Dynamics of the Norwegian Margin

Dynamics o f the Norwegia n Margi n Geological Societ y Specia l Publication s Series Editors A. J . HARTLE Y R. E . ...

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Dynamics o f the Norwegia n Margi n

Geological Societ y Specia l Publication s Series Editors A. J . HARTLE Y R. E . HOLDSWORT H

A. C . MORTO N M. S . STOKE R

It i s recommended tha t referenc e to al l or par t o f this book shoul d b e made i n one o f the followin g ways: N0TTVEDT, A . e t al . (eds ) 2000 . Dynamics o f th e Norwegian Margin. Geologica l Society , London . Special Publications , 167 . SUNDVOR, E. , ELDHOLM , O. , GLADCZENKO , T . P . & PLANKE , S . 2000. Norwegian-Greenland Se a thermal field. In: NOTTVEDT , A . e t al . (eds) Dynamics o f th e Norwegian Margin. Geologica l Society, London, Specia l Publications , 167 , 397-410 .

GEOLOGICAL SOCIET Y SPECIA L PUBLICATIO N NO . 16 7

Dynamics o f the Norwegia n Margi n

EDITED B Y

A. N0TTVED T

Norsk Hydr o Researc h Centre , Bergen , Norway

Co-editorial Boar d

BJ0RN T . LARSE N RO

Norsk Hydr o Exploration , Osl o Universit

SNORRE OLAUSSEN HARAL Saga Petroleum , Osl o Norwegia

Y H . GABRIELSE N y of Bergen

BJ0RN T0RUDBAKKE N 0RJA Saga Petroleum , Osl o Statoil

JAKOB SKOGSEI D University o f Osl o

2000 Published b y The Geologica l Society London

D BREKKE

n Petroleu m Directorate , Stavange r

N BIRKELAN D , Stavanger

THE GEOLOGICA L SOCIETY The Geologica l Societ y o f Londo n wa s founde d i n 180 7 an d i s th e oldes t geologica l societ y i n th e world . It receive d it s Roya l Charte r i n 182 5 for th e purpos e o f investigatin g th e minera l structur e o f th e Earth " an d i s no w Britain's nationa l societ y fo r geology . Both a learne d societ y and a professiona l body, th e Geologica l Societ y is recognized b y the Departmen t o f Trade an d Industry (DTI ) as the chartering authority for geoscience, abl e to awar d Chartere d Geologis t statu s upon appropriately qualified Fellows . Th e Societ y ha s a membershi p o f 8600 , o f who m abou t 150 0 liv e outsid e th e UK . Fellowship of the Societ y is open t o thos e holdin g a recognized honour s degree i n geology or cognate subjec t and wh o have a t leas t tw o years " relevan t postgraduate experience , o r wh o hav e no t les s tha n si x years' relevan t experienc e in geology o r a cognat e subject . A Fello w wit h a minimu m of five years' relevan t postgraduate experienc e in th e practice of geolog y ma y appl y fo r chartere d status . Successfu l applicants are entitle d to us e the designator y postnominal CGeo l (Chartered Geologist) . Fellow s o f th e Societ y ma y us e th e letter s FGS . Othe r grade s o f membershi p ar e availabl e to members no t ye t qualifyin g for Fellowship . The Societ y ha s it s ow n publishin g hous e base d i n Bath . UK . I t produce s th e Society' s internationa l journals. books an d maps , an d i s th e Europea n distributo r fo r publication s of th e America n Associatio n o f Petroleu m Geolo gists. (AAPG) . th e Societ y fo r Sedimentar y Geolog y (SEPM ) an d th e Geologica l Societ y o f Americ a (GSA) . Members o f th e Societ y ca n bu y book s a t considerabl e discounts . Th e publishin g Hous e ha s a n onlin e booksho p (http://bookshop.geolsoc.org.uk/. Further informatio n o n Societ y membershi p ma y b e obtaine d fro m th e Membershi p Service s Manager . Th e Geological Society . Burlingto n House , Piccadilly . Londo n W1 V OJU . UK . (Email : [email protected]: tel: +44(0)17 1 4349944) . The Society' s We b Sit e ca n b e foun d a t http://www.geolsoc.org.uk/ . Th e Societ y i s a Registere d Charity . number 210161 . Published b y Th e Geologica l Societ y from : The Geologica l Societ y Publishin g Hous e Unit 7 , Brassmil l Enterpris e Centr e Brassmill Lan e Bath BA13JN , U K (Orders: Tel . +4 4 (0)122 5 445046 Fax +44(0)122 5 442836 ) Online bookshop : http://bookshop.geolsoc.org.u k First publishe d 200 0 The publishers make no representation, express or implied . with regar d t o th e accurac y o f th e informatio n contained in this book an d canno t accep t an y lega l responsibilit y for any error s o r omission s that ma y b e made . © Th e Geologica l Societ y o f Londo n 2000 . Al l right s reserved. N o reproduction , cop y o r transmissio n o f thi s publication ma y b e made without written permission . N o paragraph o f thi s publication ma y b e reproduced , copie d or transmitte d sav e wit h th e provision s o f th e Copyrigh t Licensing Agency , 9 0 Tottenha m Cour t Road , Londo n W1P 9HE . User s registere d wit h th e Copyright Clearanc e Center. 2 7 Congres s Street , Salem , M A 01970 , USA : th e item-fee cod e fo r thi s publicatio n i s 0305-8719/00/$15.00 . British Librar y Cataloguin g i n Publication Dat a A catalogu e recor d fo r thi s boo k i s availabl e fro m th e British Library . ISBN 1-86239-056- 8 ISSN 0305-871 9 Typeset b y Aarontyp e Ltd , Bristol , U K Printed b y Th e Alde n Press , Oxfor d

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Contents

N0TTVEDT, A . Integrate d Basi n Studie s - Dynamic s o f th e Norwegia n Margin : a n 1 introduction Intra-plate riftin g an d basin formatio n CHRISTIANSSON, P., FALEIDE , J. I . & BERGE, A. M . Crusta l structur e i n the norther n Nort h 1 Sea: a n integrate d geophysica l stud y ODINSEN, T. , CHRISTIANSSON , P., GABRIELSEN , R. H. , FALEIDE , J. I . & BERGE , A . M . Th e 4 geometries an d dee p structur e o f th e norther n Nort h Se a rif t syste m TER VOORDE, M. , FARSETH , R . B. , GABRIELSEN, R. H . & CLOETHING, S. A. P. L . Repeated 5 lithosphere extensio n i n th e norther n Vikin g Graben: a couple d o r decouple d rheology ? ODINSEN, T. , REEMST , P. , VA N DER BEEK, P., FALEIDE , J. I . & GABRIELSEN , R. H . Permo - 8 Triassic an d Jurassi c extensio n i n the norther n Nort h Sea : result s from tectonostratigraphi c forward modellin g FOSSEN, H. , ODINSEN , T . FARSETH , R . B . & GABRIELSEN , R. H . Detachment s an d low - 10 angle fault s i n th e norther n Nort h Se a rif t syste m Basin fillin g RAVNAS, R. , N0TTVEDT , A. , STEEL , R . J . & WINDELSTAD , J . Syn-rif t sedimentar y 13 architectures i n th e norther n Nort h Sea NOTTVEDT, A. , BERGE , A . M. , DAWERS , N . H. , FaRSETH , R . B. , HAGER , K.-O. , 17 MANGERUD, G . & PUIGDEFABREGAS , C . Syn-rif t evolutio n an d pla y potentia l i n th e Snorre-H area , norther n Nort h Se a JORDT, H. , THYBERG , B. I. & N0TTVEDT, A. Cenozoic evolutio n of the central and norther n 21 North Se a with focu s o n differentia l vertica l movements o f the basi n floo r an d surroundin g clastic sourc e area s THYBERG, B. , JORDT, H. , BJORLYKKE , K. & FALEIDE , J. I . Relationship s betwee n sequenc e 24 stratigraphy, mineralog y and geochemistr y in Cenozoic sediment s of the northern Nort h Se a KYRKJEBO, R. , HAMBORG , M. , FALEIDE , J. I. , JORDT , H . & CHRISTIANSSON , P. Cenozoi c 27 tectonic subsidenc e fro m 2 D depositiona l simulation s of a regiona l transec t i n th e norther n North Se a basin Conjugate volcani c margin s SKOGSEID, J. , PLANKE , S. , FALEIDE , J . L , PEDERSEN , T. , ELDHOLM , O . & NEVERDAL , F . 29 NE Atlanti c continenta l riftin g an d volcani c margi n formatio n BREKKE, H. The tectoni c evolutio n of the Norwegian Se a continental margin, with emphasis 32 on th e V0rin g an d M0r e basin s MOGENSEN, T . E. , NYBY , R. , KARPUZ , R . & HAREMO , P . Lat e Cretaceou s an d Tertiar y 37 structural evolutio n o f the northeaster n par t o f th e V0rin g Basin , Norwegian Se a SUNDVOR, E., ELDHOLM , O. , GLADCZENKO , T. P . & PLANKE, S. Norwegian-Greenland Se a 39 thermal field ELDHOLM, O. , GLADCZENKO , J , SKOGSEID , J . & PLANKE , S . Atlanti c volcani c Margins : 41 a comparativ e stud y

5 1 9 3 5

3 9 9 5 3

5 7 9 7 1

VI

CONTENTS

Present stres s LINDHOLM, C . D. , BUNGUM , H. , HICKS , E . & VILLAGRAN , M . Crusta l stres s an d tectonic s 42 in Norwegia n region s determine d fro m earthquak e foca l mechanism s FEJERSKOV, M. , LINDHOLM , C . D. , MYRVANG , A . & BUNGUM , H . Crusta l stres s i n an d 44 around Norway : a compilatio n o f i n situ stres s observations FEJERSKOV, M . & LINDHOLM , C . D . Crusta l stres s in and aroun d Norway ; an evaluatio n of 45 stress generatin g mechanism s Index

9 1 1

469

Integrated Basi n Studie s - Dynamic s of the Norwegian Margin : an introduction ARVID N0TTVED T Norsk Hydro Research Centre, N-5020 Bergen, Norway Present address: Norsk Hydro Canada, lll-5th Avenue SW, Calgary AB, T2P 3Y6, Canada Several lithospheri c an d uppe r crusta l processe s interact i n th e formatio n o f rif t basin s an d evolution o f suc h basin s int o passiv e margins . Similarly, th e fillings of rif t basin s depen d o n a variety o f tectonic, morphologica l an d sedimen tary processes . Area s tha t hav e passe d throug h a complet e evolutio n fro m intra-cratoni c rift ing through breaku p an d passiv e margin forma tion offe r particula r opportunitie s t o stud y an d link th e differen t processe s o f basi n formatio n and filling . On e suc h area , tha t ha s a uniqu e combination o f a well-preserved rock recor d an d abundant data , i s th e Norwegia n Nort h Sea North Atlanti c margin , herei n referre d t o a s the Norwegia n margin . During post-Caledonia n time s th e Nort h Se a and mid-Norwegia n margi n underwen t severa l episodes o f lithospheri c extensio n (multi-phas e rifting), o f which th e lates t le d to crusta l break up an d accretio n o f oceani c crus t betwee n Norway an d Greenlan d nea r th e Paleocene Eocene transition . Prominent pre-breakup extensional episodes , i n lat e Permian-Triassic , lat e Jurassic, earl y Cretaceou s an d mid-Cretaceou s time, le d t o th e developmen t o f th e Nort h Se a rift system , th e larg e Cretaceou s V0rin g an d More sedimentar y basin s of f Norway, an d con jugate equivalent s off eastern Greenland . Lates t Cretaceous-Paleocene riftin g an d breaku p wer e accompanied b y large-scal e igneou s activity , developing th e presen t conjugat e volcani c mar gins of the North Atlantic . The Nort h Sea , i n particular , i s covered wit h an exceptionall y goo d geologica l an d geophysi cal industry an d academi c database , comprisin g both high-qualit y geophysica l profile s an d a large number o f industrial wells. About 3 0 years of active exploration i n the North Sea has le d to an advance d leve l o f understandin g o f th e geo logical evolutio n an d complexit y o f th e basin . Numerous paper s hav e bee n publishe d o n th e formation an d fillin g o f th e Nort h Se a intra cratonic rif t structure . The mid-Norwegia n margi n ha s a les s extensive, bu t increasing , industr y database . Never -

theless, th e V0rin g margi n i s amon g th e bes t studied volcani c margin s globall y du e t o a regional coverag e o f multichanne l seismi c lines, expanded sprea d profile s an d othe r geophysica l and geologica l data . I n addition , th e successfu l scientific drillin g throug h a sequenc e o f sea ward dippin g reflector s o n th e V0rin g Platea u has greatl y contribute d t o th e understandin g of th e margin . This settin g make s th e stud y regio n a n excellent laborator y i n whic h t o stud y progres sive rif t evolutio n an d it s inter-relationship with basin formatio n an d filling, including th e inter plays betwee n structura l evolution , erosion , sedimentation and magmatism, o n both regional and loca l scales. Consequently, the region allows research into fundamental earth processes, which also hav e direc t implication s fo r hydrocarbo n exploration an d assessment .

The Integrate d Basi n Studie s (IBS ) projec t The paper s an d researc h result s presente d her e have bee n prepare d a s par t o f th e Integrate d Basin Studie s projec t (Fig . l).Thi s projec t wa s funded unde r th e D G IIX , JOUL E I I pro gramme, wit h th e objectiv e t o stud y th e litho spheric an d uppe r crusta l processe s governin g the formatio n an d evolutio n o f extensiona l an d foreland basin s an d t o deciphe r th e rol e o f tectonics, sea-leve l and sedimentar y processes i n the filling of such basins. As part of this task, the project als o aime d a t studyin g the physica l laws of compactio n i n fine-graine d sediments . Base d on thes e results , a ne w generation o f descriptive as wel l a s numerica l basi n formatio n an d basi n fill models ha s bee n derived . Results fro m th e Integrate d Basi n Studie s project hav e bee n reporte d i n an extensiv e Final Report t o th e D G IIX , an d selecte d paper s prepared withi n Module s 1 an d 2 hav e bee n published i n tw o previou s Geologica l Societ y Special Publications , no s 13 4 and 156 .

From: NOTTVEDT , A . e t al. (eds ) Dynamics o f th e Norwegian Margin. Geologica l Society , London , Specia l Publications, 167 , 1-14 . 1-86239-056-8/OO/ S 15.00 © Th e Geologica l Societ y o f Londo n 2000 .

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A. N0TTVED T

This boo k ha s bee n prepare d wit h th e inten tion t o giv e a representativ e pictur e o f th e sci entific spa n an d result s o f th e projec t Modul e 3 : Dynamics o f the Norwegia n Margin . I t includes papers prepare d b y Ph D student s funde d directly throug h th e project , a s wel l a s paper s written b y researcher s i n academi a an d industry that hav e bee n workin g in , o r closel y associate d with, th e project .

IBS Modul e 3: Dynamics of th e Norwegian Margin (IBS-DNM) The Integrate d Basi n Studie s projec t Modul e 3 , Dynamics o f the Norwegian Margi n (Fig . 2), was technically par t o f th e IB S project, bu t receive d funding directl y fro m th e Researc h Counci l o f Norway (RCN). Th e module was coordinated b y Norsk Hydro , an d contractua l arrangement s were mad e betwee n Nors k Hydr o an d EU , Norsk Hydr o an d RCN , an d Nors k Hydr o an d partners. Th e modul e wa s organize d int o thre e themes an d subordinat e topic s (Fig . 3) .

Research objectives The scop e o f the IBS-DN M projec t modul e wa s to analys e and mode l th e dynamics o f the basin s off mid-Norwa y an d i n th e norther n Nort h Sea , in order to establish a better understanding of the processes controllin g basi n formatio n an d fillin g and t o develo p ne w model s fo r multiphase , intraplate riftin g an d volcani c margin formation .

Fig. 1 . D G IIX : Joule I I Geoscienc e I I Programm e

Thematically, th e projec t modul e focuse d o n three relate d themes : (i ) Intra-plat e riftin g an d basin formation , (ii ) Basi n infil l an d (iii ) Conjugate volcani c margins. The dat a coverag e an d geological settin g resulte d i n a topica l focu s o n two geological provinces, the northern North Sea and th e Norwegian-Greenlan d Se a rifte d volca nic margins . Th e M0r e Basin , o n th e volcani c More margin , link s thes e tw o provinces . Geographically, th e tw o forme r researc h theme s (Intra-plate riftin g an d basi n formatio n an d Basin infill ) focuse d o n th e norther n Nort h Se a and th e Mor e Basin , whil e the latte r (Conjugate volcanic margins) primarily dealt with the More Voring margin s and thei r conjugates. Of particula r interes t in th e intra-plat e setting is th e kinematic s an d relativ e importanc e o f deeper versu s shallow crustal processes an d how 7 these contro l overal l subsidenc e patterns : th e erosional an d provenanc e histor y o f sedimen t supply areas ; an d th e architecture , composi tion an d histor y o f basi n fill . Th e wor k con centrated o n syn-rif t basi n fill, but include d als o post-rift strata . I t ha s bee n attempte d t o estab lish genera l an d couple d model s fo r rif t basi n formation an d sedimen t fillin g i n th e norther n North Sea . The Norwegian-Greenlan d Se a represent s complete extensio n an d thinnin g o f th e litho sphere. Therefore , th e margi n basin s bea r a n imprint no t onl y o f th e sam e event s a s th e intra-plate basins , bu t hav e als o undergon e a structural , magmati c an d depositiona l his tory reflectin g th e formatio n an d subsequen t

INTEGRATED BASI N STUDIE S development o f th e presen t volcani c passiv e margin. T o understan d th e basi n evolutio n an d tectono-magmatic histor y i t i s necessar y t o assess th e timin g an d magnitud e o f the differen t tectonic, magmati c an d subsidenc e episode s fo r the variou s basins . I n orde r t o determin e th e style o f deformatio n fo r eac h tectoni c episod e and t o quantif y th e magnitud e o f crusta l an d lithospheric extension , restoratio n o f th e struc turally define d extensional deformatio n throug h time wa s particularl y emphasized .

Database and methods The projec t databas e (Fig . 4 ) consiste d o f scientific dee p reflectio n an d refractio n seismi c data, larg e volume s o f industr y regiona l 2 D reflection seismi c data , 3 D seismi c data , wel l data, potentia l fiel d data , industr y specia l studies an d reports , i n additio n t o publishe d literature.

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It is important, therefore , tha t th e project ha d access t o much o f the existing seismic, structura l and stratigraphi c framewor k i n the stud y regio n to defin e an d selec t th e optima l targe t area s fo r the detaile d studie s describe d i n th e variou s project topics . Th e industria l partner s wer e instrumental i n achievin g thi s objective. The wor k o n crusta l structur e an d basi n formation concentrate d aroun d seve n transect s from th e norther n Nort h Se a t o th e V0rin g margin (Fig . 4) , whereas th e topi c o n stres s field covers th e entir e Norwegia n continenta l shelf . The syn-rif t infil l studie s focuse d o n selecte d fault-blocks i n the northern Nort h Sea , wherea s the post-rif t (Cenozoic ) infil l studie s integrate d data fro m th e Danis h an d Norwegia n sector s of th e Nort h Sea , u p t o 62°N . Th e wor k o n erosion an d provenanc e include d th e norther n North Sea . Th e studie s o n volcani c margin s concentrated o n the mid-Norwegian margin , bu t comparative studie s wer e mad e wit h othe r vol canic margin s aroun d th e world .

Fig. 2 . IBS-DN M contractua l framewor k and involve d partners. RCN , Researc h Counci l o f Norway.

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A. N0TTVED T subsidence an d ma y therefor e contribut e significantly t o th e stretchin g facto r a s calculate d from subsidenc e analysi s o f late r rif t episodes . A majo r questio n i s the relativ e contribution t o final extensio n contribute d b y eac h o f th e mai n episodes o f extension.

Key results. Th e dee p seismi c reflectio n surve y NSDP84 ha s forme d th e basi s fo r numerou s papers o n th e crusta l structur e an d basi n evolu tion o f th e norther n Nort h Se a rif t system . However integratio n with all available geophysical an d geologica l dat a fro m bot h offshor e an d onshore area s provide d som e interestin g ne w results. Reprocessing o f th e dee p reflectio n seismi c data resulte d i n enhance d dat a quality . B y integrating seismi c refractio n data , ES P dat a and gravity/magneti c data, an d combinin g thes e transects with high-quality commercial reflectio n seismics, i t ha s bee n possibl e t o ma p ou t th e structural outlin e o f th e grabe n geometr y o f the norther n Nort h Se a (Christiansso n e t al.. Odinsen e t al. (a)) , includin g linkag e betwee n upper an d mi d crusta l faults , faul t geometrie s of th e dee p crusta l level s an d structur e o f Fig. 3. IBS-DN M projec t structur e an d organization . the lowermos t crust . O f particula r interes t i s Past member s o f th e Steerin g Committe e includ e the improve d identificatio n of top basement , the Alv Orhei m (Statoil) . Jan Volse t (Statoil ) an d Ro n confirmation o f intr a uppe r mantl e dippin g J. Stee l (UB) . reflectors an d identificatio n o f a high-velocit y lower crusta l roc k body . New informatio n on th e level s of detachmen t and genera l geometrie s o f th e differen t faul t Theme 1 : Intra-plate Riftin g an d systems ha s als o bee n gaine d (Fosse n e t al.}. Basin Formation These investigation s revea l a mor e comple x interaction betwee n faul t system s tha n antici Topics 1.1: Crustal structure, pated earlier . The kinematic s of the faultin g was 1.2: Sedimentary basin formation and studied b y the use of analogue model s (Fosse n & Gabrielsen 1995) . In addition , th e confirmatio n 1.4: Tectonic modelling of th e existenc e o f 'grabe n units / whic h ar e The dynami c framewor k fo r Mesozoi c an d characterized b y shiftin g polaritie s alon g th e Cenozoic basi n developmen t i n th e norther n basin axi s o f th e Permo-Triassi c basi n axis , i s North Se a an d o n th e mid-Norwegia n margi n important (Faerset h e t al . 19950) . has bee n wel l documented ove r th e las t decade . Forward/backward numerica l modellin g o f However, ther e i s majo r uncertaint y a s t o th e crustal geometry , palinspasti c grabe n topogra nature o f th e dee p crusta l structur e an d it s phy an d therma l developmen t hav e resulte d i n relation t o th e sedimentar y cover . I n addition , better insigh t i n th e formatio n o f th e Nort h Se a the detaile d nature , significanc e and area l exten t rift (Odinse n e t al . (b) , te r Voorde et al.). Thes e of th e faultin g an d differentia l subsidenc e i n th e investigations conclud e tha t th e Permo-Triassi c post-rift interval s ar e poorl y known . stretching even t ha s bee n underestimate d i n For a complete understanding o f evolution o f many previou s work s i n th e area . Th e Permo the norther n Nort h Se a an d Mor e Basi n i t i s Triassic even t i s responsible for a more extensive essential t o conside r th e earlie r Permo-Triassi c extension tha n th e Lat e Jurassi c even t an d th e rift, whic h generate d th e structura l framewor k former even t als o resulte d i n extensio n o f a on whic h th e late r Jurassi c an d Cretaceou s rift s wider are a tha n di d th e latter . acted (Fig . 1.3) . Compactio n o f Permo-Triassi c The results fro m th e structural mappin g o f the basin fil l an d residua l Permo-Triassi c therma l M0re Basi n particularl y emphasize th e complexanomalies ma y enhanc e Cretaceous-Cenozoi c ity o f th e basin , wher e larg e basement-involved .

INTEGRATED BASI N STUDIE S rotated fault-block s have contribute d t o th e internal compartmentalization o f the basin. As for the Nort h Sea , th e M0r e Basi n wa s subjec t to multi-phas e extensio n (Grunnaleit e & Gab rielsen 1995) . Lithospheri c stretchin g probabl y commenced i n th e lat e Permian-earl y Triassic , followed b y a second episod e in the late Jurassic. In contrast t o the North Sea, however, early-midCretaceous riftin g an d successiv e Palaeogen e breakup strongly influenced furthe r development of the basin, resulting in a cumulative beta-facto r in excess of 3. Regional-scale antiforms along the northeastern basi n margin , basi n inversio n i n the S10rebot n Sub-basi n an d th e northeaster n More Basin, compressional reactivatio n of faults as see n o n th e nea r bas e Cretaceou s level , an d reverse drag and foldin g associated wit h faults of primarily extensiona l origin , sugges t tha t pro nounced multi-phas e lat e Mesozoic-Cenozoi c inversion ha s take n plac e (Mogense n e t al., Gabrielsen e t al 1999) . Relevance t o th e petroleum industry. Th e im proved model s fo r basi n developmen t an d faul t geometries obtaine d withi n thi s topi c provid e

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new concept s t o th e industr y tha t ar e relevan t for th e understandin g o f san d distributio n i n general, an d i n searc h fo r th e subtl e tra p i n particular. I t is also expected tha t th e results will be o f importanc e i n basi n modelling , bot h o n the regiona l (crustal ) scale , an d fo r maturatio n and migratio n studie s on th e sub-regiona l scale. In addition , th e focu s o n th e pre-Jurassi c basi n development ha s opene d ne w perspective s tha t may b e o f importanc e fo r futur e exploration , particularly i n th e deepe r part s o f th e Nort h Sea Basin.

