the other hand, conclude that the K-T boundary is not characterised by dramatic changes in nannofossils and that the concept of an instantaneous mass extinction and its proposed causal connection to bolide impact may be challenged (e.g. Macleod and Keller, 1994). The strongly deformed sandstone bed at the K T boundary with brecciation and sand injection is the only deformed bed of that type in the cored succession of well 6305/5-1. A similar deformed bed has also been recorded from well 6305/7-1, also here at the K - T boundary. It is therefore tempting to believe that this deformation was a result of a catastrophic event caused by severe shock that could be related to bolide impact.

Tectonic implications In early Danian time uplift and rotation of the basin margin area to the east (Norwegian mainland) led to extensive erosion and redistribution of sandy sedi-

435

ments. These sediments prograded westwards into the MOre Basin and gave rise to deposition of amalgamated, high-density turbidites of the Egga member, that probably developed at the front of a rapidly prograding delta (the delta itself is not preserved today due to late Tertiary uplift and erosion). The seismic character of the Egga member along the eastern margin of the MOre Basin changes between different localities. In the Fr0ya High area, the internal seismic pattern is parallel or shingled with very low-gradient clinoforms (Fig. 15A), whereas a more complex and chaotic pattern with channel geometries dominates in the S10rebotn Subbasin area (Fig. 15B). The development on the FrOya High may reflect a relative shallow marine setting, whereas the pattern in the SlOrebotn Subbasin reflects a channelised and probably deformed sandy fan complex. It is tentatively suggested that delta front turbidity currents occasionally bypassed the S10rebotn Subbasin and continued into the deeper part of the MOre Basin, and hence contributed to the development of the upper reservoir interval of the Ormen Lange Field. The area west of the rotation axis (related to the Paleocene uplift) probably went through a period with overall relative sea-level rise, whereas the area to the east was uplifted with an accompanied relative sea-level fall. It is tentatively suggested here that the rotation axis was located east of the SlOrebotn Subbasin, not far away from the present-day shoreline, and that the rotation gave rise to a rapid transgression across the base-Tertiary unconformity in the SlOrebotn Subbasin and the Fr0ya High area, accompanied by rapid delta progradation westwards into the MOre Basin. A generalised palaeogeographic reconstruction of the MOre Basin and its eastern margin in late Danian time is shown in Fig. 16. The development of the base-Tertiary unconformity along the eastern area of the Norwegian continental shelf and the responding sedimentation further out in the MOre Basin and S10rebotn Subbasin imply that severe tectonic uplift of the Norwegian mainland started already in Late Cretaceous/Paleocene times. Based on correlation between offshore geology and onshore geomorphological evidence of an enveloping summit level and remnants of deep weathering in the mountain area of Scandinavia, Riis (1996) suggested that an uplift between 600 and 800 m took place during that period, and suggested that it represents a marginal uplift related to the rifting of the North Atlantic. It is tentatively suggested here that during latest Cretaceous and Palaeogene times the location of the Ormen Lange Field was an area of relative rapid subsidence (indicated by increased thickness across the dome), probably with the development of a temporary

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J. G. Gjelberg et al.

Fig. 15. The seismic character of the Egga member along the eastern margin of the MOre Basin. On the Fr0ya High area (A) the internal seismic pattern is parallel or shingled with very low angle, whereas a more complex and chaotic pattern dominates in the S10rebotn Subbasin area (B).

The Maastrichtian and Danian depositional setting, along the eastern margin of the MOre Basin

437

Fig. 16. Palaeogeographic reconstruction of the MOre Basin and its eastern margin in late Danian time. This reconstruction is based on regional sedimentological and stratigraphical studies together with detailed seismic mapping.

basin floor depression (associated with movements of underlying Upper Jurassic-Cretaceous fault blocks), where coarse-grained gravity-flow-transported sediments were deposited. Inversion and development of the Ormen Lange structure took place in latest Eocene-Early Oligocene time, probably as a result of compression related to ridge push (Vfignes et al., 1998). Fig. 17 gives an outline of the tectonostratigraphic development at the eastern margin of the Me~re Basin from the Maastrichtian to the present day. Conclusions

(1) The cored interval of well 6305/5-1 comprises both the Maastrichtian and Danian. The Maastrichtian succession (Jorsalfare Fm.) consists of grey-green strongly bioturbated, relative thick mudstone in alternation with sandy turbidites. The relative thin (<50 m) Maastrichtian succession spans a considerable time interval (>6 Ma) and long time periods were involved between deposition of each turbidite. (2) At the Cretaceous-Tertiary boundary there was a significant change in depositional environment. The degree of bioturbation and diversity of trace fossils decreased considerably, and a black laminated shale was deposited as "background" fine-grained sediments. There is also a significant change in mineral

distribution and composition across the K-T boundary, and it is suggested that significant changes in basin physiography took place. (3) The Danian Egga member is approximately 50 m thick in well 6305/5-1 and consists of amalgamated, high-density turbidites, with very good reservoir properties (more than 30% porosity and more than 1D permeability locally). (4) There is a considerable increase in the occurrence and thickness of turbidites in the Danian compared with the Maastrichtian succession. This may reflect increased tectonic activity with rotational uplift of the Fennoscandian provenance area related to the rifting and opening of the North Atlantic. It is therefore concluded that the development of the Danian basin floor fan at the Ormen Lange area mainly originated from increased erosion due to epeirogenic uplift, combined with slope instability and collapse due to increased basin margin gradient. (5) Seismic data indicate that the Ormen Lange area was a local depocentre during Paleocene and Early Eocene times, probably with the temporary development of a small depression on the sea floor capturing coarse-grained turbidity flows. Inversion of the depocentre and development of the Ormen Lange Dome took place in Late Eocene'Oligocene time. (6) Heavy-mineral composition and Nd-Sm iso-

438

J. G. Gjelberg et al.

Fig. 17. Outline of the tectonostratigraphic development for the eastern margin of the MOre Basin from the Maastrichtian to the present-day situation. (The Jurassic fault geometry illustrates the strong extensional events taking place in the Late Jurassic with rotational fault movements and large lateral displacement on fault blocks such as the Gossen High.)

tope analyses suggest that both the Jorsalfare Formation and the Egga member sands were derived from the same source area, probably strongly dominated by a metamorphic basement terrain that in provenance age corresponds to the age of the Norwegian mainland (1.6 Ga).

(7) B iostratigraphical data indicate that there has been some reworking of both Cretaceous, Jurassic and Triassic sediments (Robertson Research, unpublished data), and may suggest the former presence of Mesozoic sediments in the provenance area to the east, probably in a position corresponding to the pre-

The Maastrichtian and Danian depositional setting, along the eastern margin of the MOre Basin

sent-day metamorphic basement terrain of the M~re and Romsdal areas.

Acknowledgements We thank the license group of PL209 for permission to publish the data. We also want to thank Ian Sharp and Finn Surlyk for a very constructive review of the manuscript. Thanks are due to Inger Holmefjord at Norsk Hydro Research Centre for analyses of the samples for mineralogical composition, Rolf Birger Pedersen, University of Bergen, for Sm/Nd isotope analyses and Gry Arnesen for drafting some of the figures.

References Allen, J.R.L., 1991. The Bouma Division A and the possible duration of turbidity currents. J. Sediment. Petrogr., 61:291-295. Blystad, R, Brekke, H., Faerseth, R.B., Larsen, B.T., Skogseid, J. and T~rudbakken, B., 1995. Structural elements of the Norwegian continental shelf, Part II. The Norwegian sea region. Norw. Pet. Directorate Bull., 8, 45 pp. Bouma, A.H., 1962. Sedimentology of Some Flysch Deposits. Elsevier, Amsterdam, 168 pp. Carter, C.H., 1975. A discussion and classification of subaqueous mass transport with particular application to grain-flow, shurryflow and fluxoturbidites. Earth-Sci. Rev., 1: 145-177. Dalland, A., Worsley, D. and Ofstad, K., 1988. A lithostratigraphic scheme for the Mesozoic and Cenozoic succession offshore midand northern Norway. Nor. Pet. Directorate Bull., 4, 65 pp. Dalland, A., Mearns, E.W. and McBride, J.J., 1995. The application of samarium-neodymium (Sm-Nd) provenance ages to correlation of biostratigraphically barren strata: a case study of the Statfjord Formation in the Gullfaks Oilfield, Norwegian North Sea. In: R.E. Dunay and E.A. Hailwood (Editors), Non-Biostratigraphical Methods of Dating and Correlation. Geol. Soc. Spec. Publ., 89: 201-222. Dot6, A.G. and Lundin, E.R. Cenozoic compressional structure of the NE Atlantic margin: nature origin and potential significance for hydrocarbon exploration. Pet. Geosci., 2:299-311. Dot6, A.G., Lundin, E.R., Birkeland, 0., Eliassen, EE. and Jensen, L.N., 1997. The NE Atlantic Margin: implications of late Mesozoic and Cenozoic events for hydrocarbon prospectivity. Pet. Geosci., 3:117-131. Faure, G., 1986. Principles of Isotope Geology. Wiley, New York, 2nd edition, 557 pp. Haslett, S.K., 1994. Planktonic foraminiferal biostratigraphy and palaeoceanography of the Cretaceous-Tertiary boundary section at Bidart, South-West France. Cretaceous Res., 15(2): 179-192. Isaksen, D. and Tonstad, K., 1989. A revised Cretaceous and Tertiary lithostratigraphic nomenclature for the Norwegian North Sea. Nor. Pet. Directorate Bull., 5, 59 pp. Jongepier, K., Rui, J.C. and Grue, K., 1996. Triassic to Early Cretaceous stratigraphic and structural development of the northeastern

J.G. GJELBERG T. ENOKSEN E KJA~RNES G. MANGERUD O.J. MARTINSEN E. ROE E. VAGNES

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M~re Basin margin, off mid-Norway. Nor. Geol. Tidsskr., 76(4): 199-214. Kneller, B.C. and Branney, M., 1995. Sustained high-density turbidity currents and the development of thick massive sands. Sedimentology, 42:607-616. Lowe, D., 1979. Sediment gravity flows: their classification and some problems of application to natural flows and deposits. In: L.J. Doyle and O.H. Pilkey (Editors), Geology of Continental Slopes. Soc. Econ. Paleontol. Mineral. Spec. Publ., 27: 75-82. Lowe, D., 1982. Sediment gravity flows, II. Depositional models with special reference to the deposits of high-density turbidity currents. J. Sediment. Petrol., 52(1): 279-297. Macleod, N., Keller and G., 1994. Comparative biogeographic analysis of planktic foraminiferal survivorship across the Cretaceous/Tertiary (K/T) boundary. Paleobiology, 20(2): 143-177. Mearns, E.W., 1988. A samarium-neodymium isotopic survey of modern river sediments from northern Britain. Chem. Geol. (Isot. Geosci. Sect.), 73: 1-13. Middleton, G.V., 1967. Experiments on density and turbidity currents, III. Deposition of sediments. Can. J. Earth Sci., 4: 475505. Molina, E., Arenillas, I. and Arz, J.A., 1996. The Cretaceous/Tertiary mass extinction in planktic foraminifera at Agost, Spain. Rev. Micropaleontol., 39(3): 225-243. Morton, A.C. and Grant, S., 1998. Cretaceous depositional systems in the Norwegian Sea: heavy mineral constraints. Am. Assoc. Pet. Geol. Bull., 82(2): 274-290. Mutti, E., 19791 Turbidites et cones sous-margins profonds. In: R Homewood (Editor), S6dimentation D6trique (fluviatile, littorale et marine). Institut de G6ologie, Universit6 de Fribourg, Fribourg, pp. 353-419. Mutti, E. and Ricci Lucchi, F., 1972. Le tobiditi delt Apennino settentrionale: introduzione all'analisi di facies. Mere. Soc. Geol. Ital., 11: 161-199. Nemec, W. and Steel, R.J., 1984. Alluvial and coastal conglomerates: their significant features and some comments on gravelly mass-flow deposits. In: E.H. Kostler and R.J. Steel (Editors), Sedimentology of Gravels and Conglomerates. Can. Soc. Pet. Geol., Mere., 10:1-31. Pettijohn, F.J., Potter, RE. and Siever, R., 1972. Sand and Sandstone. Springer, Berlin, 618 pp. Pickering, K.T., Stow, D.A.V., Watson, M.E and Hiscott, R.N., 1986. Deep water facies, processes and model: a review and classification scheme for modern and ancient sediments. Earth Sci. Rev., 22: 75-174. Picketing, K.T., Hiscott, R.N. and Hein, F.J., 1989. Deep Marine Environments: Clastic Sedimentation and Tectonics. Unwin Hyman, London, 416 pp. Reading, H.G. and Richards, M., 1994. Turbidite systems in deepwater basin margins classified by grain size and feeder system. Am. Assoc. Pet. Geol. Bull., 78(5): 792-822. Riis, F., 1996. Quantification of Cenozoic vertical movements of Scandinavia by correlation of morphological surfaces with offshore data. Global Planet. Changes 12, (1-4): 331-357. Surlyk, F., 1984. Submarine fan conglomerates of the VolgianValanginian Wollaston Foreland Group, East Greenland. In: E.H. Koster and R.J. Steel (Editors), Sedimentology of Gravels and Conglomerates. Can. Soc. Pet. Geol., Mere., 10: 395-382.

Norsk ttydro Research Centre, P.O. Box 7190, N-5020 Bergen, Norway Norsk Hydro Research Centre, P.O. Box 7190, N-5020 Bergen, Norway Norsk Hydro Exploration, N-0246 Oslo, Norway Norsk Hydro Research Centre, P.O. Box 7190, N-5020 Bergen, Norway Norsk Hydro Research Centre, P.O. Box 7190, N-5020 Bergen, Norway Norsk Hydro Exploration, N-0246 Oslo, Norway Norsk Hydro Exploration, N-0246 Oslo, Norway

440 V~gnes, E., Gabrielsen, R.H. and Haremo, R, 1998. Late Cretaceous-Cenozoic intraplate contractional deformation at the Norwegian continental shelf: timing, magnitude and regional implications. Tectonophysics, 300: 29-46. Walker, R.G., 1967. Turbidite sedimentary structures and their relationship to proximal and distal depositional environments. J. Sediment. Petrol., 37(1): 25-43. Walker, R.G., 1978. Deep-water sandstone facies and ancient sub-

J. G. Gjelberg et al. marine fans: models for exploration for stratigraphic trans. Am. Assoc. Pet. Geol. Bull., 62: 932-966. Walker, R.G. and Mutti, E., 1873. Turbidite facie and facies associations. In: G.V. Middleton and A.H. Bouma (Editors), Turbidites and Deep Water Sedimentation. Pacific Section, Short Course Notes, Society of Economic Paleontologists and Mineralogists, Tulsa, OK, pp. 119-157.

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Glacial processes and large-scale morphology on the mid-Norwegian continental shelf Dag Ottesen, Leif Rise, K&re Rokoengen and Joar Sa~ttem

A regional digital bathymetric data set covering most of the mid-Norwegian continental shelf is presented and gives a unique regional view into glacial processes and ice-sheet dynamics on this part of the continental shelf during the Weichselian, indicating that forms and deposits were created by a highly dynamic ice sheet. At times, ice flow was mainly channelised through ice streams located in bathymetric depressions on the shelf areas. Glacial sedimentary processes are discussed with a focus on the marine-based part of the Scandinavian ice sheet during the last glaciation (the Weichselian). Ice sheets that grounded on the shelf edge are thought to have been responsible for depositing complex prograding sequences on the mid-Norwegian shelf during several glaciations from Late Pliocene time, reaching a maximum thickness of 1500 m on the shelf edge. During interglacials, the shelf areas were sediment-starved with little or no clastic sedimentation. On top of these prograding units, several packages of Quaternary sediments (mainly till of Weichselian age) show a more aggradational pattern. Improved knowledge about the deposition and age of the upper Cenozoic sediment wedge has proved vital for revealing the ice-sheet dynamics and may also be important in understanding the maturation and migration of hydrocarbons on the mid-Norwegian shelf.

Introduction A number of studies during the last 30 years have confirmed that the present morphology of the midNorwegian continental shelf (Fig. 1) is mainly a result of glacial processes (Holtedahl and Sellevold, 1972; Bugge, 1980; Rokoengen, 1980; Gunleiksrud and Rokoengen, 1980; Lien, 1983; Rise and Rokoengen, 1984; Rise et al., 1984; King et al., 1987; Holtedahl, 1993). The stratigraphy and age of the offshore deposits have also shown that glacial processes on the mid-Norwegian continental shelf involved sediment redistribution to a far greater extent and much faster than previously thought (Haflidason et al., 1991; Rokoengen et al., 1995; Henriksen and Vorren, 1996; Sa~ttem et al., 1996; Vorren and Laberg, 1997; Eidvin et al., 1998; Rokoengen and Frengstad, 1999). Improved models of ice-sheet dynamics within areas where the substratum shows changes on a regional scale are very important in order to understand the sediment transport from land to shelf areas, within shelf areas and onto the upper continental slope. The combination of sedimentological, geotechnical and acoustic data from the shelf areas off mid-Norway offer a unique data set to constrain such models both qualitatively and quantitatively. The purpose of the present contribution is to discuss the glacial sedimentary processes and the dynamics of the large ice sheets on the mid-Norwegian continental shelf in the

light of regional bathymetric, seismic and sedimentological data. We will focus on the marine-based part of the Scandinavian ice sheet during the last glaciation (the Weichselian), and especially the behaviour of the ice streams, which are fast moving parts of an ice sheet.

Previous investigations/geological setting In IKU's regional mapping off mid-Norway during the 1970s and early 1980s, the bedrock surface was divided into eleven units informally named I to XI and with ages of the sampled units ranging from Triassic to Pliocene. Due to basinward subsidence and glacial erosion in the inner part of the shelf, the units subcrop more or less parallel to the coast with decreasing ages westwards (Bugge et al., 1984; Rokoengen et al., 1988, 1995; Sigmond, 1992). Bedrock unit IX (Fig. 2) is found about 50 km west of the crystalline basement as a prominent ridge dominated by sand and with greater resistance to later glacial erosion than the presumably more clay-rich sediments below and above. From unit IX and landwards, the Quaternary is fairly thin. The bathymetry, especially between Fr~yabanken and Haltenbanken, reveals varying resistance to erosion of the Mesozoic and Tertiary bedrock units. An interpreted profile on the mid-Norwegian shelf (Fig. 2) illustrates the upper Cenozoic stratigraphy.

Sedimentary Environments Offshore Norway - Palaeozoic to Recent edited by O.J. Martinsen and T. Dreyer. NPF Special Publication 10, pp. 441-449, Published by Elsevier Science B.V., Amsterdam. 9 Norwegian Petroleum Society (NPF), 2001.

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D. O t t e s e n et al.

Fig. 1. Shaded relief map covering the mid-Norwegian shelf with 50 m depth contours. The data have been collected by the Norwegian Hydrographic Service with single-beam echosounder. LI = Lofoten Islands; VF = Vestfjorden; HB = Haltenbanken; SKD = Sklinnadjupet; TD = Traenadjupet; SD = Suladjupet; FB = Fr0yabanken; FR = FrCyryggen; R F -- Romsdalsfjorden; SF = Storfjorden; LG = Langgrunna; M P = Mgllc~yplatgtet; BG = Buagrunnen; BD = Breisunddjupet; OD = Onadjupet; N T = Norwegian Trench; TB -- Traenabanken; SR = Skjoldryggen.

At the present shelf edge, the extensive and complex wedge reaches a maximum thickness of about 1500 m (Rokoengen et al., 1995). A marked change in depositional pattern is observed at the regional unconformity below unit D: (1) the lower units show complex and strongly prograding sequences; (2) the upper units are subhorizontal, exhibiting both progradation and aggradation. There has been a long debate whether the different units are of glacial or non-glacial origin, especially with age estimates varying from Oligocene to Quaternary (Rokoengen et al., 1995). Eidvin et al. (1998) analysed six exploration wells on the mid-Norwegian shelf and dated the oldest parts of the sedimentary wedge on the outer continental shelf to Late Pliocene.

They correlated this with a pronounced expansion of the north European glaciers dated at about 2.6 Ma by Jansen and SjOholm (1991). Unit IX was assigned an Early Oligocene age (Eidvin et al., 1998), but a younger age can still not be excluded. The succession below the irregular base of unit D (units L-E, Fig. 2) exhibits large-scale clinoforms prograding towards the northwest and gradually building out the shelf edge. In general, the units are sheet-like with erosional boundaries in the inner part. Most of the sediments below the irregular base of unit D (units K-E) seem to have been deposited by glacial processes, e.g. deposition by grounded glaciers as various types of tills, as proximal glaciomarine sediments, or by redeposition by

Glacial processes and large-scale morphology on the mid-Norwegian continental shelf

443

Fig. 2. Composite geoseismic profile showing the upper Cenozoic stratigraphy across the mid-Norwegian shelf. Modified after Rokoengen et al. (1995). The units K-E represent Upper Pliocene/Pleistocene sediments, while the units D, B, A and U above the angular unconformity probably represent the last interglacial/glacial cycle. See Fig. 1 for location.

slumping or debris flows down the continental slope during the Late Pliocene/Pleistocene (Rokoengen et al., 1995; Henriksen and Vorren, 1996; Eidvin et al., 1998), but a younger age can still not be excluded. The sediments above the angular unconformity (units D, B, A and U) represent mainly the last interglacial/glacial cycle (Rokoengen et al., 1995; S~ettem et al., 1996). The typically irregular base of unit D is interpreted to be the result of strong glacial sculpturing, probably both constructional and erosional, in late Saalian time. The acoustically layered unit D consists of marine (Eemian) and glaciomarine sediments (Sa~ttem et al., 1996). The three topmost units (B, A and U) are dominated by unsorted material representing Weichselian tills, possibly from three major glacial advances on the continental shelf. Unit B may represent deposits from the first (early) Weichselian glaciation on the shelf, and unit A deposits from the maximum glaciation. Unit U consists of till from the last glacial advances reaching the shelf edge and glaciomarine clays from the last deglaciation (Rokoengen et al., 1995). Units B, A and U have been subdivided into a number of subunits using the till-tongue model (King et al., 1987). Knowledge about climatic changes and Scandinavian ice-sheet variations during the last glacial period (the Weichselian) have increased significantly during the last 20 years as a result of studies in the deep sea (Veum et al., 1992; Fronval et al., 1995), of Greenland ice cores (Grootes et al., 1993; Taylor et al., 1993), and from investigations of the Quaternary stratigraphy on land (Larsen and Sejrup, 1990; Olsen, 1997). Based on more than 100 ~4C AMS-datings, Olsen (1997) reported the glacial variations during the last 45,000 years, suggesting the existence of extensive ice-free areas in several intervals alternating with rapid ice-growth periods. The last glacial maximum comprises two major glacial expansion phases, dated at 22,000 years BP and 16,000 years BE Both advances probably reached the outer parts of the shelf

and are correlated with the uppermost unit A and unit U, respectively. The two youngest till units near the shelf edge (till tongues 23 and 24 in unit U) were deposited at about 15,000 and 13,500 years BP according to 14C AMSdating of shells (Rokoengen and Frengstad, 1999). Along the mid-Norwegian coast, several ~4C-dates give ages older than 12,000 years BE indicating a very rapid deglaciation of the entire shelf area. How these rapid changes in ice-sheet extent and configuration are expressed on the shelf is so far poorly known, but through the regional bathymetric data set we can better imagine the nature of these processes (Figs. 2 and 4).

Sea-bottom morphology related to ice-sheet dynamics Bathymetric data base The Norwegian Hydrographic Service collected single-beam echosounder data during the years 19651985. The regional bathymetric data set covers a large part of the Norwegian continental shelf south of 68~ with an average line spacing of 500 m. Previous bathymetric charts have also been produced (Bugge, 1975; Bugge et al., 1987). The data were gridded with a cell size of 500 m and plotted as coloured contour maps with a 5 m contour interval, and as shaded relief maps. The data set covers the areas from the outer coastal zone with crystalline bedrock to the shelf edge and parts of the continental slope, in certain areas down to 1000 m water depth. The data were collected by an Atlas Penguin echosounder (100 kHz). The positioning system used was Decca Main Chain with an absolute accuracy commonly better than 100 m, but within some areas not better than 500 m. The relative accuracy (repeatability) is, however, much better and the morphology on maps in scales of 1 : 500,000 or less will not be significantly influenced by the inaccuracy.

444 The major morphological features on the mid-Norwegian shelf between the outlet of the Norwegian Trench at about 62~ and the Lofoten Islands at about 67~ are shown in Figs. 1, 3 and 4. Between the outlet of the Norwegian Trench and northwards to 64~ the shelf is rather narrow (60-100 km wide), in contrast to areas further north. In the Skjoldryggen area (Fig. 1), the shelf reaches its greatest width, about 250 km. The shelf includes bank areas with water depths of 100-300 m north of 63~ (Tra~nabanken, Sklinnabanken, Haltenbanken and FrOyabanken). South of 63~ there are several large bank areas with water depths less than 100 m (Buagrunnen and Langgrunna, Fig. 1), in addition to Griptarane west of Kristiansund where crystalline rocks crop out at the surface. In the north, Tr~enabanken and Haltenbanken represent the largest bank areas (Fig. 1). The banks are separated by east-west-oriented depressions 350-500 m deep north of 64~ and 150300 m deep south of 64~ North of 64~ they are up to 60 km wide, while on the southern half of the shelf, between 62~ and 64~ they are narrower, and reach 10-20 km in width.