Topic 13: Present regional stress field Some rif t episode s i n th e norther n Nort h Sea More Basin-V0rin g Basi n correlate wit h break up an d initia l sea floor spreading in other basin s within th e Nort h Atlanti c rif t syste m an d the inter-relationship s betwee n pre-rif t struc tures (structura l fabric ) an d th e regiona l stres s field probabl y playe d a n importan t rol e i n structural evolution , b y rejuvenatio n o f inher ited structures .

Fig. 4. Schemati c outlin e o f the IBS-DN M stud y are a an d databas e (afte r Skogsei d e t al., this volume).

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The orientatio n an d magnitud e o f the presen t stress fiel d ar e reflecte d i n bot h wel l dat a an d earthquake foca l mechanis m solutions . Th e us e of bot h i n situ measurement s an d earthquak e based data enables mappin g o f crustal stres s as a function o f depth . Mos t importan t ar e strai n markers fo r individua l tim e intervals . Regional an d loca l variation s i n stres s direc tions an d magnitud e yiel d informatio n o n th e intra-plate forces , plat e tectonic s an d recen t geological history . Thus , i t provide s impor tant inpu t t o th e tectoni c modelling . Couplin g of non-linea r lithospher e rheolog y t o a mode l involving stres s change s actin g on a basi n offer s the prospec t o f understandin g non-therma l sub sidence an d o f solvin g discrepancie s i n presen t crustal extensio n estimates . Key results. A databas e o f unprecedente d quality an d quantit y fo r th e regio n ha s bee n established, comprisin g stres s informatio n fro m a variet y of sources. Th e projec t ha s contribute d to th e compilatio n o f a n earthquak e foca l mechanism databas e tha t i s constantl y bein g expanded, an d whic h currentl y comprise s 10 9 solutions fo r northwester n Fennoscandi a an d Svalbard (Lindhol m et al., Lindhol m et al. 1995) . In addition , a tota l o f 34 5 borehol e breakou t observations an d 10 4 overcoring measurement s have bee n collecte d an d quality-assesse d int o a high-quality database (Fejerskov & Lindholm (a), Borgerud & Suav e 1995 , Fejersko v e t al . 1995 , G01ke e t al . 1995) . All thes e dat a wer e use d i n a join t inter pretation o f th e crusta l stres s fiel d i n Norwa y and adjacen t offshor e region s (Fejersko v & Lindholm (b)) . Regionally , th e analysi s ha s revealed a crusta l horizonta l compressiv e stres s with a dominatin g WNW-ES E direction(s Hm ax) in souther n Norway , whic h graduall y rotate s into N- S compressio n i n th e Barent s Se a region. Th e observe d stres s i s largely consisten t throughout th e uppe r crust , indicatin g tha t large-scale tectoni c mechanism s ar e th e mai n source. Thi s can , wit h a hig h degre e o f con fidence, be attribute d t o crusta l spreadin g alon g the mid-Atlanti c Ridge . Th e consisten t stres s pattern show s regiona l variation s probabl y caused b y regional stress generating mechanism s like th e continenta l margin-ridg e pus h an d sediment loading . Fault s tha t pertur b an d deviate th e regiona l stres s locall y als o hav e been observed . With importan t exception s i n loca l areas , the dominan t typ e o f faultin g i s foun d t o b e contractional and , t o som e extent , strike-slip , indicating tha t th e continenta l margi n i s subjec t to compression , a s als o indicate d b y i n situ

measurements onshore . Compressiv e differen tial horizonta l stresse s ar e foun d i n al l part s of th e margi n analysed , an d th e dat a indicat e that th e stres s regim e i n th e uppe r crus t ofte n alters with depth. A stress homogeneity through out th e brittl e crust is , however, wel l documen ted b y th e homogeneou s stres s direction s obtained fro m complementar y dat a (dee p earthquakes an d shallo w boreholes) . Relevance t o the petroleum industry. O n a shortterm basis , thi s topi c ha s provide d importan t information t o th e genera l understandin g o f the stres s fiel d alon g th e Norwegia n Margin , which i s o f considerabl e importanc e whe n entering int o ne w area s lik e th e V0rin g Basin . In addition , detaile d informatio n o n i n situ stresses i n severa l petroleu m field s i s expecte d to b e directly relevant to activities in those fields. In situ roc k stres s data ar e particula r importan t for th e plannin g and drillin g o f stabl e an d saf e wells. Th e stat e o f stres s i n an d aroun d th e hydrocarbon reservoi r i s also importan t fo r field development an d reservoi r management , an d the succes s o f hydrauli c fracturin g an d injec tion depend s o n knowin g ho w fracture s an d fluid front s propagat e unde r differen t stres s systems. Th e us e o f suc h dat a ma y o n a long term basi s reduc e cost s an d hel p optimiz e dril ling an d production . Theme 2 : Basi n Infil l

Topic 2.1: Syn-rift sediment architecture The stratigraph y and architectur e o f th e Trias sic-lower Jurassi c an d Cretaceous-Cenozoi c post-rift basina l infil l o f th e Nort h Se a ar e relatively wel l documented , wherea s th e Permo Triassic an d uppe r Jurassi c syn-rif t infil l ar e known onl y i n broades t outline . Detail s o f timing, styl e an d rate s o f th e event s tha t mak e up th e syn-rif t episodes , an d o f th e sedimen tary respons e t o thes e events, hav e onl y recentl y started t o emerge. Extensio n across th e norther n Viking Grabe n i s know n t o var y betwee n th e Tampen Spur , Vikin g Grabe n proper , Hord a Platform an d Sog n Grabe n segments . Thi s variability i n space , togethe r wit h tempora l variations (Permo-Triassic , Lat e Jurassic ) ha s considerable consequence s fo r th e composi tion o f syn-rif t infill . Thes e latera l change s an d related provenanc e an d sedimen t transpor t issues hav e bee n addresse d b y analysi s o f dif ferent pattern s o f syn-rif t stackin g an d b y studies o f th e relatio n betwee n rate s an d styl e of structura l events.

INTEGRATED BASI N STUDIE S Key results. A database of seismi c examples a s well as field analogues ha s been buil t to illustrate the variabilit y of syn-rif t architecture , involvin g integration o f data on rif t topography , erosiona l and drainag e patterns , relativ e se a leve l an d it s changing position , sedimen t transpor t processe s and facies and resultan t sand body geometry an d distribution (Marjana c 1994) . I n severa l suc h examples, th e spatia l an d tempora l evolutio n of reservoir san d facie s ha s bee n semi-quantita tively linke d t o th e structura l evolutio n o f th e parent half-grabe n an d neighbourin g footwal l sediment sourc e area s (Farset h e t al . \995b; Ravnas & Bondevik 1997 ; Nettvedt et al.). A synthesi s ha s bee n compile d o f variou s aspects o f th e three-dimensiona l geometr y o f syn-rift stratigraphi c architectura l element s an d their stackin g pattern, an d th e spatia l organiza tion o f potentia l facie s tract s an d stratigraphi c surfaces presen t i n marin e syn-rif t basi n fill s (Gabrielsen e t al . 1995 ; N0ttved t e t al . 1995 ; Ravnas e t al. \991a; Ravna s & Steel 1998) . This includes a n analysi s an d compariso n o f a number o f half-grabe n sub-basin s i n th e north ern Nort h Se a (Faerseth & Ravnas 1998 ; Ravnas et al.). The studie s hav e le d t o a classificatio n o f marine rift-basin s an d syn-rif t succession s i n terms o f sedimen t suppl y int o overfilled , sediment-balanced, sediment-underfille d an d sediment-starved basin s (Ravna s & Stee l 1997) . Dependent o n whethe r th e rift-basi n wa s over filled/sediment-balanced, sediment-underfille d or sediment-starved , a three-fold , sand-clay sand package , two-fol d sand-cla y packag e o r one-fold mud-pron e packag e constitute s th e syn-rift succession , respectively . I n case s wher e the rif t episod e wa s characterize d b y repetitiv e rift phases , result s sho w tha t th e successiv e rotational til t event s ar e ofte n separate d b y periods characterize d b y les s intens e faulting , the so-calle d intra-rif t quiescenc e o r relaxa tion stages . The projec t als o include d a stud y o f th e syn rift infil l o f th e Lusitania n Basi n i n Portuga l (Ravnas e t al . \991b). Relevance t o the petroleum industry. Th e results and interpretation s obtaine d withi n thi s topi c provide new insight and idea s for analysis of rift basin developmen t an d syn-rif t infill . Th e recog nition o f th e differen t type s o f syn-rif t sedimen tary architecture s provide s a ste p forwar d i n erecting predictiv e model s fo r th e analysi s o f half-graben syn-rif t sedimen t fill . I n addition , integrated detaile d analysi s o f rif t basi n struc tural evolution , high-resolutio n biostratigraphi c zonation, reworke d biozonation , basin-fil l facie s

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and architectur e an d seismi c facies offe r a sensi tive too l fo r analysin g th e detaile d sequentia l basin development an d predictin g reservoi r san d distribution. Th e model s erecte d ma y serv e a s important reference s t o th e conceptua l trainin g of explorationists.

Topics 2.2: Post-rift sediment architecture and 2.3: Erosional episodes and provenance area Post-rift interval s commonly hav e a duratio n o f 50-100 Ma. Therma l relaxatio n followin g th e late Jurassi c rif t even t ha d almos t cease d an d the Nort h Se a basi n wa s i n a stat e clos e t o thermal equilibriu m during th e transitio n t o th e Cenozoic. Subsidenc e an d depositio n durin g the Cenozoic, therefore , wer e controlle d b y othe r factors. Th e Cenozoi c successio n i n th e Norwe gian Nort h Se a ca n b e divide d int o geneti c sequences o f 8-1 6 Ma duration , representin g major period s o f clastic wedge progradation an d retreat wit h respec t t o th e Norwegia n hin terlands. Th e clasti c wedge s ar e bounde d b y major floodin g surface s and reflec t variabl e subsidence rates . Th e characte r o f thes e variation s and o f the resultin g genetic sequences hav e bee n investigated and relate d to the uplift an d erosio n of th e surroundin g lan d areas . The sand/cla y rati o i n sedimentar y basin s i s commonly viewe d a s a functio n o f depositiona l energy. A facto r tha t tend s t o b e overlooke d is th e primar y provenanc e are a composition . A bette r understandin g o f th e processe s tha t cause change s i n th e compositio n o f th e sedi mentary sequenc e i n the Nort h Se a Basin therefore i s needed. Th e relativ e contributions o f th e Norwegian mainlan d an d Eas t Shetlan d Plat form a s sourc e area s ha s bee n evaluated , base d on analysis and descriptio n o f basin fill composition an d characteristics , mapping o f palaeodrai nage directions and transport routes , estimations of palaeo-provenanc e area s an d sourc e roc k lithologies, an d correlatio n o f basi n fill to prov enance area . Key results. Th e projec t ha s integrate d seismi c data fro m th e Danish secto r wit h the Norwegia n North Se a (Jord t e t al . 1995) . I t als o include s mineralogical an d geochemica l analyse s o f 160 0 samples o f cutting s an d core s fro m abou t 4 0 wells in th e Norwegia n secto r o f the Nort h Sea , and thu s represent s th e mos t comprehensiv e database currentl y available . Thes e dat a ar e important i n term s o f understandin g th e prove nance an d provid e a ne w basi s o n whic h t o

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interpret direction s o f sedimen t transpor t an d areas o f tectoni c uplift . The seismostratigraphi c stud y show s tha t changes i n seismi c sequenc e geometr y occurre d in-phase wit h intra-continenta l stres s variation s on th e Europea n Platfor m an d tha t sedimen t supply an d differentia l tectoni c movement s i n the basin and i n the provenance area s controlle d Cenozoic depositio n i n the centra l an d norther n North Se a (Jordt e t al.). Seismi c sequences were generated largel y independent o f marke d globa l glacio-eustatic se a leve l falls . Generatio n o f seismic onla p appear s t o hav e bee n controlle d by sedimen t suppl y and basi n floo r topography . Thinning o f sedimentary strat a towards inclined surfaces result s i n marke d seismi c onla p an d apparent unconformabl e relationships . It is further show n that seismi c velocities in the North Se a Cenozoi c mudstone s var y system atically a s a functio n o f mineralogica l composi tion. A commonly observe d velocit y inversion at the base of the Pliocene and Pleistocen e sequenc e (Reemst e t al. 1996) has been foun d t o relate to a low conten t o f smectit e an d therefor e muc h faster compactio n o f thes e sediment s compare d to th e underlyin g smectite-rich mudstones (Thy berg e t al.}. Silic a cementatio n fro m biogeni c silica, as well as alteration o f opal-A to opal-CT, have bee n show n t o caus e abrup t increase s i n seismic velocities. It has been demonstrate d that pore pressure in the Tertiar y successio n i s governe d b y miner alogical composition . Thic k sequence s o f mud stones wit h hig h smectit e content , whic h typically hav e hig h specifi c surfac e an d lo w permeability, typicall y giv e rise t o overpressure , providing there are no sandy beds causing lateral drainage. Moreover , becaus e o f their high water content smectit e ric h mudstone s usuall y hav e lower densit y than th e overlyin g mudstones an d sandstones. Thi s densit y inversio n frequentl y caused diapirism . Finally, th e projec t ha s provide d result s contributing t o th e understandin g o f th e Cen ozoic uplif t o f Norwa y (Fjeldskaa r 1994 ; Stue vold & Eldhol m 1996) . Relevance t o th e petroleum industry. Thi s topi c has contribute d t o th e understandin g o f th e Cenozoic basi n formatio n an d filling in relation to th e uplif t an d erosio n histor y o f Souther n Norway. Thi s uplif t ma y strongl y influenc e secondary an d tertiar y migratio n an d trappin g of petroleum , particularl y i n area s clos e t o th e coast o f Norway . The presen t projec t i s on e o f th e firs t i n th e North Se a wher e seismi c stratigraph y ha s bee n related t o mineralogica l compositio n i n a sys -

tematic way . Understandin g th e regiona l variations i n th e velocity/dept h functio n an d th e processes responsibl e fo r thes e trend s i s impor tant whe n dept h convertin g seismi c profiles . The dat a o n mineralog y an d diagenesi s i s als o of valu e t o calibratio n o f seismi c respons e t o lithology, i n orde r t o understan d bette r th e information tha t ca n b e extracte d fro m seismi c attributes. The demonstrated relationshi p between mineralogical composition , diagenesi s an d overpres sure is expected t o b e of great interes t to drilling , particularly horizonta l drilling, and fo r handling of rock mechanica l problems durin g production . Moreover, th e present stud y ha s show n tha t th e North Se a Cenozoi c roc k propertie s canno t b e realistically represente d a s a simpl e functio n (i.e. linea r o r exponential ) of buria l depth . Thi s has importan t consequence s t o basi n modelling, as porosit y reductio n (compaction ) i s usuall y assumed t o b e a functio n o f overburde n o r effective stress .

Topic 2.4: Stratigraphic modelling Numerical modellin g i s becomin g increasingl y important i n th e understandin g o f sedimentar y basin deposition and filling. In particular, it helps evaluate the complex interplay between tectonics, eustacy, climate , erosion and sediment transport , and bette r constrai n th e boundar y condition s of geologica l interpretation s an d models . The overal l sedimentar y architectur e o f th e upper Jurassi c syn-rif t infil l i n th e Oseber g are a as wel l a s th e Cenozoi c infil l alon g on e o f the Nort h Se a regiona l profile s hav e bee n modelled, a s a functio n o f sedimen t input , sub sidence an d se a leve l fluctuation s t o obtai n additional informatio n abou t thei r interpla y and relativ e importance . Th e modellin g ha s used forwar d process-base d simulatio n pro grams of dynamic-slope type. Comparisons hav e been made betwee n observed Stratigraphi c architecture an d syntheti c Stratigraphic models where the mos t importan t controllin g factor s hav e been considered . Key results. Th e numerica l modellin g o f th e syn-rift infil l o f th e Oseber g are a demonstrat e the interactio n betwee n structura l evolution , sedimentation an d resultin g Stratigraphi c pat tern acros s th e rotatin g fault-block s (terVoorde et al . 1997) . Th e modellin g furthe r confirm s that th e sedimentar y architectur e o f th e Nort h Sea Cenozoi c infil l canno t b e explaine d b y sedi mentary processe s and/o r eustac y alon e (Kyrk jebe e t al.). Period s o f anomalou s subsidence .

INTEGRATED BASI N STUDIE S deviating fro m th e post-rif t therma l subsidence , are require d i n th e lat e Paleocen e an d lat e Miocene. However , som e o f th e Paleocen e subsidence ma y b e relate d t o th e initia l basi n form e. g exces s wate r dept h afte r lat e Cretac eous. Th e Miocen e event , o n th e othe r hand , includes basina l subsidenc e i n th e norther n North Sea , a s wel l a s sourc e are a uplif t i n th e Norwegian mainland , an d i s believe d t o repre sent a n intraplat e effec t couple d t o th e openin g of th e Nort h Atlantic . Relevance t o th e petroleum industry. Th e itera tive proces s betwee n interpretatio n an d model ling allow s th e geologis t t o bette r constrai n th e possible geologica l model s an d t o narro w i n on a les s numbe r o f likel y interpretations . Thi s type o f modellin g i s als o a n excellen t too l t o visualize th e comple x interactio n betwee n sedi ment supply , sea-leve l and tectonic s i n fillin g o f sedimentary basins .

Theme 3 : Conjugate Volcanic Margins Topics 3.1: Rift dimensions and duration of rifting, 3.2: Geodynamic modelling and 3.3: Comparative studies It ha s lon g bee n recognize d tha t th e forma tion o f th e norther n Nort h Atlanti c conti nental margin s wa s accompanie d b y exces s magmatic activity . Thi s cause d a sub-divisio n of rifted margin s int o volcani c and non-volcani c types. Unti l recently , th e volcani c margin s wer e thought t o b e an exceptiona l case . Recen t com parative studie s involvin g othe r rifte d margin s now sugges t tha t exces s magmatis m durin g breakup i s fa r mor e commo n tha n previousl y thought. I n fact , th e volcani c margi n migh t represent th e norma l evolutionar y cas e rathe r than bein g anomalous . The volcani c signatur e o f th e Nort h Atlanti c margin ha s bee n explore d b y seismi c reflectio n data an d scientifi c drilling , whereas th e tectoni c signature, an d thu s development, hav e to a large extent bee n hidde n belo w th e volcani c rocks . The project ha s involved geophysical-geologi cal mapping o f the V0ring, Mor e and conjugat e margins, i n orde r t o obtai n ke y dimension s and timin g o f th e lat e Cretaceous-Paleocen e rift episode , b y focusin g on : (1 ) width , styl e and timin g o f syn-rif t lithospheri c extension ; (2) exten t an d timin g o f syn-rif t regiona l uplift ; (3) extent , character , tim e o f emplacemen t and dimension s o f igneou s unit s (extrusives , intrusives, lowe r crusta l high-velocit y bodies) ;

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(4) location o f continent-ocean transition ; (5) extent o f post-rift therma l margi n subsidence . Th e availability o f dat a als o of f Greenlan d offer s a possibility t o ma p an d mode l th e entir e rif t an d to compar e th e lat e Cretaceous-Paleocen e rif t dimensions wit h th e previou s rif t episodes . In addition , comparativ e studie s o f th e volcanic margi n formatio n i n th e Nort h Atlan tic, a s wel l a s globally , hav e bee n undertaken , with the objective to improv e our understandin g of tectono-magmati c volcani c margi n settin g and th e processe s governin g volcani c mar gin initiatio n an d development . I t ha s als o helped i n providin g a framewor k fo r analysi s of th e implication s o f volcani c margin s fo r erosion an d sedimentatio n o n loca l an d regiona l scales, an d fo r th e environmen t (palaeoceano graphy, palaeoclimate ) o n local , regiona l an d global scales . Key results. A numbe r o f crusta l transect s has bee n constructe d acros s th e Nort h Atlanti c margin b y us e o f dee p seismi c dat a acros s the margin . I t i s estimate d tha t th e conjugat e margins experience d som e 140k m o f crusta l stretching durin g th e Maastrichtian-Paleocen e rifting an d breakup , an d abou t 50-7 0 km of latera l displacemen t durin g Lat e Jurassic Cretaceous riftin g (Skogsei d 1994 ; Skogsei d & Eldholm 1995 ; Skogsei d e t al.). Th e Rockal l Trough has , however , experience d fa r mor e stretching, whic h i s interprete d t o b e relate d t o separate riftin g i n th e mid-Cretaceous . B y using the stretchin g estimates , palinspasti c map s hav e been constructed base d o n restorations t o 5 3 Ma pre-drift, 7 5 Ma pre-Cenozoi c breaku p an d 170 Ma Lat e Jurassi c pre-rif t plat e configura tions. Th e result s als o includ e a n evaluatio n o f the tectono-magmati c event s associate d wit h plume-lithosphere interactio n durin g rifting , with particula r focu s o n relativ e vertica l move ments an d provenanc e development . The tectoni c developmen t o f th e V0rin g an d More basin s i s controlle d b y tw o structura l trends, NE-S W an d NW-SE . I t i s suggeste d that th e NE-S W tren d wa s establishe d i n th e Paleozoic an d wa s activ e durin g al l subse quent tectoni c phases , wherea s th e NW-S E trend, probabl y reflectin g th e ol d Precambria n grain o f th e basement , controlle d th e tectoni c activity throughou t th e Cretaceou s an d Ceno zoic (Brekke) . Th e result s sho w tha t durin g th e Cretaceous an d Cenozoi c th e V0rin g Basi n wa s tectonically active , wit h repeate d phase s o f nor mal faultin g an d contractio n causin g large-scal e folding. Th e Mor e Basin , particularl y toward s the south , wa s overal l mor e tectonicall y quie t and experience d mainl y continuous subsidence .

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Comparative studie s betwee n th e N E Atlan tic, th e Namibian , th e U S Atlanti c an d th e N W Australian margin s sho w tha t al l margin s see m to hav e ha d a protracte d rif t developmen t prio r to continental separation , tha t significan t crusta l thinning i s observe d adjacen t t o th e continent ocean boundary , an d tha t th e mai n puls e o f igneous activit y coincide s wit h th e tim e o f breakup (Eldhol m e t al. 1994 ; Eldhol m e t al.Planke & Eldhol m 1994) . O n th e othe r hand , i t is als o recognize d tha t exces s volcanis m ma y occur withou t a direc t lin k t o a n activ e mantl e plume, whic h ma y explai n th e larg e variet y i n tectono-magmatic developmen t o f volcanic margins worldwide. An extensiv e Norwegian-Greenland Se a thermal field data base named HEA T has been com piled, containin g al l publi c domai n dat a i n th e region; i.e . 436 hea t flo w value s (Sundvo r e t al.}. The hea t flo w o n oceani c crus t reveal s a clear , first-order hea t flow-crusta l ag e relationship , whereas continenta l slop e maxim a o n th e Mor e and Barent s Se a margin s contras t greatl y wit h the typica l lo w hea t flo w o f ol d oceani c an d thinned continenta l crust . Relevance t o th e petroleum industry. Th e un derstanding o f basi n geometries , rif t dimension s and subsidenc e histor y i s important t o explora tion companie s workin g offshor e mid-Norway . New informatio n o n magnitud e an d tempora l development o f intra-basinal vertical movements in combinatio n wit h palinspastic reconstructio n provides a too l fo r predictin g reservoi r facie s in th e oute r margi n basins . Suc h vertica l movements, includin g th e formatio n o f a lan d bridge fro m th e Charlie Gibb s Fractur e Zon e t o the S W Barents Sea margin, may hav e led to th e establishment o f larg e drainag e system s an d the probabilit y o f prospectiv e well-sorte d sedi ment sequence s o f generall y Lat e Cretaceou s Paleocene ag e i n th e adjacen t subsidin g parts o f the basins . In addition , th e geometrica l definitio n o f units o f igneou s material s a t crusta l level s has relevanc e fo r calculatio n o f heatflo w his tory, a s bodie s o f underplate d material s a t the bas e o f th e crus t ca n spik e heatflo w an d may hav e cause d maturatio n o f organi c mat ter tha t canno t b e predicte d b y presen t heat flow measurements . Th e compilatio n o f th e heat flo w databas e i s a n importan t too l i n this respect . The comparativ e wor k betwee n volcani c margins worldwid e i s importan t fo r th e under standing o f basi c geodynami c processe s an d tectono-magmatic developmen t i n general.

Concluding remark s The result s o f th e Integrate d Basi n Studies Dynamics o f th e Norwegia n Margi n projec t emphasize th e valu e o f academi a an d in dustry working together in order t o make signifi cant scientifi c progress . I t show s ho w publi c scientific grant s an d industr y fundin g ca n b e focused t o leverage investments put into research. It als o present s a mode l fo r ho w academi a an d industry researc h ca n b e effectivel y organize d into a singl e project. The achievement s o f th e Dynamic s o f th e Norwegian Margi n projec t hav e brough t ou r understanding o f thi s regio n forward , an d i t i s our hop e tha t th e result s presented i n thi s boo k will serv e as a n importan t referenc e for th e are a in th e year s to come. Th e projec t ha s als o estab lished a databas e tha t wil l serv e as a goo d basi s for continue d research o n the Norwegian margin and severa l ne w researc h project s hav e already been initiate d tha t buil d o n th e result s o f thi s project. I n addition , it has highlighte d some new : avenues o f researc h tha t ma y furthe r increas e our understandin g o f thi s margi n an d o f multi phase rif t evolutio n in general. Project staf f an d participants Direct project participation Principal Investigator s (PI) . Doctorat e Student s (PhD), Participatin g Scientist s (PS) , Master Student s (MS) an d Researc h Associate s (RA)

Theme 1: Intra-plate Rifting and Basin Formation Topics 1.1: Crust a! structure, 1.2: Sedimentary basin formation, 1.4: Tectoniic modelling PI: Prof . J . I . Faleid e (UO) . Prof. R . H . Gabrielse n (UB). Dr. W . Fjeldskaa r (RF) PhD: M.sci . P . Christiansso n (UO). C.sci. T . Odinse n (UB), C.sci. I . Grunnaleit e (UB) RA: C.sci . K . Lokn a (UB ) Industry: Dr . A . M . Berg e (Hydro). C.real. B . T. Larse n (Hydro). Prof. R.B . Faerseth (Hydro). Dr. A . Nottved t (Hydro). Dr. H . Fosse n (Statoil ) EU: Dr . P . Reems t (VU). Dr. P . va n de r Bee k (VU). M.sci. M . te r Voord e (VU). Prof. S . Cloetingh (VU) Topic 1.3: Present regional stress field PI: Prof . A . Myrvan g (NTH) . Prof. H . Bungu m (Norsar)

INTEGRATED BASI N STUDIE S PhD: Siv.ing PS: Dr MS: T

. M . Fejersko v (NTH ) . C . Lindhol m (Norsar ) . J0rgensen (NTH) , L. Borgeru d (NTH) , E. Svar e (NTH) , M. Villgra n (Norsar/UB) , E. Hick s (Norsar/UO ) Industry: Dr . T . H . Hansse n (Hydro) , C.real. B . T. Larse n (Hydro) , C.real. R . K . Bratl i (Saga) , Dr. L . N . Jense n (Statoil ) EU: Dr . M . Golke (UKa ) Theme 2: Basin Infill

Topic 2.1: Syn-rift sediment architecture PI: "Prof . R . J . Stee l (UB) , Dr. J . Underbil l (UEd ) PhD: C.sci . R . Ravna s (UB ) PS: Dr . P . Theriaul t (UB/Statoil) , D. Meller e (UB/Statoil) Industry: C.sci . K . Bondevi k (Hydro) , Prof. R . B . Faerseth (Hydro) , Dr. A . N0ttved t (Hydro) , J. Windelsta d (Statoil ) Topic 2.2: Post-rift sediment architecture, 2.3: Erosional episodes and provenance area PI: Prof . J . I . Faleid e (UO) , Prof. K . Bjorlykk e (UO ) PhD: M.sci . H . Jordt (UO) , C.sci. B.I . Thyber g (UO ) RA: Industry: Dr . P . va n Vee n (Hydro) , Dr. L . J. Skjol d (Hydro) , Dr. A . Ryset h (Hydro) , Dr. M . Ram m (Hydro ) Dr. A . N . N0ttved t (Hydro ) Topics 2.4: Stratigraphic modelling PI: C.real . M . Hambor g (IKU ) MS: R . Kyrkjeb o (NTH ) PhD: C.sci . R . Ravna s (UB ) Industry: C.sci . K . Bondevi k (Hydro) EU: M.sci . M . te r Voord e (VU) , Prof. S . Cloetingh (VU ) Theme 3: Conjugate Volcanic Margins Topics 3.1; Rift dimensions and duration, 3.2: Geodynamic modelling, 3.3: Comparative studies PI: Dr . J . Skogsei d (UO) , Prof. O . Eldhol m (UO ) MS: B . Flakstad (UO) , U. Byrkjelan d (UO) , F. Neverda l (UO) , S. Re n (UO) , E. Alvesta d (UO ) PhD: C.sci . T . Gladszenk o (UO ) PS: Dr . S . Planke (UO) , Dr. T . Pederse n (UO) , Prof. A . M . Myhr e Industry: C.rea l B . T. Larse n (Hydro )

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Referees. Th e followin g person s kindl y serve d a s technical an d linguisti c referees fo r th e book : John Akselsen , Aril d Andresen , Kuve t Atakan , W . Scott Baldridge , Giovann i Bertotti , 0rja n Birkland , Eric Bogoslowski , Lar s Boldreel , Ros s Boutilier , Nicolas Chamot-Rooke , Sier d Cloetingh , Deni s Cou turier, Tom Dreyer , Richar d England , Doug Gardner , Rob Gawthorpe , Feli x Gradstein , Pa l Haremo , Jen s Havskov, William Helland-Hansen, Helg e Hjelmeland , Chuck Hurich, Erik P. Johannesen, Reida r Kanestrom, Ridvan Karpuz , Oddbjor n K10vjan , Joh n Knight , John Korstgard , Yngv e KristofTersen , Axe l Makurat , Ole J . Martinsen , Alai n Mascle , Joc k McCracken , Wojtec Nemec , Joha n P . Nystuen , Arvi d Nottvedt , Lars Norgard-Jensen , Nigel Platt, John Palmer , Sara h Prosser, Garr y Quinlan , Phi l D . Rice , Ja n Rivenass , Ellen Roaldset , Ala n Roberts , Yngv e Rundberg , Al f Ryseth, William Sassi, Ke n Saunders , Roge r Scrutton , Michel Seranne , Morte n Sparr e Andersen , Geral d Sullivan, Tor e Torske , Bj0r n T0rudbakken , Ja n Vol set, Erlin g Vagnes , Joh n Walsh , Marjori e Wilson , Graham Yielding and fou r anonymou s referees . Acknowledgements. Th e projec t wa s funde d b y th e European Unio n an d o f Norwa y Researc h Counci l under th e JOUL E I I researc h programm e (contrac t No. JOU2-C T 92-0110) . Nors k Hydro , Statoi l an d Saga provide d additiona l funding . Ar e B . Carlsson a t the Research Counci l o f Norway i s gratefully acknowledged fo r supportin g th e project . A s projec t leader , I am ver y gratefu l t o Bjor n T . Larse n fo r hi s involve ment i n th e projec t o n a n industr y client basis an d t o Vigdis Michelse n for he r secretaria l efforts . I als o ow e many thank s t o pas t an d recen t member s o f th e project Steerin g Committee , Ola v Eldholm , Ro y H . Gabrielsen, Snorr e Olaussen , Al u Orheim , Ton y Spencer, Ro n J . Stee l an d Ja n Volse t an d t o m y fellow colleague s o n th e Editoria l Board , Bjorn . T . Larsen, 0rja n Birkeland , Haral d Brekke , Ro y H . Gabrielsen, Snorr e Olaussen , Jako b Skogsei d an d Bjorn Torudbakken , withou t who m thi s boo k woul d never hav e com e t o light . Som e extende d thank s als o go t o th e man y scientist s wh o reviewe d th e manu scripts. However , th e principa l investigators , partici pating researcher s an d students , wh o showe d grea t efforts an d a remarkabl e spiri t o f co-operatio n throughout, o f cours e ar e th e ke y t o th e succes s of th e project . Finally , I woul d lik e t o pas s a wor d o f appreciation t o Bernar d Durand , Sier d Cloetingh , Cai Puigdefabrega s an d al l othe r scientist s i n th e Integrated Basi n Studie s project , fo r 3 year s o f stimulating co-operation .

Research contribution s Publications BLYSTAD, P. , BREKKE , H. , F^RSETH , R . B. , LARSEN , B. T. , SKOGSEID , J . & TORUDBAKKEN , B . 1995 . Structural elements of the Norwegia n continental shelf. Par t II : Th e Norwegia n Se a region . Nor wegian Petroleu m Directorat e Bulletin , 8.

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BORGERUD, L . & SVARE , E . 1995 . In-situ stress fiel d o n the Norwegia n margin . In : FEJERSKOV , M . & MYRVANG, A . (eds ) Workshop o n Rock Stresses i n the North Sea. NT U Trondheim , 13-14.02.95 , pp.165-178. CLOETINGH. S. , SASSI , W. & TAS K FORC E TEAM \994a. The origi n o f sedimentar y basins : a statu s repor t from th e tas k forc e o f th e Internationa l Litho sphere Program . Marine an d Petroleum Geologv, 11, 659-683. CLOETINGH, S. , ELDHOLM , O. . LARSEN , B . T. , GABRIELSEN, R . H . & SASSI , W . (eds ) 1994/7 . Curren t state an d perspective s o f model s fo r extensiona l and inverte d basins . Tectonophysics, Specia l Vol ume, 240 . ELDHOLM, O . & THOMAS , E . 1993 . Scratchin g th e surface: Environmenta l impact o f volcanic margin formation. Earth an d Planetarv Science Letters, 117, 319-329 . ELDHOLM, O. , MYHRE , A . M . & THIEDE , J . 1994 . Cenozoic tectono-magmati c event s i n th e Nort h Atlantic: potentia l paleoenvironmenta l implica tions. In : BOULTER , M . C . & FISHER , H . C . (eds ) Cenozoic Plants and Climates o f the Arctic. NATO, ASI Series , 127 , Springer , Heidelberg , 35-55 . ELDHOLM, O. , SKOGSEID , J. , PLANKE , S . & GLADC ZENKO. T . P . 1995 . Volcanic margin concepts . In : BANDA. E. , TALWANI , M . & TORNE , M . (eds ) Rifted Ocean. Continent Boundaries. NAT O AS I Series Volume . Kluwer , Dordrecht , 1-16 . FEJERSKOV, M. , MYRVANG , A . M. , LINDHOLM , C . & BUNGUM, H . 1995 . In-sit u roc k stres s patter n o n the Norwegian continental shelf and mainland . In: FEJERSKOV, M . & MYRVANG , A . (eds ) Workshop on Rock Stresses i n th e North Sea. NT U Trondheim, 13-14.02.95 , pp. 191-201 . FJELDSKAAR, W. 1994 . The amplitud e and deca y o f the glacial forebulg e i n Fennoscandia . Norsk Geologisk Tidsskrift, 74 , 2-8 . FOSSEN, H . & GABRIELSEN , R . H . 1995 . Experimenta l modelling of extensiona l faul t systems . Journal o f Structural Geology, 18(5) , 673-687. FROSTICK, L . E . & STEEL , R . J . 19930 . Tectoni c signatures in sedimentary basin fills: an overview . International Association of Sedimentologists Special Publication, 20, 1-9 . FROSTICK, L . E . & STEEL , R . J . 19936 . Sedimentatio n in divergen t plate-margi n basins . International Association of Sedimentologists Special Publication, 20, 111-128 . F^RSETH, R . B . & RAVNAS . R . 1998 . Th e structura l configuration o f th e Oseber g Faul t Bloc k i n th e context o f th e norther n Nort h Se a structura l framework. Marine an d Petroleum Geology, 15 , 467-490. F/ERSETH, R . B. , GABRIELSEN , R. H . & HURICH , C . A . 19950. The influenc e of basement i n structuring o f the North Se a Basin offshore wes t Norway. Norsk Geologisk Tidsskrift, 75 , 2/3 , 105-119 . F^ERSETH, R . B. , SJ0BLOM , T . S. , STEEL , R . J. , LlLJE -

DAHL, T. , SAUAR , B . E . & TJELLAND , T . 19956 . Tectonic control s o n Bathonian-Volgia n syn-rif t successions o n th e Visun d Faul t Block , norther n North Sea . In : STEEL , R . J. , FELT , V . L. , JOHAN -

NESSEN, E . P . & MATHIEU . C . (eds ) Sequence Stratigraphy on the Northwest European Margin. Norwegian Petroleu m Societ y Specia l Publica tion. 5, 325-346 . GABRIELSEN, R . H . & STRANDENES . S. 1994 . Dynamic Basin Developmen t - A complet e geoscientifi c tool fo r basi n analysis . Proceedings Worl d Petro leum Congres s 1994 . 13-2 1 GABRIELSEN, R. H., GRUNNALEITE. I. & RASMUSSEN. E. 1997. Cretaceou s an d Tertiar y inversio n i n th e Bj0rn0yrenna Faul t Complex , south-western Barents Sea . Marine an d Petroleum Geologv. 14(2) . 165-178. GABRIELSEN. R . H. , ODINSEN . T . & GRUNNALEITE . I. 1999. Structurin g of th e norther n Vikin g Grabe n and th e Mor e Basin ; th e influenc e o f basemen t structural grain , an d th e particula r rol e o f th e M0re-Tr0ndelag Faul t Complex . Marine an d Petroleum Geology, 16 . 443-465. GABRIELSEN, R. H., STEEL, R. J. & NOTTVEDT. A . 1995. Subtle traps in extensional terranes: A model wit h reference t o th e Nort h Sea . Petroleum Geoscience. 1 , 223-235 . G0LKE. M. , COBLENTZ . S. , CLOETINGH . S . & FEJERS KOV. M . 1995 . Stres s syste m o f th e Norwegia n Continental Margi n - Par t I , In-sit u roc k stres s pattern o n th e Norwegia n continenta l shal f an d mainland. In : FEJERSKOV . M . & MYRVANG . A . (eds) Workshop o n Rock Stresses i n the North Sea. NTU Trondheim . 13-14.02.95 , pp. 250-274. GRUNNALEITE. I . & GABRIELSEN . R . H . 1995 . Th e structure o f the Mor e Basin . Tectonophvsics. 252 . 221-251. JORDT, H. . FALEIDE . J. I. . BJORLYKKE . K. & IBRAHIM, M. T . 1995 . Cenozoic stratigraph y o f th e centra l and norther n Nort h Se a Basin : tectoni c development, sediment distribution and provenance areas. Marine an d Petroleum Geology. 12 , 845-879. LINDHOLM, C . D. , BUNGUM . H. . VILLAGRAN . M . & HICKS, E. 1995. Crustal stress and tectonic s in Norwegian region s determined fro m earthquak e foca l mechanisms. In : FEJERSKOV . M . & MYRVANG . A . (eds) Workshop o n Rock Stresses i n the North Sea. NTU Trondheim , 13-14.02.95 , pp. 77-91 . MELLERE. D . & Steel, R . J . 1996 . Tidal sedimentatio n in Inner Hebrides half-grabens , Scotland : th e midJurassic Bearrerai g Sandston e Formation . /// : D E BATISTE, M . & JACOBS , P . (eds ) Geology o f Siliciclastic Shelf Seas. Geologica l Society . London, Specia l Publications , 117 . 49-79 . NOTTVEDT, A. , GABRIELSEN , R . H . & STEEL . R . J . 1995. Tectonostratigraph y an d sedimentar y archi tecture o f rif t basins , wit h referenc e t o th e northern Nort h Sea . Marine an d Petroleum Geology, 12 , 881-901. PLANKE, S . & ELDHOLM , O . 1994 . Seismi c respons e and constructio n o f seawar d dippin g wedge s o f flood basalts : V0rin g volcanic margin . Journal o f Geophysical Research. 99, 9263-9278. RAVNAS. R . & BONDEVIK . K . 1997 . Architecture an d controls o n th e Bathonian-Kimmeridgia n shal low-marine syn-rif t wedge s o f th e Oseberg-Brag e area, norther n Nort h Sea . Basin Research. 9 . 197-226.

INTEGRATED BASI N STUDIE S RAVNAS, R . & STEEL , R . J . 1997 . Contrasting style s of late Jurassi c syn-rif t turbidit e sedimentation : a comparable stud y o f th e Magnu s an d Oseber g areas, norther n Nort h Sea . Marine an d Petroleum Geology, 14 , 417-449. RAVNAS, R. & STEEL, R. J. 1998 . Architecture of marine rift-basin successions. AAPG Bulletin, 82,110-146. RAVNAS, R. , BONDEVIK , K. , HELLAND-HANSEN , W. , L0MO, L., RYSETH , A . & STEEL, R . J. 19970 . Sedimentation histor y as an indicato r of rif t initiation and development : Th e lat e Bajocian-Bathonia n evolution o f th e Oseberg-Brag e area , norther n North Sea. Norsk Geologisk Tidsskrift, 77,205-232 . RAVNAS, R. , HANSEN , J . W. , MELLERE , D. , NOTT VEDT, A., SJ0BLOM , T. S. , STEEL , R . J . & WlLSON,

R. C . L . 19976 . A marin e lat e Jurassi c syn-rif t succession i n th e Lusitania n Basin , wester n Portugal - tectoni c significanc e o f stratigraphi c signature. Sedimentary Geology, 114 , 237-266 . REEMST, P. , SKOGSEID , J . & LARSEN , B . T. 1996 . Base Pliocene velocit y inversion on th e easter n Vorin g margin - cause s an d implications . Global an d Planetary Change, 12 , 201-211 . SKOGSEID, J . 1994 . Dimension s o f Lat e Cretaceous Paleocene Northeas t Atlantic rif t derived from Cenozoic subsidence. Tectonophysics, 240 , 225-247 . SKOGSEID, J. & ELDHOLM , O . 1995 . Rifted continental margins of f mid-Norway . In : BANDA , E. , TAL WANI, M . & TORNE , M . (eds ) Rifted Ocean. Continent Boundaries. NAT O AS I Series Vol ume, Kluwe r Academic PUBLISHERS , pp. 147-153 . SKOGSEID, J. , ELDHOLM , O . & PLANKE , S . 1994 . Mesozoisk kontinenta l riftin g o g Kenozoisk mar gindannelse: dypseismik k o g skorpestruktu r p a V0ringmarginen. Geonytt, 21, 3-18 . STEEL, R . J . 1993 . Triassic-Jurassi c megasequenc e stratigraphy i n th e norther n Nort h Sea : rif t t o post-rift evolution . In : PARKER , J . R . (ed. ) Petroleum Geology o f Northwest Europe. Geologica l Society, London , 299-315 . STUEVOLD, L . M . & ELDHOLM , O . 1996 . Cenozoi c uplift o f Fennoscandi a inferre d fro m a stud y o f the mid-Norwegia n margin . Global and Planetary Change, 12 , 359-386.

TERVOORDE, M. , RAVNAS , R., F^RSETH, R. B . & CLOE-

tingh, S. 1997. Tectonic modelling of middle Jurassic syn-rift stratigraph y in the Oseberg-Brage area , northern North Sea. Basin Research, 9, 133-150. THERIAULT, P. & STEEL , R . J . 1995 . Aspects o f synrif t sedimentation i n th e Uppe r Jurassi c (Helmsdal e Boulder Beds) of the Inner Moray Firt h Basin . In: STEEL, R . J. , FELT , V . L. , JOHANNESSEN , E . P . & MATHIEU, C . (eds ) Sequence Stratigraphy o n th e Northwest European Margin. Norwegia n Petro leum Societ y Specia l Publication , 5 , 365-387.

Reports Reports liste d herei n ar e publi c an d ca n b e mad e available throug h th e Researc h Counci l o f Norwa y (RCN), Universit y o f Berge n (UB) , Universit y o f Oslo (UO ) an d Norwegia n Technica l Universit y i n Trondheim (NT U = NTH) .

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FEJERSKOV, M. 1993 . Bergspenninger i Norge o g pa den norske sokkel . In : MYRVANG , A. , JOHANSEN , T. , HANSEN, A . & BERG , K . R . (eds ) Fjellsprengingsteknikk, bergmekanikk, geoteknikk. Osl o 1993 , 17pp. FEJERSKOV, M . 1994 . Breakout a s a tool fo r stress determination i n deep wellbores. NT H Repor t No. 1 , 15pp . FEJERSKOV, M. 1994 . Breakout interpretation. Methods and software used a t NTH. NT H Repor t No . 3 , 14pp. FEJERSKOV, M . 1994 . Breakout identification i n 7 wells near the Troll Field on the eastern flank of the northern Viking Graben. NTU Repor t No . 4, 39pp. FEJERSKOV, M . 1995 . Criteria for breakout identification based o n 4-arm oriented caliper logs. NT U Report No . 2 , 17pp . FEJERSKOV, M . 1995 . Breakout interpretation - 11 wells on the Visund Field, northern Viking Graben. NTU Repor t No . 5 , 39pp . FEJERSKOV, M. , i n pre p 1995 . Breakout interpretation in th e Tampen Spur area. NTU Repor t No . 6 . GLADCZENKO, T . P . & ELDHOLM , O . 1995 . LI P Database: Large Igneous Provinces - distribution and references. Geophys . Res. Group, Dep. Geol. , Univ. Oslo , Compute r Pgm./Databas e Doc . Ser . No. 15 , 5pp . HAMBORG, M. , KYRKJEBO , R . & RAVNAS , R . 1995 . Syn- and post-rift depositional modelling, northern North Sea. IK U repor t xxxxx . MARJANAC, L . T . 1994 . Reference data base for rift basins (syn-rift). Ui B Report , 41pp . NOTTVEDT, A. & IBS-DNM working group 1993 . IB S Module 3 - Dynamics of the Norwegian Margin. First Periodical Report - Project Description, June 1993, 28pp . NOTTVEDT, A . & IBS-DN M workin g grou p 1993 . Minutes of Meeting, IBS-DNM Project Seminar. Geilo, Novembe r 1993 , 57pp . NOTTVEDT, A. & IBS-DNM workin g group 1993 . IB S Module 3 - Dynamics of the Norwegian Margin. Second Periodical Report, December 1993, 24pp . NOTTVEDT, A . & IBS-DN M workin g grou p 1994 . Minutes of Meeting, IBS-DNM Project Seminar. Stavanger, Ma y 1994 , 81pp . NOTTVEDT, A. & IBS-DNM workin g group 1994 . IBSDNM Project Status Report, Ma y 1994 . NOTTVEDT, A. & IBS-DNM working grou p 1994 . IB S Module 3 - Dynamics of the Norwegian Margin. Third Periodical Report, June 1994, 32pp. NOTTVEDT, A. & IBS-DNM workin g group 1994 . IB S Module 3 - Dynamics of the Norwegian Margin. Fourth Periodical Report, December 1994, 49pp.N0TTVEDT, A . & IBS-DN M workin g group 1995 . IBS-DNM Project Status Report, February 1995. Three volumes . NOTTVEDT, A. & IBS-DNM workin g group 1995 . IB S Module 3 — Dynamics of the Norwegian Margin. Fifth Periodical Report, June 1995, 37pp . NOTTVEDT, A. & IBS-DNM workin g group 1994 . IB S Module 3 - Dynamics of the Norwegian Margin. Final Report, December 1995. PLANKE, S . 1993 . HEAT- Heat flo w data base program. Geophys . Res . Group , Dept . Geol. ,

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Univ. Oslo , Compute r program/Dat a bas e Doc . Ser. No . 2 version 3.1, 10pp . PLANKE, S . 1993 . VELO - Seismic refraction!wideangle reflection velocity data base program. Geophys. Res . Group , Dept . Geol. , Univ . Oslo , Computer program/Dat a bas e Doc . Ser . No . 1 version 3.1 , 14pp. RAVNAAS, R. . HANSEN . J . W. . MELLERE , D. . NOTT VEDT, A. , SJ0BLOM , T . S . & STEEL , R . J . 1995 .