The Traenadjupet area Tr~enadjupet is the best expressed ice-stream drainage depression in the northern part of the study area (Fig. 1). It is between 40 km and 60 km wide and generally widens towards the shelf edge. The bathymetry shows linear elements parallel to the trough axis, reflecting the flow direction of ice streams. In the eastern part of Tr~enadjupet, at least two different glacial drainage systems coalesced. Ice flow from the northeast (Vestfjorden) and southeast joined to become one major ice stream along Tra~nadjupet (Fig. 3). The bathymetry indicates that the ice stream from the southeast cuts linear features formed by an ice stream from the northeast, and thus reflects the latest erosive phase. The glacial deposits in the Tr~enabanken area are eroded in the Tr~enadjupet depression (King et al., 1987).

The Sklinnadjupet-Skjoldryggen system The Sklinnadjupet is a symmetric, U-shaped trough approximately 30 km wide, up to 470 m deep and 100 km long, presumably mainly eroded by ice streams flowing south of Tra~nabanken during the last glaciation (Fig. 3). The eastern part of Sklinnadjupet has acted as a confluence basin for drainage of ice from the onshore areas. The trough is oriented approximately east-west. The form and trend of the western part of Sklinnadjupet indicate

D. Ottesen et al.

that during the latest phase, the Sklinnadjupet ice stream was deflected towards the north, and flowed in a northwesterly direction out to the shelf edge. This is probably because the ice sheet east of Skjoldryggen was frozen to the ground or pinned by a large, supposedly mainly ice-pushed ridge, Skjoldryggen (S~ettem et al., 1996). In the eastern part of the Sklinnadjupet, the bathymetry indicates that former confluencing glaciers drained into the depressions (Fig. 3). Evidently, Sklinnadjupet received glacial ice from large inland areas. It appears that the shallow bank areas acted as barriers that partly controlled the dynamics of the ice flowing from inland areas to the continental shelves. In the TromsOflaket bank area off northern Norway, a similar observation and outline of a possible glacial mechanism was discussed by S~ettem (1990). Sklinnadjupet partly parallels another major icestream drainage route east of the Haltenbanken area and southwest of Sklinnabanken (Fig. 3). This depression is oriented N W - S E between Haltenbanken and Skinnabanken and continues westwards towards the Skjoldryggen area. Skjoldryggen (Fig. 3) has for a long time been interpreted as an end-moraine ridge at the outermost shelf edge (for further reference, see Holtedahl, 1993). It is almost 200 km long, up to 200 m high and 10 km wide and is by far the largest end-moraine on the Norwegian continental shelf. The morphology east of the Skjoldryggen is complex, comprising several depressions and ridges (Fig. 3). Sa~ttem et al. (1996) reported glaciotectonic deformation in this area, and the bathymetric data set discussed herein supports this interpretation. It seems that the displaced blocks were either transported to and incorporated into the Skjoldryggen moraine ridge, or existed as individual or complex ridges. Sa~ttem et al. (1996) suggested that, following the advance of the ice margin to Skjoldryggen, the ice lobe which deposited the ridge froze to the ground beneath. This pinned the ice, and led to a build up of stress at the ice lobe base which gave rise to glaciotectonic displacement of blocks of frozen sediments. The glacial stratigraphy in both Haltenbanken and Tra~nabanken outlines thick units of Weichselian sediments (units A, B and U, Fig. 2). King et al. (1987) described three till units, each comprising stacked till tongues with intervening glaciomarine sediments deposited during successive advances and retreats of the ice-sheet grounding line. The units generally occupy the outer and central portions of the shelf, with a thickness of up to 400 m in the Skjoldryggen area, whereas an erosional morphology dominates the central to inner shelf.

Glacial processes and large-scale morphology on the m i d - N o r w e g i a n continental s h e l f

445

Fig. 3. Colour-shaded relief map with 20 m depth contours based on a 500 m grid cell size. LI = Lofoten Islands; V F = Vestfjorden; HB = Haltenbanken; SKD = Sklinnadjupet; TD = Tramadjupet; SD = Suladjupet; FB = Fr0yabanken; FR = Fr0yryggen; TB = Tramabanken; SR = Skjoldryggen.

Ice drainage offshore Trondelag The shallow parts of Haltenbanken have in certain periods probably prevented an active westward flow of the ice sheet. Several SW-NE-trending depressions (including Suladjupet, Fig. 3) indicate that ice mainly drained southwestwards, east of Haltenbanken, turning westwards across FrOyryggen north of FrOyabanken. The Suladjupet depression is more than 500 m deep and eroded 200-300 m below the surrounding sea bottom (Figs. 3 and 4). The depression was formed by glacial erosion, mainly of the dark, Upper Jurassic/Lower Cretaceous claystone of the Spekk Forma-

tion. IKU bedrock unit IX subcrops below a thin Quaternary cover at Fr0yryggen, and it is evident that this sandy unit has been resistant to glacial erosion. Prograding glacial sequences and a wide erosional depression are seen west of Fr0yryggen (Bugge, 1980; Bugge et al., 1987; Rokoengen et al., 1995).

Ice drainage offshore More The continental shelf offshore MOre is very narrow compared to the areas further north (Fig. 4). Several WNW-ESE-trending depressions are separated by shallow bank areas. We believe that these troughs have also been drainage routes for ice streams. Out-

446

D. O t t e s e n et al.

Fig. 4. Colour-shaded relief map of the shelf area off MOre with 20 m depth contours. On Mfil0yplatgtet large curved ridges dominate, while further north, depressions which extend from land to the shelf break separate shallower bank areas. The depressions are interpreted to have been drainage routes for ice streams during maximum expansion of the late Weichselian ice sheet. SD = Suladjupet; FB = Fr0yabanken; FR = Fr0yryggen; R F = Romsdalsfjorden; SF = Storfjorden; LG = Langgrunna; M P = Mfil0yplat~et; BG = Buagrunnen; BD = Breisunddjupet; OD = Onadjupet; N T = Norwegian Trench; K = Kristiansund;/~ = ,~.lesund.

side Smr (Fig. 4), the ice drainage is directed in a southwesterly direction, towards the northern part of the Storegga slide. Another ice stream has passed south of the Griptarane highs towards the northwest, and coalesced with the ice stream off SmNa (Fig. 4). Northwest of Romsdalsfjord, a NW-SE-trending depression (Onadjupet) ends in the Storegga slide area at the shelf break. Northwest of ,~lesund, another depression extends almost to the shelf break where it coalesces with the aforementioned ice-stream eroded channel (Fig. 4). Langgrunna is a large bank area that guided icestream flow both south and north of the bank (Fig. 4). Breisunddjupet forms a narrow, elongated depression, representing the continuation of the deep Storfjord drainage system onto the open shelf, and ends in the eastern part of Langgrunna (Fig. 4). This extended fjord

feature is very uncommon on the shelf and indicates special glacial conditions during its formation; for instance, erosion by very channelised ice flow in an area where the surroundings were covered by frozen based ice could produce such a feature. A possible tectonic origin has also been discussed (Rokoengen, 1980). South of Breisunddjupet, an ice-stream drainage route from the east onto the northern part of M~10yplatfiet can be inferred (Fig. 4). MgdCyplat~et is the southernmost bank area of the mid-Norwegian shelf, located close to the outlet of the Norwegian Trench. Thus, this area probably was influenced by the large Norwegian Trench Ice Stream from time to time (King et al., 1996), as the ice dramage from the mainland either was deflected in a northerly direction and/or partly assimilated by the Norwegian Trench Ice Stream. On MfilCyplatfiet (Fig. 4), arcuate ridges

Glacial

processes

and

large-scale

morphology

on the mid-Norwegian

indicate major halts during deglaciation (Rokoengen, 1980; Rise and Rokoengen, 1984; Rise et al., 1984). Ice-flow model From the present bathymetric data set (Figs. 1, 3 and 4) and earlier investigations on the Norwegian

continental

shelf

447

continental shelf (Rise and Rokoengen, 1984; King et al., 1987; Rokoengen et al., 1995; Sa~ttem et al., 1996; Vorren and Laberg, 1997), we have reconstructed a probable flow pattern of the western part of the Scandinavian ice sheet during the late Weichselian (Fig. 5). In addition we have used investigations from Antarctica as a basis for the model.

Fig. 5. Interpreted ice-flow model during the late Weichselian with ice streams flowing along the main offshore depressions/troughs. V F = Vestfjorden; H B -- Haltenbanken; S K D = Sklinnadjupet; T D - Tramadjupet; S B = Sklinnabanken; S D - Suladjupet; F B = Fr0yabanken; M P = Mfil0yplatfiet;N T - Norwegian Trench; T B - Tr~enabanken;L G - Langgrunna; S K - Skagerrak; T = Trondheim.

448

Extensive research has been carried out in West Antarctica during recent years in order to understand the ice-sheet dynamics of large, marine-based ice sheets (e.g. Shabtaie and Bentley, 1987). The emphasis has been on the large ice streams which drain about 90% of the West Antarctic ice sheet. These ice streams are fast moving parts of ice sheets, normally 300-500 km long, 50-80 km wide and with speeds of 300-700 m/year, whereas the surrounding ice sheet may have a speed of less than 10 m/year (Bindschadler et al., 1996). Generally, the ice streams are located in overdeepened troughs, often eroding several hundred metres below the surrounding seafloor. The glaciological setting of West Antarctica today can partly be compared to the situation on the midNorwegian shelf during late Weichselian time. Studies on the shelf areas in Antarctica have outlined both prograding and aggrading glacial sequences (e.g. Cooper et al., 1991; Latter and Cunningham, 1993), comparable to what we find on the mid-Norwegian shelf. In the Ross Sea, Shipp and Anderson (1997) have described glacial megaflutes and trough forms, both related to palaeo-ice streams across the Ross Sea. The largest ice stream followed the Norwegian Trench along the southern and western coast of Norway (Fig. 4), and ended where the ice calved in the Norwegian Sea west of Mfil~yplatfiet (King et al., 1996). The idea of an immense Skagerrak glacier flowing along the Norwegian coast was introduced by Helland (1885). For some years this theory was generally accepted, but later became more controversial or was even rejected (Andersen, 1964; Holtedahl, 1993). Investigations both in the northern North Sea (Rise and Rokoengen, 1984) and in the Skagerrak (Longva and Thorsnes, 1997), however, have demonstrated the ice movements along the trench and proven the existence of the ice stream. In the Vestfjorden/Tra~nadjupet area another major ice stream has flowed out to the shelf edge several times (Fig. 5). On the mid-Norwegian shelf, the location of the ice-stream drainage routes are mainly located between the shallow bank areas, such as Tra~nabanken, Haltenbanken, FrOyabanken, Buagrunnen and Langgrunna (Figs. 3-5). Grounded ice sheets are thought to have been responsible for depositing the prograding sequences. During the initial advance of the grounded ice, the inner shelf would have been heavily eroded and gently dipping glacial strata were probably deposited on the shelf. Ice streams carved broad depressions across the shelf and carried sediments directly to the continental shelf edge, thereby creating trough-mouth fans (Vorren and Laberg, 1997) and sheet-like prograding

D. Ottesen et al.

sequences (King et al., 1987). During interglacial periods, the shelf areas were starved of sediment and thus received little or no clastic sedimentation. A ckn owled g ements

We gratefully acknowledge the Norwegian Hydrographic Service for the regional bathymetric data set and IKU for access to seismic data. This paper benefited from reviews by Tom Bugge and Tore Vorren. The English language has been improved by David Roberts. References Andersen, B.G., 1964. Har Ja~ren v~ert dekket av en Skagerrakbre? Er "Skagerrakmorenen" en marin leire? Nor. Geol. Unders., 228: 5-11. Bindschadler, R., Vornberger, E, Blankenship, D., Scambos, T. and Jacobel, R., 1996. Surface velocity and mass balance of Ice Streams D and E, West Antarctica. J. Glaciol., 42(142): 461-475. Bugge, T., 1975. Kart med kystkontur og dybdekoter for den norske kontinentalsokkel. Cont. Shelf Inst. (IKU), Publ. 55, 21 pp. Bugge, T., 1980. Ovre lags geologi pfi kontinentalsokkelen utenfor MOre og Tr0ndelag. Cont. Shelf Inst. (IKU), Publ., 104, 44 pp. Bugge, T., Knarud, R. and M0rk, A., 1984. Bedrock geology on the mid-Norwegian continental shelf. In: A.M. Spencer, S.O. Johnsen, A. M0rk, E. Nys~ether, E Songstad and A. Spinnanger (Editors), Petroleum Geology of the North European Margin. Graham and Trotman, London, pp. 271-283. Bugge, T., Rise, L. and Rokoengen, K., 1987. Dybdekart over midtnorsk kontinentalsokkel. Mfilestokk 1" 1,000,000. Cont. Shelf Inst. (IKU), Publ., 115. Cooper, A.K., Barrett, P.J., Hinz, K., Traube, V., Leitchenkov, G. and Stagg, H.M.J., 1991. Cenozoic progradation sequences of the Antarctic continental margin: a record of glacio-eustatic and tectonic events. In: A.W. Meyer, T.A. Davies and S.W. Wise (Editors), Evolution of Mesozoic and Cainozoic Continental Margins. Mar. Geol., 102: 175-213. Eidvin, T., Brekke, H., Riis, F. and Renshaw, D., 1998. Cenozoic stratigraphy of the Norwegian Sea continental shelf, 64~176 Nor. Geol. Tidsskr., 78:125-151. Fronval, T., Jansen, E., Bloemendal, J. and Johnsen, S., 1995. Oceanic evidence for coherent fluctuations in Fennoscandian and Laurentide ice sheets on millennium timescales. Nature, 374: 443-446. Grootes, EM., Stuiver, M., White, J.W.C., Johnsen, S. and Jouzel, J., 1993. Comparison of oxygen-isotope records from the Gisp2 and Grip Greenland ice cores. Nature, 366: 552-554. Gunleiksrud, T. and Rokoengen, K., 1980. Regional geological mapping of the Norwegian continental shelf with examples of engineering applications. In: D.A. Ardus (Editor), Offshore Site Investigations. Graham and Trotman, London, pp. 23-35. Haflidason, H., Aarseth, I., Haugen, J.E., Sejrup, H.E, LOvlie, R. and Reither, E., 1991. Quaternary stratigraphy of the Draugen area, Mid-Norwegian Shelf. Mar. Geol., 101: 125-146. Helland, A., 1885. Om jordens 10se afleiringer. Meddelelse fra Den naturhistoriske Forening i Christiania, pp. 27-42. Henriksen, S. and Vorren, T., 1996. Late Cenozoic sedimentation and uplift history on the mid-Norwegian continental shelf. Global Planet. Change, 12: 171-199. Holtedahl, H., 1993. Marine geology of the Norwegian continental margin. Nor. Geol. Unders. Spec. Publ., 6, 150 pp. Holtedahl, H. and Sellevold, M.A., 1972. Notes on the influence of glaciation on the Norwegian continental shelf bordering on the Norwegian Sea. Ambio Spec. Rep. 2, 31-38.

Glacial processes and large-scale morphology on the mid-Norwegian continental shelf Jansen, E. and Sj0holm, J., 1991. Reconstruction of glaciation over the past 6 My from ice-borne deposits in the Norwegian Sea. Nature, 349: 600-603. King, E.L., Sejrup, H.R, Haflidason, H., Elverh0i, A. and Aarseth, I., 1996. Quaternary seismic stratigraphy of the North Sea Fan: glacially fed gravity aprons, hemipelagic sediments, and large submarine slides. Mar. Geol., 130: 293-315. King, L., Rokoengen, K. and Gunleiksrud, T., 1987. Quaternary seismostratigraphy of the Mid-Norwegian Shelf, 65~176 A till tongue stratigraphy. Cont. Shelf Inst. (IKU), Publ., 114, 58 pp. Larsen, E. and Sejrup, H.R, 1990. Weichselian land-sea interactions: Western Norway-Norwegian Sea. Quat. Sci. Rev., 9: 85-97. Larter, R.D. and Cunningham, A.R, 1993. The depositional pattern and distribution of glacial-interglacial sequences on the Antarctic Peninsula Pacific margin. Mar. Geol., 109: 202-219. Lien, R., 1983. P10yemerker etter isfjell p~ norsk kontinentalsokkel. Cont. Shelf Inst. (IKU), Publ., 109, 147 pp. Longva, O. and Thorsnes, T. (Editors), 1997. Skagerrak in the past and at the present an integrated study of geology, chemistry, hydrography and microfossil ecology. Nor. Geol. Unders. Spec. Publ., 8, 100 pp. Olsen, L., 1997. Rapid shifts in glacial extension characterise a new conceptual model for glacial variations during the Mid and Late Weichselian in Norway. Nor. Geol. Unders. Bull., 433: 54-55. Rise, L. and Rokoengen, K., 1984. Surficial sediments in the Norwegian sector of the North Sea between 60~ and 62~ Mar. Geol., 56:287-317. Rise, L., Rokoengen, K., Skinner, A. and Long, D., 1984. Nordlige Nordsj0. Kvart~ergeologisk kart mellom 60~ og 62~ og 0st for 1~ M 1: 500,000. Continental Shelf Institute (IKU) in cooperation with British Geological Survey (BGS). Rokoengen, K., 1980. De 0vre lags geologi p~ kontinentalsokkelen utenfor MOre og Romsdal. Beskrivelse til kvarta~rgeologisk kart 6203 i mfilestokk l" 250,000. Cont. Shelf Inst. (IKU), PUN., 105, 49 pp. Rokoengen, K. and Frengstad, B., 1999. Radiocarbon and seismic

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evidence of ice-sheet extent and the last deglaciation on the midNorwegian continental shelf. Nor. Geol. Tidsskr., 79: 129-132. Rokoengen, K., Rise, L., Bugge, T. and Sa~ttem, J., 1988. Berggrunnsgeologi pfi midtnorsk kontinentalsokkel. Kart i mfilestokk 1 1,000,000. Cont. Shelf Inst. (IKU), Publ., 118. Rokoengen, K., Rise, L., Bryn, R, Frengstad, B., Gustavsen, B., Nygaard, E. and S~ettem, J., 1995. Upper Cenozoic stratigraphy on the Mid-Norwegian continental shelf. Nor. Geol. Tidsskr., 75: 88-104. Sa~ttem, J., 1990. Glaciotectonic forms and structures on the Norwegian continental shelf: observations, processes and implications. Nor. Geol. Tidsskr., 70: 81-94. Sa~ttem, J., Rise, L., Rokoengen, K. and By, T., 1996. Soil investigations, offshore mid-Norway: a study of glacial influence on geotechnical properties. Global Planet. Change, 12:271-285. Shabtaie, S. and Bentley, C.R., 1987. West Antarctic ice streams draining into the Ross ice shelf: configuration and mass balance. J. Geophys. Res., 92(B2): 1311-1336. Shipp, S. and Anderson, J., 1997. Paleo-ice stream boundaries, Ross Sea, Antarctica. In: T. Davies, T. Bell, A. Cooper, H. Josenhans, L. Polyak, A. Solheim, M. Stoker and J. Stravers (Editors), Glaciated Continental Margins. An Atlas of Acoustic Images. Chapman and Hall, London, pp. 106-109. Sigmond, E.M.O., 1992. Berggrunnskart, Norge med havomrfider. Mfilestokk 1"3 millioner. Norges Geologiske Unders0kelse, Trondheim. Taylor, K.C., Lamorey, G.W., Doyle, G.A., Alley, R.B., Grootes, RM., Mayewski, RA., White, J.W.C. and Barlow, L.K., 1993. The "flickering switch" of Late Pleistocene climate change. Nature, 361: 432-436. Veum, T., Jansen, E., Arnold, M., Beyer, I. and Duplessy, J.C., 1992. Water mass exchange between the North-Atlantic and the Norwegian Sea during the past 28,000 years. Nature, 356: 783785. Vorren, %0. and Laberg, J.S., 1997. Trough mouth fans - - paleoclimate and ice-sheet monitors. Quat. Sci. Rev., 16(8): 865-881.

Geological Survey of Norway, N-7491 Trondheim, Norway Geological Survey of Norway, N-7491 Trondheim, Norway Norwegian University of Science and Technology, N-7034 Trondheim, Norway SINTEF Petroleum Research, N-7465 Trondheim, Norway Present address: Sauherad Kommune, N-3812 Akkerhaugen, Norway

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451

Late Quaternary sedimentary processes and environment on the Norwegian-Greenland Sea continental margins Tore O. Vorren and Jan Sverre Laberg

Our aim is to explain the morphology and late Quaternary sedimentary processes and palaeoenvironment of the Norwegian-Greenland Sea continental margin. In particular we concentrate on the North Norwegian and Barents Sea continental margin. Three main groups of morphological features will be discussed: slides, trough mouth fans and channels. Most of the present paper is an extended summary of Vorren et al. (1998). However, some new data concerning the Tra~nadjupet Slide are added.

Slides Several smaller and larger slides have been identified on the Norwegian continental margin (Fig. 1). Table 1 summarises data on dimensions of the largest slides according to our most recent results. The age of the slides varies from more than 200,000 years (Bj~rn~yrenna Slide) to probably less than 5000 years (Traenadjupet Slide). Some slide features are illustrated by the Tra> nadjupet Slide located on the continental slope just northeast of the VOring Plateau (Fig. 2). The area affected by the slide extends from the shelf break to more than 3000 m water depth in the Lofoten Basin, implying a run-out distance of ca. 200 km. The slide headwall is about 150 m high and 20 km long. The slide scar (area of evacuation) can be followed downslope to about 2400 m water depth, and covers

an area of about 5000 km 2. The slide deposits cover an area of about 9100 k m 2 and total slide-affected area was estimated to be about 14,100 km 2 (Table 1). Some of the morphological features like the escarpment, sediment blocks and flow features from the upper slide area are illustrated in Fig. 3. The large slides (Fig. 1) occurred in areas characterised by high sediment input. Thus, relatively high sediment supply leading to unstable sediments may have been important for the initiation of slope failures in these areas. Another important factor may be the presence of shallow gas, suggested by Knutsen et al. (1992) as important for the instability of sediments found on the central part of the Bear Island trough mouth fan. Studies of the present seismicity in the eastern Norwegian-Greenland Sea have shown relatively high activity along older fault systems (Kvamme and

TABLE 1 Dimensions of large slides on the Norwegian continental margin

Continental slope gradient Run-out distance (km) Maximum thickness (m) Height/length ratio b Slide scar area (kin 2) Total slide influenced area (km 2) Volume (km 3)

Traenadjupet Slide

Storegga Slide a

And0ya Slide

Bj0rnOyrenna Slide

1.25 ~ 200 150 0.0125 5000 14,500 900 d

0.6 ~ 850 430 0.004 34,000 112,500 5580

7~ 190 c

0.6 ~

0.0126 3630 c 9700 c

0.0063 6200 23,000 1100

400

a From Bugge (1983) and Bugge et al. (1987). b Height is the elevation difference between the top of the failed mass at the point of initiation and the top of the failed mass in the depositional zone, and length is the distance from the origin (Hampton et al., 1996 and references therein). c Based on Dowdeswell and Kenyon (1995). d Average thickness 100 ms (= 100 m assuming a seismic velocity of 2000 m/s). Sedimentary Environments Offshore Norway - Palaeozoic to Recent edited by O.J. Martinsen and T. Dreyer. NPF Special Publication 10, pp. 451-456, Published by Elsevier Science B.V., Amsterdam. 9 Norwegian Petroleum Society (NPF), 2001.

452

T.O.

Vorren and J.S. Laberg

Fig. 1. Bathymetric map showing location and extent of slides, small slides and large channels on the Norwegian-Greenland Sea continental margin. The figure is compiled from several sources such as Bugge et al. (1987); Mienert et al. (1993); Laberg and Vorren (1993) and Dowdeswell et al. (1996).

Hansen, 1989; Bungum et al., 1991). From other continental margin areas, earthquakes have been recognised as one of the most likely triggering mechanisms for submarine slides (e.g. Hampton et al., 1996). Thus, for the triggering of the large slides along the eastern Norwegian-Greenland Sea continental margin, earthquakes have probably been an important mechanism (Bugge, 1983; Bugge et al., 1987; Kenyon,

1987; Knutsen et al., 1992; Laberg and Vorren, 1993, 1996a; Evans et al., 1996). Decomposition of gas hydrates was also suggested as a possible triggering mechanism for the Storegga Slides (Bugge et al., 1987). Andreassen and Hansen (1995), using a phase boundary diagram for a methane hydrate system, infer that gas hydrates may be present under the Norwegian/Barents Sea continental slope

TABLE 2

Dimensions and slope gradients of trough mouth fans along the Norwegian-Greenland Sea continental margin

Radius (km) Width upper (km) Width lower (kin) Depth upper (kin) Depth lower proximal (kin) Depth lower distal (km) Area (kin 2) Gradient (upper) Gradient (middle) Gradient (lower)

Kongsfjorden

Isfjorden

Bellsund

Storfjorden

Bear Island

North Sea

Scoresby Sund

TMF

TMF

TMF

TMF

TMF

TMF

TMF

55 40 60 0.2

50 45 75 0.25

70 55 85 0.15

2.0 2700

3.0 3700

2.3 6000

3.2 ~

1.8 ~

590 250 550 0.5 3.0 3.2 215,000 0.8 ~ 0.4 ~ 0.2 ~

560 165 300 0.4 2.7 3.5 142,000 0.6 ~ 0.8 ~ 0.3 ~

110 180 240 0.3 1.5 1.5 19,000

1.9 ~

190 130 210 0.4 2.4 2.7 35,000 1.8 ~ 1.0 ~ 0.2 ~

2~

Late Quaternary sedimentary processes and environment on the Norwegian-Greenland Sea continental margins

453

Eig. 2. Map of the Traenadjupet Slide area. The dotted line within the slide separates the upslope slide scar area and the downslope area of slide deposition.

today. Increased bottom water temperature and/or lowering of sea level would cause destabilisation of the hydrate zone (McIver, 1982). To summarise, high sedimentation rates, which in turn may have led to a build-up of excess pore water pressure, and perhaps with additional pressure caused by gas bubbles, probably led to unstable or metastable

sediments within relatively large parts of the eastern Norwegian-Greenland Sea continental margin. Destabilising and triggering may have been prompted by earthquakes or perhaps by decomposition of gas hydrates.