A marine to continental svn-rift succession: the Kimmeridgien Abadia and Lourinha Formations of the Santa Cruz area, Lusitanian Basin, Portugal. University o f Berge n Report . VANVEEN, P., ' SKJOLD . L . J . & RYSETH , A . E . 1994. A High-Resolution Stratigraphic Framework for the Paleogene i n th e Northern North Sea. Norsk Hydro Repor t R-055419 .

Theses BORGERUD, L . 1995 . Relations between Rock Stresses and Pore Pressure on the Norwegian Margin, 62-67 north - a study based on borehole breakouts. Diplom a thesis , Dep. of Geolog y an d Mineral Resource s Engineering , Norwegian Technical University, Trondheim . FEJERSKOV. M . 1996 . Determination o f in-situ rock stresses related to petroleum activities on the Norwegian continental shelf. Dr . ing . thesis . Department o f Geolog y an d Minera l Resource s Engineering, Norwegia n Technica l University , Trondheim, Norway . 162pp . FINSTAD, A . G . 1995 . Main structural elements o n th e Norwegian continental margin, 68-72"N and adja-

cent lan d areas : tempora l an d spatia l develop ment. Cand . scien t thesis , Dep . o f Geology . University o f Oslo . HOLMSEN. C . 1994 . Vndersokelse av kenoioiske o g mesoioiske sedimenter i Vikinggraben i relasjon til de n tertiare landhevingen. Cand . scient . thesis. Dep. o f Geology . Universit y o f Oslo . IBRAHIM. M . T . 1993 . Post-Jurassic basin fill i n th e northern North Sea. Cand. scien t thesis , Dep. of Geology, Universit y o f Oslo . JORDT, H . 1995 . The Cenoioic geological evolution o f the central and nor them North Sea based on seismicsequence stratigraphy. Dr . scien t thesis . Depart ment o f Geology, Universit y o f Oslo . JORGENSEN. T . 1994 . Determination an d Evaluation o f In-situ Rock Stresses a t th e Snorre Field. Diploma thesis, Dep . of Geolog y an d Minera l Resource s Engineering. Norwegia n Technica l University . Trondheim. KARLBERG. T . 1995 . En geofysisk undersokelse a v Hovgaardryggen. Cand . scien t thesis . Dep . o f Geology. Universit y o f Oslo . RAVNAS. R . 1996 . Variability o f syn-rift sedimentary architecture in marine rift-basins: examples from the middle-late Jurassic of the northern liking Graben, North Sea, and the Lusitanian Basin, western Portugal. Dr . scient . thesis . Department of Geology . Universit y o f Bergen . 210pp. SVARE, E . 1995 . Relations between Rock Stresses an d Pore Pressure on the Norwegian Margin, 62 67 north - a study based on leak-off tests and formation pressure. Diplom a thesis . Dep . of Geolog y and Minera l Resource s Engineering . Norwegian Technical University , Trondheim.

Crustal structur e i n the norther n North Sea : a n integrated geophysical stud y P. CHRISTIANSSON, 1 2 J . I . FALEIDE 1 & A . M . BERGE 3 1

2

Department of Geology, P.O. Box 1047 Blindern, N-0316 Oslo, Norway Present address: Norsk Hydro ASA, P.O. Box 200, 1321 Stabekk, Norway 3 Norsk Hydro Research Centre, P.O. Box 646, 5020 Sandsli, Norway Abstract: Thi s study focuses o n the deep structure o f the Viking Graben and adjacen t area s of th e norther n Nort h Se a (60-62°N) , an d it s implication s fo r th e amount , timin g an d nature of lithospheric extension. Two regiona l transects have been constructed base d o n a n integrated analysis of deep seismic reflection and refractio n data , gravity and magnetic data, and correlation s betwee n offshore an d onshor e geology. The shallow interpretation is based on high-qualit y conventional seismi c reflectio n dat a calibrate d agains t a larg e numbe r o f exploration wells . Th e ne w and partl y reprocesse d seismi c data , combine d wit h th e othe r geophysical data, make possible a better documentation of the crustal configuration, such as the pre-Jurassi c sedimen t distribution , basement an d Moh o relief , an d dee p faul t geom etries. A lower-crusta l bod y characterize d b y a n 8+kms" 1 velocit y an d a n averag e bul k density of 2.95 gcm~3 is present beneath the Horda Platform. This body probably represents a deep crustal root of partially eclogitized rocks that formed during the Caledonian orogeny. Heterogeneities withi n this bod y giv e rise to th e non-typica l velocity-density relation . Th e crust-mantle boundary is located a t th e bas e o f this body a t a depth o f 30-35 km and doe s not coincid e with the seismically defined Moho . Th e geometry of crustal thinning reflects th e cumulative effec t o f severa l post-Caledonian rif t phases . Result s sho w tha t Permia n riftin g affected a wid e area, fro m th e 0ygarde n Faul t Comple x t o th e Hutto n Fault .

Deep seismi c reflectio n dat a hav e forme d th e basis fo r numerou s paper s throug h th e las t decade, focusin g o n th e crusta l structur e an d basin evolutio n i n th e norther n Nort h Se a rif t system. However , poo r dat a quality , especiall y the lo w S/ N rati o a t depth , ha s le d t o man y model-driven interpretations. The NSDP84-line s were firs t describe d b y Gibb s (1987<2 , b) an d Klemperer (1988) . Additiona l interpretation s have been presented by Harrison (1987) , Kusznir & Matthew s (1988) , White & McKenzie (1988), Klemperer & White (1989) , Pine t (1989) , White (1989, 1990), Klemperer &Hurich( 1990), Reston (1990) an d Bru n & Tro n (1993) . A commercia l deep seismi c profile , Britoi l NNS83-22 , tha t i s located close to NSDP84-1 (Fig. 1 ) was described by Beac h (1985 , 1986 ) and Beac h e t al. (1987). The eastern ends of the NSDP84 lines are tied by two north-south profiles, ILP-10 and -1 1 (Fig. 1) , recorded b y the Norwegian Internationa l Lithosphere Projec t (Huric h & Kristofferse n 1988 ; Klemperer & Huric h 1990 ; Fasrseth e t al . 1995 ; Hurich 1996) . Interpretations an d combine d model s o f th e NSDP84 reflectio n data an d gravit y data hav e been publishe d b y Hollige r (1987) , Hollige r & Klemperer (1989 ) and Fichle r & Hospers (1990) ;

and Zervo s (1987 ) modelle d th e gravit y field along si x regiona l seismi c profile s take n fro m Ziegler (1982 ) an d Glenni e (1984) . Hosper s & Ediriweera (1988 , 1991 ) publishe d a ma p o f depth t o th e crystallin e basement , base d o n a n integrated analysi s o f magnetic , gravit y an d seismic data . Most paper s agre e i n Moh o depth s rangin g between 9.5 and 10. 5 s twt identified as the base of lower-crustal reflectivity . However , n o moder n deep seismi c refraction data hav e been available to constrai n th e dept h conversio n an d con firm th e publishe d crusta l thicknes s an d dept h to Moh o maps . Furthermore , poo r definitio n o f the to p o f th e crystallin e basemen t make s i t difficult t o construc t basemen t thicknes s map s that reflec t th e cumulativ e crusta l thinnin g i n response t o severa l post-Caledonian rif t phases . Several models have been proposed t o explain the crusta l thinnin g an d basi n formatio n i n th e northern Nort h Sea, an d the same regiona l dee p seismic line s have bee n interprete d accordin g t o both th e simpl e an d pur e shea r models . Th e symmetrical pur e shea r mode l (McKenzi e 1978 ) was favoure d for th e Vikin g Graben b y Giltne r (1987) an d Badle y e t al . (1988) . Alternatively, the asymmetrica l simple shear mode l (Wernick e

From: NOTTVEDT , A . e t al . (eds ) Dynamics o f th e Norwegian Margin. Geologica l Society , London , Specia l Publications, 167 , 15-40 . 1-86239-056-8/00/S15.0 0 © Th e Geologica l Societ y of Londo n 2000 .

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P. CHRISTIANSSON , J . I . FALEID E & A . M . BERG E

Fig. 1 . Mai n structura l element s withi n th e stud y area . Locatio n o f th e tw o regiona l transect s an d othe r dee p seismic reflectio n an d refractio n profiles .

1985) ha s bee n use d t o explai n the asymmetry o f the grabe n structur e (Beac h 1986 ; Beac h e t al. 1987; Scott & Rosendahl 1989) . Tectonic model s combining simpl e an d pur e shear , characterize d by structura l asymmetr y o f th e brittl e uppe r crust an d symmetrica l thinnin g o f th e lowe r ductile crust , hav e gathere d increasin g suppor t (Kusznir e t al . 1991 , 1995 ; Kuszni r & Ziegle r 1992). Odinsen e t al. (19990) have also favoure d such a combined model , involvin g separation o f the crust at differen t detachmen t level s caused b y rheology an d pre-existin g zones o f weakness .

The ai m her e i s t o presen t a n integrate d geophysical stud y of th e crusta l structure acros s the norther n Nort h Sea . Th e dee p crusta l structure wil l b e relate d t o th e onshor e geolog y of wester n Norwa y wherea s th e geometr y o f crustal thinnin g wil l b e relate d t o th e post Caledonian evolutio n includin g severa l rif t phases. Fo r thi s purpose, tw o regiona l transects have bee n constructe d base d o n high-qualit y conventional seismi c reflectio n dat a an d repro cessed dee p seismi c reflectio n profile s (NSD P 84-1 an d -2 ; Fig. 1) . The crusta l configuration is

CRUSTAL STRUCTUR E I N TH E NORTHER N NORT H SE A further constraine d b y integratio n o f dee p seis mic refraction , gravit y an d magneti c data . Furthermore, seismi c mappin g o f th e stud y area includin g som e 15000k m o f high-qualit y conventional seismi c dat a ha s bee n carrie d ou t to establis h a regiona l structura l an d strati graphic framework . Geological framewor k The Nort h Se a rif t system , post-datin g th e Caledonian orogeni c extensiona l collapse , wa s affected b y tw o lithospheri c extensio n event s i n Permian-?earliest Triassi c time (in the literature often referre d t o a s th e Permo-Triassi c event ) and late Mid-Jurassic to earliest Cretaceous time (here referre d t o a s Lat e Jurassic) , respectively. The rif t axi s fo r th e Permo-Triassi c rif t i s thought t o li e beneath th e presen t Hord a Plat form wherea s th e Lat e Jurassi c rif t wa s centre d beneath th e presen t Vikin g Grabe n (Fig . 1) . Each rif t phas e wa s followe d b y a therma l cooling stage , characterize d b y regiona l subsi dence i n th e basi n area s (Ziegle r 1982 ; Gabrielsen 1986 ; Giltne r 1987 ; Badle y e t al. 1988 ; Gabrielsen e t al. 1990) . The norther n North Se a rift syste m i s bounde d b y th e Eas t Shetlan d Platform i n th e wes t an d th e 0ygarde n Faul t Zone i n th e eas t (Fig . 1) . Structures within thi s area ar e characterize d b y larg e rotate d faul t blocks wit h sedimentar y basin s i n asymmetri c half-grabens associate d wit h lithospheri c extension an d thinnin g of th e crust . Devonian strata , deposite d i n response t o th e extensional collapse of the Caledonides, are present i n onshor e wester n Norwa y (Hossac k 1984 ; Steel e t al . 1985 ; Norton 1986 , 1987 ; Seranne & Seguret 1987 ; Osmundsen 1995) , Shetland (Allen 1981; Flin n 1985 ; Serann e 1992 ) an d easter n Greenland (Hart z & Andresen 1995) . Althoug h Devonian sediment s i n th e norther n Nort h Se a have bee n reache d i n only a fe w wells, there ar e reasons t o believ e tha t Devonia n sediment s are presen t regionall y i n th e deepe r part s o f the pre-Triassic half-graben s beneath th e Hord a Platform, Vikin g Grabe n an d Eas t Shetlan d Basin. Recen t studie s o f seismi c reflectio n data on th e Eas t Shetlan d Platfor m (Hollowa y e t al. 1991; Plat t 1995 ) indicat e larg e sedimentar y basins though t t o contai n Uppe r Palaeozoi c (Devonian-Carboniferous) rocks . Th e presenc e of Uppe r Palaeozoi c rocks , o f bot h Devonia n and Lowe r Permian ag e (Rotliegendes), has also been confirme d b y drilling o n th e Eas t Shetlan d Platform (Johnso n e t al . 1993 ; Duncan & Buxton 1995) . Upper Palaeozoi c rock s hav e not yet been encountere d i n th e easter n par t o f th e

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northern Nort h Se a basin. However , th e seismi c data o n th e Hord a Platfor m indicat e tha t a s much a s 3- 4 km o f strat a exis t i n th e deepes t parts o f th e half-grabens , whic h ar e no t teste d by drilling. These sediment s probably dat e bac k into Permia n an d olde r times . Permo-Triassic riftin g i s bes t documente d i n areas margina l t o th e Vikin g Graben , e.g . th e Horda Platfor m an d th e Magnu s Basin-Uns t Basin region , becaus e thes e were les s influence d by Jurassi c deformation . Here , seismi c dat a display Triassi c an d olde r sediments , whic h show wedge-shape d geometrie s indicatin g tha t these area s wer e technicall y activ e i n th e earl y rift phase . O n th e Hord a Platform , seismi c data calibrated agains t wel l dat a (e.g . 31/6-1 ; Fig . 1 ) suggest tha t thi s phas e terminate d i n earl y Triassic times . Th e initiatio n o f thi s phas e i s not date d i n th e stud y are a bu t i s generall y assumed t o hav e bee n durin g Permia n tim e (Fcerseth e t al . 1995 ; Odinse n e t al . 19990) . Palaeomagnetic datin g o f th e Nordfjord-Sog n detachment, boundin g th e Devonia n basin s o f Western Norway , indicate s a tectoni c even t a t around 26 0 Ma (Torsvi k e t al . 1992) . Dolerit e dykes i n th e coasta l Sunnfjor d regio n an d immediately wes t o f th e Devonia n basin s ar e also o f Permian ag e (250-270 Ma; Torsvi k e t al. 1997). Followin g th e rif t structure s belo w th e Horda Platfor m southward , tilte d faul t block s are observe d jus t belo w th e Uppe r Permia n Zechstein sal t (Ditch a 1998) . Th e north-sout h trending structures that dominate d th e early rif t phase were inherited from earlie r deformation of Precambrian basemen t (Faerset h e t al . 1995) . The Triassi c t o Middl e Jurassi c successio n reflects a pattern o f repeated outbuildin g of clastic wedge s fro m th e Norwegia n an d Eas t Shet land hinterland s withi n a generall y evolvin g post-rift basi n (Lervi k e t al . 1989 ; Stee l & Ryseth 1990 ; Stee l 1993) . A broadl y simila r geometry o f th e megasequence s i n bot h con tinental Triassic an d marine Jurassic succession s was relate d b y Stee l (1993 ) t o subsidence-rat e variations. Differentia l subsidenc e acros s fault s throughout Triassi c time has als o bee n reporte d (Steel & Ryset h 1990) . Th e 0ygarde n Faul t Zone, formin g the eastern margi n of the Permo Triassic basin, was active throughout most of the time interval. The lates t tectonic event o f major importanc e in th e norther n Nort h Se a i s th e Lat e Jurassi c stretching event, characterized b y the northeastsouthwest trendin g Vikin g Graben . Th e initia tion o f th e Jurassi c rif t even t i n th e norther n North Se a is dated t o lat e Mid-Jurassic time by most workers , wit h a clima x i n Lat e Jurassi c time (Ziegle r 1982 ; Beac h e t al . 1987 ; Badley

Fig. 2 . Dee p seismi c reflectio n profile s acros s th e norther n N o r th Sea. (a) NSD P 84- 1 (Transec t I ) ; (b ) NSD P 84-2 (Transect 2) . Locatio n o f profile s i s shown i n Fig . I.

CRUSTAL STRUCTUR E I N TH E NORTHER N NORT H SE A et al 1988 ; Gabrielse n e t al. 1990) . Th e initia l stages o f Jurassi c riftin g influence d a broade r area tha n di d th e fina l stage s o f grabe n forma tion (Gabrielse n e t al . 1990) . Continue d exten sion wa s concentrate d o n fewe r fault s alon g th e graben margins . Simultaneously , th e interna l graben relie f becam e mor e pronounced , an d th e rift syste m develope d a matur e grabe n topogra phy wit h platforms , sub-platforms , platfor m marginal high s an d a grabe n featur e wit h complex centr e o f subsidenc e alon g it s axi s (Gabrielsen e t al . 1990 ; Nottvedt e t al . 1995) . During Earl y Cretaceou s tim e continue d dif ferential subsidenc e wa s accompanie d b y th e gradual infillin g o f th e grabe n relie f wit h deep water shale s an d mino r pelagi c carbonates . Regional subsidenc e combine d wit h a genera l rise i n eustati c se a leve l resulte d i n th e progres sive oversteppin g o f basi n edge s an d a gradua l decrease i n clasti c influ x int o th e Nort h Se a basin. Th e hig h Lat e Cretaceou s eustati c se a level provided condition s suitabl e for widesprea d deposition o f chalk. A thick Uppe r Cretaceous Danian chal k successio n cover s mos t o f th e central an d souther n Nort h Sea , wherea s terri genous mudstone s dominat e i n th e norther n North Sea. The northward chang e fro m chalk t o mudstone i n th e centra l Vikin g Graben i s gra dational, bot h laterall y an d vertically . Even a t the norther n limit s of th e Vikin g Graben, ther e are a fe w metres o f Maastrichtia n chal k (Han cock 1990) . The Cenozoic basi n fill of the northern Nort h Sea reflect s a comple x patter n o f regiona l ver tical movement s controllin g basi n geometr y an d provenance o f th e sediment s (Jord t e t al . 1995 , 1999). Paleogene sediment s were mainly derived from area s i n th e wes t an d north , wherea s southern Norwa y becam e a n importan t sourc e area from Oligocene an d throughout Lat e Ceno zoic times . Depocentr e location s an d sequenc e geometries indicat e tha t differentia l vertica l movements o f th e basi n floo r wer e closel y related t o boundarie s o f underlyin g Mesozoi c and Palaeozoi c structure s i n th e Nort h Se a Basin (Jord t e t al . 1999) . Data analysi s For a bette r understandin g o f th e crusta l structure an d it s implication s fo r basi n forma tion an d evolution , w e have integrate d al l available geophysica l data . Th e uppe r crus t i s bes t displayed o n high-qualit y conventiona l seismi c reflection data (0-7 s twt) whereas the middle and lower crus t i s bes t studie d o n dee p seismi c reflection dat a (0-15 s twt) (Figs 2 and 3) . Dee p

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seismic refractio n dat a (expande d sprea d pro files, ESPs ) provid e crusta l velocitie s tha t ar e used i n th e dept h conversion . Gravit y an d magnetic data constrain the deep basin geometry and crusta l structure . Ou r integrate d approac h also include s correlation s betwee n th e offshor e and onshor e geology . The regiona l profile s i n Fig . 4 wer e con structed b y combining traditional interpretation of th e mai n sequenc e boundarie s wit h th e pat tern o f crusta l reflectivit y o n th e dee p seismi c lines (Fig. 3) . Transect 1 is oriented NW-S E an d crosses th e Magnu s Basin , Eas t Shetlan d Basin , Viking Grabe n an d Hord a Platform , an d Transect 2 run s E- W acros s th e Eas t Shetlan d Platform, Vikin g Grabe n an d Hord a Platfor m (Figs 1 and 4) .

Conventional seismic reflection data (0-7 s twt) Despite th e stron g focu s o n th e tw o crusta l transects, w e have interprete d a relativel y dens e grid of conventional regional seismic lines covering th e Norwegia n sector , bu t onl y a fe w line s have bee n availabl e o n th e U K side . Ke y line s for th e regiona l transect s hav e bee n th e repro cessed version s of SG8043-40 3 (Transec t 1 ) an d SG8043-101 (Transect 2) , which show improve d resolution o f th e dee p basi n stratigraph y afte r reprocessing. Th e seismi c stratigraph y alon g the transect s ha s bee n calibrate d agains t selec ted wells , which , i n combination , cove r th e stratigraphic interva l from Triassi c t o Quatern ary time . The shallo w par t o f th e profile s (Fig . 4 ) i s characterized b y a relativel y flat-lying Cenozoic and Cretaceou s post-rif t sequenc e overlyin g faulted Mesozoi c an d Uppe r Palaeozoi c strata . The axi s o f th e Vikin g Grabe n display s th e thickest depocentre s o f post-rif t sequences . Jurassic strat a ar e als o thickes t i n th e Vikin g Graben, particularl y on Transec t 2 (Fig. 4) , and thin a t th e basi n flanks. A Permo-Triassic basi n is bes t visualize d o n th e Hord a Platform . Here , growth sequence s toward s th e majo r fault s indicate pre-Jurassi c faul t activit y giving rise t o a syste m of half-grabens. Th e mos t pronounce d difference i n structura l styl e betwee n th e tw o transects i s th e chang e i n polarit y o f th e tilte d fault block s beneat h th e easter n par t o f th e Horda Platfor m (Fig . 4) . Th e polarit y shif t i s probably connecte d t o transfe r faults runnin g in an east-west direction. A strong reflecto r tied t o basement drille d i n wel l 31/6- 1 (Fig . 1 ) can b e traced fo r som e distanc e o n th e centra l Hord a Platform, bu t i s difficult t o interpre t regionally.

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P. CHRISTIANSSON , J . I . FALEID E & A . M . BERG E

Fig. 3. Seismi c dat a example s fro m th e sam e par t o f Transec t 1 . (a) Dee p seismi c reflectio n lin e NSD P 84-1 : (b) conventiona l seismi c reflectio n lin e NVGTI-92-102 .

Deep seismic reflection data (0-15 stwt) The dee p seismic reflection lines (NSDP84, Fig s 1 and 2 ) were acquired an d processed i n 1984-198 5 by GEC O o n behal f o f BIRP S an d severa l oi l companies. The y wer e furthe r reprocesse d i n 1991 by BIRPS (Klempere r & Hobbs 1991) . Th e dataset was also reprocesse d b y Norsk Hydr o i n 1994 fo r th e IB S projec t t o enhanc e th e lower crustal reflectivit y an d Moh o definition . Important elements in the post-stack reprocessin g were correct scalin g an d balancin g o f section s t o preserve dynami c range , an d prope r migra tion. Reprocessing improve d definitio n of lowercrustal structure , an d th e Moh o ca n no w b e observed a s a contras t i n reflectiv e characte r along mos t o f th e dee p seismi c line s (Fig s 2 and 3) . Migration, using the velocity fields established alon g th e crusta l transects , ha s provide d more correc t geometrie s an d faul t definition . However, th e dat a ar e stil l contaminate d b y long-period multiples . The genera l outlin e o f th e transect s i s similar to tha t o f man y extende d area s (e.g . Blundel l 1990), wit h a reflectiv e sequenc e o f sediment s

near the surface, a transparent middl e crust, an d a reflectiv e lowe r crust (Figs 2 and 4) . The uppe r reflective crus t ha s alread y bee n describe d base d on the conventional seismic reflection data. Here , we wil l focu s on th e middl e an d lowe r crust. Although th e middl e crus t i s characterize d by poo r reflectivity , severa l importan t observa tions ca n b e mad e o n th e reprocesse d dee p seismic dat a whe n thes e ar e combined wit h con ventional data . Som e graben-margi n maste r faults hav e bee n trace d throug h th e transparen t middle crus t t o th e lowe r reflectiv e crust , where faults see m t o merg e a t c . 18km dept h (Fig . 4) (Odinsen e t al. \999a). Th e Hutto n Faul t o n Transect 1 i s mainl y a Permo-Triassi c featur e whereas the f c Western Margi n Fault' of the Viking Graben o n Transec t 2 is mainly a Lat e Jurassi c feature (Fig . 4). The larg e east-dippin g faul t boundin g th e horst structur e at th e central Horda Platform on Transect 2 seems to correlate with a zone of eastdipping mid-crusta l reflection s below th e 0ygarden Faul t Zon e (Fig s 2 an d 4) . Thes e dippin g reflections migh t b e linke d t o Caledonia n struc tures beneat h wester n Norway . East-dippin g

Fig. 4 . Lin e drawing of the tw o regiona l crusta l transects . Th e crusta l reflectivit y patter n i s interpreted fro m th e reprocesse d dee p seismi c reflectio n profiles in Fig . 2 , whereas the basin stratigraphy is based o n conventional seismic reflection data. HF, Hutto n Fault ; WMF, 'Wester n Margin Fault'. Location o f profiles is shown in Fig. 1 .