T.O. Vorren and J.S. Laberg

454

Fig. 3. Mosaic of TOBI side-scan sonar data covering the western, upper Tra~nadjupet Slide. Escarpments and sediment blocks are indicated. For location, see Fig. 2.

Trough mouth fans and debris flows

On the continental margins surrounding the Norwegian-Greenland Sea, fan- or delta-like protrusions

occur in front of many of the glacial troughs or channels crossing the continental shelf and ending on the shelf break (Fig. 1). Nansen (1904) noted many of these protrusions. Vogt and Perry (1978) pointed

TABLE 3 Dimensions of individual debris flows on trough mouth fans in the Norwegian-Greenland Sea

Width (kin) Thickness (m) Length (kin) Area (km 2) Volume (km 3)

Isfjorden TMF

Storfjorden TMF

Bear Island TMF

North Sea TMF

Scoresby Sund TMF

2-5 10-30 10-20 < 100 0.5-1

1-5 15 50-100 <500 3-8

3-24 5-50 100-200 850-1880 10-32

2-40 (5) <60 (15-30)

0.5-2 5-15

10-50

Late Quaternary sedimentary processes and environment on the Norwegian-Greenland Sea continental margins

out that these protrusions are probably prograded deltas and attached fans. Vorren et al. (1988, 1989) proposed naming these features, which also occur on other glaciated margins, "trough mouth fans" (TMF). The TMFs vary in size and shape (Table 2). In recent years the fans have been investigated with regard to their architecture, sedimentological processes and origin, and their potential as palaeoclimatic archives. At least during the late Quaternary, the TMFs have been the sites of intense debris flow activity. Damuth (1978) was the first to indicate their presence. Vorren et al. (i988, 1989) found that the debris flows, seen in cross-section, are bundled in sets of lenses separated by high-amplitude reflections, and pointed to their potential as a palaeoclimatic proxy. They suggested that in most cases each of the lobes in the lens-set represents a single debris flow originating on the upper slope, emplaced during periods when the glacier grounding line was near the shelf break. The high-amplitude reflections between each lens-set probably represent periods of low sediment input/erosion during interstadials or interglacials. Later mapping by seismic and by side-scan sonars has confirmed that the debris flows indeed are the main building blocks, and palaeoclimatic recorders, of the younger part of the TMFs in the NorwegianGreenland Sea (Vogt et al., 1993; Laberg and Vorren, 1995, 1996a,b; King et al., 1996; Sejrup et al., 1996; Dowdeswell et al., 1996, 1997). The dimensions of the debris flows vary from fan to fan. The width varies between 0.5 and 40 km, thickness between 5 and 60 m, length from less than 10 km up to 200 km, the areas covered are up to 1880 km 2 and the volumes from 0.5 to 50 km 3 (Table 3). There is a clear tendency towards larger fans having the larger and more voluminous debris flows (Table 3). Channels The observed channels on the Norwegian margin can be grouped into two types, according to their size and morphology: gullies on the upper and middle continental slope, and large channels on the continental slope, rise, and in the deep sea. Two large deep-sea channel systems have been observed along the Norwegian continental margin, the Lofoten Basin Channel and the Inbis Channel. Along the East Greenland margin, several channel systems have been identified (Fig. 1). The deep-sea channels may have been formed by dense water originating from cooling, sea-ice formation and brine rejection close to the glacier margin or they may originate from small slides on

455

the upper slope transforming into debris flows and turbidity currents. References Andreassen, K. and Hansen, T., 1995. Inferred Gas Hydrates Offshore Norway and Svalbard. Dr. Scient. Thesis, University of Troms0. Bugge, T., 1983. Submarine slides on the Norwegian continental margin, with special emphasis on the Storegga area. Cont. Shelf Pet. Res. Inst. Publ., l l0, 152 pp. Bugge, T., Belting, S., Belderson, R.H., Eidvin, T., Jansen, E., Kenyon, N.H., Holtedahl, H. and Sejrup, H.R, 1987. A giant threestage submarine slide off Norway. Geo-Mar. Lett., 7: 191-198. Bungum, H., Alsaker, A., Kvamme, L.B. and Hansen, R.A., 1991. Seismicity and seismotectonics of Norway and nearby continental shelf areas. J. Geophys. Res., 96(B2): 2249-2265. Damuth, J.E., 1978. Echo character of the Norwegian-Greenland Sea: relationship to Quaternary sedimentation. Mar. Geol., 28: 1-36. Dowdeswell, J.A. and Kenyon, N.H., 1995. Cruise Report RRS James Clark Ross Cruise 08 22 July to 1 September 1994. University of Aberystwyth, 50 pp. Dowdeswell, J.A., Kenyon, N.H., ElverhN, A., Laberg, J.S., Hollender, EJ., Mienert, J. and Siegert, M.J., 1996. Large-scale sedimentation on the glacier-influenced polar North Atlantic margins: long-range side-scan sonar evidence. Geophys. Res. Lett., 23: 3535-3538. Dowdeswell, J.A., Kenyon, N.H. and Laberg, J.S., 1997. The glacierinfluenced Scoresby Sund Fan, East Greenland continental margin: evidence from GLORIA and 3.5 kHz records. Mar. Geol., 143: 207-221. Evans, D., King, E.L., Kenyon, N.H., Brett, C. and Wallis, D., 1996. Evidence for long-term instability in the Storegga Slide region off western Norway. Mar. Geol., 130: 281-292. Hampton, M.A., Lee, H.J. and Locat, J., 1996. Submarine landslides. Rev. Geophys., 34: 33-59. Kenyon, N.H., 1987. Mass-wasting features on the continental slope of north west Europe. Mar. Geol., 74: 57-77. King, E.L., Sejrup, H.R, Haflidason, H., Elverhoi, A. and Aarseth, I., 1996. Quaternary seismic stratigraphy of the North Sea Fan: glacially-fed gravity flow aprons, hemipelagic sedimentation, and submarine sliding. Mar. Geol., 130:293-315. Knutsen, S.-M., Richardsen, G. and Vorren, T.O., 1992. Late Miocene-Pliocene sequence stratigraphy and mass-movements on the western Barents Sea margin. In: T.O. Vorren, E. Bergsager, O.A. Dahl-Stamnes, E. Holtar, B. Johansen, E. Lie and T.B. Lund (Editors), Arctic Geology and Petroleum Potential. Nor. Pet. Soc. Spec. Publ., 2: 573-606. Kvamme, L.B. and Hansen, R.A., 1989. The seismicity in the continental margin areas of Northern Norway. In: S. Gregersen and RW. Basham (Editors), Earthquakes at North-Atlantic Passive Margins: Neotectonics and Postglacial Rebound. Kluwer, Dordrecht, pp. 429-440. Laberg, J.S. and Vorren, T.O., 1993. A Late Pleistocene submarine slide on the Bear Island Trough Mouth Fan. Geo-Mar. Lett., 13: 227-234. Laberg, J.S. and Vorren, T.O., 1995. Late Weichselian submarine debris flow deposits on the Bear Island Trough Mouth Fan. Mar. Geol., 127: 45-72. Laberg, J.S. and Vorren, T.O., 1996a. The Middle and Late Pleistocene evolution of the Bear Island Trough Mouth Fan. In: A. Solheim, F. Riis, A. ElverhN, J.L. Faleide, L.N. Jensen and S. Cloetingh (Editors), Impact of Glaciations on Basin Evolution: Data and Models from the Norwegian Margin and Adjacent Areas. Global Planet. Change, 12: 309-330. Laberg, J.S. and Vorren, T.O., 1996b. The glacier fed fan at the mouth of Storfjorden Trough, western Barents Sea: a comparative study. Geol. Rundsch., 85: 338-349.

T.O. Vorren and J.S. Laberg

456 McIver, R.D., 1982. Role of naturally occurring gas hydrates in sediment transport. Am. Assoc. Pet. Geol. Bull., 66: 789-792. Mienert, J., Kenyon, N.H., Thiede, J. and Hollender, F.J., 1993. Polar continental margins: studies off East Greenland. EOS, Trans. Am. Geophys. Union, 74(20): 225-236. Nansen, E, 1904. The bathymetrical features of the north Polar Seas, with a discussion of the continental shelves and previous oscillations of the shore-line. In: F. Nansen (Editor), The Norwegian North Polar Expedition 1893-1896 Scientific Results, IV. pp. 1232. Sejrup, H.E, King, E.L., Aarseth, I., Haflidason, H. and Elverhr A., 1996. Quaternary erosion and depositional processes: Western Norwegian fjords, Norwegian Channel and North Sea Fan. In: M. DeBatist and E Jacobs (Editors), Geology of Siliciclastic Shelf Seas. Geol. Soc. Spec. Publ., 117:187-202. Vogt, ER. and Perry, K.K., 1978. Post rifting accretion of continental

T.O. VORREN J.S. LABERG

margins in the Norwegian-Greenland and Labrador Seas: morphologic evidence. EOS, Trans. Am. Geophys. Union, 59: 120. Vogt, ER., Crane, K. and Sundvor, E., 1993. Glacigenic mudflows on the Bear Island submarine fan. EOS, Trans. Am. Geophys. Union, 74: 449; 452-453. Vorren, T.O., Hald, M. and Lebesbye, E., 1988. Late Cenozoic environments in the Barents Sea. Paleoceanography, 3:601-612. Vorren, T.O., Lebesbye, E., Andreassen, K. and Larsen, K.-B., 1989. Glacigenic sediments on a passive continental margin as exemplified by the Barents Sea. Mar. Geol., 85:251-272. Vorren, T.O., Laberg, J.S., Blaume, F., Dowdeswell, J.A., Kenyon, N.H., Mienert, J., Rumohr, J. and Werner, F., 1998. The Norwegian-Greenland Sea continental margins: morphology and late Quaternary sedimentary processes and environment. Quat. Sci. Rev., 17: 273-302.

Department of Geology, University of TromsO, N-9037 TromsO, Norway Department of Geology, University of TromsO, N-9037 TromsO, Norway

457

Accretionary, forced regressive shoreface sands of the Holocene-Recent Skagen Odde spit complex, Denmark-- a possible outcrop analogue to fault-attached shoreface sandstone reservoirs Lars Henrik Nielsen and Peter Niels Johannessen

Spit systems are often important components of Recent deltas and coastal plains, where waves redistribute fluvial-derived sediments and they show large variations in terms of size, volume and grain size. However, spit systems are likely to form along any wave-influenced coastline and may form where faulted terrains of poorly to moderately lithified rocks are exposed to wave erosion. In such areas large spit systems are likely to be formed down-current from fault escarpments, along fault crests and on fault ramps, but appear to have been overlooked in the geological record. The Holocene-Recent Skagen Odde spit complex has developed within the last 8000 years and is still growing. The most important factors that have governed its growth are well-known, including climate, sea-level changes, palaeo-relief, free fetch, tidal range, progradation and sediment transport rates. The spit system has developed down-current from a steep slope comparable to a fault escarpment. It shows a triangular form with sides of 30-40 km and it contains 5-10 km s of sand. The main part of the spit system consists of five sedimentary units that constitute a shallowing-upward shoreface succession: (lower) storm sand; bar-trough sand; beach sand; peat and (upper) modern aeolian sand. The succession was formed during a forced regression as uplift outpaced eustatic rise. As the average basin-floor gradient has been steeper than the shoreline trajectory, the accommodation increased during basinward progradation of the spit system, and the resulting succession represents an accretionary, forced regression. The spit-system model may serve as an analogue to ancient fault-related shoreface sandstones without associated fluvial or deltaic deposits.

Introduction

Geologica! models are mandatory for focusing exploration efforts and optimisation of exploitation of proven hydrocarbon fields. The occurrence of potential exploration targets and the architecture of deeply buried reservoirs is, however, in many cases difficult to predict or interpret in basins with poor constraints on the distribution of sea and land, palaeorelief, sea-level fluctuations, climate, sediment transport rates, and duration of formation of the reservoir. The shoreface sands of the raised Holocene-Recent Skagen spit complex is compared to the Late Pleistocene Lyngsfi spit system which developed on a submerged glacial fault ridge (Nielsen et al., 1988), and together the two spit systems may serve as a well-constrained, full-scale analogue to deeply buried reservoirs developed down-current from partly submerged faulted terrains exposed to wave erosion. The Skagen spit complex forms the northernmost part of Jylland, and the Lyngs~ spit system is located on a raised marine plateau approx. 10 km south of the baseline of the Skagen spit (Fig. 1).

A spit system is attached to a land mass at one end and terminating in open water at the other, and is younger than the land mass (Evans, 1942; Meistrell, 1972; Nielsen et al., 1988). The baseline defines the line along which a spit system is attached to a land mass, and the spit point is the distal termination of the spit system. A spit system consists of a spit platform which is a large-scale primary sedimentary structure formed by sediment transport along the coast and rises above the shelf but lies below mean sea tide, and a spit which is a sediment ridge on the spit platform, partly elevated over mean low tide, and follows the formation of the spit platform. Spit systems develop where waves and wave-induced currents favour shoreparallel sediment transport. They are most common in microtidal settings as strong tidal currents favour sediment transport perpendicular to the coast. They occur both along low-energy and high-energy coasts and can display grain sizes varying from very finegrained muddy sand to pebbles depending on energy and available material. Commercial amounts of hydrocarbon are commonly found in shallow marine sandstones, and the spit-system model may offer an

Sedimentary Environments Offshore Norway - Palaeo~.oic to Recent edited by O.J. Martinsen and T. Dreyer. NPF Special Publication 10, pp. 457-472, Published by Elsevier Science B.V., Amsterdam. @ Norwegian Petroleum Society (NPF), 2001.

458

L.H. Nielsen and P.N. Johannessen

Fig. 1. Map showing the location of the Skagen Odde spit complex at the northernmost tip of Denmark, its present-day coastline, its baseline, four reconstructed positions of the palaeo-coastline, and the position of the Skagen 3/4 core well (modified from Hauerbach, 1992).

alternative interpretation of the genesis, orientation and geometry of fault-attached shoreface sandstones, where associated fluvial or deltaic deposits are missing (Nielsen and Johannessen, 1999).

Geological setting Large areas in North Jylland were transgressed by the sea as the last glacial ice cover melted resulting in an archipelago of small glacial islands at ca. 14,650 BP (Jessen, 1899, 1936; Tauber, 1966; Krog and Tauber, 1974; Knudsen, 1978; Nielsen et al., 1988; Richardt, 1996). Subsequently, the area was uplifted owing to glacial rebound at a faster rate than the contemporaneous eustatic rise and a general sea-level fall followed. The fall was punctuated by short periods of stable relative sea levels when uplift was balanced by eustatic rise. Several spit systems were formed within a few hundred years during such a period in eastern Vendsyssel (ca. 13,100-12,600 BP), and sharp-based regressive shoreface sands were deposited in northern Vendsyssel (Zirphae beds; ca. 12,500-12,000 BP). The regression probably reached a temporary maxi-

mum at ca. 10,000 BP when large parts of the southern North Sea emerged. In North Jylland the coastline was situated a short distance basinward of the baseline of the Skagen spit complex (Fig. 1). The general regression was quickly interrupted by the rapid Early Atlantic transgression that peaked at ca. 8000 B P (Petersen, 1991), and lagoons and fjords were established close to the baseline behind partly submerged barriers. After this short event, the general regression continued and the large regressive Skagen spit complex was formed during the last ca. 8000 years (Jessen, 1899, 1936; Schou, 1949; Hauerbach, 1992).

The Skagen Odde spit complex The Holocene-Recent Skagen spit complex in northernmost Denmark is among the largest in the world (Fig. 1). It is attached to the main land along a steep slope up to 70 m high comparable to fault escarpments, and the spit terminates in deep water at the spit point. Due to glacial rebound the proximal part of the spit complex is raised up to 13 m above present sea level. The subaerial part of the spit

Accretionary, forced regressive shoreface sands of the Holocene-Recent Skagen Odde spit complex, Denmark

complex covers approx. 400 k m 2, and the thickness of the spit-platform succession ranges from a few metres to more than 25 m (excluding the variable aeolian cover) and it forms a marine sand-dominated reservoir body with a volume of 5-10 km 3. The spit complex has a large triangular form with a W N W ESE-trending baseline of 30 km, a west side of 40 km and an east side of 35 km (Fig. 1). The baseline along the steep slope forms the boundary between the spit complex and the Pleistocene glacial and late glacial landscape. The two other sides are the current coastlines of which the western is a retreating coastal cliff exposing the raised spit-complex deposits and the eastern is a prograding coastal plain. Both sides pass northward into the active spit coast. Preserved beach ridges, C-14 dates from shells and from peats formed in lows between successive ridges, archaeological remains and old maps allow a detailed reconstruction of various stages of the spit complex development (Fig. 1) (Hauerbach, 1992). Its initial formation at ca. 8000 B P was supported by partly submerged barriers of older sediments, and it developed as a cuspate foreland receiving sediments from both the west and east coasts of Jylland (Jessen, 1936; Schou, 1949; Fredericia, 1988; Petersen, 1991). At this stage lagoons, small spits and beach ridges were formed in its western and southern part close to the baseline at U ggerby and Tversted (Fig. 1). As a result of the northeasterly progradation of the spit system and the overall regression forced by the relative sealevel fall, a large lagoon was formed to the east. Further lowering of sea level caused progradation of curved sandy beach ridges that eventually enclosed the lagoon transforming it into a lake (Gftrdbo SO, Fig. 1). During the following 4000 years the spit system reached its present shape with extensive marine foreland with sand ridges in the eastern part. The spit complex thus includes several coastal elements that have developed and amalgamated during the last ca. 8000 years. The present paper mainly deals with the large northeast-trending recurved spit that constitutes the medial to distal part of the spit complex formed within the latest ca. 5400 years (Fig. 1).

Controlling factors of spit-complex development Since the Skagen spit complex began to form at ca. 7600 BP (Hauerbach, 1992; Knudsen et al., 1996), sea-level change has been one of the most important factors governing its development. The spit complex prograded seaward while the area experienced isostatic uplift that outpaced eustatic rise. Based on the gradients of raised beach deposits formed during the highest Holocene sea level (Mertz, 1924) it is esti-

459

Fig. 2. Reconstruction of average palaeo-slopes. Holocene beach ridges occur ca. 13 m above present sea level at the baseline of the spit complex, the core well at the present distal point of the Skagen spit complex shows 100 m of sediments younger than ca. 7600 years (Knudsen et al., 1996), and the sea-level fall at the position of the present point of the spit is estimated to ca. 25 m based on the shoreline gradient from Mertz (1924).

mated that during the lifetime of the spit complex, the relative sea level has fallen ca. 13 m along the baseline of the spit system and ca. 25 m at the position of the present distal point of the spit due to differential uplift. The spit complex has therefore experienced a forced regression during its formation. However, the loss of accommodation space caused by the relative sea-level fall was compensated as the spit system prograded into increasingly deep water on a steeply inclined sea bottom, that precluded sediment bypass across an erosional surface. The palaeo-water depth at the present spit point at the time of initial spit formation is estimated to ca. 125 m based on the presence of 100 m of mud and sand deposited within the last 7600 years shown by cores from the Skagen 3/4 well (Knudsen et al., 1996) plus 25 m of sea-level fall. The average gradient of the sea bottom at the time of spit initiation was thus ca. 1 : 280, while the gradient of the spit coastline trajectory is 1:2700 (Fig. 2). The present circulation pattern in the Skagerrak which greatly influences the sediment transport pattern along the spit system consists primarily of the northeastward flowing Jylland Current and the outflow of less-saline water from the Baltic Sea following the Swedish and Norwegian coasts (Fig. 3). The two currents combine into the southwest flowing Norwegian Coastal Current. This circulation pattern was possibly established shortly after the rapid Holocene sea-level rise at ca. 7600 BP due to opening of the English Channel (Van Weering, 1975; Nordberg, 1992; Knudsen et al., 1996). The present interference between the North Sea amphidromic systems reduces the tidal range to less than 0.3 m at Skagen. A change in faunas and lithology at ca. 5500 BP shown by the Skagen 3/4 cores suggests that the large shallow grounds of glacial de-

460

L.H. Nielsen and RN. Johannessen 10~

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in the North

Sea. Widths

of arrows

indicate

magnitude

of volume

transport

(modified

from

Turrell

et al., 1 9 9 2 ) .

posits, the Jylland Bank west of the present coastline of Jylland (Fig. 3), became submerged at this time allowing the development of a more pronounced Jylland Current, which probably reduced the tidal range to the present amount (Knudsen et al., 1996; Leth, 1996). Since this time tidal changes have not influenced the development of the spit complex as much as meteorological events. Strong westerly storms may rise the sea level up to 1.5 m or more. The change from interbedded silt and clay to sand and silt at ca. 5500 BP in the Skagen 3/4 cores probably reflects that a continuous coastline from the mainland toward the spit point was now established, allowing a high rate of sediment transport along the west coast toward the spit point (Fig. 1).

The spit system has developed in a warm temperate and maritime climate. The wind system has been closely monitored since 1860 and is dominated by (1) the passage of cyclones with winds from the south and southwest followed by winds from the west and northwest, and (2) high-pressure cells situated over the Scandinavian Peninsula or the European continent causing winds from the north, northeast, east or southeast (Kristensen and Frydendahl, 1991; Anthonsen et al., 1996). The growth of the spit system has especially been facilitated by southwesterly and westerly winds setting up strong waves that cause longshore drift of sand toward the northeast. The free fetch from these directions has been 800-900 km. The transport capacity of sand along the shore of the

Accretionary, forced regressive shoreface sands of the Holocene-Recent Skagen Odde spit complex, Denmark

461

Fig. 4. Air photo showingthe northern tip of Skagen Odde with various stages in the location of the north cost indicated. New light houses were built as the spit coast prograded northeastward. The averagediameter of pebbles along the present coastline is indicated.

present-day spit has been monitored since 1954 and is calculated to vary between 950,000 and 1,450,000 m3/year along the northwestern coast of the spit showing a marked decrease to one fourth or fifth at the point where the coastline changes from a S W NE trend to a W - E trend (Hansen and Langballe, 1993). The reduced transport capacity causes rapid deposition along the north coast as reflected by rapid progradation of the accretional northern spit coast estimated to average to more than 8 m/year during the latest 300 years (Fig. 4) (Foster and Jensen, 1990; Hauerbach, 1992). Estimates based on C-14 dates indicate an average progradation rate of the spit complex of 4-5 m/year since its start.

Sedimentary facies A sedimentological investigation of the cliff sections (loc. 1, 2, 4 and 5 Fig. 1) exposing the medial part of the large regressive spit system has revealed four principal sedimentary units occurring in consistent stratigraphic order forming a shallowing-upward sandy shoreface succession terminated by aeolian sand (Figs. 5 and 6). The units termed 1-4 are described in ascending order. The shoreface sands are overlying muddy facies which occur below the exposed sections and are thus only known from wells. The muddy facies and the aeolian sand are not dealt with here.

Unit 1: storm sand beds

The lowest exposed unit consists of fine-grained, occasionally medium-grained sand beds, 12-60 cm thick, typically 20-30 cm with marine shells (Fig. 7a). The bed boundaries are sharp or erosional, occasionally lined with scattered pebbles. The beds show a parallel-laminated and weakly bioturbated lower part grading upward into an upper totally bioturbated part. The majority of the beds are non-graded; few beds show a weak fining-upward trend. The uppermost bioturbated 5-15 cm, however, often shows a slightly upward-increasing content of mud up to a few percentages. Occasionally, the beds are terminated by mud drapes up to a few centimetres thick. The sand laminae in the lower part of the beds are typically 13 cm thick, non-graded and well-sorted occasionally with few scattered marine shells and small pebbles. Cross-laminated and wave-rippled sand layers and thin mud layers occur sporadically within the sand beds. Skolithos and escape traces are seen occasionally. The upper fully bioturbated part shows primarily burrows of heart urchins, occasionally mud-lined. Skolithos and other vertical to oblique burrows are also observed. In addition, large compound burrows, 10-15 cm in diameter and typically 5-25 cm long, but occasionally up to 60 cm long, with steep to vertical walls occur in most of the beds. The compound burrows contain

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Fig. 5. Sedimentological section from loc. 1 (Fig. 1).

Accretionary, forced regressive shoreface sands of the Holocene-Recent Skagen Odde spit complex, Denmark

Fig. 6. Sedimentological section from loc. 4 (Fig. 1).