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P. CHRISTIANSSON , J . I . FALEID E & A . M . BERG E

mid-crustal reflections are also observe d beneat h the Magnu s Basi n o n Transec t 1 (Fig. 1) . The reflectiv e lowe r crus t i s see n a s a n undulating transition zone between the overlying transparent middl e an d uppe r crystallin e crus t and th e underlyin g upper mantle . Th e thicknes s of this reflectiv e band normall y varies between 3 and 6 km but locall y increases t o 1 0 km under th e basin margins . Th e reflectivit y is strongest unde r the basi n flanks , clos e t o th e Norwegia n coast , and become s weake r unde r th e Vikin g Graben , which is also the most extende d area . Th e Moho is well displayed, eithe r a s a single reflector or a s the bas e o f th e lower-crusta l reflectivity , an d shallows belo w th e mos t strongl y extende d par t of th e rif t syste m (Fig s 2 an d 4) . Separation o f th e lowe r crus t int o a n uppe r and lowe r reflectiv e zon e ca n b e see n i n bot h transects beneat h th e Horda Platform , althoug h the lower-crustal reflection s are weaker o n Transect 2 giving rise to tw o Moho candidates in this area. A similar geometry was observed b y Beac h (1986) on a nearby deep seismi c reflection profile (profile 3 i n Fig . 1) . W e wil l retur n t o a dis cussion o f th e bod y bounde d b y th e tw o Moh o candidates afte r integratio n wit h velocit y an d density informatio n fro m analyse s o f othe r geophysical data . Reflections belo w th e Moh o ar e recorde d i n both transect s (Fig s 2 an d 4) . Thes e features , dipping awa y fro m th e grabe n axis , ar e locate d close t o th e transitio n betwee n th e thinne d crus t and th e thicker crust beneat h th e platform areas . This patter n o f intra-mantl e reflection s has als o been describe d b y previou s wor k (Beac h 1986 ; Beach e t al. 1987 ; Klempere r 1988 ; Hollige r & Klemperer 1989 ; Bru n & Tro n 1993) . Th e interpretation o f these structure s offer s severa l solutions, bu t th e coincidenc e betwee n th e dipping reflector s an d th e di p gradien t o f th e Moho ma y sugges t tha t th e reflection s represent faults o r shea r zone s withi n the mantle .

Deep seismic refraction data The dee p seismi c refractio n dat a (ESPs ) wer e collected b y Elf Aquitaine Norg e i n co-operatio n with th e Universitie s o f Osl o an d Bergen . During th e researc h programme , fro m 198 3 t o 1986, 10 3 ESPs an d 17 3 heat-flow measurements were acquire d fro m th e Norwegia n continenta l margin, th e northern Nort h Sea , th e Jan Maye n Ridge an d i n th e Barent s Sea . The ESP s were acquired with a two-ship tech nique (Stoff a & Buh l 1979 ) suc h tha t th e sho t and receive r were moved apar t a t equa l intervals maintaining th e sam e midpoint . Th e tw o ships .

with on e firin g it s airgun s (6000in. 3 ) whil e th e other wa s recordin g wit h it s multichanne l streamer (2.4k m long) , starte d fro m initia l positions a t th e en d point s o f th e ESP , usuall y between 7 0 an d 100k m apart , steamin g t o th e opposite en d points and crossing the midpoint at a separatio n o f abou t 1 km. I n thi s manner, tw o ESPs ar e collecte d i n on e experiment , on e a t closing range s an d anothe r a t expandin g ranges . Asymmetry whe n comparin g th e lef t an d righ t side o f th e ES P reflect s two-dimensiona l effect s related t o dip-topography . The dat a wer e processe d b y Institu t Frangai s du Petrol e (IFP ) i n co-operatio n wit h El f Aqui taine. Th e trace s wit h commo n source-receive r offsets wer e sorte d an d stacke d t o on e singl e trace. Th e dat a wer e the n transforme d fro m th e distance-travel tim e (x-t) domai n t o th e intercept time-ray parameter (tau-/; ) domain, accord ing t o th e slant-stac k procedur e o f Stoff a e t al . (1981). On e advantag e o f th e ESP s commo n midpoint geometr y i s tha t dippin g interface s have mino r effec t o n determinatio n o f interva l velocities (Diebol d & Stoff a 1981) , i n contras t to sonobuo y dat a fo r whic h commo n receive r geometry yield s a velocit y tha t migh t b e sig nificantly affecte d b y dippin g layer s (Olafsso n et al . 1992) . The seve n norther n Nort h Se a ESP s (Fig . 1) . from 1983-1984 , hav e no t bee n publishe d be fore, althoug h Pine t (1989 ) referre d t o th e IFP-Elf interpretation . Bot h distance-trave l time (x-t. Fig . 5a ) an d intercep t time-ra y parameter (tau-/; . Fig. 6a ) domain presentation s were availabl e fo r al l profiles . T o extrac t th e maximum velocity-dept h information , a dat a reduction procedur e simila r t o tha t o f Eldhol m & Mutte r (1986 ) an d Jackso n e t al . (1990 ) wa s used wit h thre e differen t method s applie d i n the x- t a s wel l a s on e i n th e tau-/ ? domain . Th e travel tim e computation s wer e carried ou t usin g in-house compute r softwar e develope d b y Kit terod (1986 ) and Hegn a (1989) . Th e interpreta tion procedur e involve d th e followin g steps : • Al l seismi c arrival s identifie d a s refracte d events wer e firs t analyse d wit h a slope intercept metho d (Ewin g 1963). The assump tions o f thi s reductio n techniqu e ar e plan e layers an d a constan t velocit y withi n each . If thes e constraint s ar e met , th e dat a wil l consist o f straight-lin e segments where hea d or criticall y refracted wave s are produced . • Next , th e seismi c arrival s wer e examine d as divin g wave s an d a flat-Eart h Herglotz Wiechert inversio n wa s performe d (Ak i & Richards 1980) . Th e assumptio n her e i s a continuous velocit y increas e wit h depth :

CRUSTAL STRUCTUR E I N TH E NORTHER N NORT H SE A

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Fig. 5 . ES P 5 1 in x-t domain , (a ) Data example ; (b ) synthetic ray-tracing model . (Se e Fig 1 for locatio n of ES P 51.) thus, ra y path s wer e curve d s o tha t o n th e time v. distance plot the seismic arrivals form a curv e also . • A wide-angle reflection preogram based on the t2-x2 procedure of LePichon e t al. (1968) was used fo r determinin g interval velocities. The near-vertical deep seismic reflection dat a

•

were use d t o identif y wide-angl e reflection s in th e ES P data. Finally, a tau-sum interpretatio n (Diebol d & Stoffa 1981 ) wa s carrie d ou t o n al l ESP s using th e primary tau-p curv e as input. Thi s is equivalen t t o th e standar d first-arriva l slope-intercept metho d except that it is valid

Fig. 6 . ES P 5 1 i n ta n p domain , (a ) Dat a example ; (b ) syntheti c ray-tracin g model . (Se e Fig . 1 tor locatio n o f ES P 51. )

CRUSTAL STRUCTUR E I N TH E NORTHER N NORT H SE A for al l critical an d post-critica l arrival s without distinctio n betwee n reflection s an d refractions. The velocity-dept h curve s derive d fro m th e various techniques are similar (Fig. 7) . However, the slope-intercep t an d t 2-x2 solution s provid e a step-shape d roughe r curv e tha n th e others . Computed (synthetic ) travel time curves (Figs 5b and 6b ) were produced b y ra y tracin g t o exam ine th e accurac y o f the inversions . After severa l iterations w e arrived a t a bes t fit final velocitydepth functio n (Fig . 7) . Th e interpretatio n procedure i s summarized i n Fig . 8 and th e final velocity-depth curve s ar e plotte d o n a depth converted Transec t 1 in Fig. 9 . The result s of the analyses o f al l ESP s hav e bee n presente d b y Christiansson e t al. (1999a), wh o als o discussed

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the vertica l an d latera l velocit y distributio n along Transec t 1 in mor e detail . Th e reductio n of th e ESP s i n combinatio n wit h dee p seismi c reflection dat a ha s improve d ou r knowledg e o f the crustal velocity distribution across the northern North Sea, especially for the mid- and lowercrustal levels . Wide-angl e reflectio n hyperbola s from th e lower-crusta l interface s an d possibl y the Moh o adde d mor e constraint s t o th e dee p crustal configuration. The velocity field used in the depth conversio n of the crustal transect s wa s constructed fro m the ESP result s an d additiona l velocit y dat a suc h as informatio n fro m well s an d stackin g veloc ities. Typica l velocitie s withi n th e sedimentar y sequences are : Tertiar y 2.0-2. 5 km s"1; Cretac eous 3.0-4. 1 kms^1 ; Jurassi c 3.3—4.6kms~ 1 ; pre-Jurassic 4.0-5. 8 km s"1. Th e uppermos t

Fig. 7. Velocity-dept h curves from the analysis of ESP 51. The velocity-depth curves are plotted agains t two-wa y time fo r direc t comparison wit h th e dee p seismi c panel.

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Fig. 8 . Interpretatio n o f deep seismi c refraction an d wide-angl e reflection data, (a) ESP 50 data example ; (b ) dee p seismic reflectio n data clos e t o ES P 50 ; (c) velocity-depth curve s fro m th e analysi s of ES P 50 ; (d) synthetic ray-tracing model . (Se e Fig. 1 for locatio n o f ES P 50.)

Fig. 9. Th e final velocity-depth solutio n for all ESPs plotted o n a depth converted Transect 1 . Velocities in the range 5.8-6.1 km s~! are interpreted a s typical for the uppe r part o f the crystallin e basement. The high-velocit y lower-crusta l bod y (8.1-8.4kms M) detected i n ESP s 51 , 52 and 5 3 beneath th e Hord a Platfor m i s outlined.

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Fig. 10 . Crustal-scal e gravit y models o f Transect s 1 and 2 . The geometrie s an d densitie s are constraine d b y th e deep seismi c reflectio n and refractio n data . crystalline crus t show s velocitie s rangin g fro m 5.Skins" 1 i n area s wit h shallo w basement , e.g. eas t o f th e 0ygarde n Faul t Zone , t o 6.1 6.2kms' 1 i n mor e centra l area s (Fig . 9) , which is i n accordanc e wit h typica l basemen t veloc ities fro m mor e globa l studie s (Christense n & Mooney 1995) . The to p o f the crystalline crust is

not alway s eas y t o interpre t o n th e seismi c reflection data , especiall y i n th e deepe r part s o f the Vikin g Graben, whic h ar e infille d wit h sedi ments o f 12k m thicknes s a t th e most . A typica l velocity fo r th e uppe r transparen t crystallin e crust i s 6.4-6.5kms~ 1 . Highe r value s withi n the middle-lowe r crus t ar e observe d unde r th e

CRUSTAL STRUCTUR E I N TH E NORTHER N NORT H SE A eastern par t o f th e Hord a Platfor m i n ES P 53 , where wide-angl e reflection s giv e a P-wav e velocity o f 6.8kms~ 1 fo r th e dept h interva l between 1 3 an d 29km . A simila r velocit y wa s reported b y Deeme r & Hurich (1991 ) base d o n wide-aperture (WA ) recordin g a t lan d station s during acquisitio n o f th e dee p seismi c reflectio n line ILP-1 1 (profil e 5 i n Fig . 1) . Lower-crusta l velocities o f 6.6-7. 0 km s"1 wer e als o derive d from a n extensiv e ocean-botto m seismograp h (OBS) refractio n experimen t conducte d o n a 150km profil e alon g th e Sognefjor d (Fig . 1 ) (Iwazaki e t al 1994) . The ESPs located o n the Horda Platform sho w strong wide-angl e reflection s correspondin g t o the tw o Moh o candidate s i n th e dee p seismi c reflection data , confinin g a bod y wit h interva l velocities rangin g betwee n 8. 1 an d S^kms^ 1 (Fig. 9) . This high-velocity , lens-shaped bod y i s c. 100km wide and 5- 1 Okm thick , with the bas e of th e bod y locate d i n th e southeaster n par t o f Transect 1 at depth s of 30-35 km. Firs t arrival s from thi s bod y hav e no t bee n obtaine d becaus e of limitation s i n offse t o f th e ESPs . I n th e W A data, however , firs t arrival s fro m thi s bod y ar e recorded a t offset s > 130-140 km giving a velocity o f S.l-S^kms- 1 (Deeme r & Hurich 1991) . Velocities of 8-hkms" 1 ar e normall y assigne d to mantl e rocks , pointin g t o th e shallo w Moh o candidate i n th e dee p seismi c reflectio n dat a a s the bas e o f th e crust . However , a s wil l b e discussed below , thi s is not necessaril y th e case .

Gravity and magnetic data In additio n t o th e seismi c dat a w e hav e use d gravity an d magneti c dat a t o constrai n th e crustal structur e alon g Transect s 1 and 2 . Bot h regional an d residua l anomalie s hav e bee n use d in th e gravit y an d magneti c modelling , an d estimates o f dept h t o magneti c basemen t sho w generally goo d agreemen t wit h the seismi c data . Based o n th e seismi c constraints, a crustal-scale polygon mode l fo r eac h transec t wa s con structed. Densities wer e assigned t o al l polygons (Fig. 10 ) accordin g t o a standar d velocity density relatio n (Barto n & Woo d 1984 ; Chris tensen & Mooney 1995) . The observe d Bougue r gravit y anomalie s along th e tw o transect s (Fig . 10 ) reflec t th e large-scale structure s associated wit h crusta l thinning. Th e regiona l fiel d show s a positiv e gravity anomal y a t th e centr e o f th e Vikin g Graben, an effec t cause d b y the elevatio n o f th e Moho an d the denser upper mantle . The shallow basement area s wes t o f th e Vikin g Grabe n an d east of the 0ygarden Fault Zone are also show n as gravit y highs.

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The gravit y respons e o f th e crustal-scal e models alon g Transect s 1 and 2 wa s calculate d and compared with the observed regiona l anom alies (Fig . 10) . The gravit y model o f Transec t 1 shows clearl y tha t bes t fi t wit h th e observe d gravity anomaly i s achieved whe n the base of the crust i s place d a t th e bas e o f th e high-velocit y body, giving it a bulk density of 2.95 gcm~3 . The calculated gravit y response show s a deviation o f c. 40mgal i f w e assig n mantl e densitie s t o th e high-velocity body . Simila r result s wer e obtained fo r Transec t 2 (Fig . 10) . Thus , i t i s necessary t o explai n th e compositio n an d origi n of a roc k bod y characterize d b y mantl e velo cities but densitie s intermediate betwee n norma l crust an d mantle , befor e we can decid e where t o put th e crust-mantl e boundar y beneat h th e Horda Platform . W e wil l retur n t o thi s below , when al l result s fro m th e dat a analysi s ar e integrated. The density contrast betwee n crust and mantl e ranges fro m 0.4 7 t o 0.50gcm~ 3 alon g bot h transects, excep t fo r th e lower-crusta l high velocity body , whic h has a crust-mantle density contrast of 0.25 gem"3. We also had to introduce a slightl y higher basemen t densit y (2.73gcm~ 3 ) in th e wester n part s o f th e tw o transect s t o obtain th e bes t fi t betwee n th e observe d an d calculated gravit y anomalies. Unde r th e centra l part o f Transec t 2 wher e thinnin g i s mos t prominent, bot h gravit y an d magneti c dat a indicate a dens e an d magneti c bod y a t lower crustal leve l wit h a densit y o f 3.0gcirr 3. This high-densit y bod y ma y represen t mag matic rock s intrudin g an d underplatin g th e crust beneat h th e souther n Vikin g Grabe n i n Jurassic tim e (Dixo n e t al . 1981 ; Underbil l & Partington 1993) . Residual gravit y and magneti c anomalies als o indicate a Palaeozoic basi n on the East Shetlan d Platform, whic h is given a density of 2.45gcm~ 3 for th e upper par t an d 2.55gcm~ 3 fo r th e lower part. Recen t result s from seismi c reflection dat a on th e Eas t Shetlan d Platfor m jus t sout h o f Transect 2 sho w a dee p sedimentar y basi n of possible Palaeozoi c age (Holloway e t al. 1991 ; Platt 1995) . Gravit y modellin g an d magneti c depth estimate s suppor t th e interpretatio n o f this Palaeozoi c basin , whic h ha s als o bee n confirmed b y drillin g (Johnso n e t al . 1993 ; Duncan & Buxton 1995) . Crustal structur e Based o n th e integrate d analysi s o f al l geophy sical dat a we hav e constructe d the crusta l transects i n Fig . 11 . I n thi s pape r w e wan t t o

Fig. 11 . Fina l crusta l model s fo r Transect s 1 and 2 , base d o n integratio n o f al l availabl e geophysica l an d geologica l data .

CRUSTAL STRUCTUR E I N TH E NORTHER N NORT H SE A focus o n th e geometr y o f crusta l thinnin g an d relate tha t t o severa l phases o f post-Caledonia n rifting. However , befor e w e ca n d o tha t w e must decid e wher e th e Moh o i s locate d be neath th e Horda Platfor m an d wher e the to p of the crystallin e basemen t i s locate d beneath'th e Viking Graben .

Moho The Moh o discontinuit y i s define d b y a rathe r abrupt increas e o f th e P-wav e velocit y fro m about Tkms" 1 o r les s t o typicall y i n exces s o f S.Okms" 1 , and i s expected t o b e associated wit h a densit y contrast o f about 0.5gcm~ 3 (Steinhar t 1967). Bot h geologist s an d geophysicist s typi cally regar d th e seismicall y define d Moh o a s the bas e o f th e crust . However , th e Moh o doe s not necessaril y correspon d t o th e petrologica l boundary betwee n crus t an d mantl e (Griffi n & O'Reilly 1987 ; Menge l & Ker n 1992) . Wher e crustal rock s ar e transforme d int o eclogite s a t the bas e o f orogenicall y thickene d crust , th e position of the seismic Moho (defined as the firstorder velocit y discontinuity ) i s no t identica l t o the petrologica l Moh o (define d a s th e bound ary betwee n non-peridotiti c crusta l rock s an d olivine-dominated mantl e rocks) . A s a conse quence, refractio n seismi c field studies ma y no t detect eclogite s a s crusta l rocks . Th e tw o con trasting base s fo r Moh o definitio n hav e le d t o misunderstanding concernin g th e natur e o f th e crust-mantle transition . In th e dee p crus t belo w th e Hord a Platfor m we ar e probabl y facin g thi s problem . ES P dat a along Transect 1 demonstrate velocities o f more than Skms" 1 fo r th e body , whic h i s bounde d by th e two Moh o candidate s i n the deep seismic reflection dat a (Fig . 9) , suggesting that th e bod y is no t a par t o f th e crus t bu t represent s mantl e rocks. Gravit y modelling , o n th e othe r hand , indicates a densit y o f 2.95gcm~ 3 fo r thi s body , intermediate betwee n norma l densitie s for crys talline basement an d mantle. Introducing mantle densities i n th e high-velocit y body give s ris e t o significant discrepancy between the observed an d calculated gravit y anomalie s (Fig . 10) . Further more, result s fro m tectonostratigraphi c model ling ar e no t compatibl e wit h th e shallo w Moh o candidate (Odinse n e t al. 19996) . This high-velocity , high-density body i s interpreted to represent a deep crustal root of partially eclogitized rock s tha t ar e related t o th e onshor e Western Gneiss Regio n followin g the conceptua l model of Andersen et al. (1991) for the Late Cale donian extensiona l collaps e i n western Norway . In their model th e crustal thicknes s was doubled during th e Caledonia n collisio n betwee n th e

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Laurentian an d Balti c plate s i n Mid-Siluria n time, followe d b y decouplin g o f th e therma l boundary laye r an d extensiona l collaps e o f the orogen . Thi s mode l permit s th e presenc e of eclogites within the lowe r crust a s an effec t o f pro- an d retrograd e metamorphis m durin g th e subduction an d eductio n o f th e lowe r crust . These rock s wer e brough t t o th e surfac e in th e Western Gneis s Regio n i n respons e t o rapi d uplift an d exhumation . The y ar e characterize d by hig h velocitie s (u p t o S.Skms" 1 ) an d den sities (u p t o 3.6gem" 3). A lowe r crus t wher e large segment s underwen t partia l eclogitizatio n as a resul t o f flui d infiltratio n an d deformatio n may explai n th e increase d velocitie s observe d toward th e bas e o f the crust . Partial eclogitizatio n lead s t o a mixtur e o f granulite an d eclogit e a s in th e cas e o f th e eclo gite breccia s (Austrhei m 1990 ) and create s roc k bodies wit h intermediat e densitie s betwee n th e protolith an d th e product . Eclogite s foun d a t the surfac e sho w variabl e degree s o f back reaction varyin g fro m ni l t o complet e o n th e scale o f a n outcro p o r eve n o f a thi n section . The variabl e degre e o f back-reactio n i s a ques tion o f flui d availability , rock permeabilit y an d kinetics rathe r tha n a n indicatio n o f differen t uplift path s (Austrheim 1994) . Detailed mappin g of onshor e wester n Norwa y (Austrhei m 1987 , 1990; Austrhei m & M0r k 1988 ; Bound y e t al . 1992) reveal s tha t eclogite s wer e forme d alon g anastomosing shea r zone s up t o 15 0 m thick and continuous ove r severa l kilometres. I n a n are a of mor e tha n 50km 2 th e amoun t o f eclogite s formed wa s estimate d a s betwee n 3 0 an d 40 % (Austrheim 1987) . I n a smalle r are a o f 1km 2 , eclogites accoun t fo r mor e tha n 50 % o f th e exposed rock s (Austrhei m & M0rk 1988) . A predominanc e o f basi c rock s i s foun d i n the Berge n Arc s wherea s th e Wester n Gneis s Region i s muc h mor e heterogeneou s wit h a large amoun t o f quartzo-feldspathi c rocks . Th e gabbro-eclogite relation s studie d s o fa r hav e been limite d t o rathe r smal l bodie s o f a fe w hundred metre s thickness. In extrapolating these observations t o a lowe r crus t o f predominantl y gabbroic composition , however , on e woul d expect t o fin d larg e volume s o f incompletel y reacted rocks . Interna l shea r zone s o r origina l fracture zone s i n suc h massif s shoul d pla y a n important rol e i n metamorphi c reactions . Thu s the physical properties o f such a crust would , t o a large extent, be governed b y the distribution of original zone s o f weaknes s an d th e generatio n of fracture s and shea r zones . Th e field evidence presented b y Austrhei m e t al . (1997 ) demon strates th e importanc e o f fluid s i n promotin g reactions an d henc e controllin g petrophysica l

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properties. Thi s fluid-controlle d eclogitizatio n can resul t i n a partiall y eclogitize d crust , wit h a mixture o f roc k type s an d metamorphi c facies . These observation s suppor t indication s o f compositional heterogeneit y i n th e lowe r crus t suggested b y th e geophysica l data . The seismi c velocit y o f a partl y eclogitize d crust i s difficult t o evaluate . Austrhei m & M0r k (1988) assume d tha t th e velocit y woul d b e a linear functio n o f th e volumetri c rati o o f th e two rock types . However, Austrheim et al. (1997) suggested tha t i t i s possibl e tha t a n abrup t change i n velocit y ma y occu r a t a certai n percentage o f eclogitization , whic h depend s o n seismic wavelength and the spatial distribution of eclogites an d granulite s with depth. I f a marke d increase i n velocit y occur s a t lo w degre e o f eclogitization, th e resul t wil l b e a crus t wit h a high velocit y an d relativel y lo w density . Thi s argument fits well with our geophysica l observa tions and modellin g results of Transects 1 and 2 . Brittan & Warner (1996 ) examined th e impor tance o f heterogeneit y i n biasin g wide-angl e seismic velocit y measurements . Thei r syntheti c seismic modellin g indicate s tha t velocitie s measured fro m wide-angl e experiment s ma y depen d significantly upo n th e geometri c arrangemen t and scal e o f th e heterogeneities . Heterogeneitie s whose correlatio n lengt h i s large r tha n th e dominant wavelengt h o f th e investigatin g wave will produc e th e greates t bia s i n seismi c velo cities; the seismi c wave will alway s tend toward s the fastes t possibl e path . Thus , a t larg e offsets , the inferre d velocit y will b e large r tha n th e tru e average velocit y o f crus t traverse d (Britta n & Warner 1996) . Different geometri c arrangement s of the sam e heterogeneou s compositio n ca n lea d to a wid e rang e o f velocitie s measure d b y th e surface observer . Thes e result s ar e importan t to conside r whe n inferrin g materia l composi tions fro m seismi c velocitie s measured b y wide angle experiments . We thu s interpre t th e bas e o f th e crus t t o b e at o r nea r th e bas e o f th e deepes t reflection s at 30-35 km depth , i n contras t t o mos t o f th e earlier publishe d interpretation s o f dee p seismi c data i n th e norther n Nort h Se a (Beac h 1986 ; Hurich & Kristoffersen 1988 ; Deemer & Huric h 1991; Faerset h e t al . 1995) . O n Transec t 2 n o ESPs ar e available , s o th e high-velocit y bod y from Transec t 1 cannot b e detected b y refraction and wide-angl e reflectio n seismi c methods . Instead, dee p seismi c reflectio n dat a (Fig s 2 and 4 ) an d gravit y modellin g (Fig . 10 ) provid e the necessar y constraint s fo r th e bas e o f th e crust (Fig . 11). Hynes & Snyde r (1991 ) als o discusse d th e potential fo r crusta l rock s belo w th e Moh o i n

the Scottis h Caledonides . Th e rocks betwee n the Moho a t abou t 30k m an d a stron g reflecto r (W; McGear y & Warne r 1985 ) at abou t 50k m have been interprete d a s crustal, but wit h mantle velocities an d densities . Hyne s & Snyde r (1991 ) investigated th e mineralog y an d seismi c char acteristics of such rocks . Thei r calculation s indicated tha t continenta l crus t o f uniforml y mafi c composition coul d hav e seismic properties indistinguishable fro m thos e o f mantl e rocks . If , o n the other hand , th e continental crust containe d a significant proportio n o f more felsi c rock s (a s in the Wester n Gneis s Region ) i t woul d no t ex hibit th e hig h seismi c velocitie s commonl y attributed t o th e uppe r mantle . Our results , supported b y the wor k o f Brittan & Warne r (1996 ) an d Austrhei m e t al . (1997) , show tha t a partially eclogitized lower crust bes t fits the observe d seismi c an d gravit y responses . The non-typica l velocity-densit y relatio n with in th e lower-crusta l bod y ma y b e explaine d b y large-scale anisotrop y relate d t o th e distribution of eclogites . Th e amoun t o f eclogite s an d t o what exten t th e lower-crusta l rock s ar e mor e mafic i n compositio n ar e difficul t t o constrain , making downwar d extrapolation s fro m th e surface distributio n an d compositio n o f eclogite s uncertain. To illustrat e the offshore-onshor e correlatio n of the crustal unit s containing eclogites, w e have constructed tw o regiona l transect s acros s th e northern Nort h Se a an d th e Mor e Basi n continuing 200k m t o th e southeas t ont o th e Norwegian mainlan d (Fig . 12) . The Nort h Se a transect (X-X' ) i s base d o n Transec t 1 an d onshore geolog y modifie d fro m Fosse n (1992) . Its onshor e continuatio n crosses th e Berge n Arcs, th e allochthonou s nappes , intersect s th e Hardangerfjord Shea r Zon e an d finall y end s a t the Balti c Shield (Fig . 12). The crusta l structur e onshore i s based o n work by Sellevoll & Warrick (1971), wh o sho t tw o seismi c profile s acros s southwestern Norway . The y foun d tha t th e Moho discontinuit y was a t depth s o f 28-3 0 km beneath th e coastlin e increasin g eastwar d t o 36-38 km beneat h centra l part s o f souther n Norway. Th e dept h t o th e Moh o wa s deter mined fro m bot h Moh o reflection s (using t 2-.\2) and b y travel times for the refracted wave , which gave a mantl e velocit y of 8.0 5 km s"1. Profile X-X ' (Fig . 12 ) show s ho w th e crus t gradually thin s toward s th e Vikin g Graben , th e depth t o th e bas e o f th e crus t bein g reduce d from c . 40 km t o c . 20 km. I t is important t o not e that th e 8-t-kms" 1 velocitie s are confine d to th e lower-crustal body , indicatin g a differen t com position and/o r amoun t o f eclogite s compare d with th e outcroppin g geolog y onshore .

CRUSTAL STRUCTUR E I N TH E NORTHER N NORT H SE A

33

Fig. 12 . Of f shore-onshore correlation o f majo r crusta l features . Transect 1 is extended 200k m acros s th e Caledonides o f Western Norwa y ont o th e Balti c Shield (X-X'). A simila r profile (Y-Y' ) is constructed t o th e north acros s th e M0re-Tr0ndela g Faul t Zone . (Not e th e relationshi p betwee n the lower-crustal high-velocity body offshor e an d th e Wester n Gneis s Regio n expose d onshore. ) Reference s are give n in figure .

The M0r e transec t (Y-Y' ) i s base d o n a n offshore crustal-scal e transec t fro m Olafsso n et al . (1992) , wit h th e onshor e continuatio n modified fro m Anderse n e t al . (1991) . Profil e Y-Y' (Fig . 12 ) crosse s th e M0r e Basi n an d

intersects th e souther n par t o f th e M0re Trondelag Faul t Zon e a t a n almos t perpendi cular angle . I t continue s int o th e Norwegia n mainland passin g th e Wester n Gneis s Regio n north o f the Devonian basins , the allochthonous

34

P. CHRISTIANSSON , J . I . FALEID E & A . M . BERG E

nappes, th e Laerdal-Gjend e Faul t an d end s a t the Balti c Shiel d i n th e southeast . A high velocity bod y (8.5kms -1 ) wa s detecte d b y Olafsson e t al. (1992 ) beneath th e easter n flan k of the M0r e Basi n i n a simila r geological settin g to ou r norther n Nort h Se a transec t (Fig . 12) . However, th e densit y o f thi s bod y ha s no t been teste d b y gravit y modelling . Wit h onl y one ES P i t i s difficul t t o constrai n th e shap e o f this body .

Basement The to p o f th e crystallin e crus t i s identifie d i n parts o f th e stud y are a a s a stron g seismi c reflector, whic h i s correlate d wit h exploratio n well 31/6- 1 drille d int o crystallin e basemen t on th e Hord a Platfor m (Fig . 1) . However, i t i s not possibl e t o trac e thi s reflecto r ove r larg e distances. Well s penetrating crystalline basement are als o know n fro m th e northeaster n par t o f the norther n Nort h Se a (e.g. i n block s 35/ 3 an d 35/9) an d o n th e Shetlan d Platform , wher e about 3 0 U K well s hav e reache d crystallin e basement rock s (Johnso n e t al . 1993) . Further constraint s o n th e to p crystallin e basement interpretatio n ar e base d o n high quality filtere d gravit y an d magneti c data . A fairly goo d matc h betwee n magneti c basemen t depth estimate s an d interpretatio n fro m seis mic dat a i s achieved , excep t i n th e centr e o f the Vikin g Graben , wher e th e interpretatio n i s less constraine d becaus e o f poore r seismi c resolution. Here , th e gravit y an d magneti c filtered residua l anomalie s i n combinatio n wit h the 6+kms" 1 velocitie s i n th e ES P dat a ar e used i n ou r interpretatio n o f to p crystallin e basement. The transect s i n Fig . 1 1 revea l larg e thick nesses o f Uppe r Palaeozoi c rocks . Magneti c depth estimate s alon g th e easter n par t o f Transect 2 coincid e wit h th e deepes t reflecto r interpreted o n lin e SG8043-101 , indicatin g a considerable thicknes s o f sedimentar y strat a (Devonian-?Carboniferous) beneat h th e wedg e of syn-rift (Permian ) origin preserved in the tilted half-grabens. Th e basi n beneat h th e Tertiar y cover of the East Shetlan d Platfor m in Transect 2 probably contain s Uppe r Palaeozoi c rocks . Devonian rock s wer e penetrate d i n U K wel l 3 / 29-2 locate d o n Transec t 2 o n th e terrac e separating th e Eas t Shetlan d Platfor m an d th e Viking Graben (Fig . 11) ; and i n severa l well s o n the Platfor m (Johnso n e t al . 1993) . Hollowa y e l al. (1991 ) published a seismi c lin e coverin g th e East Shetlan d Platfor m jus t sout h o f Transect 2 reflecting a deep (a t least 6 km) Upper Palaeozoi c

basin. Unfortunately , w e canno t differentiat e between Devonia n an d Permia n rock s i n Trans ect 1 as wa s possibl e i n Transect 2 (Fig. 11) .

Crustal thinning and rifting Basement an d Moh o relie f alon g th e transect s (Fig. 11 ) allow s u s t o stud y th e geometr y o f crustal thinning. The basemen t thicknes s reflect s cumulative crustal thinning in response t o several post-Caledonian rif t phases , assumin g constan t crustal thicknes s befor e th e onse t o f crusta l stretching an d thinning . Stretchin g estimate s based o n crusta l thinnin g hav e bee n compare d with estimates based o n other technique s suc h as subsidence analysi s (Christiansso n 2000 ; Fjelds kaar e t al. 1999 ) or forwar d tectonostratigraphi c modelling (Odinsen et al. 19996) . The crusta l thickness varies from c . 30 km an d 36 km belo w th e platform s o n th e grabe n flank s to 21-24 km below the axis of the Jurassic Viking Graben (Fig . 11) . Th e transect s sho w maximu m basin dept h wher e crystallin e basement i s thin nest an d Moh o shallowest , a s predicte d b y th e McKenzie model . Furthermore , th e syn - an d post-rift basi n fil l ar e superimposed , suggest ing tha t th e crusta l thinnin g an d lithospher e temperature fiel d perturbation s wer e laterall y coincident. A basement relie f of > 10 km is indicated along Transect 1 (Fig. 11) . Th e crus t i s thickest i n th e east, clos e t o th e Norwegian coast , an d become s considerably thinne r west of the 0ygarden Faul t Zone. Th e Hord a Platfor m wa s relativel y littl e affected b y th e Lat e Jurassi c rifting , s o mos t o f this thinnin g mus t b e relate d t o Permia n an d earlier (Devonian? ) stretching . Thus , th e Per mian riftin g affecte d a wider area i n the norther n North Se a than did the Jurassic event. Transect 1 clearly show s tha t th e Permia n stretchin g mus t have bee n extensive , affectin g a n are a fro m th e 0ygarden Faul t Comple x t o the Hutton Faul t in the Eas t Shetlan d Basi n (Fig s 4 an d 11) . Thi s is supported b y correlatio n acros s pre-Jurassi c faults wit h mor e tha n 3 km o f vertica l displacement in some cases, and a basin much wider than the presen t Vikin g Grabe n underlai n b y a thinned crystallin e crust . However , whe n esti mating th e widt h o f th e Permia n rif t w e mus t correct fo r th e crusta l stretchin g relate d t o th e younger (Lat e Jurassic) event. Crustal thinnin g along Transec t 2 follow s th e same patter n a s fo r Transec t 1 , wit h maximu m thinning o f th e crus t unde r th e presen t Vikin g Graben (Fig . 11) . However , th e pre-Jurassi c basin geometr y i s her e uncertain . Ther e i s hardly roo m fo r bot h a thic k Permo-Triassi c

CRUSTAL STRUCTUR E I N TH E NORTHER N NORT H SE A and Devonian-?Carboniferous sequence beneath the deeply burie d Jurassi c strata . A thick Devo nian sequenc e i s probably presen t i n vie w o f it s existence bot h o n th e Eas t Shetlan d Platfor m in th e wes t an d th e Hord a Platfor m i n th e east. I f th e dee p basi n fil l consist s mainl y o f Devonian-?Carboniferous strata , th e main par t

35

of th e Permia n rif t basi n i s located furthe r east , on th e Hord a Platform . However , Permia n movements o n th e 'Wester n Margi n Fault ' (Figs 4 and 11 ) cannot b e ruled out . Another pronounce d differenc e i n crusta l structure betwee n the tw o transect s (Fig . 11 ) i s the shif t i n polarit y o f som e o f th e Lat e

Fig. 13 . (3 curves fo r crusta l Transects 1 and 2 . 1 , From crusta l thinnin g assuming an initia l crustal thicknes s of 36km; 2 , from revers e modelling (Transec t 1 Fjeldskaar e t al. 1999 ; Transect 2 , Christiansson 2000) ; 3, fro m forwar d modelling (Odinse n e t al . 19996) .

36

P. CHRISTIANSSON , J . I . FALEID E & A . M . BERG E

Palaeozoic rif t structure s beneat h th e easter n Horda Platform . Th e polarit y shif t i s probabl y connected t o transfe r faults running in east-west direction. Furthermore , Lat e Jurassi c riftin g appears t o b e more focuse d o n Transec t 2 compared wit h Transect 1 which gav e ris e t o a very thick sequenc e o f Uppe r Jurassi c strat a i n th e Viking Grabe n aroun d 60° N (Fig s 4 an d 11) . Most faul t geometrie s see m t o b e planar , although som e fault s sho w listri c geometrie s a t depth. Th e Western Margi n Fault , confinin g th e Viking Grabe n t o th e wes t o n Transec t 2 show similarities i n curvatur e an d di p t o th e Hutto n Fault an d it s deeper continuatio n o n Transec t 1 (Figs 4 and 11) . ,/3-curves derive d fro m crusta l thinnin g hav e been calculated fo r the two transects assumin g a constant initia l crustal thicknes s of 36 km before Permian riftin g (Fig . 13) . Ther e are , however , many uncertaintie s related t o thes e curves . Th e uncertainties i n identificatio n o f th e Moh o an d top crystallin e basement hav e bee n discusse d in previous chapters . The Moh o is identified along most o f the transects bu t uncertaintie s exis t with respect t o velocitie s use d i n th e dept h conver sion. To p crystallin e basemen t i s uncertai n particularly beneat h th e Vikin g Graben. W e d o not have control on the distribution of Devonian strata al l along th e transects . Therefore w e used the thicknes s of th e crystallin e part o f th e crust in ou r estimates . Strictl y speakin g thi s i s no t correct becaus e th e Devonian strat a shoul d hav e been considere d a s part o f the crus t involve d in Permian an d younge r stretching . Furthermore , the assumptio n o f a unifor m crus t o f constan t thickness i s questionable . Latera l thicknes s variations wer e probabl y presen t followin g th e extensional collapse of the Caledonides. We used the present-da y thicknes s o f 36k m eas t o f th e 0ygarden Faul t Zon e a s th e initia l crusta l thickness assumin g tha t Permia n an d younge r stretching di d no t affec t thi s area . Thi s i s no t quite true , a s w e know tha t structure s onshor e were reactivate d i n Permia n an d Jurassi c time s (Torsvik e t ai 1992) . All thes e uncertaintie s must b e kep t i n min d when analysing the /3 curves derived from crustal thinning. Despit e th e uncertainties , th e curve s reflect th e cumulativ e effec t o f severa l post Caledonian rif t events . I n Fig . 1 3 we compar e our j3 value s wit h thos e obtaine d b y forwar d modelling (Odinse n e t al. \999b) an d revers e modelling (Christiansson 2000 ; Fjeldskaa r e t al . 1999) of both transects . The curves are similar in shape bu t th e modelle d value s ar e sys tematically lowe r tha n thos e estimate d fro m crustal thinning . A carefu l analysi s o f th e vari ous /3 curve s i s beyon d th e scop e o f thi s paper .

For detail s an d discussion s w e therefore refe r t o the modellin g papers cite d above .

Summary an d conclusions Two regiona l transect s acros s th e norther n North Se a hav e bee n constructe d base d o n a n integrated analysis of deep seismic reflection an d refraction data , a s wel l a s gravit y and magnetic data. Th e shallo w part s ar e base d o n high quality conventiona l seismi c reflectio n dat a calibrated agains t a large number o f exploration wells. Th e ne w an d partl y reprocesse d seismi c data, combine d wit h the other geophysica l data, make possibl e a bette r documentatio n o f th e crustal configuration , such a s th e pre-Jurassi c sediment distribution, basement and Moho relief , and dee p faul t geometry . A lower-crusta l bod y characterize d b y a n 8+kms" 1 velocit y and a n averag e bul k density of 2.95gem" 3 i s presen t beneat h th e Hord a Platform. Thi s bod y probabl y represen t a dee p crustal roo t o f partially eclogitized rocks, which formed i n respons e t o Caledonia n collisio n and crustal thickenin g followed b y extensiona l collapse o f th e orogen . Th e non-typica l velocity density relatio n withi n th e lower-crusta l bod y may b e explaine d b y heterogeneitie s relate d t o the distributio n o f eclogites , whic h cause a bia s in th e seismi c velocitie s measure d fro m wide angle data. The crust-mantle boundar y is located at th e bas e o f thi s bod y an d doe s no t coincid e with th e seismically defined Moh o a t th e top . The geometr y o f crusta l thinnin g reflects th e cumulative effec t o f severa l post-Caledonia n rift phases . Stretchin g estimate s base d o n crus tal thinnin g have bee n compare d wit h estimates based o n othe r technique s suc h a s subsi dence analysis and forwar d tectonostratigraphic modelling. The interpretatio n presented her e woul d no t have bee n possibl e withou t th e integratio n o f various geologica l an d geophysica l dat a fro m both offshor e an d onshor e areas . This wor k wa s funde d b y th e Commissio n o f th e European Union and the Norwegian Research Council in th e framework o f the DGXII - Joul e Programme, sub-programme: Energ y fro m fossi l sources : Hydrocarbons, Integrate d Basin Studie s - Th e Dynamics o f the Norwegia n Margin . Nors k Hydr o ASA , Sag a Petroleum a.s . an d Statoi l a.s . provide d data fo r thi s study. W e ar e especiall y indebte d to J . E . Lie , who reprocessed th e NSD P lines. The paper also benefite d from discussion s with T . Andersen , P. T. Osmundsen, H. Austrheim , S . Plank e and J . Skogseid . The clos e co-operation withi n th e IB S project, particularly wit h R. Gabrielsen , A. N0ttved t and T . Odinsen , is highly

CRUSTAL STRUCTUR E I N TH E NORTHER N NORT H SE A appreciated. W e woul d lik e t o than k N . Chamote Rooke, Y . Kristofferse n an d C . Huric h fo r thoroug h and constructive review s of the manuscript. G . Farro w and S . Thompson improve d th e English .

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BLUNDELL, D . J . 1990 . Seismi c image s o f continenta l lithosphere. Journal o f th e Geological Society, London, 147 , 895-913 . BOUNDY, T . M. , FOUNTAIN , D . M . & AUSTRHEIM , H . 1992. Structura l development an d petrofabric s o f eclogite facie s shea r zones , Berge n Arcs , wester n Norway: implication for deep crustal deformatio n processes. Journal o f Metamorphic Geology, 10 , 127-146. BRITTAN, J . & WARNER , M . 1996 . Wide-angle seismi c velocities i n heterogeneou s crust . Geophysical Journal International, 129 , 269-280 . BRUN, J . P . & TRON , V . 1993 . Developmen t o f th e North Vikin g Graben: inference s from laborator y modelling. Sedimentary Geology, 86, 31-51. CHRISTENSEN, N . I . & MOONEY , W . D . 1995 . Seismic velocity structur e an d compositio n o f th e con tinental crust : a globa l view . Journal o f Geophysical Research, 100 , 9761-9788 . CHRISTIANSSON, P . 2000 . Crustal structure an d basin evolution o f th e Northern North Sea. D r Scien t Thesis, Universit y o f Oslo. , BERGE, A. M . & FALEIDE, J. I. 19990 . Analysis o f Expanded Spread Profiles in the northern North Sea. Departmen t of Geology, Universit y o f Oslo , Internal Publication , 68 . DEEMER, S . J. & HURICH , C . A . 1991 . Comparison o f coincident high-resolutio n wide-apertur e an d CDP profilin g alon g th e southwes t coas t o f Norway. In : Continental Lithosphere: Deep Seismic Reflections. America n Geophysica l Union , Geodynamic Series , 22, 435-442. DIEBOLD, J . B . & STOFFA , P . L . 1981 . Th e trave l equation, tau-; ? mappin g an d th e inversio n o f common midpoin t data . Geophysics, 46, 238-254. DITCHA, E . 1998 . Pre-Tertiary evolution o f th e Stord Basin, northern North Sea. Cand . scient . thesis , University o f Oslo . DIXON, J . E. , FITTON , J . G . & FROST , R . T . C . 1981 . The tectoni c significanc e o f post-Carboniferou s igneous activit y i n th e Nort h Se a Basin . In : Petroleum Geology of the Continental Shelf. Heyden, London , 121-123 . DUNCAN, W . I . & BUXTON , N . W . K . 1995 . Ne w evidence fo r evaporiti c Middl e Devonia n lacus trine sediments with hydrocarbon sourc e potential on the East Shetlan d Platform, North Sea . Journal of th e Geological Society, London, 152 , 251-258 . ELDHOLM, O . & MUTTER , J . C . 1986 . Basi n structur e on th e Norwegia n margi n fro m analysi s o f digitally recorde d sonobuoys . Journal o f Geophysical Research, 91 , 3763-3783. EWING, J. 1963 . Elementary theory of seismic reflection and refraction measurements. In: HILL, M. N. (ed.) The Sea. Wiley-Interscience, New York, 3-19 . FICHLER, C . & HOSPERS , J . 1990 . Dee p crusta l structure o f th e norther n Nort h Vikin g Graben : results fro m dee p reflectio n seismi c an d gravit y data. Tectonophysics, 178 , 241-254 . FJELDSKAAR, W. , JOHANSEN , H. , TE R VOORDE , M. , CHRISTIANSSON, P . & FALEIDE , J . I . 1999 . Modelling o f crusta l thinnin g an d th e effect s o f possible intr a plat e stres s i n th e norther n Vikin g Graben. In : Basin Modelling. AAP G Memoir .