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Fig. 7. (a) Sharp-based storm sand bed. The lower part is parallel-laminated, the upper part is completely bioturbated by sea urchins. A large compound burrow probably formed by a lobster is also seen. Lens cap for scale. (b) Parallel-laminated to low-angle cross-bedded beach sand with scattered pebbles, diffuse bioturbation and Skolithos burrows. Ruler, 20 cm long for scale. (c) Bar-trough deposits showing planar and trough cross-bedding. Shovel, 1 m long for scale.

mostly fine-grained sand, but the concave-up bottom of the burrows is often mantled with small pebbles and mud clasts. Within the compound burrows a variety of mud-lined burrows occur, including Skolithos, which make the fill of the compound burrows more muddy than the surrounding sand. The sand beds are interpreted to represent storm, waning storm and fair weather deposits. The sharp bases represent a sudden change in the environment, occasionally including reworking caused by storms. The overlying laminated sand was mainly deposited from heavily loaded suspension clouds of sand subjected to traction currents. The scattered wave tipples and cross-lamination were formed by relatively weak oscillatory and unidirectional currents. The escape traces reflect the upward movements of organisms trying to adjust to the aggrading seafloor. During fair weather conditions after the storms, the upper parts of

the waning storm deposits and overlying fair weather deposits were bioturbated by heart urchins and other animals. Similar burrows are described from contemporaneous shoreface sand from the Island of L~esO in the Kattegat east of northern Jylland (Fig. 3; Hansen, 1977). The large compound burrows were probably formed by large crustaceans (e.g. lobster burrows), which during the storms tried to sustain their burrows as testified by the lining of pebbles along the bottom part of the burrows. After the death or disappearance of the crustacean, the burrow was filled with sand and mined by other organisms forming the numerous mud-lined burrows within the compound burrow.

Unit 2: bar-trough deposits The storm sand beds are overlain by bar-trough deposits with some interfingering. Unit 2 is 4.35-4.6

Accretionary, forced regressive shoreface sands of the Holocene-Recent Skagen Odde spit complex, Denmark m thick and consists primarily of planar and trough cross-bedded, fine-grained sand beds, 10-50 cm thick with erosive bases. The lower part of the cross-beds consist occasionally of medium- to coarse-grained sand, with small pebbles in the foresets and on the bases (Fig. 7c). Several cross-beds show preserved topsets; planar cross-beds show occasionally well-preserved bottomsets and topsets. Especially the planar cross-beds contain occasionally trough crosslamination, wave ripples and back-flow ripples within the foresets and bottomsets. The cross-beds in the lower part of the unit are totally bioturbated in the top by heart urchins (Figs. 5 and 6). Skolithos are also common and burrows of Diplocraterion and escape burrows occur occasionally. Large compound burrows similar to those in the underlying storm sand beds are also seen. The degree of bioturbation increases upward in each cross-bed, but shows an overall decrease upward in the unit. The cross-beds in the upper ca. 3-4 m of the unit contain very small diffuse burrows, giving the cross-beds a vague appearance. The burrows are 1-2 mm in diameter, and less than 1 cm long. Shells are only found in the lower part of the unit. The planar cross-beds represent straight-crested to slightly curved bars. Measurement of palaeo-current directions mainly indicate migration towards the west-southwest suggesting a slightly more curved shoreline at the spit point at approx. 5000 B P than shown on the reconstruction by Hauerbach (1992) in Fig. 1. The cross-beds separated by pebble-strewn surfaces resemble structures formed by recent coastattached oblique bars as shown in box cores samples (Hunter et al., 1979). During pauses in bar migration, current ripples and wave ripples were formed on the bar surfaces by weak unidirectional and oscillatory currents. The back-flow ripples show that back-flow eddies occasionally were established. The preserved bottom- and topsets indicate that large amounts of sediment were transported in suspension and only minor erosion of the bar tops occurred. The burrows show that the bars were inhabited. The trough cross-beds represent 3D mega-ripples, which primarily migrated towards the east, southeast and northeast as shown by palaeo-current directions measurements. The mega-ripples were driven by longshore currents in wide troughs between the bars. The intensively bioturbated cross-beds indicate periods of decreased current velocities, where the mega-ripples were relatively stationary allowing colonisation of burrowing animals. Escape traces show that some burrowing animals were able to move up through the sediment, while the mega-ripples were migrating. The diffuse bioturbation may represent activity of amphipods.

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Unit 3: beach sand Unit 3 overlies the bar-trough sediments with some interfingering and is 2-3.6 m thick. It consists primarily of low-angle cross-bedded and parallellaminated, fine- to medium-grained sand (Fig. 7b). Scour-fill structures and wide, flat troughs, 2 m wide and 10 cm deep, occur in the middle part of the unit. Scattered pebbles occur. Indistinct bioturbation similar to that in the uppermost part of the bar-trough deposits is seen. A large pine trunk has been found in the unit (Hauerbach, 1992). Roots are present in the upper 1-1.2 m of the unit showing an upwardincreasing abundance and obliteration of the primary structures below the overlying peat. The low-angle cross-bedded and parallel-laminated sand was deposited in the swash-backwash zone and on modified swash bars (e.g. Thompson, 1937; Hunter et al., 1979). The scour fills, flat troughs and cross-lamination were formed in shallow scours and runnels between beach berms. The trunk represents drift wood deposited by waves on the beach. The uppermost 0.5-1.0 m of the unit may represent aeolian reworked sand or small dunes on the beach, similar to the recent thin aeolian sand cover occurring below the peat swamps on the active spit coast to the north. Unit 4: peat Unit 4 overlies beach sand and is composed of dark-brown organic-rich material with thin sand layers (Figs. 5 and 6). It forms isolated lenses, up to 1-1.5 m thick and up to ca. 100 m wide. Trunks and branches of birch, up to 10 cm in diameter, are common. Roots from the peat protrude down into the beach sands. The peats represent in situ accumulations of organic matter in low-lying areas on the coastal plain presumably between successive beach ridges. Today, small peat swamps are established about 200 m from the active spit coast to the north.

Dynamics of spit systems Based on wave tank experiments (Meistrell, 1966, 1972), descriptions of Recent spit systems (e.g. Kumar and Sanders, 1974; Komar, 1976; Hine, 1979) and well-exposed raised Pleistocene spit systems, Nielsen et al. (1988) developed a three-dimensional morphological-sedimentological model for spit systems. When waves set up longshore drift of a sufficient amount of sediment, a mainland beach will be extended by the formation of a platform structure, which basically is composed of foresets and

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topsets beds. The platform is prograding by deposition of steep avalanche foresets beds or gently dipping shoreface beds depending on available grain sizes and energy. Tidal currents in inlets in front of the platform may modify the platform foresets. During progradation of the platform, the water depth above the platform will remain constant irrespective of irregularities in shelf topography, and a partially subaerial spit ridge is formed on the platform. In the constructive phases of the development of a spit system, obliquely incoming waves create an oblique bar-trough system on the platform in the surf zone along the seaward side of the spit (Fig. 8). Sand is transported forward to the edge of the platform by onshore migration of bars and alongshore migration of dunes in the troughs driven by currents. Pebbles

are transported in the swash-backwash zone along the spit coast. Deposition will mainly occur where the coast turns away from the sea and refraction of the waves results in reduction of wave energy per unit length of the coast. At the same time, the currents running parallel to the coast will expand over the platform, as the controlling effect of the subaerial spit gradually ceases. This combination causes high sedimentation rates on the platform along the end of the spit, and the refracted bars will become wider and longer and the interjacent troughs shallower. This modification of bars and troughs will gradually increase as the spit coast further bends away from the sea. From the point on the spit coast where the bartrough system begins to bend around the tip of the spit and to the point where the bar-trough system is

Proximal bar-trough system Bar has steep slipface

3D mega-ripples in trough

Swash bar

Distal bar system Bar has low angle slipface / /

"-

Shoreline of spit

Bar slipface migration direction

0

500

I

I

/ Beach ridges

1000 m i

A Beach ridge Marsh

0

100

Sho rehne'

200 m

Proximal bar-trough system Bar has steep slipface

Distal bar system Bar has low angle slipface I

B

Wave-ripples

.....

Fig. 8. Schematic map view of the present Skagen spit point showing welding of oblique bar-trough systems to the spit coast (upper). Schematic cross-section A - B showing emerged beach ridges with marsh formation in adjacent lows and active submarine bar-trough systems approaching the coast (lower). Internal structures are inferred.

Accretionary, forced regressive shoreface sands of the Holocene-Recent Skagen Odde spit complex, Denmark losing its significance, successive bars will migrate towards and weld onto the spit coast and contribute to the progradation of the spit point. When bars migrate up the shoreface, they will emerge and become swash bars with swash-backwash lamination formed on the seaward side. The preservation potential of bars is thus very large on recurved spits in contrast to linear-barred coast, where longshore migration of the interjacent troughs will remove the bar deposits resulting in shoreface successions without bar deposits (Hunter et al., 1979; Nielsen et al., 1988).

Interpretation of the sections The depositional facies of units 1-4 represent an upward-shallowing succession of shoreface sands deposited by the prograding Skagen spit complex. A marine shell from the bar-trough deposits in section 4 indicates an age of 6245 BR which is in accordance with an age of 5330 BP of the peat layer topmost in the section (Hauerbach, 1992). The shoreface sands show some important differences from typical shoreface successions from prograding deltaic or linear coasts. The storm sand beds in the lower part are unusually thick, bars are well-represented, the succession shows only a very weak upward-coarsening trend, pebbles are small and occur only scattered in the beach sand, and gravel beach ridges are missing. The thick storm sand beds and the pervasive bioturbation of their upper parts suggest that they were deposited in an area that was relatively protected and received large amounts of sediments. During westerly storms very large amounts of sand are transported in suspension along the spit coast to the point of the spit where it is deposited because the refracted waves lose their energy, and the shore-parallel current becomes unconfined as the spit coast turns abruptly toward the south. Bars are only likely to be well-preserved along the recurved part of the spit coast where they weld onto the beach and add to its progradation (Fig. 8; Hine, 1979; Nielsen et al., 1988). The measured directions indicate that the palaeo-spit coast had a N W - S E trend as the bars migrated toward the westsouthwest, while the dunes in the troughs migrated toward the east, southeast and northeast In contrast to the preserved sections, pebbles are abundant on the northern spit coast where they occur on pebble-strewn surfaces and as pebbly beach ridges. The pebbles originate from erosion of the mainland coast more than 40 km to the southwest and are transported during westerly storms in the swash-backwash zone along the coast to the northern spit coast, where they are deposited onto pebble beach ridges formed during storms. The average size of the largest pebbles shows a lateral decrease toward the east along the north-

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ern spit coast from ca. 7 cm to ca. 4 cm reflecting the lateral decrease in wave energy as the waves are refracted along the spit coast as it turns away from the sea (Fig. 4). At the point where the active spit coast turns toward the southeast, the average pebble size decreases significantly to ca. 2.4 cm and their abundance similarly decreases. These evidences all indicate that the measured sections represent a position on the palaeo-coast of the early Skagen spit system, where the recurved coast had turned away from the sea and had a NW-SE trend. This implies that the reconstruction in Fig. 1 made by Hauerbach (1992) showing the 5400 BP situation needs to be slightly modified by having the seaward facing coastline located more westward and a more pronounced curvature of the coast along the palaeo-spit point.

Sequence stratigraphy The medial to distal part of the regressive spR system Based on regional context it is evident that the Skagen spit complex was formed during a forced regression. In most coastal settings with a stabilised or equilibrium shoreface-shelf profile, the slope of the sea bottom basinward of the shoreface toe is very gentle, typically showing a gradient of 1 : 20001:3000 (Walker and Plint, 1992; Nummedal et al., 1993). Therefore, shoreface successions formed during falling sea level typically show evidence of reduced accommodation space during their formation by being compressed and sharp-based overlying a regressive surface of marine erosion (Plint, 1988; Posamentier et al., 1992; Hunt and Tucker, 1992, 1995; Hamberg and Nielsen, 2000). For the Skagen spit complex the thickness of the preserved bar-trough deposits and the beach sands is, however, within the expected range deduced from the present-day bathymetry along the active spit, and the preserved successions do not show signs of the fall of sea level (Fig. 9). Likewise, the Skagen-3/4 well shows that the sandy spit complex is thicker at the present spit point than further south, although the area at the point of the Skagen spit has been subjected to a falling sea level within the latest 7600 years of ca. 25 m. This apparent paradox is related to an inherited glacial and possibly fault-induced topography that shows a very marked basinward increase of water depth up to several hundred metres in the Norwegian trench between Denmark and Norway. As discussed earlier, the average gradient of the basin floor at the initial spit formation was ca. 1:280 while the resulting gradient of the spit coast trajectory is 1:2700, which means that the spit system has experienced an increase in

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Fig. 9. Schematic model of the Skagen Odde spit complex. The main part consists of an accretionary, forced regressive systems tract.

accommodation during its basinward progradation in spite of the sea-level fall (Fig. 2). Thus the medial to distal part of the spit system constitutes an accretionary, forced regressive systems tract.

The medial-proximal part of the regressive spit system At locality 4 in the medial-proximal part of the regressive spit system, a conglomerate occurs, 1550 cm thick which can be followed over a distance of ca. 1 km (Figs. 1, 6 and 10). It is sheet-like and overlies bar-trough deposits with an erosional unconformity that shows a basinward down-cutting of ca. 0.4-0.5 m/100 m to the northeast. Basinward the conglomerate splits up into 3-4 thin pebble beds that pinch out into coarse-grained bar-trough deposits. In most places the basal erosion surface is even and sharp, but in places it shows a highly irregular appearance with erosion holes up to 15 cm deep and filled with pebbles that occasionally are cross-bedded. The pebble-filled erosion holes show some similarities with the gravelly gutter casts described from nearshore deposits by Chiocci and Clifton (1991); however, it has not yet been possible to determine their 3D morphology in detail. The conglomerate is typically clast-supported, non-graded, without imbrication and consists of rounded to sub-angular clasts, 2.5-9.5 cm in diameter, with an average of ca. 5 cm. Parallel bedding and foresets are occasionally seen. The conglomerate contains large reworked lumps of peat and occasionally abundant marine shells and consists of the coarsest grain size seen in the profiles. Its planar appearance, coarse grain size and the way it splits up and thins basinward, indicate that it was deposited on a beach by amalgamation of pebble-strewn surfaces (similar to the stone horizons on

the active spit coast) and gravel beach ridges formed by storm waves. The seaward-dipping foresets may represent progradation of the beach toe or gravely beach cusps, and landward-dipping foresets represent modified swash bars. The relief of the erosive base indicates that the conglomerate was deposited simultaneously with erosion of the sand to hinder collapse of the loose sand. The basinward down-cutting of ca. 5 m, the coarse grain size and planar appearance suggest that the conglomerate was formed during a fall in sea level of ca. 5 m and marks a sequence boundary. After the short-lived regression, a transgression of a similar amount re-established the accommodation space at the spit point and normal spit progradation continued (Fig. 9). Age determination from a well-preserved shell immediately below the conglomerate has yielded 6245 BP (reservoir-corrected, sample AAR-4474), while reworked shells from the conglomerate have yielded 7030 BP (reservoir-corrected, sample K-6844). As these ages are comparable to ages from the Storregga Slide (Haflidason et al., 1999), the conglomerate may alternatively be interpreted as a tsunami deposit formed by waves triggered by the slide similar to those observed in Norway and Scotland (Bondevik et al., 1997). Further work including comparison to sea-level curves, further age-datings and a detailed study of the 3D nature of the conglomerate and its sedimentary structures is needed to evaluate the alternatives. Discussion

The coasts of the present-day Skagen spit complex show a full range of coastal processes ranging from cliff erosion and formation of a ravinement surface along the western coast (transgression), rapid active

Accretionary, forced regressive shoreface sands of the Holocene-Recent Skagen Odde spit complex, Denmark

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Fig. 10. Locality 4. (a) Close-up of the base of the conglomerate where it is highly irregular. SB = sequence boundary. (b) Notice that the beach conglomerate is cutting the gently northeastward-dipping spit deposits in the lower half of the section. Persons for scale.

accretion along the northeastern coast (the spit point) due to amalgamation of swash bars with the spit coast and formation of beach ridges (mainly depositional regression), and formation of marine foreland along the eastern coast due to uplift and deposition (combined forced and depositional regression). Similar suites of coastal processes are likely to occur along the coastlines of active and partly submerged fault blocks. The Skagen spit complex is different from spits developed as part of a delta (1) by being sourced from wave erosion of coastal cliffs and not fluvialderived sediment, (2) and by being developed during uplift, which has influenced its form by (passive) formation of marine foreland to the east and reduction of the coastal retreat along the proximal part of spit, which normally takes place as the mainland coast

is eroded. The geological setting of the Skagen spit complex may thus be comparable to terrains under active faulting. Large spit systems are formed where strong winds from a dominant direction with a sufficiently long free fetch cause wave erosion of poorly consolidated to moderately lithified rocks and establishment of strong currents that follow a mainland coast until it bends away from the sea. Such scenarios may occur both in extensional and compressional settings. In extensional settings fault blocks may develop with varying topography and gradients, and it is likely that parts of some fault blocks will be subjected to wave erosion with the potential formation of spit systems (Fig. 11A). In basins where compression causes inversion of fault blocks, some blocks may be uplifted

L.H. Nielsen and P.N. Johannessen

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Fig. 11. Formation of a shallow marine sandstone reservoir by spit-system progradation. (A) An extensional setting with wave erosion of faulted and rotated blocks. (B) A compressional setting with wave erosion of inverted fault blocks.

above sea level while others may remain below, and thus being potential sites for deposition of shoreface sands (Fig. l lB). In both types of settings spit systems may be formed, and the spit-system model may be an alternative to "standard" interpretations of shallow marine sandstones. Especially rotated extensional fault blocks are widely represented within the Jurassic-Cretaceous of the large North Sea basin, and spit systems may be formed on both hanging and foot wall blocks (Fig. 12). Conglomeratic layers are reported from the Upper Jurassic shelf sediments from the Danish Central Graben (Johannessen and Andsbjerg, 1993; Johannessen et al., 1996; Johannessen, 2001). As pebbles are not very likely to be transported on a flat shelf, the occurrences of the conglomerates are somewhat puzzling. However, as described here, pebbles are transported along the Skagen spit coast for at least 40 km in the swash-backwash zone. If a spit system never makes it into the geological record because of continued erosion of the seaward side of the fault block, it is likely that a ravinement surface with pebbles will be left. The spit-system model may thus provide an alternative interpretation of pebble lags also, instead of the "classic" model with regression accompanied by fluvial deposition of pebbles that later during ensuing transgression are reworked by

Fig. 12. Hypothetic map view of possible sites of spit-system development in extensional settings depending on dominant wind direction and available accommodation space. (A) Spit systems formed on foot-wall slope. (B) Spit systems formed on hanging wall fault block. (C) Spit systems formed along structural strike.

shoreface erosion (Posamentier et al., 1992; Johannessen, 2001).

Acknowledgements Norsk Hydro, represented by T. Dreyer, G. Mangerud and O. Martinsen is thanked for financial support and stimulating discussions. E Andreasen, J. Fredericia, R Hauerbach and J.O. Leth introduced us to the geology of Skagen. The referees H.E. Clifton and E Livbjerg are thanked for valuable suggestions. J. Heinemeyer and K.L. Rasmussen conducted the C-14 dates; Gurli E. Hansen made the figures.

References Anthonsen, K.L., Clemmensen, L.B. and Jensen, J.H., 1996. Evolution of a dune from crescentic to parabolic form in response

Accretionary, forced regressive shoreface sands of the Holocene-Recent Skagen Odde spit complex, Denmark to short-term climatic changes: R~bjerg Mile, Skagen Odde, Denmark. Geomorphology, 17: 63-77. Bondevik, S., Svendsen, J.I. and Mangerud, J., 1997. Tsunami sedimentary facies deposited by the Storegga tsunami in shallow marine basins and coastal lakes, western Norway. Sedimentology, 44:1115-1131. Chiocci, F.L. and Clifton, H.E., 1991. Gravel-filled gutter casts in nearshore facies indicators of ancient shoreline trend. In: R.H. Osborne (Editor), From Shoreline to Abyss: Contribution in Marine Geology in honor of Francis Parker Shepard. Soc. Econ. Paleontol. Mineral., Spec. Publ., 46: 67-78. Evans, O., 1942. The Origin of Spits, Bars, and Related Features. Dowden, Hutchinson and Ross, Stroudsburg, PA, pp. 53-72. Foster, T. and Jensen, J., 1990. Coastal development on the southeast coast of the Skaw Spit. In: E Bruun and N.K. Jacobsen (Editors), Proceedings, Skagen Symposium. J. Coastal Res. Spec. Vol., 9: 693-723. Fredericia, J., 1988. Den hydrogeologiske kortl~egning af Nordjyllands Amtskommune. Dan. Geol. Unders., Intl. Rep. 22, 231 pp. Haflidason, H., Gravdal, A., Sejrup, H.P., Bryn, R, Lien, R. and Mienert, J., 1999. TOBI Imagery side-scan sonar and seismic data of the northern escarpment of the Storegga Slide off Mid-Norway: evidence of long term instability. In: O. Martinsen and T. Dreyer (Editors), Sedimentary Environments Offshore Norway - - Palaeozoic to Recent. Norwegian Petroleum Society, Bergen, Norway, 3-5 May, Extended Abstracts, pp. 205-208. Hamberg, L. and Nielsen, L.H., 2000. Shingled, sharp-based shoreface sandstones: depositional response to stepwise forced regression in a shallow basin, Upper Triassic Gassum Formation, Denmark. In: D. Hunt and R. Gawthorpe (Editors), Sedimentary Responses to Forced Regressions. Geol. Soc., London, Spec. Publ., 172: 69-89. Hansen, J.M., 1977. Sedimentary history of the island LassO, Denmark. Bull. Geol. Soc. Denm., 26: 217-236. Hansen, L.V. and Langballe, I., 1993. Kystlinjevariationer og sedimenttransport pfi Skagens Oddes vestkyst. Unpublished M.Sc. Thesis, Geographic Institute, Copenhagen University, 156 pp. and 10 enclosures. Hauerbach, E, 1992. Skagen Odde-Skaw Spit. An area of land created between two seas. Folia Geogr. Danica, 20, 119 pp. Hine, A.C., 1979. Mechanisms of berm development and resulting beach growth along a barrier spit complex. Sedimentology, 26: 333-351. Hunt, D. and Tucker, M.E., 1992. Stranded parasequences and the forced regression wedge systems tract: deposition during baselevel fall. Sediment. Geol., 81: 1-9. Hunt, D. and Tucker, M.E., 1995. Stranded parasequences and the forced regressive wedge systems tract: deposition during baselevel fall - - reply. Sediment. Geol., 95: 147-160. Hunter, R.E., Clifton, H.E. and Phillips, R.L., 1979. Depositional processes, sedimentary structures, and predicted vertical sequences in barred nearshore systems, southern Oregon coast. J. Sediment. Petrol., 49:711-726. Jessen, A., 1899. Beskrivelse til geologisk kort over Danmark. Kortbladene Skagen, Hirtshals, Frederikshavn, Hjorring og Lokken. Dan. Geol. Unders. I, Ra~kke 3,368 pp. Jessen, A., 1936. Vendsyssels Geologi. Dan. Geol. Unders. V, Ra~kke 2, 195 pp. Johannessen, EN., 2001. Depositional environments and sequence stratigraphy of paralic and shallow marine Upper Jurassic reservoir sandstones in the northern part of the Danish Central Graben. In: E Surlyk and J.R. Ineson (Editors), The Jurassic of Denmark and Greenland. Geol. Denm. Surv. Bull. (in press). Johannessen, EN. and Andsbjerg, J., 1993. Middle to Late Jurassic basin evolution and sandstone reservoir distribution in the Danish Central Trough. In: J.R. Parker (Editor), Petroleum Geology of Northwest Europe: Proceedings of the 4th Conference. Geological Society, London, pp. 271-298. Johannessen, EN., Dybkjam K. and Rasmussen, E.S., 1996. Sequence stratigraphy of Upper Jurassic sandstones in the northern

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Geological Survey of Denmark and Greenland (GEUS), Thoravej 8, DK-2400 Copenhagen NV, Denmark Geological Survey of Denmark and Greenland (GEUS), Thoravej 8, DK-2400 Copenhagen NV, Denmark

473

References index

Aarseth, I., 5, 448, 449, 455, 456 Aasheim, S., 40, 41, 45, 49, 50 Aasheim, S.M., 200, 210 Abbate, E., 416, 418 Adams, R.D., 472 Affleck, L.G., 403 Agdestein, T., 84, 195 Agterberg, F.P., 34, 143, 196, 231, 363

Ahlberg, A., 197 Ainsworth, N.R., 73, 84 Aitken, J.F., 90, 95, 98, 102 Allen, G.E, 143, 237, 239-241, 244-246, 255, 471 Allen, H.W., 144 Allen, J.R.L., 109, 142, 246, 255, 269, 277, 426, 439 Alley, R.B., 449 Allison, EA., 367, 380 Alsaker, A., 455 Alsen, E, 319 Alve, E., 369, 381 Alvestad, E., 35 Amedro, F., 291 Anastas, A.S., 109, 140, 142 Andersen, B.G., 448, 448 Andersen, C., 196, 197 Andersen, M.S., 31, 33 Andersen, ER., 196 Andersen, S.B., 368, 381 Andersen, T., 2, 4 Andersen, T.B., 13, 33, 34, 49 Anderson, D.S., 40, 49 Anderson, J., 448, 449 Anderton, R., 8, 13, 16, 17, 31, 33 Andreassen, K., 452, 455, 456 Andriessen, EA.M., 36 Andsbjerg, J., 2, 4, 176, 178, 179, 183-185, 187, 189, 192, 194, 195-197, 470, 471 Anthonsen, K.L., 460, 470 Archard, G.M., 4 Arenillas, I., 439 Arhus, N., 35, 225,231,232 Armstrong, L.A., 400, 402 Arnaud, N.O., 34 Arnesen, D.M., 102, 256 Arnold, M., 449 Arz, J.A., 439 Ashton, M., 34, 84, 173, 210