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FLINN, D. 1985 . The Caledonide s o f Shetland. In: GEE , D. G . & STURT , B . A . (eds ) Th e Caledonide Orogen - Scandinavia an d Related Areas. Wiley , Chichester. FOSSEN, H . 1992 . The rol e o f extensiona l tectonics i n the Caledonide s o f sout h Norway . Journal o f Structural Geology, 14 , 1033-1046 . F^ERSETH, R . B. , GABRIELSEN , R. H . & HURICH , C . A . 1995. Influenc e of basemen t i n structurin g o f th e North Se a Basin , offshor e southwes t Norway . Norsk Geologisk Tidsskrift, 75 , 105-119 . GABRIELSEN, R . H . 1986 . Structura l element s i n gra ben system s an d thei r influenc e o n hydrocarbo n trap types . In: SPENCER , A. M . e t al. (eds) Habitat of Hydrocarbons on the Norwegian Continental Shelf. Graha m an d Trotman , London , 55-60 . , F^RSETH , R. , STEEL , R . J. , IDIL , S . & KLOVJAN , O. S . 1990 . Architectural styles of basinfil l i n th e northern Vikin g Graben. In : BLUNDELL . D . J . & GIBBS. A . D . (eds ) Tectonic Evolution o f th e North Sea Rifts. Clarendon , Oxford , 158-179 . GIBBS, A . 1987r/ . Linke d tectonic s o f th e norther n North Se a basin . In: BEAUMONT , C . & TANKARD , A. J . (eds ) Sedimentary Basins an d Basin Formation Mechanisms. Canadia n Societ y of Petroleu m Geologists, Memoir , 12 , 163-171 . 19876. Dee p seismi c profile s i n th e norther n North Sea . In : BROOKS , J . & GLENNIE . K . (eds ) Petroleum Geology of North West Europe. Graham an d Trotman , London , 1025-1028 . GILTNER, J . P . 1987 . Applicatio n o f extensiona l models t o th e norther n Vikin g Graben . Norsk Geologisk Tidsskrift, 67 , 339-352 . GLENNIE, K . W . 1984 . The structura l framewor k an d the pre-Permia n histor y o f th e Nort h Se a Area . ///: GLENNIE , K . W . (ed. ) Introduction t o th e Petroleum Geology o f th e North Sea, Blackwell Scientific, Oxford , 17-39 . GRIFFIN. W . L . & O'REILLY , S . Y . 1987 . I s th e continental Moh o th e crust-mantl e boundary ? Geology, 15 , 241-244. HANCOCK, J. M. 1990 . Cretaceous. In : GLENNIE, K . W . (ed.) Introduction t o th e Petroleum Geology o f th e North Sea. Blackwel l Scientific, Oxford , 61-73 . HARRISON, M . G . 1987 . A n interpretation o f deep seismic reflection profiles in the northern North Sea: implications for extension models. MS c thesis. University o f London , Roya l Hollowa y and Bed ford Ne w College . HARTZ, E . & ANDRESEN , A . 1995 . Caledonia n sol e thrust of East Greenland: a crustal scale Devonia n extensional detachment? Geology, 23 , 637-640. HEGNA. S . 1989 . Kinematiske metoder for bestemmelse av hastigheter fra marin-seismiske data med anvendelse i ostlige deler a v Loppahoyden. Cand . scient. thesis , Universit y of Oslo. HOLLIGER, K . 1987 . Crustal structure o f th e North Sea. Diploma thesis , ET H Zurich . & KLEMPERER , S . 1989 . A compariso n o f th e Moho interprete d fro m gravit y dat a an d fro m deep seismi c reflection dat a i n the northern Nort h Sea. Geophysical Journal, 97 , 247-258 . HOLLOWAY, S. . REAY , D . M. , DONATO , J . A . & BEDDOE-STEPHENS, B . 1991 . Distribution o f gran-

ite and possibl e Devonian sediment s in part o f the East Shetlan d Platform , North Sea . Journal o f the Geological Society, London, 148 , 635-638 . HOSPERS, J . & EDIRIWEERA , K. W . 1988 . Mappin g o f the to p o f th e crystallin e continental crust i n th e Viking Graben area , Nort h Sea . /// : KRISTOFFER SEN. Y . (ed. ) Progress i n Studies o f th e Lithospheres i n Norway. Norge s Geologisk e Under sokelse, Specia l Publication , 3 , 21-28. & 1991 . Dept h an d configuratio n o f th e crystalline basemen t i n th e Vikin g Graben area , northern Nort h Sea . Journal o f th e Geological Society, London, 148 , 261-265 . HOSSACK, J . R . 1984 . The geometr y o f listri c growth faults i n th e Devonia n basin s o f Sunnfjord , wes t Norway. Journal o f th e Geological Society, London, 141 , 629-638. ' HURICH, C . A . 1996 . Kinematic evolution of the lower plate during intracontinental subduction: an example from the Scandinavian Caledonides. Tectonics. 15,1248-1263. & KRISTOFFERSEN , Y . 1988 . Dee p structur e o f the Caledonid e oroge n i n souther n Norway : new evidenc e fro m marin e seismi c reflectio n profiling. In : KRISTOFFERSEN . Y . (ed. ) Progress in Studies o f th e Lithosphere i n Norway. Norge s Geologiske Undersokelse . Specia l Publication . 3 . 96-101. HYNES, A . & SNYDER . D . B . 1991 . Deep-crusta l mineral assemblage s an d potentia l fo r crusta l rocks belo w th e Moh o i n th e Scottis h Caledo nides. Geophysical Journal International. 123 . 323-339. IWAZAKI. T. , SELLEVOLL . M . A. . KANAZAWA . T. . VEGGELAND, T. & SHIMAMURAH, H. 1994 . Seismic refraction crusta l stud y alon g th e Sognefjord . south-west Norway , employin g ocean-botto m seismometers. Geophysical Journal International. 119,791-808. JACKSON, H . R. , FALEIDE . J . I . & ELDHOLM . O . 1990 . Crustal structur e o f th e sheare d southwester n Barents Se a continental margin. Marine Geology. 93, 119-146 . JOHNSON, H. , RICHARDS , P . C. , LONG . D. & GRAHAM . C. C . 1993 . United Kingdom Offshore Regional Report: the Geology of the Northern North Sea. British Geologica l Survey . HMSO. London . JORDT, H. , FALEIDE , J . L , BJORLYKKE . K . & IBRAHIM . M. T. 1995 . Cenozoic sequenc e stratigraph y o f the Central an d Norther n Nort h Se a basin. Tectoni c development, sedimen t distributio n an d prove nance areas . Marine and Petroleum Geology. 12. 845-879. ,, THYBERG , B . I . & NOTTVEDT , A . 1999 . Cenozoic evolutio n o f th e centra l an d norther n North Se a wit h focu s o n differentia l vertica l movements o f th e basi n floo r an d surroundin g clastic sourc e areas . This volume. KITTEROD, N.-O . 1986 . Hastighets variasjoner me d dypet, tolkningsprosedyrer for ESP og sonarboye data. Cand . scient . thesis. Universit y of Oslo . KLEMPERER, S . L . 1988 . Crustal thinnin g and natur e of extension in the norther n North Se a from dee p seismic reflectio n profiling . Tectonics, 1, 803-821.

CRUSTAL STRUCTUR E I N TH E NORTHER N NORT H SE A & HOBBS , R . 1991 . Deep Seismic Reflection Profiles around the British Isles, the BIRPS Atlas. Cambridge Universit y Press, Cambridge . & HURICH, C . A. 1990 . Lithospheri c structur e o f the Nort h Se a fro m dee p seismi c reflectio n profiling. In : BLUNDELL , D . J . & GIBBS , A . (eds ) Tectonic Evolution o f th e North Se a Rifts. Oxfor d University Press, Oxford , 37-6 3 & WHITE , N . 1989 . Coaxia l stretchin g o r litho spheric simpl e shea r i n th e Nort h Sea ? Evidenc e from dee p seismi c profilin g an d subsidence . In : TANKARD, A . J . & BALKWILL , H . (eds ) Extensional Tectonics and Stratigraphy of the North Atlantic Margins. America n Associatio n o f Pet roleum Geologists , Memoir , 46 , 511-522. KUSZNIR, N . J. & MATTHEWS, D. H. 1988 . Deep seismi c reflections an d th e deformationa l mechanis m o f the continenta l lithosphere . Journal o f Petrology, Special Lithosphere Issue , 66-87. & ZIEGLER , P . A . 1992 . Th e mechanis m o f continental extensio n an d sedimentar y basin for mation: a simpl e shear/pur e shea r flexural cantilever model . Tectonophysics, 215 , 117-131 . , MARSDEN , G . & EGAN , S . S . 1991 . A flexura l cantilever simpl e shear/pur e shea r mode l o f continental lithospher e extension : applicatio n t o the Jeann e d'Ar c basi n an d Vikin g Graben . In : ROBERTS, A . M. , YIELDING , G . & FREEMAN , B. (eds) Th e Geometry o f Normal Faults. Geologi cal Society , London , Specia l Publications , 56 , 41-60. , ROBERTS , A . M . & MORLEY , C . 1995 . Forwar d and revers e modelling o f rif t basi n formation . In: LAMBIASE, J . (ed. ) Hydrocarbon Habitat i n Rift Basins. Geologica l Society , London , Specia l Pub lications, 80 , 33-56. LEPICHON, X. , EWING , J . & HOUTZ , R . E . 1968 . Deep-sea sedimen t velocit y determinatio n mad e while reflectio n profiling . Journal o f Geophysical Research, 73, 2597-2614. LERVIK, K . S. , SPENCER , A . M . & WARRINGTON , G . 1989. Outline of Triassic stratigraphy and structur e in central an d northern Nort h Sea. In: COLLINSON, J. D. (ed.) Correlation in Hydrocarbon Exploration. Graham and Trotman, London, 173-189 . McGEARY, S . & WARNER , M . 1985 . Seismic profilin g the continental lithosphere . Nature, 317, 795-797 . MCKENZIE, D . P . 1978 . Som e remark s o n th e devel opment o f sedimentar y basins . Earth an d Planetary Science Letters, 40, 25-40. MENGEL, K . & KERN , H . 1992 . Evolutio n o f th e petrological and seismic Moho - implication s for the continenta l crust-mantl e boundary . Terra Nova, 4 , 109-116 . NORTON, M . 1986 . Lat e Caledonia n extensio n i n western Norway : a respons e t o extrem e crusta l thickening. Tectonics, 5, 195-204 . 1987. The Nordfjord-Sogn Detachment , W . Nor way. Norsk Geologisk Tidsskrift, 67 , 93-106. N0TTVEDT, A. , GABRIELSEN , R . H . & STEEL , R . J . 1995. Tectonostratigraphy an d sedimentar y architecture o f rif t basins ; with reference to th e north ern Nort h Sea . Marine an d Petroleum Geology, 12, 881-901.

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ODINSEN, T. , CHRISTIANSSON , P., GABRIELSEN , R . H. , FALEIDE, J . I . & BERGE, A . M . 19990 . The geom etries an d dee p structur e o f th e norther n Nort h Sea rif t system . This volume. , REEMST , P. , VA N DER BEEK, P. , FALEIDE , J . I . & GABRIELSEN, R . H . 1999/7 . Permo-Triassi c an d Jurassic extensio n i n th e norther n Nort h Sea : results fro m tectonostratigraphi c forwar d model ling. This volume. OLAFSSON, I. , SUNDVOR, E., ELDHOLM, O. & GRUE, K . 1992. M0r e Margin : crusta l structure s fro m analysis of expande d sprea d profiles . Marine Geophysical Research, 14 , 137-162 . OSMUNDSEN, P . T . 1995 . Late-orogenic structural geology and Devonian basin formation in western Norway: a study from the hanging wall of the Nordfjord-Sogn detachment in the Sunnfjord region. Dr . scient . thesis, Universit y of Oslo . PINET, B . 1989 . Dee p seismi c profilin g an d sedimen tary basins . Bulletin d e l a Societe Geologique d e France, 8, 749-766. PLATT, N . H . 1995 . Structure s an d tectonic s o f th e northern Nort h Sea : ne w insight s fro m dee p penetration regiona l seismi c data . In : LAMBIASE , J. J . (ed. ) Hydrocarbon Habitat i n Rift Basins. Geological Society, London, Special Publications, 80, 103-113 . RESTON, T . J . 1990 . Shea r i n th e lowe r crus t durin g extension, not s o pure and simple. Tectonophysics, 173, 175-183 . SCOTT, D . L . & ROSENDAHL, B. R. 1989 . North Viking Graben: a n Eas t Africa n perspective . AAPG Bulletin, 73 , 155-165 . SELLEVOLL, M . & WARRICK , R . E . 1971 . A refractio n study of the crustal structure in southern Norway. Bulletin of the Seismological Society of America, 61, 457-471. SERANNE, M . 1992 . Devonia n extensiona l tectonic s versus Carboniferou s inversio n i n th e norther n Orcadian basin . Journal of th e Geological Society, London, 149 , 27-37 . & SEGURET , M . 1987 . Th e Devonia n basin s o f western Norway : tectonic s an d kinematic s o f a n extending crust . In: COWARD, M . P. , DEWEY , J . F . & HANCOCK , P . L . (eds ) Continental Extensional Tectonics. Geologica l Society , London , Specia l Publications, 28 , 537-548. STEEL, R . J. 1993 . Triassic-Jurassic megasequence stra tigraphy in the northern Nort h Sea: rift to post-rif t evolution. In : PARKER, J. R . (ed. ) Petroleum Geology of Northwest Europe. Proceedings of the 4th Conference. Geologica l Society , London, 299-315 . & RYSETH , A . 1990 . The Triassic-earl y Jurassi c succession i n th e norther n Nort h Sea : megase quence stratigraph y an d intra-Triassi c tectonics . In: HARDMAN , R . F . P . & BROOKS , J . (eds ) Tectonic Events Responsible for Britain's Oil and Gas Reserves. Geologica l Society , London , Spe cial Publications , 55 , 139-168 . , SIEDLECKA , A . & ROBERTS , D . 1985 . The Ol d Red Sandston e basin s o f Norwa y an d thei r deformation: a review . In : GEE, D. G . & STURT , B. A . (eds) Th e Caledonide Orogen - Scandinavia and Related Areas. Wiley , New York , 293-315 .

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The geometries and deep structure of the northern Nort h Sea rif t system TORE ODINSEN, 1'4 PETE R CHRISTIANSSON, 2'5 RO Y H . GABRIELSEN, JAN ING E FALEIDE 1

2

& ANKE R M . BERGE

1

3

Department of Geology, University of Bergen, Allegaten 41, N-5007 Bergen, Norway 2

Department of Geology, University of Oslo, P.O. Box 1047 Blindern, N-0316 Oslo, Norway 3

Norsk Hydro Research Centre, N-5020 Bergen, Norway 4

5

Present address: Statoil, N-5020 Bergen, Norway

Present address: Norsk Hydro ASA, P.O Box 200, N-1321 Stabekh, Norway Abstract: Th e enormou s quantit y o f commercia l reflectio n seismi c line s across th e Nort h Sea Basin have made th e area on e of the most thoroughl y studied continental setting s in the world. Furthe r insight in the deep architectur e of the crust is provided by c. 10000km deep reflection seismi c data . Unfortunately , thes e uniqu e database s hav e rarel y bee n combine d systematically to constrain possible tectonic models for the area. This paper i s built on a ful l integration o f high-qualit y commercia l line s (7stwt ) an d th e dee p (15stwt ) NSDP84- 1 and - 2 lines . Th e dee p line s hav e bee n post-stac k reprocesse d an d depth-converted . A numbe r o f dee p well s hav e provide d stratigraphi c contro l alon g th e lines . The overal l reflective pattern i n the lines divides the crust i n three, wit h a reflective upper an d lower crus t separated b y a les s reflectiv e middl e crust . Th e latera l change s i n reflectivit y matche s th e observed variatio n i n crusta l thickness , wher e th e thinnes t crust coincide s wit h th e Viking Graben are a wit h a tota l crusta l thicknes s o f 21-2 4 km, increasin g t o 30-3 6 km i n th e platform areas . Th e lowe r crust i s seen as a n undulatin g 4-1 Okm thic k ban d wit h shallow dipping reflections , with a Moh o tha t consist s o f reflection s with variable lateral thickness and amplitude , rathe r tha n on e single stron g reflection . The structura l analysi s show s tha t the crus t i s cut b y a numbe r o f larg e norma l fault s wit h varying geometries. I t i s assume d that som e o f thes e majo r fault s ar e long-live d feature s roote d i n ol d basemen t grains . The mos t spectacula r norma l fault s develope d durin g th e Permo-earl y Triassi c exten sional phase , bu t wer e ofte n reactivate d durin g th e Jurassi c extensiona l phase , an d wit h continued minor faul t movemen t int o the Cretaceous therma l cooling period. Integration of commercial an d dee p reflectio n seismi c section s show s tha t thre e detachmen t level s ar e present within the crust. These levels, which control changes in fault geometries , are believed to represen t latera l rheologica l interfaces combined wit h or intersecte d b y long-lived zones of weaknesses . Th e uppermos t leve l i s represente d b y supra-basemen t low-angl e norma l faults controlled by gravity and/or lithologica l changes during extension. An intra-basement (middle crust) leve l between 5 and 7 s (twt) coincides with decreasing dip o f the large r basi n bounding faults . Th e lowe r crus t i s th e deepes t detachmen t level , whic h probabl y exert s control o n th e geometri c change s o f th e upper-mantl e shea r zone s an d th e larges t crusta l normal faults .

The structura l framewor k for Mesozoi c t o Ter - nificanc e an d area l extent of the faultin g and it s tiary basi n developmen t i n th e norther n Nort h relatio n t o differentia l subsidence an d sedimen Sea (Fig. 1 ) has becom e increasingl y well docu - tatio n ar e stil l unde r debate . I n particular , th e mented durin g th e las t decad e (McKenzi e 1978 ; pre-Triassi c evolutio n i s debatable, a s th e thic k Sclater et al. 1986; Beach et al. 1987; Giltner 1987 ; sedimentar y layer s observe d withi n th e dee p Badleyetal. 1988 ; Gabrielsen e/#/. 1990; Roberts half-graben s o n reflectio n seismi c line s ar e no t et al. 1990) . Discrete phases of crustal extension, penetrate d b y wells. Data from th e southwestern with fault bloc k rotatio n an d sub-basi n develop- par t o f Norwa y an d sequenc e stratigraph y ment, ar e separate d b y therma l coolin g an d studie s indicat e tha t riftin g starte d i n Permia n broad basina l subsidenc e fro m Permia n t o Cre - tim e (Faerset h 1978 ; Torsvi k e t al . 1992 ; Steel taceous time . However , th e detaile d nature , sig - 1993) . I t i s als o acknowledge d tha t Devonia n

From: N0TTVEDT , A . e t al . (eds ) Dynamics o f th e Norwegian Margin. Geologica l Society , London , Specia l Publications, 167 , 41-57. 1-86239-056-8/OO/ S 15.00 © Th e Geologica l Societ y of Londo n 2000 .

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T. ODINSE N E T AL .

Fig. 1 . Ke y ma p an d mai n structura l elements , norther n Nort h Se a (referenc e leve l i s base Cretaceous) . Area s that are primarily affected b y mid-Jurassic-early Cretaceous stretching in grey. Locations of transects 1 and 2 and semi-regional commercia l line s use d i n thi s paper ar e displayed .

(-Carboniferous?) sediments , whic h ar e wel l known fro m the UK sector , may cover basemen t in the deeper part s o f the northern Nort h Sea , a s suggested b y Platt (1995) , Fserseth e t al (19950 ) and Fsrset h (1996) . For a mor e complet e understandin g o f th e evolution an d interpla y betwee n th e differen t extensional phases , basi n geometr y an d subsi dence history , i t i s essential t o analys e thi s pre Jurassic evolution , whic h also encompasse s Caledonian an d post-Caledonia n structures . Compaction o f Permo-Triassi c basi n fil l an d residual Permo-Triassi c therma l anomalie s ma y enhance Triassi c an d Jurassi c subsidenc e an d may therefor e cause overestimatio n o f stretching estimates derive d fro m subsidenc e analysi s

assuming onl y on e rif t phas e (i n mid-Jurassic early Cretaceou s time) . T o succee d i n thi s evaluation, i t i s essential to perfor m a structural and geometrica l analysi s of sufficien t detai l an d completeness t o provid e a bas e fo r modelling . The presen t wor k i s an attemp t t o provid e thi s database, whic h ha s bee n used i n th e modellin g of tw o crusta l transects . Th e modellin g result s are presente d separatel y (Odinse n e t al . 2000) . To achiev e thi s aim , w e hav e combine d an d correlated th e dee p reflectio n seismi c line s NSDP84-1 an d - 2 ( I S s t w t ) (Fig s 2 a an d 3a ) with conventional (7 s twt) reflection seismi c lines to stud y th e structur e at differen t crusta l levels . This analysi s o f th e dee p crusta l structur e als o allows u s t o recogniz e th e pre-Jurassi c structure

Fig. 2 . (a ) Time-section o f crustal transec t 1 . Boxe s show the location s of figures referred to i n the text , (b ) Depth-converted crusta l transec t 1 . (Fo r location , se e Fig. 1. ) No vertica l exaggeration.

Fig. 3 . (a ) Time-section o f crustal transec t 2 . Boxe s sho w th e location s of figure s referre d t o i n th e text , (b ) Depth-converted crusta l transec t 2 . (For location , se e Fig. I. ) No vertica l exaggeration .

NORTH-SEA GEOMETRIE S AN D DEE P STRUCTUR E and particularl y th e architectur e o f th e lowe r crust. Lines NSDP84-1 and -2 (hereafter referred to a s transect 1 and transect 2 respectively) hav e been post-stac k reprocesse d an d depth-con verted (Fig s 2 b an d 3b) . Intersectin g well s have been use d t o establis h stratigraphi c contro l an d have provided velocit y data in the upper 3- 4 km of th e crust , wherea s analyse s o f expande d spread profile s (ESP ) hav e give n velocitie s a t deeper crusta l level s (Christansson e t al. 2000) . In th e first stage o f the analysi s th e geometr y o f the upper, middl e and lowe r crust, as well as the reflection Moho , hav e been determine d for transects 1 and 2 . Th e secon d stag e o f th e analysi s used th e observation s derive d fro m commercia l reflection seismi c lines.

Extensional models Following th e publicatio n o f dee p seismi c reflection dat a acquire d i n 198 3 an d 198 4 (Beach 1986 ; Beac h e t al . 1987 ; Gibb s 1987 ; Klemperer 1988) , three principal models (Fig. 4) for th e dee p structur e o f th e Nort h Se a Basi n have coexisted . Bot h th e symmetrica l pur e shear model , favoure d fo r th e Vikin g Grabe n by Giltne r (1987 ) and Badle y e t al . (1988) , an d

45

the inhomogeneou s shea r model , propose d b y Klemperer (1988 ) for thi s area , ma y b e see n a s developments o f th e McKenzi e (1978 ) model . The asymmetrica l simpl e shea r model , congru ent wit h th e mode l o f Wernick e (1985) , wa s proposed fo r th e Vikin g Grabe n b y Beac h (1986), Beac h et al. (1987) , Gabrielse n (1989 ) and Scot t & Rosendahl (1989) . Finally, differen t types of delamination models (Lister et al. 1986), which allo w fo r differentia l depth-dependen t extension (Royde n & Kee n 1980) , hav e bee n proposed fo r differen t stage s of the developmen t of th e norther n Nort h Se a b y Cowar d (1986) , Fossen et al. (2000) and Ter Voorde et al. (2000). Extensional models are generally based o n the interpreted geometrie s o f majo r crusta l faults . However, suc h model s als o hav e t o conside r rheology an d th e genera l layerin g o f th e conti nental crus t (Dunba r & Sawye r 1989) . Suc h layering result s in larg e jumps i n yiel d strength at laye r interfaces, and thereb y a vertical alteration o f stronger an d weake r intra-crustal zones . In a simpl e two-layere d mineralogica l model , where felsi c rock s dominat e th e upper crus t an d mafic rock s dominate th e lower crust, one would expect norma l fault s t o detac h a t th e weake r interface betwee n thes e tw o layers . However , a s shown b y Kuszni r & Par k (1987) , dependin g

Fig. 4. Thre e models for continenta l extension (afte r Liste r e t al . 1986 , fig. 1).

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on th e crusta l composition , crusta l thickness , geothermal gradien t an d strai n rate , extensiona l faults ma y als o penetrat e throug h th e lowe r crust. I n a settin g suc h a s th e norther n Nort h Sea, whic h ha s undergon e severa l phase s o f tectonic activity , one would als o expect that pre existing weaknesse s withi n th e crust , suc h a s faults o r intrusions , pla y a majo r rol e durin g fault evolutio n (Dunba r & Sawye r 1989) . Thus, th e uppe r crus t i s deforme d b y brittl e localized faulting , wherea s th e lowe r crus t extends b y distribute d ductil e deformatio n con trolled b y non-Newtonia n power-la w cree p (Kusznir & Park 1987) . Evidenc e fro m expose d lower-crustal rock s an d recen t seismi c reflection studies sho w tha t ductil e deformatio n i s ver y heterogeneous an d localize d (Blundel l e t al. 1989; Blundel l 1990 ; Vilcott e e t a l 1993) . Thi s scale-relationship implie s tha t on e ha s t o dif ferentiate betwee n th e bul k ductil e (pur e shear) and localize d modes o f deformation when describing extensiona l mechanism s o f th e lower crust .