Astin, T.R., l l7, 142 Atkinson, K., 379 Aubert, J., 380 Augustson, J., 231

Bertram, G.T., 321, 323, 344, 347,

Baartman, J.C., 194, 195 B~ickstr6m, S.A., 4, 31, 34, 231,321, 325, 345, 347, 351, 359, 363, 367-369, 377-379, 380, 381 Badescu, M.O., 3, 4, 389, 402 Badley, M.E., 69, 84, 85, 147, 172, 174, 193, 195, 345 Bailey, H.W., 34, 380 Baker, E.K., 116, 120, 142 Baker, R.A., 367-369, 381 Bakke, S., 231,232 Baldwin, C.T., 413,418 Banks, N.L., 269, 277 Banner, J.A., 400, 402 Barker, M.J., 243,244, 256 Barlow, L.K., 449 Barnes, K., 291,364, 381 Barr, D., 193, 196 Barrett, P.J., 448 Barrett, T.J., 389, 402 Barton, M.D., 234, 256 Barwis, J.H., 89, 102 Basov, V.A., 223,224, 232 Batist, M., 210 Batten, D.J., 81, 84 Beamish, G.W.J., 291,418 Beaudoin, B., 389, 402-404 Beckly, A., 255 Beckmann, J.E, 380 Befring, S., 4, 455 Beicip Franlab, 323, 344, 348, 363, 369, 379 Belderson, R.H., 4, 455 Bentley, C.R., 448, 449 Bergan, M., 22, 24, 33 Berge, A.M., 33 Berge, C., 85 Bergfjord, E., 291 Berggren, W.A., 321, 344, 345, 347,

Beyer, I., 449 Bhatia, S.B., 380 Bhattacharya, J.B., 256 Bhattacharya, J.E, 134, 142 Bidstrup, T., 195, 196 Bindschadler, R., 448,448 Birkeland, 0., 34-36, 439 Birkelund, T., 206, 210 Bjerke, M.D., 375,377, 380 Bj0rlykke, K., 4, 35, 345, 363, 380 Bj0rnseth, H.M., 29, 33, 36 BjorCy, M., 216, 231 Blackbourn, G.A., 4 Blanche, J.E, 403 Blankenship, D., 448 Blaume, F., 418, 456 Blevens, M., 256 Bloch, R.B., 291,418 Blodgett, R.H., 74, 84 Bloemendal, J., 448 Blum, M.D., 136, 142 Blystad, E, 9, 12, 22, 26, 28, 31, 33, 282, 291, 421,439 BOen, F., 5, 35, 364, 381 Boersma, J.R., 116, 121, 124, 136,

363, 364

Besley, B.M., 95, 102 Besly, B.M., 4, 9, 14, 16-18, 33, 74, 84

363, 380, 381

Berglund, L.T., 228, 231 Bergsager, E., 5 Bergslien, D., 403, 418 Bergstr6m, J., 180, 181,197 Bertelsen, F., 79, 84, 194, 196

142

Bohacs, K., 96, 100, 102 Bojesen-Koefod, J., 196 Bojesen-Koefoed, J.A., 197 Boldreel, L.O., 31, 33 Bonde, N., 380 Bondevik, S., 468,471 Borisov, A.V., 35, 231 Bortolotti, V., 418 Bose, M.N., 232 Bosence, D.W.J., 367, 380 Boulter, M.C., 37 Bouma, A.H., 52, 54, 56, 64, 383, 384, 389, 402, 405, 409, 418, 423, 439

Bourgeois, J., 115, 143 Bourillet, J.-F., 210 Bowen, J.M., 405 Bowen, R.N.C., 365,380 Bowman, M.B.J., 29, 30, 33, 173,

474 398, 399, 402 Boyd, R., 143, 144, 172, 255, 257 Bradshaw, M.J., 225, 226, 231 Branney, M., 426, 439 Branney, M.C., 409, 418 Breivik, A., 34 Breivik, A.J., 29, 33 Brekke, H., 2, 3, 4, 9, 24-31, 33, 34, 69, 84, 211,231,291,439, 448 Brett, C., 455 Breusers, H.N.C., 116, 124, 144 Brewster, J., 399, 400, 402 Bridge, J.S., 79, 84 Bridges, EH., 33 Brierley, G.J., 79, 84 Bristow, C.S., 13, 33 Britze, E, 193, 196 Bro, E.G., 35, 231 Brodie, J., 319 Bromley, R.G., 218, 223, 231 Brooks, EW., 35, 231 Brunel, M., 50 Bruun Christensen, O., 35 Bryhni, I., 40, 49 Bryn, E, 4, 5, 449, 471 Bubb, J.N., 232, 345, 364, 405 Buch, A., 194, 197 Bue, B., 231 Bugge, T., 3, 4, 105, 142, 232, 441, 443, 445,448, 449, 451,452, 455 Bukovics, C., 106, 107, 142 Bungum, H., 33, 452, 455 Burchell, M.T., 35, 84 Burns, B.A., 48, 49 Burton, C.A., 380 Busi, C., 418 Butt, A., 380 By, T., 449 Bylund, G., 180, 196 Callomon, J.H., 23, 33 Cameron, T.D.J., 17, 33 Campion, K.M., 50, 102, 232, 256 Candace, O.M., 36 Cannon, S.J.C., 148, 172 Cant, D.J., 42, 49, 79, 84 Carman, G.J., 399, 400, 402 Carmichael, S., 36 Carr, I., 314, 318 Carter, C.H., 423,439 Cartier, E.G., 142 Cartwright, J.A., 287, 291,354, 364 Castellucci, E, 418 Cazzola, C., 389, 397,403 Cecchi, M., 3, 4, 35, 418 Chadwick, R.A., 236, 244, 255 Chalmers, J.A., 36 Chan, M.A., 389, 403 Chapin, M.A., 387,403 Chappell, J., 144 Charnock, M., 5 Charnock, M.A., 2, 4, 35, 73, 81, 84, 364, 367, 368, 375,380, 381

References index

Chatellier, J.Y., 402 Chauvet, A., 50 Chen, C., 56-58, 64, 375, 380 Cheredeev, S.I., 35 Cherry, S.T.J., 354, 363 Chiocci, EL., 468, 471 Christensen, O.B., 194, 195 Christiansen, F.G., 36, 52, 64, 65, 278, 296, 318, 319 Christiansson, E, 11, 33, 364 Christie, EA.E, 348, 364 Claoue-Long, J., 85 Clausen, O.R., 65, 278 Clayton, C.J., 173 Clemmensen, L.B., 298, 316, 318, 319, 470

Cliff, R.A., 13, 33 Clifton, H.E., 124, 125, 136, 143, 468, 471 Cloetingh, S., 11, 30, 33, 35, 36, 342, 344

Cohen, A.D., 96, 102 Coleman, J.M., 73, 84, 95-98, 102, 140, 143, 389, 402 Coleman Jr., J., 418 Coleman, M.L., 128, 143 Collins, M.B., 255 Collinson, J.D., 1, 2, 4, 16, 33, 35, 90, 92, 96, 102, 243,255, 405 Conradsen, K., 471 Contescu, L., 389, 403 Cook, EJ., 246, 255 Cooper, A.K., 448, 448 Copestake, E, 85, 169, 172, 173, 21 O, 231,380

Corbett, P.W.M., 256 Corfield, S., 2, 4, 200, 201,210 Corfield, S.M., 9, 14, 17, 18, 33 Corfu, F., 17, 33 Corliss, B.H., 375,380 Cornaggia, F., 418 Cossey, S.EJ., 387, 403 Coward, M.E, 9, 16, 33 Crane, K., 456 Crawford, R., 401,402, 403 Cronin, B.T., 389, 403 Cunningham, A.E, 448, 449 Curtis, C.D., 128, 143 Cutts, EL., 401,403 D' Agostino, A.E., 418 D'Heur, M., 401,402, 403 D'Iorio, M., 345, 363 Dahl, C.R., 195 Dahl, N., 26, 33 Dahl-Stamnes, ~).A., 5 Dahlgren, S., 17, 33, 84 Daley, B., 236, 255 Dalland, A., 3, 4, 22, 24-27, 29, 31, 33-35, 69, 79, 84, 87, 102, 104, 107-109, 125, 134, 136-140, 143, 173, 199, 200, 205, 210, 211,

218-220, 222, 223, 225, 228, 229, 231,232, 282, 291,363, 423,427,

431-433,439 Dallmann, W., 232 Dalrymple, R.W., 49, 142-144, 156, 172, 210, 233, 234, 236, 237, 243-245,255, 257 Dam, G., 65, 108, 138, 140, 143, 319 Damtoft, K., 182, 193, 196 Damtofte, K., 197 Damuth, J.E., 291,418, 455,455 Dangree, J.B., 345 Danyushevskaya, A., 231 Davies, E, 403 Davies, S.J., 102 Davis, R.J., 404 De Graciansky, E-C., 161, 166-170, 172, 173, 280, 282, 285,286, 291 De Vries, M.B., 402, 418 Deegan, C.E., 280, 286, 291 Deeks, N., 196 Dehls, J.F., 31, 33 den Dulk, M., 381 Deroo, G., 291 Dever Jr., G.R., 84 Di Liegro, G., 65 Dickinson, W.R., 389, 403 Digranes, E, 35 Dimitrov, E, 36 Dixon, R.J., 233,255 Dobinson, A., 36 Dodd, C., 255 Donovan, D.T., 259, 260, 277, 294296, 298, 303, 305,308-311, 314, 316,318,318

Donselaar, M.E., 402 Dor6, A.G., 8, 9, 11, 22-28, 31, 34, 37, 67, 69, 79, 83, 84, 85, 105-107, 143, 145, 147, 148, 151, 166, 172, 173, 199, 205, 210, 283, 291,294, 319, 421,439 Dorn-Lopez, D., 36 Dott Jr., R.H., 115, 143, 389, 403, 409, 418 Dowdeswell, A., 418 Dowdeswell, J.A., 451,452, 455,455, 456

Downie, C., 232 Downie, R.A., 19, 34 Doyle, G.A., 449 Doyle, J.D., 255 Drewery, S.E., 33 Dreyer, T., 1, 4, 5, 34, 102, 103, 116, 143, 146, 151, 160, 167-171,172, 210, 231,234, 255

Duchaufour, E, 95, 102 Duijnstee, I.A.E, 381 Duplessy, J.C., 449 Dybkja~r, K., 4, 79, 84, 176, 178, 179, 187, 189, 192-194, 195, 196, 471 Dyer, R., 231,380 Dypvik, H., 232

475

References index

Edvardsen, A., 363 Edwards, A.S., 381 Egan, S.S., 363 Egeberg, T., 172 Eger, J.D., 403 Ehrenberg, S.N., 106, 108, 143, 146, 171,172, 199, 200, 203,210 Eide, E.A., 13, 33, 34 Eide, E, 71, 84 Eidvin, T., 4, 31, 34, 378, 380, 441-443, 448, 455 Einchcomb, C.C.G., 255 Ekdale, A.A., 223,231 Ekern, O.F., 103, 143, 363 E1 Maherssi, C., 402 Eldholm, O., 30, 35, 36 Eliassen, EE., 34, 439 Elliott, T., 43, 45, 49, 93, 95, 97, 98, 102

Elvebakk, G., 35, 231 Elverhcfi, A., 5, 319, 449, 455, 456 Elvsborg, A., 231 Embry, A., 220, 221,231 Embry, A.F., 12, 14, 18, 20, 21, 23, 24, 27, 34, 35, 221,231 Emeleus, C.H., 278 Enay, R., 23, 34 Engkilde, M., 206, 209, 210, 319 Englehorn, J.A., 403 Enoksen, T., 4 Erlstr6m, M., 180, 196, 197 Ernst, S.R., 381 Errat, D., 9, 11, 22, 34 Escher, J.C., 12, 34 Ethridge, F.G., 50 Ettensohn, F.R., 74, 84 Eugeno-S Working Group, 196 Evans, D., 455 Evans, G., 255 Evans, O., 457, 471 Eynon, G., 148, 172, 175, 178, 193, 196

Faccini, U.E, 49 Fairbanks, R.G., 245,255 Faleide, J.I., 4, 28, 29, 31, 33-37, 345, 363, 364, 380

F~ilt, L.M., 234, 255 Fanavoll, S., 231 Farley, M.B., 173 Faure, G., 431,439 Fa~rseth, R.B., 4, 9, 11, 19, 22, 26, 33-35, 68, 69, 71, 84, 147, 172, 231,282, 291,350, 363, 439

Felt, V.L., 5 Fenton, J.EG., 4, 84 Feurere, J.R., 402 Fichler, C., 34 Fielding, C.R., 43, 49, 74, 84, 93, 95, 96, 102 Finney, B.E, 348, 364 Fisher, M.J., 22, 23, 34 Fjeldskaar, W., 31, 36, 345, 363

Fjellanger, E., 148, 172 Flach, ED., 79, 85 Flanagan, K.R, 400, 402, 404, 407, 418

Fleet, A.J., 173 Flesche, H., 49 Flint, S.S., 90, 95, 98, 102 Folkestad, A., 2, 4, 40, 41, 49, 49 Fonnesu, F., 403 Forrester, M.M., 256 Fossen, H., 13, 34, 40, 49 Foster, T., 461,471 Fowler, N.M., 403 Fralick, EW., 389, 402 Frandsen, N., 85 Fraser, A.J., 33, 35, 84, 255 Fraser, A.R., 401,405 Fredericia, J., 459, 471 Frengstad, B., 5, 441,443,449 Fries, G., 389, 398, 403, 404 Fritsen, A., 28, 34 Fronval, T., 443,448 Frydendahl, K., 460, 471 Funnell, B.M., 144 Fumes, H., 2, 5 Gabaldon, V., 404 Gabel, S.L., 256 Gabrielsen, R.H., 4, 9, 1l, 12, 34, 35, 37, 143, 211, 231, 279, 291, 321, 323, 325, 337, 344, 345, 348, 350, 363, 364, 380, 381,440

Gage, M., 33 Gage, M.S., 69, 83, 84, 148, 172 Gagliano, S.M., 102, 143 Gale, A.S., 255 Gale, B., 256 Gallagher, K., 25, 35 Gallagher, L., 34 Galloway, W.E., 54, 64, 148, 173, 220, 231,350, 363, 367, 380 Garber, J.L., 363, 380 Gardner, M.H., 40, 47, 49 Gardner, T.W., 95, 102 Gary, M., 47, 49 Gastaldo, R.A., 116, 143 Gawthorpe, R.L., 33, 103, 144, 201, 206, 210 Gelati, R., 389, 403 Gellein, J., 35 Geroch, S., 375, 380 Gersib, G.A., 96, 102 Gibbs, A.D., 348, 363 Gibson, J.L., 403 Gibson, T.G., 325,345, 367-370, 375, 380

Giles, M.R., 172 Gillmore, G.K., 3, 4, 322, 323, 325, 327, 333, 336, 337, 345, 347, 348, 352, 362, 363, 364, 365,380 Gingras, M.K., 124, 125, 136, 143 Gjelberg, J., 3, 4, 25, 34, 35, 37, 47, 48, 49, 87, 92, 98, 102, 103, 107,

143, 199, 205, 206, 209, 210, 228, 231,277, 277 Gjerstad, H.M., 143, 210 Glennie, K.W., 8, 14, 19-22, 34 Gloppen, T.G., 40, 41, 49, 50, 232 Gluyas, J., 173 Gnaccolini, M., 389, 403 Goesten, M., 256 Golden, M., 85 Goldring, R., 243,255, 256 Goldsmith, P.J., 82, 84 Good, T.R., 256 Goodchild, M.W., 277, 277 Gorfir, N., 36 Gostin, V.A., 117, 144 Gradstein, F.M., 4, 27, 31, 33, 34, 36, 129, 143, 178, 196, 221,224, 231, 321,323, 325,344, 345, 347, 351, 359, 363, 367-369, 377-379, 380 Graham, J.R., 78, 84

Granholm, E-G., 294, 319 Granholm, EG., 291 Grant, A.C., 33 Grant, D., 403 Grant, S., 316, 319, 433,439 Grant, S.M., 33 Graue, E., 24, 34, 82, 84, 148, 156, 170, 171,173, 206, 210 Gravdal, A., 4, 471 Grayson, R.E, 16, 34 Green, S.C.H., 291 Griffin, E G., 401,402, 403 Grinde, P., 399, 401,403 Grogan, E, 34 Grootes, EM., 443,448, 449 Grow, J.S., 84 Grue, K., 35, 439 Grunnaleite, I., 34, 291 Guargena, C., 4 Gudlaugsson, S.T., 9, 11-13, 16, 17, 19, 21, 33, 34, 37 Gulbrandsen, A., 277, 277, 280, 291 Gunleiksrud, T., 441,448, 449 Gupta, S., 201,210 Gustavsen, B., 5, 449 Haas, S., 256 Hfibrekke, H., 35 Hadler-Jacobsen, F., 143, 210 Haflidason, H., 3, 4, 5, 441,448, 449, 455, 456, 468,471 H~iger, K.-O., 4 Hfikansson, E., 16-18, 20, 21, 23, 24, 26, 27, 29,, 34 Hald, M., 456 Hald, N., 278 Haldorsen, H.H., 233, 251,255 Hallam, A., 79, 84, 106, 108, 128, 136, 143, 178, 193, 196, 224-228, 231,232

Halland, E.K., 67, 85 Haller, J., 261, 278, 295, 310, 314, 319

476 Hallsworth, C.R., 403 Halvorsen, E., 180, 196 Hamar, G.E, 172, 280, 282, 291 Hamberg, L., 35, 194, 196, 278, 467, 471

Hamborg, M., 364 Hamp, R., 403 Hampson, G., 102 Hampson, G.J., 90, 98-100, 102 Hampton, M.A., 34, 314, 319, 384, 403, 409, 416, 418, 451,452, 455 Hancock, J.M., 367, 380 Hanisch, J., 29, 34 Hanken, N.-M., 144 Hansen, C.F., 319 Hansen, E.K., 33, 35, 319 Hansen, J.M., 464, 471 Hansen, J.W., 216, 228, 231,232 Hansen, K., 318 Hansen, L., 4 Hansen, L.V., 461,471 Hansen, R.A., 452, 455 Hansen, T., 452, 455 Hanslien, S., 400, 402, 403 Haq, B.U., 108, 128, 136, 143, 173, 193, 196, 335, 342, 345 Hardenbol, J., 143, 148, 160, 173, 196, 231,291,345

Harding, A.W., 400, 402, 403 Harding, I.C., 256 Haremo, E, 37, 440 Harker, S.D., 291 Harland, W.B., 28, 34 Harms, J.C., 143 Harper, H., 210 Harris, N.B., 143 Harris, ET., 142 Hart, M.B., 368, 375,380 Haslett, S.K., 435,439 Hatelid, W.G., 232 Hatleid, W.G., 405 Hatlelid, W.G., 345, 364 Hatton, I.R., 404, 418 Hauerbach, E, 458, 459, 461, 465, 467,471 Haugen, J.E., 448 Hawkes, P.W., 235,236, 244, 255 Hayward, A.B., 148, 173, 350, 364 Hedgpeth, J.W., 367, 380 Hedley, R.J., 85 Hein, F.J., 319, 404, 439 Heinburg, C., 206, 210 Heir-Nielsen, S., 471 Helgesen, G., 5, 256 Helland, A., 448,448 Helland-Hansen, W., 24, 34, 47, 48, 49, 82, 84, 148, 173, 206, 210 Heller, EL., 49, 389, 403 Hemmens, ED., 108, 143 Henderson, E.W., 472 Henk, F.H., 65 Henriksen, S., 3, 4, 441,443,448 Henry, K.L., 277

References index

Herbin, J.-P., 291 Hesjedal, A., 280, 282, 291 Hesselbo, S., 166, 173 Hesselbo, S.E, 172 Heum, O.R., 199, 200, 210 Hickin, E.J., 79, 84 Hicks, E.C., 33 Hine, A.C., 465,467, 471 Hines, F.M., 145, 146, 148, 170, 173 Hinkley, R.J., 277 Hinz, K., 36, 448 Hiscott, R.N., 56-58, 64, 319, 389, 395,403, 404, 439 Hodgkinson, R.J., 291,319, 418 HCgseth, K., 49, 282, 291 Hr A., 34, 102, 143, 210, 231 Holbrook, EW., 102 Hole, A., 143 Hollender, F.J., 455, 456 Hollingsworth, R.R., 35, 319 Holtedahl, H., 4, 441,444, 448, 448, 455

Holter, E., 5 Hooper, R., 210 Hossack, J., 36 Hossack, J.R., 33, 40, 49 Howard, J.D., 102 Howell, D.G., 389, 403, 405 Huang, Z., 196 Huang, Z.-H., 143 Hubbard, R.N., 37 Huc, A.-Y., 143 Humpherey, T.J., 403 Hunt, D., 47, 48, 49, 206, 210, 467, 471

Hunter, R.E., 465,467, 471 Hurich, C.A., 4, 291 Hurst, J.M., 65, 278 Hutt, S., 236, 255 Ibrahim, M.T., 4, 35, 345, 363, 380 Idil, S., 49 Ignatenko, E.A., 35 Imbus, S.W., 277 Ineson, J.R., 176, 193, 196 Insole, A.N., 236, 255 Isaksen, D., 3, 4, 280, 282-284, 291, 350, 351,363, 367, 380, 423, 439 Jackson, J.A., 314, 319, 364 Jackson, J.S., 201,210 Jackson, M.D., 5, 234, 255 Jacobel, R., 448 Jacobsen, T., 34, 84, 173, 255, 363 Jacobsen, V.W., 22, 34 Jacquin, T., 161, 166-170, 172, 173, 280, 282, 285,286, 291 Jakubowski, M., 225,231 James, A.V., 85 James, D.E, 144, 471 James, N.E, 142 Jannink, N.T., 381 Jansen, E., 4, 34, 381,442, 448, 449,

455

Japsen, P., 195, 196 Jenkyns, H.C., 166, 169, 173 Jennette, D.C., 234, 255, 256 Jensen, J., 461,471 Jensen, J.H., 470 Jensen, J.L., 256 Jensen, L.N., 34, 231,439 Jenssen, A.I., 400, 403, 407, 413, 414, 418

Jervey, M.T., 50, 232, 342, 345 Jessen, A., 458, 459, 471 Johannessen, E.E, 5, 24, 34, 82, 84, 148, 171,173, 221,231,232, 350, 363

Johannessen, EN., 3, 4, 5 176, 178, 179, 189, 192, 196, 458,470, 471 Johansen, B., 5 Johansen, R., 232 Johansen, S.E., 13, 16, 20, 22, 24, 26, 27, 34, 35 Johnsen, J.R., 84, 173, 210 Johnsen, S., 448 Johnson, B., 169, 172 Johnson, H.D., 5, 25.5, 257, 402, 413, 418

Johnson, J., 34 Johnson, J.G., 220, 231 Johnson, K., 25, 35 Johnson, S.Y., 73, 79, 84 Jolley, E.J., 35, 84 Jones, A.D.W., 234, 255 Jones, B., 34 Jones, B.G., 73, 84 Jones, C.M., 4 Jones, C.R., 234, 255, 256 Jones, G.A., 36 Jones, G.D., 321,325, 345, 347, 363, 368, 380 Jones, R.W., 367, 368, 375,380 Jongepier, K., 25, 35, 422, 439 Jordan, D.W., 404, 405, 418 Jordan, T.E., 389, 403 Jorde, K., 85 Jordt, B., 36, 65, 278, 319 Jordt, H., 3, 4, 31, 35, 321,323, 325, 337, 345, 350, 351,359, 363, 364, 367, 375, 377, 378, 380 JCrgenv~g, S.H., 102, 256 Joseph, E, 403 Jouzel, J., 448 Joy, A.M., 321,345, 347,363 Jutson, D., 34 Kaas, I., 5, 256 Kalheim, J.E., 231 Kaminski, M.A., 33, 34, 345, 363, 368, 375, 380 Kaminski, M.J., 380 Karlsson, W., 103,143 Kazimieras, S., 36 Ke, X., 241,255 Keene, J.B., 142

References index

Keller, G., 435,439 Kelly, S.R.A., 37, 259, 260, 262-264, 268, 269, 272, 276, 278, 318, 319 Kemper, E., 284, 291 Kenoulty, N., 255 Kent, R., 319 Kenyon, N.H., 4, 119, 140, 143, 418, 451,452, 455, 456 Kessler, L.G., 401,402, 403 Kihle, O., 35 King, A.D., 380 King, C., 368-370, 380 King, E.L., 5, 448, 449, 455,455, 456 King, L., 441,443,444, 446-448, 449 Kirk, M., 73, 84 Kittilsen, J.E., 28, 35, 316, 319 Kja~refjord, J.M., 5, 98, 102, 256 Kj~ernes, E, 4 Kjennerud, T., 3, 4, 323-325, 333, 336, 337, 340, 344, 345, 363, 364, 365,369, 378, 380 Kjonsvik, D., 255 Klapper, G., 231 Kleverlaan, K., 389, 395,403 Kligfield, R.M., 405 Klingspor, I., 180, 196 Knarud, R., 22, 24, 33, 85, 142,448 Kneller, B.C., 384, 403, 409, 418, 426, 439 Knott, S.D., 29, 35, 69, 84 Knox, W.O'B., 166, 173 Knudsen, K.L., 458-460, 471 Knutsen, S.-M., 451,452, 455 Koch, J.-O., 199, 200, 210 Koch, J.O., 194, 196 Koch, L., 259, 262, 278, 295, 319 Kodaira, S., 35 Komar, ED., 465,471 Kooi, H., 33 Koraini, A.M., 37, 278, 319 Korstgfird, J.A., 65, 182, 196, 278 Kortekaas, T.F.M., 234, 255 Koutsoukos, E.A.M., 368, 375,380 Kouwe, W.F.E, 179, 194, 197 Kouwenhoven, T.J., 381 Kovach, W.L., 368, 375,380 Krabbe, H., 34, 318 Kreiner-M011er, M., 2, 4, 52, 60, 61, 64

Kristensen, L., 460, 471 Kristensen, S.E., 35, 37, 232 Kristiansen, I.L., 4, 84, 345, 363 Krog, H., 458, 471 Krystinik, L.F., 115, 143, 209, 210 Kuenen, EH., 384, 403 Kuhnt, W., 380 Kumar, N., 465,471 Kusznir, N.J., 36, 348, 363 Kuznir, N.J., 364 Kvamme, L.B., 451,455 Kyrkjebr R., 3, 4, 345, 347, 348, 350, 353, 357-359, 362, 363, 364, 365, 376, 38O

477

Laberg, J.S., 3, 5, 416, 418, 419, 441, 447, 448, 449, 452, 455,455, 456 Lake, L.W., 255 Lamorey, G.W., 449 Landrum, W.R., 143 Langballe, I., 461,471 Langlais, V., 102, 256 Larsen, B.T., 5, 33, 35, 291,439 Larsen, E., 443,449 Larsen, F., 85 Larsen, H.C., 13, 27, 29, 35, 318, 319 Larsen, K.-B., 456 Larsen, L.M., 318 Larsen, M., 2, 3, 5, 27-30, 35, 65, 262, 272, 277, 278, 294, 319 Larsen, O., 180, 197 Larsen, V., 40, 49 Larsen, V.B., 24, 25, 35, 36 Larssen, G.B., 232 Larter, R.D., 448, 449 Lasseter, T.J., 234, 255 Latham, A., 403 Laursen, G.V., 368, 381 Lawrence, D.A., 173 Leach, ER.L., 143 Lebesbye, E., 456 Leckie, D.A., 115, 143 Lee, H.J., 314, 319, 455 Leeder, M.R., 16, 24, 33, 35, 193, 196, 201,210, 418 Leereveld, H., 36, 232 Lehtonen, L.R., 319, 418 Leitchenkov, G., 448 Leith, D.A., 5, 256 Leith, T.L., 23, 27, 35, 225, 226, 231 Lerche, I., 196 Lericolais, G., 210 Lervik, K.S., 71, 85 Leth, O.L., 460, 471 Lie, E., 5 Liegro, G.D., 65 Lien, R., 4, 441,449, 471 Lilleng, T., 34, 102, 143, 172, 210, 231

Lindholm, C.D., 33 Lindholm, R.M., 403, 418 Link, M.H., 389, 395,403 Linsay, D.R., 102 Lippard, S., 11, 35 Lippard, S.J., 4, 345, 380 Littlefair, R.W., 403 Liu, G., 11, 35 Livbjerg, F., 35, 145, 173 Livsic, J.J., 29, 35 Locat, J., 455 Loeblich, A.R., 367, 381 Lofaldli, M., 223,232, 381 Lohmann, G.E, 381 Lomo, L., 34, 84, 173, 210 Long, D., 449 Longman, M.W., 418 Longva, O., 448, 449 LOnOy, A., 35

Lord, A.R., 170, 173 Los, A., 255 L0seth, H., 4, 5, 211-213, 225, 231, 232, 345, 380

Lottes, A.L., 12, 27-29, 36 Loutit, T.S., 173 L0vlie, R., 448 Lowe, D., 426, 439 Lowe, D.R., 52, 54, 57, 58, 64, 383, 384, 389, 403, 409, 418 Lund, J.J., 73, 85 Lund, T.B., 5 Lundin, E.R., 9, 31, 34, 36, 421,439 Lunsford, M.K., 403 Lyck, J., 380 MacCarthy, I.A.J., 78, 85 MacDonald, A., 67, 85 Mackensen, A., 378, 381 Mackertich, D., 193, 196 MacKinnon, T.C., 389, 403 Macleod, N., 435,439 Macnaughton, R.B., 42, 43, 49 Magniez-Jannin, F., 291 Magnus, C., 4, 33, 84 Maguregui, J., 120, 143, 233,255 Mangerud, G., 4, 5, 35, 364, 381 Mangerud, J., 471 Mangold, C., 23, 34 Manure, S.B., 223, 232 Marcussen, C., 36, 65, 278, 318, 319 Margulis, L.S., 35 Marjanac, T., 145, 146, 151, 166-170, 173, 233, 245,256 Markussen, S., 49 Marsden, G., 363 Marssets, T., 210 Marten, R.F., 35, 319 Martin, C.A.L., 269, 278 Martin, S., 172 Martinius, A.W., 2, 5, 235,256 Martinsen, O.J., 1, 3, 5, 30, 31, 35, 48, 49, 148, 173, 269, 278, 351, 359, 364, 378, 381 Martinsen, R.S., 209, 210 Martirosjan, V.N., 35 Marzo, M., 49 Masson, A.G., 78, 85 Mathieu, C., 5 Mayewski, EA., 449 Maync, W., 259, 262, 278, 294, 319 Mayo, W., 246, 255 M~ehle, S., 50 McAffce Jr., R., 49 McBride, B.C., 405 McBride, J.J., 84, 143, 439 McCabe, EJ., 39, 47, 48, 50, 74, 85, 96, 102 McCarthy, EJ., 48, 49 McGilvery, T.A., 54, 64 McGovney, J.E., 344, 345 McIver, R.D., 453,456

478 McKenzie, D.E, 314, 319, 348, 364 McMahon, A.H., 4 Mearns, E.W., 69, 83, 84, 85, 143, 433, 439 Meisling, K.E., 210 Meistrell, F.J., 457,465, 471 Mellere, D., 233, 256 Mertz, E.L., 459, 471 Miall, A.D., 42, 43, 49, 71, 73, 79, 85, 241,246, 256, 257 Michelsen, O., 79, 85, 176, 179, 184, 185, 187, 192-195, 196 Midboe, ES., 35 Middleton, G.V., 384, 389, 395, 403, 409, 416, 418, 426, 439 Mienert, J., 4, 418, 452, 455, 456, 471 Migliorini, X.X., 384, 403 Miller, K.G., 380 Miller, R.G., 225, 227, 232 Mills, S.J., 36 Milnes, A.G., 37 Milton, N.J., 4, 321, 323, 344, 347, 351,363, 364 Mitchell, R.W., 143 Mitchell, S.M., 291,418 Mitchener, B.C., 85, 148, 173, 210 Mitchener, J., 36 Mitchum Jr., R.M., 50, 102, 173, 232, 256, 345, 364, 405 Mitlehner, A.G., 381 Mjelde, R., 11, 25, 35 MjCs, R., 34, 84, 145, 173, 363 Moe, A., 34 Mogensen, T.E., 194, 196

Mohrig, D., 300, 305,319 Moiola, R.J., 280, 282, 291,383, 384, 389, 404, 409-411,418 Molina, E., 435,439 Moller, J.J., 176, 187, 189, 192, 193, 196

Montadert, L., 291 Monteil, E., 36, 232, 380 MCrk, A., 5, 35, 36, 142, 221, 223, 231,232, 448

MOrk, M.B.E., 5, 22, 35, 231,232 Morton, A.C., 69, 85, 398, 403, 433, 439

Morton, N., 81, 83, 85 Morton, R.C., 316, 319 Moskalenko, V., 36 Mossop, G.D., 79, 85 Mould, A.S., 36 Moullade, M., 380 Mound, D.G., 401,404 Mudge, D.C., 22, 23, 34 Muggeridge, A.H., 5, 234, 255 Mulder, T., 418 Muller, C., 291 Mute, E., 399-401,404 Murray, J.W., 368-370, 375,378, 381 Mutter, J.C., 36 Mutterlose, J., 257 Mutti, E., 55, 58, 59, 64, 384, 385,

References index

394, 396, 403-405, 418, 423, 426, 439, 440

Nagy, J., 223, 224, 232, 368, 381 Nansen, F., 454, 456 Narbonne, G.M., 49 Naylor, D., 37 N~ess, A., 5, 255, 256 N~ess, O., 35 Nedkvitne, T., 5, 319 Nelson, C.S., 142 Nemec, W., 27, 35, 305, 319, 426, 439

Netland, A., 210 Neuhaus, D., 402 Neverdal, E, 36 Newman, M.S.J., 400, 402, 404 Newman, M.St., 407, 418 Newton, S.K., 400, 402, 404, 407, 418

Nielsen, E.B., 34, 262, 278 Nielsen, L.H., 3, 4, 5, 79, 85, 176, 178, 179, 184, 187, 192-194, 196, 457, 458,465,467, 471 Nilsen, H.R., 50 Nilsen, T.H., 389, 395,403, 404 Nio, S.D., 234, 241, 245, 256, 269, 278

Nipen, 0., 172 Noe-Nygaard, N., 2, 3, 5, 65, 259, 260, 277, 278, 294, 316, 319 N0hr-Hansen, H., 36, 65, 260, 262, 263, 278, 296, 298, 300, 303, 319 Nordberg, K., 459, 471 Nordgard Bolas, H.M., 400, 404 NCrgaard-Pedersen, N., 278 Norling, E., 180, 181, 192, 197 Normark, W.R., 58, 59, 64, 384, 385, 404

Norton, M.G., 40, 49 NOttvedt, A., 5, 13, 16, 23, 24, 26, 27, 29, 34, 35, 84, 107, 143, 173, 210, 232, 350, 364, 367, 381 Nummedal, D., 256, 467,471 Nygaard, E., 5, 449 Nyland, B., 33, 84 Nyong, E., 368, 381 Nystuen, J.E, 67, 71, 82, 85, 285, 291,342, 345

Nytoft, H.E, 197

Oldham, L., 16, 34 Olesen, O., 33 Olesen, O.G., 13, 35 Olsen, L., 443,449 Olsen, R.R., 319 Olsen, T., 48, 49, 100, 102, 235, 256 Olsen, T.R., 172 Olsson, R.K., 368, 381 Oomkens, E., 89, 102 Orton, G.J., 141,143 Oschmann, W., 225, 227, 232 Osmundsen, ET., 33, 40, 49 Ostisty, B.K., 35 Ostvedt, O.J., 401,404 Ottesen, D., 5 Overli, EE., 34 Oxnevad, I.E., 34, 35 Paola, C., 49 Parize, O., 389, 402, 404 Parker, G., 319 Parker, J.R., 398, 404 Parkinson, D.N., 145, 146, 148, 170, 173

Partington, M.A., 23, 24, 27-30, 35-37, 73, 82, 85, 145, 148, 151, 160, 161, 166-170, 173, 175, 176, 178, 180, 193, 194, 197, 201,210 Pascoe, R., 201, 210 Paternoster, B., 403 Paterson, S., 244, 256 Payne, R., 472 Pchelina, T.M., 35, 231,232 Pearce, G.J., 369, 381 Peattie, D.K., 399, 400, 405 Pedersen, G.K., 176, 197 Pedersen, T., 36, 103, 108, 143 Peet, R.K., 368, 375,379, 381 Pegrum, R., 37 Pegrum, R.M., 20, 35 Peijs-van Hilten, M., 233,234, 256 Perry, K.K., 454, 456 Petersen, H.I., 195, 197 Petersen, K.S., 458, 459, 471 Petersen, N.W., 34 Pettijohn, F.J., 430, 439 Pettingill, H.S., 403 Phillips, R.L., 471 Phleger, F.B., 367, 381 Piasecki, S., 36, 52, 59, 61, 63, 64, 65, 278, 318, 319

O'Connor, S.J., 416, 418 O'Neill, M., 84 Oakman, C.D., 27-30, 35 Odinsen, T., 11, 34, 35, 291 Ofstad, K., 4, 33, 102, 143, 210, 231, 291,439

Ogg, J., 323,345 Ogg, J.G., 143, 231 Ogg, O.G., 196 Olafsson, I., 11, 35 Olaussen, S., 1-3, 5, 17, 18, 20, 21, 35, 36, 228,232, 278, 319

Pickering, K.T., 314, 319, 363, 380, 386, 404, 426, 439 Pickup, G.E., 233, 256 Pielou, E.C., 368, 381 Pinault, M., 402, 404 Pitman III, W.C., 36 Planke, S., 30, 35, 36 Platt, N.H., 354, 364 Please, EM., 172 Plint, A.G., 467, 471,472 Plint, G., 49 Plummer, ES., 117, 144

479

R e f e r e n c e s index

Pollard, J.E., 243,255 Posamentier, H.W., 48, 50, 140, 144, 173, 220, 221,232, 237, 239-241, 244-246, 255, 342, 345, 384, 390, 404, 467, 470, 471 Potter, EE., 439 Poulsen, N.E., 176, 179, 184, 187, 192, 197 Pratt, B.R., 117, 144 Prestholm, E., 35 Prestvik, T., 9, 36 Preuss, T., 35, 278 Price, H.S., 251,256 Price, J.D., 84, 195 Price, S.M., 278 Price, S.E, 37, 259, 278, 294, 298, 314,316,319

Printzlau, I., 180, 197 Prior, D.B., 95-98, 102 Pritchard, D.W., 236, 256 Proust, J.-N., 210 Provan, D.M.J., 209, 210 Pulham, A.J., 43, 50 Pulvertaft, T.C.R., 12, 34 Radley, J.D., 241,243,244, 256 Radovich, B.J., 344, 345 Rahmani, R.A., 124, 144 Rahmanian, D., 50 Rahmanian, V.D., 102, 232, 256 Rambech Dahl, C., 84 Ramberg-Moe, H., 231 Ramirez, A., 418 Ramm, M., 67, 71, 78, 82, 85, 148, 173

Rampone, G., 403 Ramsbottom, W.H.C., 16, 35 Rasch, S., 256 Rasmussen, A., 35, 37 Rasmussen, E.S., 35, 176, 187, 189, 196, 471

Rasmussen, J.A., 52, 64 Rattey, R.E, 148, 173, 350, 364 Raunsgaard Pedersen, K., 73, 85 Ravenne, C., 389, 404 Ravnas, R., 396, 404 Ravnfis, R., 2, 5, 64, 64 Rawson, EF., 26, 35 Raestad, N., 85 Reading, H.G., 141, 143, 235, 250, 256, 384, 386, 396, 404, 405, 430, 439

Reeder, M.L., 404 Reemst, E, 11, 35 Rees, G., 37 Reid, B.E., 143 Reineck, H.E., 116, 119, 120, 124, 139, 144, 256 Reither, E., 448 Remacha, E., 404 Renshaw,, 84 Renshaw, D., 34, 173, 448 Renshaw, D.K., 34

Reshaw, D., 363 Retallack, G.J., 95, 96, 102 Reynaud, J.-Y., 206, 210 Rhodes D., 4 Rhodes, D.N., 405 Rhodes, R.N., 109, 125, 139, 143, 245,255 Ricci Lucchi, F., 384, 385, 389, 391, 404, 423,439 Rich, B., 84 Richards, M., 384, 386, 396, 404, 430, 439 Richardsen, G., 455 Richardt, N., 458,471 Rider, M., 171,173 Riding, J.B., 169, 173, 184, 197 Rigazio, G., 389, 403 Riis, F., 9, 22, 26, 28, 29, 31, 33-36, 211,231,380, 423,435,439, 448 Riley, G.W., 256, 471 Riley, L.A., 26, 35, 284, 291 Ringrose, E, 255 Ringrose, ES., 234, 256 Risdal, D., 369, 381 Rise, L., 5, 231, 441, 447, 448, 448, 449

Ritchie, J.D., 194, 197 Ritter, U., 232 Roberts, A., 4 Roberts, A.M., 11, 30, 36, 173, 323, 325,345, 348, 364 Roberts, D.G., 8, 22, 23, 33, 36 Roberts, G., 85 Robertson, G., 233, 256 Robertson, I.D., 404 Rochow, K.A., 359, 364 R~e, S.-L., 49, 50, 67, 85 Roest, W.R., 289, 291 Rofheart, D.H., 418 Rogers, D.A., 117, 142 Rohrman, M., 22, 36 Rokoengen, K., 3, 5, 441-443, 445-448, 448, 449 Rolle, F., 36, 65 Rolle, F. Scholle, EA., 278 R~nning, K., 34, 84, 173, 210 Rosell, EJ., 389, 404 Rosvoll, K.J., 102, 256 Rowan, M.G., 405 Rowley, D.B., 12, 27-29, 36 Rubino, J.L., 172, 389, 404 Ruffell, A.H., 234, 236, 237, 244, 246, 256, 257 Ruhmor, J., 418 Rui, J.C., 35, 439 Rumohr, J., 456 Rundberg, Y., 31, 34, 36, 350, 364, 367, 380, 381 Rusciadelli, G., 291 Russell, L.R., 210 Rust, B.R., 73, 78, 79, 84, 85 Rutherford, M.M., 84 Rutledge, D., 36, 232

Ryan, W.B.F., 22, 36 Rye-Larsen, M., 403, 418 Rykkelid, E., 40, 49 Ryseth, A., 2, 4, 5, 34, 49, 67, 69, 71, 78, 82, 84, 85, 147, 148, 173, 210 Sagri, M., 418 Said, R., 367, 381 Saknic, M., 36 Sandberg, C.A., 231 Sanders, J.F., 465,471 Sandsdalen, C., 102, 256 Sandulescu, M., 389, 404 Sandvik, K.O., 4 Sangree, J.B., 232, 364, 405 Sarg, J.F., 173 Savoye, B., 418 Sawyer, D., 34 S~ettem, J., 5, 441,443,444, 447, 449 Scambos, T., 448 Schaaf, A., 291 Schlager, W., 47, 50 Schmidt, W.J., 107, 144 Schmitz, B., 381 Schnapper, D.B., 403 Scholle, P.A., 52, 63, 64, 64, 65 Scholz, C.A., 348, 364 Schou, A., 458, 459, 471 Schumm, S.A., 42, 50 Sclater, J.G., 348, 364 Scotese, C.R., 14, 36 Scull, B.J., 280, 286, 291 Seguret, M., 40, 50 Seidenkrantz, M.-S., 471 Seidler, L., 65, 278 Sejrup, H.P., 3, 4, 5, 381, 443, 448, 449, 455,455, 456, 471 Sellevold, M.A., 441,448 Selley, R.C., 244, 256 Sellwood, B., 34 Sellwood, B.W., 33, 124, 125, 144, 178, 193, 196 Seranne, M., 40, 50 Sgavetti, M., 389, 403 Shabtaie, S., 448, 449 Shanley, K.W., 39, 47, 48, 50 Shanmugam, G., 280, 282, 285, 291, 316, 317, 319, 364, 381,383, 384, 389, 404, 407,409-411,416, 418 Sharp, I.R., 4, 200, 201, 210 Shaw, N.D., 142, 172 Shields, K.E., 291 Shimamura, H., 35 Shiobara, H., 35 Shipp, S., 448,449 Short, S.A., 142 Siedlecka, A., 1, 5 Siegert, M.J., 455 Siever, R., 439 Sigal, J., 291 Sigmond, E.M.O., 441,449 Sinclair, H.D., 389, 404 Sivhed, U., 196, 197

480

References index

Sj0holm, J., 442, 449 Sjulstad, H.I., 4 Skar, T., 49 Skarpnes, 0., 172 Skibeli, M., 280, 282, 291, 319, 350, 353,364, 367, 375, 377, 381,413, 418

Skilbrei, J.R., 35 Skinner, A., 449 Skjervold, R., 291 Skjold, L.J., 37 Skogseid, J., 11, 28, 30, 33, 36, 291, 439

Skott, P.H., 35 Skuce, A.G., 36 Slatt, R.M., 383, 384, 389, 404, 405, 410, 418 Slesser, G., 472 Sliter, W.V., 367-369, 381 Sloan, B.J., 321,345, 347,364 Smalley, P.C., 144 Smedvig Technologies, 256 Smelror, M., 2, 5, 26, 27, 36, 213, 216, 219, 221,225,231,232, 380 Smethurst, M.A., 13, 35, 36 Smith, A.G., 105, 144 Smith, D.G., 144 Smith, K., 194, 197 Smith, R.M., 85 Smith, W.G., 143 Smniders, F., 256 Smythe, D.K., 30, 36 Solberg, P.O., 2, 4 Solli, T., 26, 33 Sonnino, M., 403 Sorbie, K.S., 256 SCrenes, N., 35 Sorgenfrei, T., 194, 197 Spalding, T.D., 418 Speijer, R.P., 381 Spencer, A.M., 85 Spencer, R.S., 367,381 Spinnanger, A., 50 Srivastava, S.P., 11, 12, 28, 36, 289, 291

Stadtler, A., 257 Stagg, H.M.J., 448 Standring, J., 84 Stanley, D.J., 389, 405 Stanley, K.O., 85 Staub, J.R., 96, 102 Stauffer, EH., 389, 405 Steel, R.J., 1, 2, 4, 5, 12, 13, 16-21, 23, 27, 29, 31, 34-36, 39-41, 45, 48, 49, 49, 50, 64, 64, 67, 69, 71, 82, 84, 85, 143, 145-148, 151, 161, 166-170, 173, 210, 233, 245, 256, 277, 277, 364, 381,396, 404, 426, 439 Stemmerik, L., 2, 4, 12-14, 16-18, 20-24-29, 35, 36, 34, 51-53, 59, 61, 63, 64, 65, 259, 262, 277, 278, 294, 296, 314, 318, 319

Stephen, P.J., 81, 85 Stewart, D.J., 79, 85, 236, 243, 244, 255, 256

Stewart, K.R., 35, 231 Stiberg, J.P., 212, 213,232 Stockbridge, C.E, 85 Str T., 37 Stollhofen, H., 102 Str H.-H., 130, 144 Stone, C., 418 Stone, C.G., 402, 405, 418 Stoneley, R., 236, 244, 256 Storhaug, K., 210 Stouge, S., 64 Stow, D.A.V., 26, 36, 439 Stow, D.A.W., 386, 405 Strand, T., 34 Straume, T., 291,319, 364, 381,418 Stride, A.H., 413,418 Strider, M.H., 403 Stuiver, M., 448 Styles, R, 364 Sundvor, E., 35, 456 Surlyk, F., 2, 3, 5, 12, 13, 20-29, 36, 51, 52, 59, 63, 65, 108, 128, 136, 138, 140, 143, 144, 196, 206, 209, 210, 221-224, 232, 259-261,277, 278, 294, 295, 297, 307, 310, 316, 319, 426, 439, 471 Suter, J., 96, 100, 102 Svela, K.E., 2, 5, 102, 256 Svendby, A.K., 49 Svendsen, J.I., 471 Syvertsen, S.E., 291, 319, 364, 381, 418

Syvitski, P.M., 418 Talwani, M., 30, 36 Tappan, H., 367, 381 Tapscott, C.R., 11, 12, 36 Tauber, H., 458, 471,472 Taylor, A.M., 103, 144, 206, 210 Taylor, D.R., 256 Taylor, J.C.M., 21, 22, 36 Taylor, K.C., 443,449 Teln~es, N., 64 Templet, P.L., 471 ten Have, A., 402 Terwindt, J.H.J., 116, 120, 121, 124, 125, 128, 136, 140, 142, 144, 250, 256

Tessier, B., 210 Tesson, M., 144, 471 Therkelsen, J., 65, 278, 319 Thiede, J., 456 Thierry, J., 143, 173, 196 Thorn, B.G., 144 Thomas, G.M., 34 Thomas, S.A., 196 Thompson, D.B., 243,255 Thompson III, S., 232, 345 Thompson, M., 33, 36 Thompson, S., 364, 405

Thompson, W.O., 465,472 Thomsen, E., 36, 65, 195, 196, 197, 278

Thomsen, R.O., 196 Thomson, M., 282, 291 Thon, A., 210 Thorsnes, T., 448, 449 Thusu, B., 223,232 Timbrell, G., 400, 402, 405, 407, 418 Tjalsma, R.C., 381 Tjelland, T., 34, 102, 143, 210, 231 Todd, R.G., 232, 342, 345, 364, 405 Tonkin, P.C., 401,405 Tonstad, K., 3, 4, 280, 282-284, 291, 350, 351,363, 367,380, 423, 439 Tooby, K.M., 143 Torkildsen, G., 49 Torske, T., 9, 36 Torsvik, T.H., 33, 34 TCrudbakken, B., 33, 291,439 Tralau, H., 180, 197 Traube, V., 448 Trewin, N.H., 84 Tucker, M.E., 47, 48, 49, 206, 210, 467,471 Tudhope, A.W., 84 Tunbridge, I.E, 45, 50 Turner, B.R., 269, 278 Turrell, W.R., 460, 472 Tveiten, B., 4 Tveten, E., 211-213, 225,232 Tyler, N., 120, 143, 233,255 Tyson, R.V., 225,226, 232 Ujetz, B., 368, 375,381 Ulmer, D.S., 64, 65 Ulmer-Scholle, D., 65 Underhill, J.R., 4, 8, 14, 19, 20, 22-27, 34, 36, 37, 82, 85, 145, 148, 173, 175, 176, 178, 180, 193, 194, 197, 210, 244, 256 Upton, B.G.J., 262, 278 V~genes, E., 31, 34, 35, 37 Vagle, G.B., 291 V~gnes, E., 422, 437,440 Vail, E, 291 Vail, ER., 50, 143, 173, 196, 220, 232, 342, 345, 349, 364, 384, 390, 404, 405

Valmori, E., 389, 391,404 van den Berg, J.H., 256 van der Beek, EA., 22, 24, 25, 35-37 van der Zwaan, G.J., 381 van Veen, E, 22, 34, 143, 196 van Veen, EM., 23, 35 Van Adrichem Boogaert, H.A., 179, 194, 197 Van Buchem, ES.P., 166, 173 Van de Weerd, A.A., 103, 120, 144 Van den Berg, J.H., 121, 124, 136, 144

Van der Zwaan, C.J., 16, 37

481

References index

Van der Zwaan, G.J., 365, 367, 369, 381

Van Morkhoven, F.P.C.M., 369, 381 Van Schaak, M., 35 Van Straaten, L.M.J.U., 116, 144 Van Veen, RM., 37 Van Vliet, A., 389, 398,405 Van Wagoner, J.C., 48, 50, 87, 98, 100, 102, 151, 173, 220. 232, 234, 236, 256 Van Weering, T.E.C., 459, 472 Vann, I.E., 364 Vejba~k, O.V., 176, 197 Verba, M.L., 35, 231 Verdenius, J.G., 231,232 Verhoef, J., 28, 36 Verral, E, 348, 364 Veum, T., 443,449 Vially, R., 389, 404 Vigna, B., 403 Vigran, J.O., 223,231,232 Vincent, A.J., 375,381 Vischer, A., 261, 262, 278, 294, 297, 316,319

Visser, C.A., 402 Vogt, ER., 454, 455,456 Vollset, J., 26, 67, 69, 85, 172, 173 Vornberger, E, 448 Vorren, T.O., 1-3, 5, 416, 418, 419, 441,443,447,448, 449, 451,452, 455,455,455, 456 Vosgerau, H., 278, 314, 319 Wach, G.D., 234, 236, 237, 239, 244, 246, 256 Wack, G., 236, 256 Waggoner, J.R., 255 Walker, D., 416, 418 Walker, I.M., 36 Walker, R.G., 42, 49, 54, 65, 79, 84, 134, 142, 384, 385, 389, 395, 405,

423,426, 440, 467,472 Wall, G.R.T., 34 Wallensky, E., 144 Wallis, D., 455 Wallis, R.J., 404 Warren, J.E., 251,256 Warrington, G., 85 Watson, M.E, 439 Watt, W.S., 318 Webb, J.E., 102 Weimer, E, 3, 4, 396, 405, 418 Weimer, R.J., 89, 102 Weiss, H.M., 5, 35, 213, 216, 231, 232

Wells, J.T., 141,144 Welton, J.E., 389, 403 Wenk, A.D., 403 Wennberg, O.P., 13, 37 Werner, F., 418, 456 West, R.R., 367,381 Wezel, F.C., 389, 405 Whitaker, M.F., 172 White, C.D., 234, 256 White, J.W.C., 448, 449 White, N.J., 348, 364 White, R.S., 30, 37 Whiteman, A., 24, 37 Whitham, A.G., 28, 37, 259, 261, 264, 268, 269, 276, 277, 278, 294, 298, 314, 316, 319 Whitley, EK., 108, 144, 235,256 Whittaker, A., 255 Whyatt, M., 401,402, 405 Widmark, J.G.V., 381 Widmier, J.M., 232, 345, 364, 405 Wignall, EB., 224-228, 232 Wiig, M., 146, 151, 160, 167-171, 172

Wilkinson, G.C., 403 Williams, E.G., 102 Williams, R., 4

Willis, B.J., 48, 50, 137, 138, 144, 234, 256, 257 Wills, J.M., 399, 400, 405 Wilson, M.S., 405 Wilson, R.C., 232 Winsvold, I., 37 Withjack, M.O., 201,210 Wolf, C.L., 49 Wonham, J.P., 234, 257 Wood, L.J., 47, 50 Woodroffe, C.D., 136, 144 Woodruff, A.H.W., 404, 418 Woolam, R., 173 Worsley, D., 4, 12, 13, 16-21, 23, 27, 29, 31, 33, 36, 102, 143, 210, 213, 215, 219, 220, 225, 228, 229, 231, 232,291,439

Wrang, E, 36, 65 Wunderlich, F., 116, 119, 120, 139, 144, 250, 256 Xijin, L., 363, 380 Yang, C.-S., 140, 144 Yang, C.S., 234, 245,256, 269, 278 Yielding, G., 36, 69, 71, 85, 147, 173, 345, 364

Yoshida, S., 2, 5, 234, 255, 257 Young, R., 399, 400, 402 Ytice, H., 36 Zaitlin, B.A., 134, 136, 137, 143, 144, 172, 236, 255-257 Zakharov, V.A., 225,232 Zamorano, M., 404 Zang, R.D., 403 Ziegler, EA., 8, 13, 14, 18-21, 23, 24, 27, 29, 37, 79, 83, 85, 142, 148, 174, 175, 176, 178, 193, 194, 197, 294, 319, 350, 359, 364

This Page Intentionally Left Blank

483

Subject index 2D seismic, 280, 407 3D seismic, 90, 199, 280, 407, 411,413-415 Aalenian, 23, 24, 145, 148, 151, 156, 170, 175, 179-181, 183, 184, 194, 201,205 Accommodation, 48, 87, 93, 136-138, 145, 194, 206, 273, 342, 365,457,468 Accommodation space, 39, 48, 87, 93, 96, 98, 100, 136-138, 140, 142, 175, 184, 187, 194, 286, 321, 325, 336, 342, 359, 365,459, 467,468, 470 Aeolian sand 22, 457, 461,465 Agat Field 277, 279, 280, 316 Agat Formation 284 Aggradation 47, 48, 54, 93, 98, 134, 140, 213,442 Alba 400, 402 Albian 28, 29, 31, 213, 216, 225, 259, 261, 263, 264, 267, 270, 276, 277, 279, 282, 284, 285, 289, 290, 293, 296, 298-301, 303, 304, 313, 315-318, 327, 333, 336, 337, 340, 341 Alge Member 211, 215, 216, 219, 223, 225-229 Alluvial 7, 13, 16-18, 20, 22, 39-41, 43, 47, 48, 67, 69, 71, 72, 74, 79, 82, 83, 98, 100, 145, 148, 166, 201 Alluvial fan 16, 17, 31, 41, 81, 105 Alluvium 41 Amplitude map 203,205 AndOya 22, 25, 27-29, 211-213, 218-223, 225-227, 229, 451 Anisian 22, 23, 82 Ankerite 215,220 Annot 384, 388, 390, 393,395 API gravity 100 Aptian 7, 27, 31, 216, 221, 225, 228, 236, 259, 261, 263, 267, 273, 275-277, 280, 282-285, 289, 290, 293, 296, 317, 333, 340, 341,354-356 Aquaplaning 414, 416 Are Formation, 2, 25, 87, 107, 109, 116, 117 Arbroath 401,402 Asselian 19, 20 Australia 246 Bajocian 24, 25, 145, 148, 175, 179-181, 184, 194, 195, 201, 205, 206, 208, 209, 211, 213, 221-223, 227, 231, 293-295, 314, 316, 317, 350 Balder 376, 400, 402, 423 Baltic Sea 459 Baltic Shield 9, 17, 22, 179, 187, 193-195 Barents Sea 1, 2, 7-9, 11, 13, 16, 17, 19-29, 31, 33, 51,213, 221,223-226, 229, 451,452 Barremian 26, 27, 211, 213, 216, 220, 221, 225, 227, 228, 236, 259-261, 263, 264, 271-273, 275-277, 279,

281-285, 289, 290, 293, 315, 317, 340, 341 Basalt 180, 194, 262, 296 Base-level fall 47, 103, 134, 136 Bashkirian 16-18 Basin physiography 1, 3, 148, 199, 201,205, 427, 430, 435, 437 Basin-floor topography 383, 386, 387, 390, 395-398, 402 Bat Group 107 Bathonian 23, 25, 26, 148, 175, 179, 181, 183-185, 194, 195, 201, 205, 207-209, 211, 216, 222, 223, 227, 229-231, 316, 350, 354 Bathymetric data 441,443,444, 447, 448 Bear Island 451,452, 454 Beardmore-Geraldton 388, 395 Bergen area 12 Bergen-Sunnhordland Arcs 13 Bernbjerg Formation 295,305 Billefjorden 13, 17 Biostratigraphic data 81, 87, 89, 104, 108, 128-131, 134, 140, 146, 199, 201,203, 206, 209, 216, 218, 280, 351 Biostratigraphic zonation 148, 149, 160, 279, 284 Biostratigraphy 87, 160, 220, 262, 263, 389 Bioturbated 28, 52, 54, 61, 63, 74, 79, 108, 110, 111, 128, 151, 153, 155-157, 160, 187, 192, 205, 213, 216, 219, 237, 239, 267, 313,426-431,435,437, 461,464, 465 Bioturbation 39, 49, 98, 109-112, 117, 119, 124, 126-128, 130, 137, 160, 189, 206, 213, 215, 218-220, 226, 243, 269, 270, 313,426-428, 430, 437,464, 465,467 Bittern 407 Bjarmeland Platform 13 Bjcrn0ya 13, 14, 16, 17, 19, 31 Bj0rnOyrenna 451 Blodoks Formation 286, 287 Blossville Coast 12, 13 Boreal Ocean 7, 8, 103, 106, 227 Bornholm 124 Bosporus 22 Botryococcus 130, 167, 168, 216, 219, 226, 227 Bouma sequence 55, 384, 409, 426 Breathitt Group 90, 92-95, 98 Bredehorn Member 51-53, 59-61, 63, 64 Breisunddjupet 442, 446 Brent delta 23, 24, 82, 148 Brent Group 2, 83, 147, 148, 151, 171, 172, 205, 350 Bristol Elv Formation 295 Bryne Formation 182-184, 194 Buagrunnen 442, 444, 446, 448 Burrows 97, 110-113, 127, 128, 216, 239, 264, 313, 461, 464, 465 Burton Formation 172

484 C-14 dates 459, 461,470 Calcrete 74 Caledonian 12, 40, 316 Caledonian orogeny 69, 261,350 Caledonide foldbelt 69 California 369, 388 Callovian 23, 175, 179, 183-185, 187, 189, 194, 195, 216, 223, 226, 228-230, 295 Cambrian 1 Campanian 28, 29, 282, 287, 289, 296, 317, 327, 333, 336, 337, 340, 341,363, 378, 421,433 Canyon 311,393 Carbonate 7, 13, 16, 17, 19, 21, 22, 30, 51, 52, 54, 55, 59, 61, 63, 64, 74, 216, 219, 220, 223, 224, 282, 298, 412, 426, 427 Carbonate cement 426, 428 Carbonate platform 16, 17, 19, 21, 22 Carboniferous 1, 2, 7-9, 11-13, 16-20, 25, 32, 51, 61, 63, 90, 99, 176, 179, 194, 297, 298, 388 Carnian 22, 23 Cengio 388, 390, 393-397 Cenomanian 28, 29, 276, 280, 286, 287, 290, 293, 296, 311, 313, 318, 321, 325, 336, 337, 340-342, 344, 350, 354, 356, 357, 363,376, 378 Cenozoic 3, 11, 108, 211,228, 296, 298, 314, 315, 317, 321, 347, 350, 359, 362, 365, 367, 375, 376, 421, 422, 441, 443 Central Graben 21, 82, 148, 179, 183, 184, 194, 399, 401, 408-413 Central Trough 25 Chalk 28, 30, 175,407, 408, 426-428 Channel 30, 31, 42, 57-59, 61, 62, 71, 73, 74, 78, 79, 81, 87, 90, 92-95, 98, 100, 101, 109, 121, 123-125, 127, 128, 130, 134, 139, 183, 184, 187, 233, 234, 241, 243-245, 268-271,273,393, 394, 399, 414, 435,446, 455 Chlorite 146, 171,225, 411,426 Climate 2, 7, 8, 16, 21, 63, 64, 106, 175, 221,227, 396, 430, 457, 460 Clinoform 39, 48 Clinothem 39, 48 Coal 16, 18, 73, 74, 78, 89, 90, 92, 95-97, 100, 113, 127, 128, 184, 185, 216, 218, 347, 348, 359, 369, 426 Coal beds 17, 74, 98, 187, 191 Coal seams 16, 107, 184 Coaly 113, 127, 184, 193, 216, 219 Coastal onlap 79, 323, 349 Coastal plain 2, 7, 17, 25, 26, 183, 185, 193, 195,459, 465 Cod 401,402 Condensed section 48, 286 Coniacian 28, 279, 287, 289, 290, 293, 296, 301,303, 304, 313, 315,317, 318, 340, 350, 357, 358, 363,376 Continental margin 7, 8, 11, 12, 15, 23, 28, 30, 31, 212, 451-453, 455 Continental shelf 3, 7, 8, 435,441-445,447,448, 454 Continental slope 441,443,451,452, 455 Cook Formation 146, 151,167-169, 171, 172, 233, 245 Coriolis force 227 Correlative conformity 48, 145 Cretaceous 2, 3, 7-9, 11, 13, 24-32, 124, 147, 175, 194, 201, 203, 205, 211, 213, 215, 216, 225, 227-229, 233-236, 244, 253, 259-264, 270, 272, 274-277, 279-287, 289, 290, 293-298, 305, 313-318, 321, 323, 325, 327, 333, 336, 337, 340-342, 344, 347, 348, 350, 352, 354-356,

Subject index 359, 361-363, 365-367, 369, 370, 375, 376, 378, 379, 408,421-423,427,432, 433,435,437, 438, 445,470 Cretaceous-Tertiary boundary 325, 327, 333, 336, 337, 340, 341 350, 352, 421,432, 435,437 Cyclicity 39, 48, 116, 127 Danish Central Graben 175, 176, 178, 179, 181-185, 187, 189, 191-195,470 De Geer Zone 28, 29, 31 Debris avalanches 293, 300 Debris flow 193, 216, 286, 308, 409411, 415, 416, 455 Decompaction 325, 348, 350, 361,362 Deep marine 28, 31, 134, 145, 151, 156, 160, 170, 282, 283, 289, 347-350, 352, 361,367, 411 Deep-water deposits 2, 293,298, 316, 384 Deep-water fan 407 Delta 7, 16, 17, 24-28, 31, 41, 87, 95, 97, 116, 120, 121, 134, 138-141, 166, 193, 236, 244, 245, 259, 261, 268, 271-273, 435,469 Delta front 31, 134, 142, 264, 267, 270, 272, 273, 435 Delta plain 13, 25, 26, 31, 87, 89, 95-98, 100, 105, 107, 128, 134, 142, 270, 272, 273 Deltaic 7, 13, 16, 17, 31, 79, 81, 87, 97, 145, 148, 170, 205, 259, 261,268, 270-273, 276, 277, 350, 457, 458, 467 Denmark 1, 64, 67, 79, 84, 176, 194, 228, 277, 457, 458, 467 Desiccation cracks 45, 95, 116, 241 Devonian 1, 2, 8, 13, 16, 19, 39, 40, 42, 69, 78, 298, 350 Diagenesis 100, 389 Dinocysts 168, 296, 318 Dinoflagellates 218, 220, 267, 269, 273 Disconformity 206, 208, 223 Distributary 79, 81, 95, 96, 98, 100, 119, 123, 127, 131,139, 140, 160, 183, 396 Distributary channel 95, 118, 130, 134, 140, 160 Dolomite 19, 215, 430 Domal structure 107, 175 Downlap 287 Drake Formation 145, 169-171 Draugen Field 209 Draupne Formation 147, 282 Dunlin Group 2, 67, 69, 71, 73, 81, 145-148, 150-152, 161, 171, 172 East Greenland 1-3, 7, 12, 13, 16, 17, 20, 22, 23, 25, 26, 28, 30, 32, 51, 52, 63, 64, 108, 206, 209, 223-225,259-261, 276, 277, 293-297, 307, 314, 316, 317, 433,455 East Shetland Basin 322, 351,354, 360 East Shetland Platform 69, 322, 344, 351,354, 360, 361 Ebbadalen Formation 17 Echosounder 442, 443 Egga member 421,422, 428-431,433-438 Eirikson Member 67, 69, 71, 73 Ellesmere Island 12, 29 Embayment 108, 109, 139, 141,223,233, 237, 243,244 England 2, 16, 27, 227, 228, 233-236, 238, 241,244, 253 English Channel 225,227, 459 Eocene 3, 7-9, 12, 25, 29-31, 106, 120, 213, 233, 234, 261, 262, 287, 321, 327, 333, 336, 337, 340-344, 351, 359, 363, 376, 378, 383, 388, 398, 399, 421,423,437 Ephemeral 42, 45 Erosional surface 240, 270, 271,276, 423, 459 Estuarine 103, 111, 112, 119, 123, 130, 131, 133, 135-137, 140, 141, 145, 146, 151, 156, 160, 166-168, 170-172,

Subject index 184, 187, 189, 192, 233-235,237,239, 244, 246, 253 Estuary 119, 125, 130, 134, 136-138, 140, 184, 187, 233, 237, 241-246, 253 Eustasy 175, 398 Evaporites 7, 16, 17, 22, 52, 241,298 Expanding puddle model 226, 227 Extension 7-9, 11-13, 16, 19, 20, 22, 25, 32, 40, 69, 107, 140, 147, 181,201,226, 244, 262, 279, 290, 314, 350 Facies architecture 142, 237, 389, 402 Facies tract 98 Faeroe-Shetland Basin 261,277 Fan delta 31, 41, 67 Farsund Formation 191, 193, 195 Fault crest 61 Fault escarpment 457 Fault-created physiography 199 Feda Graben 189 Fennoscandia 69, 148 Fennoscandian 40, 79, 83, 107, 148, 175, 179, 421,433,437 Fennoscandian Border Zone 175, 179, 187 Finnmark 1 Fischschiefer 284 Fisher Bank 194 Fission track 25 Flexural unloading 348 Flood basalts 261,262 Floodbasin 39-41, 43-46 Flooding surface 100, 136, 137, 240, 241,243,246 Floodplain 14, 16, 17, 67, 71, 73, 78, 82, 87, 90, 92, 93, 95, 96, 98, 100, 183, 187, 236 Fluvial 2, 7, 13, 16-18, 20-22, 26, 28, 30, 39-43, 45, 51, 52, 61, 63, 67, 71-74, 76-79, 81-83, 87-90, 92-95, 97-100, 121, 123, 127, 128, 134, 136, 137, 139, 141, 142, 151, 166, 167, 183-185, 187, 192-195, 222, 223, 233, 234, 237, 239, 244, 246, 269, 272, 276, 277, 295, 457, 458, 470 Fluvial channel 41, 42, 46, 69, 71, 73, 76, 78, 137, 183, 184, 236 Fluvial sand 100 Foraminifera 169, 215, 216, 226, 365, 367, 369, 375, 376, 378, 427, 435 Forced regression 167, 171,206, 273,457, 459, 467 Forced regressive systems tract 48, 468 Foresets 89, 119, 126, 246, 247, 253, 264, 268, 269, 465, 466, 468 Foreshore 218, 222, 229 Forties 399, 400 Fosdalen Member 263, 267, 270, 276 Fram Field 281 Frigg 30, 399, 400 Fr0ya High 107, 433,435,436 FrOyabanken 441,442, 444-448 FrOyryggen 442, 445,446 Fuglen Formation 211, 213, 215, 216, 222, 223,226-229 Garn Formation 199-20 l, 203, 205,206, 208, 209 Gas hydrates 452, 453 Gassum Formation 79, 83, 194 Genetic stratigraphic sequence 220 Geographical Society 0, 259, 261,297, 308, 318 Geomorphological evidence 435 Gipsdalen Group 16 Gjallar Ridge 259, 260

485 Gj0a Field 279 Glacial processes 441,442 Glaciomarine sediments 442-444 Glauconite 234, 411, 412, 426 Graben system 13, 20 Grane 407 Gravity flow 52, 59, 193, 293, 294, 305, 308, 309, 316, 317, 411,415,418 Great Britain 1 Greenland 1, 2, 7-9, 11-13, 16, 17, 20, 25-30, 32, 51, 52, 64, 84, 138, 140, 142, 176, 222, 227,261,277, 283,294, 317,398, 421,443, 451,452 Greensand 235 Griptarane 444, 446 Growth fault 140 Guipuzcoa-Jaizkibel 388, 397 Gulf of Mexico 410 Gullfaks 69, 79, 151,161, 164, 165, 167-172, 233,354 Gypsum 117 Half-graben 18, 19, 60, 61, 175, 185, 189, 191, 195, 201, 282, 396, 397 Halten Terrace 25, 31, 87, 103-109, 131, 141, 199-201,206, 209, 316, 433 Haltenbanken 103,235,441,442, 444, 445,447,448 Hanging wall 140, 201, 203-205, 208, 296, 298, 300, 310, 317, 325,470 Hanging-wall slope 178, 187, 189, 192, 193, 195 Harding 400, 402 Harstad Basin 211 Hauterivian 26, 211, 216, 220, 221,225,227, 228, 263,269, 272, 277, 279, 283, 289, 293, 315, 317, 325,341 Heavy-mineral composition 433,437 Heidrun 2, 87-90, 92, 93, 99-105, 108, 109, 114-142, 209, 235 Heimdal 399, 401,407,408, 410, 413-417 Hekkingen Formation 211, 215, 216, 219, 223,225, 227-229 Helgeland Basin 25 Hemipelagic 193, 279, 282, 284-286, 289, 290, 426, 427, 429-431,433 Heno Formation 189, 191,195 Hettangian 71, 79, 81-83, 107, 148 Hiatus 7, 24, 25, 29-31, 170, 179, 183, 194, 223, 226, 227, 237, 264, 317, 327, 333, 336, 337, 340, 341, 351, 359, 376, 378 High-density turbidity current 54 High-resolution biostratigraphy 212 Highstand systems tract 48, 141,142, 211,221,273, 276 Hild-Alwyn Alignment 336, 354, 360 Hinterland rejuvenation 30, 32, 67, 148 Hold with Hope 259, 261-263, 272, 275-277, 318 Holocene 1, 3,245,457-459 Home Forland Formation 259, 261,263, 264, 276, 318 Horda Platform 26, 67, 69, 71, 73, 74, 76, 78, 79, 81-83, 145, 147, 151, 155-157, 160, 161, 166-170, 172, 289, 322, 341,344, 354, 358, 360, 361 Hornelen Basin 1, 39-42, 47, 48 Hornsund 13 Hornsund Basin 17 Hummocky cross-bedding 97, 98 Hydrocarbon column 100 Hydrocarbon discoveries 146, 175,280 Hydrocarbon discovery 279 Hydrocarbon recovery 233,234

486 Hydroplaning 293, 300, 305 Ice flow 441,444, 446 Ice sheet 441,444-448 Ice stream 444, 446, 448 Ice-sheet dynamics 441,443,448 Ile Formation 201 Inbis Channel 455 Incised valley 87, 90, 95, 98-100, 134, 166, 184, 187, 233, 237,253,269-271,273 Injection 100, 399, 401,402, 416, 427, 429, 435 Interdistributary 139 Interglacial 443,448 Intraplate extension 8 Intraslope basin 421 Intrusions 296, 390 IRAP RMS, 250, 255 Ireland 23, 24 Isle of Wight 233-242, 244-247 Isotope data 69, 107, 432, 433 Italy 286, 388, 416 Jackfork Group 384 Jameson Land 12, 25-27, 51, 52, 61, 63, 64, 108, 138, 140, 223 Jameson Land Basin 27, 51,297 Jan Mayen Lineament 9, 12, 422, 423 Janusfjellet Formation 26 Johansen Formation 145, 151,166, 172 Jorsalfare Formation 289, 421,427, 428,433, 438 Jotun 407 Jurassic 1-3, 7-9, 11-13, 22-26, 32, 33, 51, 63, 67, 69, 71, 79, 81-83, 103, 104, 106-108, 124, 128, 136, 145-152, 160, 169, 171, 172, 175, 176, 178-185, 187, 191, 193-195, 199, 201, 203, 205-207, 209, 211, 213, 215, 216, 223-227, 229, 233-236, 245, 259, 261, 262, 264, 270-272, 276, 279, 280, 282, 283, 289, 290, 294-298, 305, 310, 314-317, 344, 354, 355, 363, 365, 422, 437, 438, 445,470 Jylland 457-460, 464 Jylland Current 459, 460 Kangerlussuaq 12, 277 Kazanian 21, 22, 52, 63 Kentucky 90, 92, 94, 98 Kerogen 213, 215, 216, 218-220, 224 Kimmeridgian 26, 27, 175, 187, 189, 191-193, 195, 219, 223-230, 295, 316 Klippfisk Formation 27, 216, 219, 220, 225 Knurr Formation 220, 228 Kolje Formation 211, 216, 220, 225 Kong Karls Land 27, 223 Kvitebjcrn 354, 355, 360 Kviting Formation 29 Kyrre Formation 287, 289 K-T boundary 423,426, 427,429-433,435,437 Lacustrine 7, 16, 17, 20, 21, 41, 43-45, 78, 81, 82, 183-185, 187, 192, 195,216 Ladinian 22, 23 Lagoonal 184, 185, 187, 191, 192, 195, 213, 222, 233, 236, 237 Lake Malawi 348 Laminar flow 383, 410

Subject index Langgrunna 442, 444, 446-448 Lava 30 Lava deltas 30 Lava plateau 13 Lista Formation 408, 423 Liverpool Land 51 Lobe 51, 54-57, 59, 61-64, 111, 115, 117, 121, 122, 141, 268, 285,394, 414, 444 Lofoten 211-213, 225, 228, 316, 442, 444, 445, 451 Lofoten Basin Channel 455 Lola Formation 187, 189, 191 Lomre Terrace 145, 151, 155, 156, 161, 166-170, 172, 322, 340, 351,358, 360 Lower Greensand Group 234, 236 Lowstand fan 51 Lowstand systems tract 98, 275, 276 Lulu Formation 185, 194 Lyr Formation 27 Magnus Basin 322, 351,354, 360 Mandal High 189 Marflo Ridge 283 Marries Bleues 384, 388, 390, 393, 395, 397, 398 Marnoso-Arenacea 388, 390, 391,393-395, 397 Mass movement 307, 414 Maximum fooding surface 48, 98, 100, 128, 130, 137, 161, 221,223,224, 276 Maximum transgressive surface 156, 166, 170 Meandering 16, 79, 82, 87, 89, 90, 92-94, 99, 100, 124, 125, 136, 137, 236, 243 Meandering channel 414 Megasequence 161, 166-170 Melke Formation 199, 203, 205-209, 229 Mesozoic 2, 8, 51, 108, 147, 194, 211, 213, 222, 223, 228, 259, 261, 262, 273, 276, 277, 294-296, 298, 314-316, 354, 398, 438, 441 Meteorite 431 Micropalaeontological analysis 321,325, 340, 347, 370, 378 Micropalaeontological data 168, 321, 325, 327, 333, 336, 337, 340-342, 344, 365, 376, 379 Micropalaeontology 160, 167, 168, 325, 348, 350, 362, 377, 378 Mid-Cimmerian Unconformity 23, 24 Mid-Norway 1-3, 32, 51, 64, 87, 103, 108, 199, 203, 209, 234, 235, 282, 283, 367, 441 Midland Valley 16 Midland Valley-Ling Depression 13, 20 Millstone Grit 16 Miocene 3, 31,287, 321,323, 327, 333, 336, 337, 340, 342, 344, 351,359, 363, 376, 378, 379, 388, 421 Modern systems 141 Moine Thrust 433 Montrose 401,402 Moray Firth 17, 25, 81, 82, 148 Morphogroups 368, 375 Moscovian 16-18 M~l~y Terrace 153, 166, 169 M~lcyplat~et 442, 446-448 M~nedal Formation 293,296, 301,303,304, 306-310, 318 Mgtsnykan Formation 211, 213, 216, 218, 219, 223, 226, 228, 229 Mere Basin 13, 25, 30, 31,279, 286, 378, 421-424, 433-438 MCre-TrCndelag Fault Complex 13

Subject index Namurian 14, 16, 17 Nansen Member 67, 69, 71, 73, 81, 161 Neill Klinter Group 108, 109, 138, 140 Nelson 401,402 Neogene 7, 30-32, 195, 261, 327, 340, 347, 359, 363, 365, 367 Neritic 367 New Zealand 388 Nordkapp Basin 13 Nordland 107, 211-213, 216-219, 221-230 Nordland Ridge 107, 201, 316 Nordmela Formation 26 Norian 82 North Atlantic 3, 7-9, 11, 30, 51, 84, 229, 244, 260, 277, 283,286, 289, 294, 317, 367, 435,437 North Atlantic Rift System 69, 83, 357 North Sea 1-3, 7-9, 11, 13, 16-18, 20-28, 30, 31, 33, 39, 49, 51, 67-69, 71, 82, 145-149, 151, 168, 169, 172, 175, 176, 178-180, 189, 193-195, 201, 205, 206, 233, 259, 276, 279, 280, 282, 286, 287, 289, 290, 294, 316, 321-324, 335, 342-344, 347, 348, 350, 351, 354, 359, 362, 363, 365, 367, 369, 375, 376, 378, 379, 383, 384, 398-400, 407, 411, 415, 416, 418, 423, 427, 432, 448, 452, 454, 458-460, 470 North-East Greenland 259-261,276, 277, 316, 317 Norwegian Coastal Current 459 Norwegian Sea 7, 8, 12, 13, 17, 20-22, 25, 27, 28, 31, 69, 84, 261,375,448 Norwegian-Danish Basin 81-83, 175, 176, 178, 179, 181-185, 187, 192-195 Norwegian-Greenland Sea 7, 8, 11, 25-30, 32, 106, 225, 317, 421,451-453,455 Not Formation 199, 205, 206, 208 Old Red Continent 16 Old Red Sandstone 78 Oligocene 31, 287, 327, 333, 336, 337, 342-344, 351, 359, 363,376, 378, 379, 388, 421,437, 442 Olistostromes 407, 409-411, 413, 414, 416 Onadjupet 442, 446 Onlap 40, 61, 67, 176, 179-181, 187, 189, 194, 199, 204-206, 208, 259, 282, 283, 285-287, 289, 290, 323, 348, 349 Oosterschelde 121, 124, 136 Ordovician 388 Organic matter 128, 213, 215, 216, 218-220, 225,226, 465 Organic-rich shales 52, 63, 193, 195,225-227 Ormen Lange 421-423,427, 430, 431,433,435,437 Oseberg Formation 156, 170-172 Oslo Graben 2, 17 Oslo region 1 Outcrop analogue 233, 234, 457 Outcrop studies 233,235,254, 383, 387, 390, 395,397, 398 Oxfordian 23, 26, 27, 147, 175, 184, 185, 187, 189, 191, 194, 195,205, 211, 212, 219, 221-230, 295,350 Oygarden Fault Complex, 12, 13, 69 Oygarden Fault Zone, 282 P : B % (planktonic : benthonic ratio), 369, 376 Palaeobathymetry 3, 347-349, 352, 357, 360, 362, 363,365368, 377, 379 Palaeocene 3,350, 351,358, 359, 363 Palaeogene 29, 194, 259, 261,262, 270, 321,333, 347, 363,

487 365, 367, 369, 421,433,435 Palaeoslope 67, 69, 71, 78, 79, 82, 301 Palaeosols 71, 74, 95, 98 Palaeowater depth 279, 284, 365,367, 369, 375-377, 379 Palaeozoic 1, 3, 8, 9, 11, 13, 19, 20, 51, 106, 194, 261, 315 Palaeocurrent 140 Palynology 160, 167, 168, 180 Palynomorph 115 Pangea 23, 32 Paralic 109, 182, 184, 194, 277 Parasequence 98, 100 Parasequence set 206, 209 Passive margin 31, 107, 390 Peat 74, 90, 96, 98, 100, 457,465,467, 468 Pechora 13 Peneplanation 20, 32 Permeability 87, 99-101, 119, 171,233, 251-254, 383, 394, 395, 399-402, 437 Permian 2, 7-9, 18-22, 32, 51-53, 61, 63, 64, 69, 147, 176, 179, 194, 236, 259, 293, 297, 298, 300, 305, 309, 310, 314, 315,317, 318, 350 Plate tectonics 8 Pleistocene 3, 8, 31,410, 443,457,459, 465 Pliensbachian 73, 81, 82, 87, 105, 107, 108, 136, 141, 145-148, 151, 156, 160, 161, 166-169, 171, 172, 179, 183, 194 Pliocene 3, 31,321,327, 333, 336, 337, 340-342, 344, 351, 359, 363,365, 378, 379, 410, 441-443 Porosity 93, 100, 119, 171,383,400-402, 437 Processes 1, 45, 71, 73, 87, 88, 93, 98, 100, 109, 115, 117, 124, 127, 134, 139, 141, 178, 233, 237, 241, 244, 276, 279, 290, 291, 293, 298, 308, 316, 317, 344, 347, 363, 365, 379, 383, 384, 398, 407, 411, 413, 415, 418, 441, 443,451,455,468, 469 Progradation 7, 23-25, 27, 30, 43, 47, 48, 63, 82, 83, 137, 141, 145, 148, 153, 156, 166, 168, 171, 172, 187, 189, 192, 193, 195, 206, 209, 213, 226, 259, 261, 268, 270, 273, 276, 351, 353, 359, 435, 442, 457, 459, 461, 466-468, 470 Provenance 16, 17, 22, 28, 104, 106, 107, 128, 130, 225, 298, 300, 305, 314, 316, 317, 387, 411, 421, 430-433, 437, 438 Ramp 140, 184, 185, 193, 259, 273,277, 386, 389 Rannoch Formation 156, 201 Raude Member 67 Raufite Member 211, 213, 219, 226, 228, 229 Ravinement surface 137, 160, 189, 468, 470 Ravnefjeld Formation 51-53, 57, 59, 63, 305 Recent 1, 3, 22, 25, 30, 40, 47, 67, 69, 145, 206, 259, 260, 279, 316, 348, 384, 385, 407, 448, 451, 455, 457, 458, 465 Recovery factor 87 Red beds 16, 17 Regression 17, 23, 47, 151, 155, 187, 206, 284, 287, 458, 459, 468-470 Regressive erosion surface 206 Regressive systems tract 221,228 Reservoir 2, 3, 51, 87, 88, 95, 99-101, 103, 108, 115, 131, 133, 142, 145, 146, 151, 169, 171, 172, 175, 176, 194, 199, 209, 233, 234, 248, 251, 253-255, 259, 276, 383, 387, 389, 398-402, 421, 422, 429, 431, 433, 435, 437, 457, 459, 470

488 Reservoir architecture 103,246, 402 Reservoir characterization 233, 253, 255 Reservoir geometry 199 Reservoir modelling 250 Reservoir rocks 1-3, 51,147, 175, 194 Reservoir zonation 108, 151 Reservoir zonation scheme 209 Retrogradation 47, 48, 141,145 Rhaetian 67, 69, 71, 79, 81, 82, 87, 194 Ribban basin 106, 107, 211, 213 Rift 11-13, 16, 20, 22, 23, 25, 26, 29, 32, 51, 64, 84, 106, 107, 147, 175, 195, 201, 209, 213, 228, 236, 259, 261, 276, 277, 282, 293-296, 314-317, 344, 347, 350, 352, 354, 356-359, 363,367 Rift basin 22, 51,235,295 Rifting 2, 3, 7-9, 11, 12, 16, 24, 25, 28-32, 64, 69, 106, 147, 178, 201, 227, 244, 261, 262, 272, 279, 282, 283, 289, 290, 293, 294, 296, 314, 316, 317, 344, 350, 355, 398, 435,437 Ringkcbing-Fyn High 175, 176, 178-181,187, 192-194 River 16, 23, 73, 78, 79, 82, 116-118, 120, 134, 137-141, 222, 234, 236, 263, 267, 269, 273, 305, 318, 433 River planform 79 Rock-Eval 213, 215 Rockall Plateau 28 Rogn Formation 26 Rold Bjerge Formation 293, 296, 298-301, 317 Rona Ridge 69 Ror Formation 107, 130, 131,142 Ross Sea 448 Rotliegende 18, 19, 21 Ryazanian 26, 191-193, 195, 281-283, 290, 316, 322, 323, 325, 327, 333, 336, 337, 340-343, 350-352, 354, 357 Rodby Formation 284 Salt 17, 107, 195, 201,396, 398 Salt Dome Province 181 Sandy debris flow 383, 410, 411, 416, 417 Santonian 28, 29, 261, 263, 279, 280, 282, 289, 290, 313, 336, 340 Schuchert Dal Formation 51, 59, 63 Scoresby Sound, 12, 13 Scoresby-Bergen lineament 12, 13 Scythian, 82 Seafloor spreading, 7, 8, 28, 32, 290 Sediment supply, 31, 39, 40, 47, 48, 69, 87, 138, 141, 145, 166, 175, 189, 221,342, 344, 365,451 Seismic attribute 199, 203, 209 Seismic attribute analysis 288, 415 Seismic data, 3, 4, 18, 201, 205, 208, 209, 216, 279, 280, 282-286, 288-290, 321~ 323, 340, 349, 351, 352, 407, 415, 437, 448 Seismic expression, 201 Seismic sequence, 348 Seismic stratigraphy, 201,323, 325 Sele Formation, 423 Selje High, 283, 289, 433 Senja Fracture Zone, 9, 28 Sequence, 2, 7, 12, 16-18, 22, 27-31, 48, 52, 57, 87, 90, 95, 97-99, 103, 128-130, 134, 136-138, 141, 142, 145, 146, 148, 150, 151, 155-158, 160-162, 164, 166-172, 176, 184, 191, 192, 199, 201, 211, 215, 217, 218, 220, 221, 224, 225, 227, 228, 233-236, 240, 246, 253, 261, 265,

Subject index 266, 272, 273, 281, 282, 284, 285, 323, 325, 348, 350, 351,359, 378, 379, 384, 390, 407 Sequence boundary, 47, 128, 156, 160, 167, 169, 175, 184, 187, 211, 221, 223, 227, 228, 233, 237, 239-241, 246, 267, 276, 468,469 Sequence stratigraphy, 1, 2, 47, 48, 145, 147, 148, 161,220, 236, 237, 259, 272, 347, 350, 367, 369, 379, 385, 389, 467 Shale Grit, 388, 390, 393, 395 Shallow marine, 2, 7, 13, 16, 17, 19, 22, 24, 26-28, 32, 52, 67, 79, 81, 83, 103, 151, 160, 175, 184, 187, 189, 191-193, 199, 205, 206, 213, 223, 226, 229, 259, 261, 268, 269, 273, 277, 279, 282, 283, 289, 295, 298, 347, 348, 351,359, 369, 435,457, 470 Shallow water, 17, 28, 120, 323, 368 Sheets, 31, 45, 72, 393,402, 441,448 Shelf, 2, 3, 7-9, 17, 20, 22-25, 28, 29, 31, 51, 64, 73, 79, 115, 122, 136, 145, 146, 151, 155, 156, 166, 167, 172, 176, 179, 187, 193, 199, 203, 205, 206, 209, 211, 213, 215, 219-226, 228, 229, 259, 260, 270, 276, 277, 282, 283, 293, 294, 316, 344, 369, 376, 377, 396, 398, 421, 423, 427,433,441-444, 446, 448,457, 466, 467, 470 Shelf break, 28, 31,446, 451,454, 455 Shore-parallel current, 467 Shoreface, 3, 79, 81,109, 115, 137, 151,155, 156, 160, 167, 168, 170-172, 179, 183-185, 187, 189, 191, 192, 194, 195, 199, 203, 205, 206, 208, 213, 218, 222, 259, 264, 270-273, 276, 457, 458, 461,464, 466, 467, 470 Shoreline, 17, 23, 30, 48, 67, 78, 83, 115, 136, 140, 145, 151, 153, 155, 156, 160, 166, 168, 171, 172, 209, 325, 398, 435,459, 465 Shoreline trajectory, 457 Siderite, 73, 74, 110, 111, 113, 121,215, 220, 426 Silurian, 40, 350 Sinemurian 67, 69, 71, 73, 79, 81-83, 107, 145, 148, 161, 171, 194 Sinuosity, 79 Sinuous, 126, 264, 413,414 Siri, 407-414, 417, 418 Skagen Odde, 3,457, 458, 461,468 Skagerrak-Kattegat Platform, 192 Skjoldryggen, 442, 444, 445 Sklinnabanken, 444, 447 Sklinnadjupet, 442, 444, 445,447 Slide, 3, 279, 286, 293, 294, 296, 300-304, 316, 451,453, 454, 468 Slide headwall, 451 Slide scar, 451,453 Sliding, 286, 293, 301,303, 308, 309 Slope, 30, 78, 79, 82, 175, 178, 183, 187, 189, 216, 276, 279-281, 283, 285-288, 300, 301, 307, 309, 310, 314, 318, 367, 369, 376-378, 384, 395, 414, 422, 437, 451, 452, 455,457-459, 467, 470 Slope apron, 293, 296, 303, 304, 309, 310, 318, 386 Slump, 268, 293,294, 296, 300-304, 316, 414 Slumping, 57, 58, 280, 285, 286, 293, 301, 303, 308, 309, 313,395, 399, 443 Slurry, 395,397 SlCrebotn Subbasin, 421,422, 433,435, 436 Sm01a, 446 SmCrbukk, 103-105, 108, 109, 115-117, 119, 123, 126-130, 132, 134, 138-142, 199-201,203-209, 235 SmCrbukk Fault, 199-201,203,205,208

Subject index Soft-sediment deformation, 39, 41-45, 48, 49, 93 Sogn Graben, 279-283,285,286, 289, 290, 322, 344 Sognefjord Formation, 26 Sola Formation, 283 Sorgenfrei-Tornquist Zone, 179-181, 183, 184, 187, 194 Source rock, 7, 22, 26, 51,193, 195 Spekk Formation, 209, 445 Springar Formation, 427 Statfjord Field, 67 Statfjord Formation, 2, 67, 69, 71-74, 76, 78, 79, 81, 82, 145, 151, 157, 161, 166, 172 Stauning Alper Fault, 297 Steensby Bjerg, 259, 261-264, 266, 267, 269, 270, 272-274, 276, 277 Stochastic reservoir models, 248 Stord Basin, 13, 18, 21, 22, 322, 341,344, 351-355, 361 Storegga Slide, 3,446, 451 Stratified basin model, 226 Stratigraphic play, 51,277 Strike-slip, 9, 12, 13, 16, 40, 235,388, 390 Structural restoration, 321, 323-325, 327, 333, 336, 337, 340-342, 344, 347, 350, 352, 360, 365,376-379 Sto Formation, 211, 213,222, 227, 228 Subaerial, 52, 117, 148, 209, 241, 259, 272, 347, 350, 369, 376, 378, 423,458, 466 Subaerial unconformity, 48 Submarine canyon, 398 Submarine fan, 30, 289, 398, 430, 433 Svalbard, 7, 12-14, 16, 17, 22, 26-29, 31,223,277, 286 Sverdrup Basin, 23 Swamp, 90, 96, 98 Syn-rift, 26, 51, 69, 71, 199, 201,236, 316, 317, 350 Synaeresis, 116, 117 T-R sequence, 221-223, 226 Tail End Graben, 178, 181, 185 Tampen Spur, 67, 69, 71, 74, 76, 78, 79, 81-83, 156, 160, 161, 166-168, 325,327, 357 Tang Formation, 423,427 Tarbert Formation, 208, 209 Tatarian, 21, 22 Tau values, 368, 369, 376, 378 Tectonic influence, 221 Tectonics, 8, 9, 17, 19, 40, 67, 227, 229, 387, 389, 390, 398 Tectonism, 2, 3, 7, 28, 29, 290 Tertiary, 7-9, 11, 13, 29, 32, 147, 235, 282, 289, 290, 321, 323, 325, 327, 333, 336, 337, 340, 341, 344, 347, 350, 354, 360, 361, 367, 388, 393-396, 398-400, 421, 423, 432, 435,441 Tethys Ocean, 103, 106 Thermal dome, 23, 82, 83 Thermal maturity, 216 Thermal relaxation, 8, 22, 25, 32 Tholeiitic flood basalts, 30 Tidal channel, 124, 243,268, 269, 271,272, 276 Tidal delta, 244 Tidal flat, 245, 246, 269, 270, 276 Tidal ravinement surface, 136, 246 Tilje Formation, 2, 103-109, 114-116, 127-130, 136-138, 140-142, 234, 235,250 Till, 26, 441,443,444 Toarcian, 26, 82, 83, 107, 108, 136, 141, 145, 147, 148, 151, 156, 167-172, 179, 180, 183, 185, 194, 213

489 TOC (total organic carbon), 195, 213, 215, 216, 218-220, 223-225,228, 423 Tofte Formation, 107 Tourelle, 388, 390, 393,395 Tournaisian, 13, 16 Trace fossil, 156, 160, 264, 270 Trace-fossil suite, 109, 116, 117, 121 Traill f3, 60, 61,259, 261,293-297, 314-318 Transgression, 16, 17, 19, 21, 22, 24, 27-29, 51, 79, 81, 82, 95, 115, 134, 136-138, 142, 151, 161, 166, 169, 171, 172, 184, 185, 187, 189, 192-195, 209, 221-227, 236, 245, 270, 272, 273, 276, 277, 282, 284, 286, 287, 435, 458, 468, 470 Transgressive surface, 81, 136, 140, 169, 185,221,222, 237, 239-241 Transgressive systems tract, 48, 142, 221,226, 228, 273,276 Triassic, 2, 7-9, 11, 13, 19, 22-26, 32, 61, 63, 67-69, 71, 72, 79, 81, 82, 87, 106, 107, 147, 148, 175, 179, 180, 182, 201, 259, 261, 262, 273, 276, 282, 293, 295, 298-300, 305, 314, 315,317, 318, 342, 350, 438, 441 Troms, 17, 26-28, 31,211-213, 215,216, 218, 220-230, 444 Tromsr Basin, 31 Tryggvason Formation, 287, 290 Tra~nabanken, 229, 442, 444, 445,447, 448 Tra~nadjupet, 442, 444, 445,447, 448, 451,453,454 Tr~ndelag Platform, 11, 22, 25, 26, 31, 105, 107, 316, 423, 433 Tsunami, 468 Tub~en Formation, 25, 26 Tuffites, 180 Turbidite, 3, 30, 51, 58-64, 187, 193, 259, 283, 285, 289, 301, 303, 311-313, 316, 317, 384, 385, 407, 409-411, 414-4 17, 423,426-430, 433, 437 Turbidite system, 59-64, 301,396 Turbidity current, 57, 58, 384, 394, 427, 428 Turonian, 28, 29, 263, 276, 279, 284, 286, 287, 289, 290, 293, 296, 301, 303, 304, 309, 313, 315, 317, 318, 325, 327, 337, 340-342, 350, 376 Uer Terrace, 322, 340, 360 Unconformity, 18, 20, 22, 24-26, 28, 31, 145, 148, 151,175, 179-185, 187, 193-195, 212, 221, 222, 225, 229, 259, 262-264, 267, 270-273, 275, 276, 282, 289, 340, 351, 421,423,433,435,442, 443, 468 Underplated magmatic bodies, 11 Uplift, 3, 7, 16, 17, 20, 22, 24, 25, 27-32, 51, 69, 82, 141, 148, 151, 170, 175, 176, 178-180, 193-195, 201, 205, 208, 212, 259, 261, 277, 279, 283, 286, 287, 289, 290, 293, 294, 296, 300, 301, 305, 315, 317, 321, 344, 351, 357, 359, 365, 379, 398, 421, 423, 433, 435, 437, 457-459, 469 Upwelling, 227 Uralian orogeny, 19, 21, 22 USA, 1, 90, 92, 94, 98, 201,234 Utsira Formation, 31,378 Utsira High, 18-20, 67, 69, 71-74, 76-79, 82, 322, 351,354, 357, 361 Valley fill, 166, 167, 184, 187 Vardeklc~fl Formation, 206, 209 Variscan orogeny, 19, 20 Vectis Formation, 233-237, 239, 245, 246, 250, 253 Vega Sund Formation, 293,296, 311-313, 317, 318

Subject index

490 Vendsyssel, 458 Venezuela, 120, 233 Veslefrikk Field, 146, 171 Vester~len, 211-213, 228 Vestfjorden, 213,442, 444, 445,447, 448 Viking Graben, 2, 17, 21, 25, 26, 67, 69, 71, 73, 79, 81-83, 145-148, 157, 160, 161, 166-168, 171, 194, 279-283, 289, 322, 327, 333, 335-337, 344, 351, 352, 354, 357, 358, 361,365,401,408, 413, 414, 416, 417 Viking Group, 199 Visean, 13, 14, 16, 17, 32 Volcanic, 18-20, 107, 194, 270 Volcanism, 7, 148, 180 Volgian, 26, 147, 175, 192, 193, 195, 211, 224, 226, 227, 230, 272, 307, 314, 316, 352, 353 V~le Formation, 421-423,427-429, 433 V0ring Basin, 3, 25, 27, 28, 30, 31,260, 282, 287, 316, 317 V0ring Escarpment, 25 Vcring Marginal High, 11, 25

Wandel Sea, 28, 29 Water escape, 74, 287, 288, 426, 427 Wave erosion, 457, 469, 470 Wave-influenced coastline, 457 West Antarctica, 448 Western Norway, 2, 12, 13, 39 Westphalian, 17 Willapa Bay, 124, 136 Wollaston Forland, 259, 261,307, 310, 316 Wordie Creek Formation, 61, 63, 262, 293, 298, 299, 314, 317 Yoredale cycles, 16 Yorkshire, 145, 160, 166, 169, 170, 172 Ypresian, 11, 30, 400, 422 Zechstein Group, 18, 21, 22 Zircons, 316

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