Geological setting The presen t crusta l architectur e i n th e norther n North Se a is generally accepte d t o b e a resul t of Permian-early Triassi c an d mid-Jurassic-earl y Cretaceous stretching , separate d b y therma l subsidence (Eyno n 1981 ; Gabrielse n e t al . 1990; Faerset h et a l 1995# ; Robert s e t a l 1995 ; Faerseth 1996) . I t i s als o anticipate d tha t bot h Precambrian an d Caledonia n structures , a s well as extensiona l collapse o f th e Caledonides , hav e influenced late r extensio n an d crusta l reconfi guration i n th e are a (Fros t 1987 ; Huric h & Kristoffersen 1988 ; Klempere r & Huric h 1990 ; Faerseth e t al \995a). Although th e detail s o f timing , significanc e and latera l exten t o f th e Permo-earl y Triassi c stretching ar e stil l a matte r o f debat e (Giltne r 1987; Gabrielsen e t al. 1990; White 1990 ; Faerseth et al . 1995a ; Robert s e t al . 1995) , recen t dat a suggest tha t thi s rif t phas e wa s mor e significan t than th e mid-Jurassic-earl y Cretaceou s exten sional event . Larg e tilte d faul t block s wit h throws o f th e orde r o f severa l kilometre s formed a 150k m wid e N- S oriente d basi n i n late Palaeozoi c time . Durin g th e therma l sub -

sidence tha t followe d th e rifting , faultin g occurred o n bot h margin s (Stee l & Ryset h 1990) a s a consequenc e o f interactio n o f latera l variations in thermal subsidence , sedimen t load ing, compactio n an d flexur e (Gabrielse n 1986 ; Badley e t al. 1988) . Several studie s hav e empha sized that thermal subsidence was still continuing when the mid-Jurassic rifting started . Hence, this must b e taken into account i n stretching calculations from subsidence analysis of the Cretaceou s post-rift sequenc e (Giltne r 1987 ; Gabrielsen et al. 1990; Roberts ^ al. 1995). The Jurassi c extensio n i n th e norther n Nort h Sea i s well constrained . Rotatio n o f majo r faul t blocks show s tha t riftin g wa s initiate d i n lat e Bajocian-early Bathonia n time , an d terminate d in earl y Ryazania n tim e (Ziegle r 1982 ; Leede r 1983; Badle y e t al . 1988 ; Ratte y & Haywar d 1993; Faerset h e t al . \995b). Th e Permo-earl y Triassic maste r fault s wer e partl y reactivated i n the Jurassic rifting , influencin g th e general structural patter n o f th e entir e basin, an d promotin g segmentation an d subsidenc e wit h opposin g polarities i n som e area s (Faerset h 1996) . Later , during th e rif t clima x i n lates t Jurassi c time , fault activit y wa s concentrate d a t fewe r fault s along th e grabe n margin . A s a result , th e in ternal grabe n relie f became mor e pronounce d a s the syste m developed a matur e grabe n topogra phy wit h platforms , sub-platforms , platfor m marginal high s an d a grabe n featur e sensu stricto, wit h a comple x centr e o f subsidenc e along it s axis . Durin g th e therma l coolin g tha t followed th e rifting , a n earl y to mid-Cretaceou s rapidly subsidin g basi n developed , accompanie d by mino r faul t movemen t alon g som e o f th e master fault s (Gabrielse n 1986) . Late Cretaceou s to Tertiar y modes t subsidenc e affectin g th e northern Nort h Se a i s believe d t o hav e bee n affiliated wit h openin g o f th e Nort h Atlanti c in early Tertiar y times .

General crusta l architecture, northern North Se a In bot h dee p reflectio n seismi c transect s (Fig s 2 and 3) , the geometry o f the upper reflectiv e crust is dominate d b y th e Cenozoic , relativel y flat lying an d unfaulte d post-rif t sequence , whic h

Fig. 5 . Reflectivit y patter n i n th e lowe r crus t fro m reprocesse d NSDP84- 1 an d 2 . (See Figs 2 a an d 3 a fo r location.) (a) Portion belo w Gullfaks faul t bloc k (transec t 1 ) showing a highl y reflectiv e lowe r crust wit h reflecto r sets arranged a s lenticula r zone s an d a discontinuou s Moh o reflectio n (arrows) , (b ) Portion belo w th e Hord a Platform (transec t 2 ) showing a diffus e lowe r crus t reflectivit y patter n an d n o well-define d Moh o (arrows) , (c) Portion belo w th e Hord a Platfor m (transec t 2 ) showing a discret e Moh o reflectio n (arrows ) an d a thi n reflective lowe r crus t above .

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unconformably overlie s the heavily faulted base ment an d it s Permia n t o lat e Jurassi c sedimen tary cove r (Jord t e t al. 1995) . Bot h transect s reveal asymmetrica l large-scal e extensiona l geo metries, wit h th e deepes t part s o f th e basi n situated i n th e westernmos t hangin g wal l o f the inne r wester n maste r faul t o f th e Vikin g Graben. Th e presen t depth-converte d dat a suggest tha t th e faul t geometrie s i n th e uppe r reflective crus t var y wit h positio n withi n th e basin an d tha t mor e tha n on e leve l o f detach ment ma y b e identified . I t i s particularly note d that a loca l detachmen t a t c . 6-7 km (4- 5 s twt) is observe d i n th e reflectio n seismi c data . Th e existence o f thi s supra-basemen t detachmen t on th e wester n shoulde r o f th e grabe n i s als o confirmed b y detaile d studie s o f Gullfak s faul t block (Fosse n 1989 ; Fosse n e t al . 2000) . Fo r the crustal-scal e structures , man y maste r fault s have planar o r slightl y curviplanar geometries in their uppe r parts , bu t see m t o flatte n a t a dept h of 12-1 5 km. It shoul d als o b e note d tha t ther e ar e differ ences betwee n th e tw o transects . Th e majo r easterly thro w a t basemen t level is spread across five major fault s i n the more northerly transect 1 (Fig. 2b) , whereas thi s thro w i s concentrated o n only on e maste r faul t i n the southerl y transect 2 (Fig. 3b) . Bot h transect s clearl y illustrat e pre Jurassic fault movement s as well as fault polarity shift (transec t 2), confirming the existence of preJurassic developmen t o f separat e N- S trendin g graben segment s (Gabrielsen et al. 1990 ; Faerseth et al . 19950) . Although th e middl e crus t i s characterized b y poor reflectivity , structura l observation s o f major significanc e ca n b e mad e i n th e repro cessed dee p seismi c reflectio n data , whe n thes e are combine d wit h an d correlate d wit h com mercial reflection seismi c lines. Thus, i t has bee n possible t o trac e som e maste r fault s int o th e more transparen t middl e crus t and , i n a fe w cases, t o th e to p o f th e muc h mor e reflectiv e lower crust . Althoug h les s clearl y displayed , similar geometrie s ar e als o indicate d belo w th e central par t o f th e mid-Jurassic-earl y Cretac eous Vikin g Graben . Th e fe w strong reflection s that ar e see n i n th e middl e crus t ar e primaril y low-angled. Mor e steepl y dippin g reflection s have bee n mappe d i n th e coast-paralle l dee p reflection seismi c line s by Faerset h et al. (1995a), who relate d thes e t o th e pre-lat e Mesozoi c structuring. The lowe r reflectiv e crus t i s see n a s a n undulating reflectiv e 4-1 0 km thic k ban d wit h primarily gentl y dippin g reflections , bot h i n sections transvers e t o th e strik e o f th e N- S oriented Permo-Triassi c an d Jurassi c basi n axi s

(Figs 2 and 3) , and alon g its strike in the Hord a Platform. Th e to p o f th e lowe r crus t shallow s from mor e tha n 20k m belo w th e platfor m margins t o 17k m beneat h th e presen t Vikin g Graben axis . Thi s reflectiv e zon e i s les s pro nounced belo w th e grabe n axis . I t i s also note d that th e reflection s i n th e lowe r crus t ar e no t always very prominent, an d ca n also appear a s a diffuse pattern . A t a large r scale , th e reflectiv e lower crus t i s undulating i n a long-wavelength low-amplitude mode , i n whic h th e shallowes t part reflect s th e mos t thinne d crust . I n detail , this undulatin g mod e i s repeate d b y reflecto r sets, ofte n arrange d a s lenticula r zones . Thi s is especially pronounced wher e the lowe r crust is most reflectiv e below the Gullfaks area (Fig . 5a). The genera l expressio n o f thes e zone s i s similar to th e anastomosin g shea r band s describe d b y Blundell e t al . (1989 ) and Blundel l (1990), wh o suggested tha t extensio n i n th e uppe r crus t i s accommodated b y norma l faultin g wherea s ductile shea r dominate s withi n the lowe r crust. The reflectiv e Moho i s generall y wel l defined in bot h transects . I n th e area s wher e th e lowe r crust i s more transparent , th e Moho is seen only as a diffus e patter n o f reflection s (Fig . 5b) . Ou r data sho w tha t th e Moh o consist s o f a se t o f reflectors wit h variabl e latera l thicknes s an d amplitude, rathe r tha n on e singl e stron g reflec tor. Thus , th e Moh o reflectio n i s displaye d partly wit h abrup t cut-off s o f th e reflectiv e band (Fig . 5a) , and partl y as a discrete reflection (Fig. 5c) . Below th e Hord a Platfor m th e reflection s o f the lowe r crus t spli t int o tw o separat e band s (Fig. 6) . A dedicated velocit y study has reveale d that th e body delineated by these two reflection s is characterized b y velocities typical for the uppe r mantle (S+kms" 1 ) (Christiansso n e t al . 2000) . However, result s from forwar d modellin g o f th e basin developmen t (Odinse n e t al . 2000 ) an d gravity modellin g (Christiansso n e t al . 2000), a s well a s th e thicknes s an d distributio n o f th e associated sedimentar y accumulatio n indicat e that th e bas e o f th e crus t woul d b e locate d at th e bas e o f th e lowe r reflectiv e band . Acknowledging thes e observations , th e entir e crustal thicknes s i s interprete d t o var y fro m c. 30km an d 36k m belo w th e platform s o n the grabe n flank s t o 21-2 4 km belo w th e axi s of th e bas e o f th e Jurassi c sequenc e i n th e Viking Graben . Beneath th e reflectiv e Moh o intra-mantl e reflectors ar e recorde d i n both transect s (Fig. 7). These features , dippin g 30-45 ° awa y fro m th e graben axis, are focused at the transition between the basin an d th e platform areas, an d follo w th e general patter n fo r intra-mantl e reflection s

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Fig. 6. Portio n o f the lower crust below the Horda Platfor m (see Fig. 2a for location) showin g how the reflection s split int o tw o separat e band s (arrows ) envelopin g a mor e transparen t zone . (Fo r discussion of this pattern , see text.)

Fig. 7 . Portio n o f th e lowe r crust an d uppe r mantl e (see Fig. 2a for location ) showin g intra-mantle reflection s (the tw o lowe r arrows) .

described b y Beach (1986) and Klemperer (1988) . The interpretation o f the mantle structures offer s several alternative solutions, bu t th e coincidenc e between th e dippin g reflection s an d th e di p gradient o f th e Mon o ma y sugges t tha t th e re flections represent mantl e fault s o r shea r zones , similar t o thos e observe d fo r offshor e Britai n in the BIRP S dee p seismi c dat a (McGear y e t al. 1987; Blundel l 1990 ; Reston 1990) .

Large structura l feature s i n the norther n North Se a Although th e relationshi p betwee n th e forma tion o f th e reflectiv e pattern s an d th e tectoni c events in the norther n Nort h Se a is equivocal, a generally hel d assumptio n i s tha t repeate d crustal stretchin g ma y hav e modifie d the reflec tive architectur e substantiall y (Klemperer 1988 ;

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Fig. 8 . Portio n o f transec t 1 (see Fig . 2 a fo r location ) showin g the fault-plan e reflectio n of th e principa l east-dipping faul t i n th e rif t syste m (se t o f arrow s t o th e left) . Part s of th e fault-plan e reflection of th e supra-crustal detachmen t belo w th e Gullfak s faul t bloc k (Dl ) ar e als o see n (se t o f arrows t o th e uppe r right).

Fig. 9 . Portio n o f transect 2 (see Fig. 3 b for location ) showing the fault-plan e reflection (arrows ) of the principal east-dipping faul t i n th e rif t system .

NORTH-SEA GEOMETRIE S AN D DEE P STRUCTUR E Blundell 1990 ; Klemperer & Hurich 1990) . It i s also accepte d tha t th e reflectio n Moh o corre sponds to the refraction Moho, whic h is the base of th e crus t i n th e are a (Barto n 1986) . One o f th e ke y factor s i n understandin g ho w the lowe r crus t deform s i n respons e t o litho spheric stretching , an d thereb y i n constrainin g possible tectonic models for the area, is to stud y the reflectio n from larg e norma l fault s from th e upper crus t an d it s change s wit h depths . I n previous studies , n o dee p line s crossin g th e North Se a Basi n sho w a singl e reflecto r identi fied a s a norma l fault , whic h ca n b e tracke d from th e uppe r crus t dow n t o o r belo w Moh o (Klemperer & White 1989 ; Blundel l 1990) . Ou r approach wa s to loo k specificall y a t th e deepes t penetrating norma l fault s an d t o analys e thei r fault plan e reflection , bot h o n th e commercia l lines an d o n th e dee p reprocessed lines . The geometr y o f th e principa l east-dippin g fault i n th e rif t syste m belo w Gullfak s i s illustrated i n Fig . 8 . Its wester n upper 3.5- 7 km (3-5 s twt) i s interpreted i n transec t 1 using cutoffs o f th e horizon s a s see n i n th e hangin g wall and footwall . From 7 km depth (5 s twt) the fault is seen a s a n almos t continuou s reflectio n down to c . 18km (8 s twt). It s uppe r par t i s character ized by a sub-planar, smoot h faul t plan e dippin g c. 55° (measure d i n depth-converte d section s with n o vertica l exaggeration) . It s dip graduall y decreases, definin g a sub-listri c geometr y whe n entering th e lowe r crust . Whe n th e di p o f th e fault reflectio n becomes les s tha n 10 ° it ca n n o

51

longer b e distinguishe d fro m othe r reflection s within th e lowe r crust . A simila r east-vergen t master faul t o f th e Mesozoi c Vikin g Graben i s observed i n transect 2 . But in contrast to what is seen i n transec t 1 , th e maste r faul t i n transec t 2 can only be tracked continuously fro m the upper crust int o the middle crust a t c . 8 s (twt) (Fig. 9). Assuming th e sam e geometr y fo r the tw o faults, the detachment-dept h i n transec t 2 may also b e c. 20 km. I t shoul d her e b e note d tha t th e dee p stratigraphy an d basemen t leve l i n transec t 2 is more uncertai n tha n i n transec t 1 . The interferenc e betwee n th e east-dippin g master faul t i n transect 1 and th e reflectiv e ban d of th e lowe r crus t ha s bee n a subjec t o f severa l papers (Beac h et al 1987 ; Klemperer 1988 ; Pinet 1989; Brun & Tron 1993) . In this study, the intra lower-crustal reflection s are sometime s found t o be aligne d wit h th e continuatio n o f th e east dipping maste r faul t (Fig . 8) . Althoug h thes e lower-crustal reflection s cannot b e confirmed t o represent a continuation o f the faul t i n the form of shea r zone s dow n t o th e Moho , ther e i s th e possibility tha n th e entir e crus t ma y hav e bee n offset b y th e maste r fault . West-dipping fault s dominat e th e easter n platform o f th e stud y area , especiall y t o th e north (transec t 1 ; Fig. 2b) . Th e mos t prominen t structure o f this set is the 0ygarden Faul t Zon e (Fig. 10) , whic h i s characterize d b y multipl e reactivation, an d whic h define s th e limi t o f th e rift syste m sensu stricto (Gabrielse n 1986 ; Gab rielsen et al. 1995 ; Nottvedt et al. 1995). The faul t

Fig. 10 . Portio n o f th e lin e SG8043-403 (see Fig. 1 for location ) showing the west-dippin g fault-plane, reflectio n (indicated wit h arrows ) o f the 0ygarden Fault.

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Fig.11. Portion of the line SG8043 .01in the Horda Platform ( seeFig.1 for location )showing the low-angle

east-dipping fault-plane reflection (indicated with arrows).

Fig. 12 . Portio n o f th e lin e NVGTI-92-105 i n th e Hord a Platfor m (se e Fig . 1 for location ) showing three west-dipping fault s with different dips . Part s o f th e fault-plan e reflections are visibl e (indicated wit h arrows).

is clearly seen as an almost continuou s reflectio n from 2 to 8 s (twt) depth (Fig. 10) . The upper part of the fault , whic h cuts through th e sedimentary column, typicall y dip s a t 40-50° . I n th e base ment th e di p decrease s t o 20-30° . Th e detach ment leve l of this faul t i s seen t o b e near th e to p of the lower crust. I t is noted that discontinuou s westerly dippin g reflection s ar e see n belo w th e 0ygarden Faul t Comple x i n bot h transect s a t mid-crustal level s (Figs 2 an d 3) . Also, a se t o f east-dipping extensiona l fault s ar e see n i n th e Horda Platfor m are a i n transec t 2 . I n general , the di p o f th e fault s i n th e Hord a Platfor m varies, goin g fro m low-angle , curviplana r t o steep planar (Fig s 1 1 an d 12) . On th e wester n grabe n margin , th e fault s are curviplana r an d listri c wit h detachment s

situated shallowe r tha n th e major crusta l faults , so that th e shallower faults ar e transected by , or merge with , th e uppe r par t o f th e large r faults . This i s particularl y well documente d belo w th e Gullfaks an d Snorr e field s (Fosse n e l al 2000 ) (Figs 8 and 13) . Constraints on extensional model s fo r the northern Nort h Se a Tectonic model s ar e generall y base d o n th e interpreted geometrie s o f majo r norma l faults . The differen t model s proposed fo r th e northern North Se a have therefore le d to a debate o n th e geometric relation s an d th e natur e o f rift s i n general: planar rotationa l (Yieldin g et al. 1991) .

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Fig. 13 . Portio n o f line NVGTI-92-105 across the Snorr e fault bloc k (see Fig. 1 for location ) showin g the fault-plan e reflectio n o f tw o east-dippin g intra-basemen t fault s wit h decreasin g di p a t depth s (indicate d with arrows) .

listric (Beac h e t al 1987 ; Gibb s 1987 ) an d combinations thereo f (Gabrielse n 1986 ; Spec k snijder 1987 ; Gabrielsen e t al . 1995 ) have bee n proposed. An y discussio n an d applicatio n o f such model s ha s t o tak e int o accoun t tha t th e area ha s experience d repeate d extensio n an d crustal reconfiguration , as i s seen i n th e presen t basin configuration , whic h show s considerabl e differences betwee n th e tw o transects . Beneat h the Hord a Platfor m i n transec t 1 , mos t o f th e Permo-early Triassi c maste r fault s di p toward s the mid-Jurassic-earl y Cretaceou s Vikin g Gra ben axis . In transec t 2 , it is seen that th e easter n part o f th e Hord a Platfor m i s influence d b y master fault s tha t di p awa y fro m th e grabe n axis, wherea s th e wester n par t i s characterize d by a serie s o f fault s wit h shiftin g polarities . It i s assume d tha t thi s contras t i n polarit y reflects th e primar y Permo-earl y Triassi c basi n configuration (Faerset h e t al . 19950 ; Faerset h 1996). Thi s ha s been confirme d b y restoration s recently published by Faerseth (1996). The width of th e Permo-earl y Triassi c basi n wa s c . 150km (transects 1 an d 2) . Thi s correspond s t o th e location o f th e are a o f thinne d crus t a s identi fied i n transec t 1 an d wher e modelle d stretch ing estimate s displa y mor e extensio n i n th e Permo-early Triassi c tha n i n th e mid-Jurassic early Cretaceou s (Robert s e t al . 1995 ; Odinse n et a l 2000) . In spit e o f som e difference s i n lower-crusta l reflectivity a s see n i n transect s 1 and 2 beneat h the platfor m areas , th e genera l impressio n i s

that th e signatur e o f th e lowe r crus t i s simila r between th e tw o transects . Thus , transec t 1 is characterized b y eastwar d crusta l thickenin g immediately belo w th e oute r limi t o f the Hord a Platform, wherea s a simila r thickenin g i s found further eas t in transect 2. The two transects show similarities bot h i n th e widt h o f th e are a wit h Moho shallowing , an d th e depth s t o th e Moh o below th e central par t o f the Viking Graben. The intra-mantl e reflection s observe d belo w the Hord a Platform , a s see n particularl y i n transect 1 , hav e a di p angl e o f 30-45° . Suc h features are frequently recorded withi n the upper mantle i n othe r extende d areas , on e exampl e being offshor e Britai n (McGear y e t al . 1987 ; Blundell 1990 ; Resto n 1990) . Althoug h i t ha s been propose d tha t the y may represen t magma , concentration o f fluids , o r diffractio n fro m th e lower crust , th e curren t assumptio n i s that the y are localized mantle faults (Blundell 1990; Klemperer & Hurich 1990 ; Reston 1993) . In transect 1 it i s noted tha t thes e mantl e reflection s coincide with shif t i n th e lower-crusta l reflectivity , whic h is a t th e transitio n betwee n th e thinne r crus t below th e Vikin g Grabe n an d th e thicke r crust below the Horda Platform (Fig. 2). This, as well a s th e observe d geometr y o f th e feature , makes i t reasonabl e t o sugges t a mantl e fault . The coincidenc e betwee n th e locatio n o f th e mantle fault, th e lateral thinning of the crust an d the high-velocit y (8+kms" 1 ) bod y belo w th e Horda Platfor m (transec t 1 ) ma y hav e signifi cance. Althoug h thi s mus t b e subjecte d t o

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further study , this observation ma y indicate tha t the high-velocit y bod y i s a long-live d (Caledo nian?) featur e (Christiansson e t al. 2000) . Figure 1 4 summarize s th e mai n structura l observations fro m th e seismi c reflectio n line s in the area. This conceptual mode l i s based mainl y on transect 1 but also has relevance to transect 2. The model show s that, although th e most o f the faults ar e linea r t o curvilinear , listri c an d mor e complex geometrie s ar e also observed. Th e latter geometries ma y b e a resul t o f pre-existing base ment structure s influencin g renewe d faulting . In th e Hord a Platfor m area , fo r example , i t i s indicated tha t listri c faults are linke d t o o r situ ated abov e stee p plana r norma l fault s (Badle y et al. 1988 ; Gabrielsen 1986) . It is acknowledged that th e shiftin g polarit y o f th e Permo-earl y Triassic (o r older? ) basi n syste m beneat h th e Horda Platfor m impart s geometri c problem s that hav e no t ye t bee n solve d (Fig s 2 an d 3) . The resolution in the present data is not sufficien t to solv e th e proble m o f cross-cuttin g relation s between th e major variably dipping faults . Three level s of detachments (Dl, D2 and D3 ) are present within the crust, as shown in Fig. 14.

The supra-basemen etachment (Dl ) i s found within th e sedimentar y column , exemplifie d within th e Gullfak s faul t bloc k (Fig . 8) , bu t i s also observe d elsewher e i n th e are a (Gabrielse n et al . 1995) . Accordin g t o Fosse n e t ai (2000) , Dl i s situate d directl y abov e basement , an d i s probably linke d to the larger Gullfaks fault com posing a ramp-flat-ramp geometr y a t depth . The intra-basemen t detachmen t leve l (D2 ) i s frequently see n i n th e commercia l line s within the middl e crystallin e crust . Th e di p relation s and th e fac t tha t thes e feature s ca n b e easil y followed alon g strik e sugges t tha t the y ar e no t side-reflections o r nois e fro m processing . The y also generate deeper multiple s in the commercial lines. The D 2 level is similar on bot h side s of the Viking Graben . As state d above , continuatio n betwee n th e master fault s o f th e uppe r crus t an d structure s in th e lowe r crust-uppe r mantl e i s viable . Therefore th e possibilit y of a n easterl y dipping simple shea r zon e exists . Thi s i s supporte d b y recent studie s showin g tha t change s betwee n coupled an d decouple d mode s o f deformatio n are i n best agreemen t wit h the observe d patter n

Fig. 14 . Conceptua l model wit h special reference t o transec t 1 summarizes the mai n structura l observations from th e dee p an d commercia l lines. (Note th e thre e detachment levels (Dl, D 2 an d D3 ) an d th e possibl e lin k between th e mantle faul t an d th e principa l east-dipping normal faul t throug h th e lowe r crust.) Two possibl e scenarios are illustrated for the deformation o f the lower crust during movement of the largest normal fault. Als o displayed ar e stee p planar fault s i n th e footwal l o f th e rif t system . (See tex t for furthe r information.)

NORTH-SEA GEOMETRIE S AN D DEE P STRUCTUR E

55

of subsidenc e (Te r Voord e e t al. 2000) . How - by rheological interface s within the crust, but are ever, a t leas t tw o type s o f contacts betwee n th e probably als o influence d b y pre-existin g loca upper-crustal faul t an d tha t o f the upper mantle lized zone s o f weaknesses , suc h a s fault s an d can b e depicted . I n eithe r case , th e intr a lowe r shear zones. Hence, the fault geometrie s suppor t crust wil l the n b e th e deepes t detachmen t leve l • tha t decoupling occurre d a t mor e that on e level, (D3) a s show n i n Fig . 14 . and tha t th e principal decouplin g surfac e can be The principal flattenin g levels D2 and D 3 can traced int o the lower crust, and possibly throug h easily b e explained a s a consequenc e o f changes the lowe r crust an d int o th e upper mantle . in rheolog y a t depth , wher e flattenin g occur s The deformatio n o f th e uppe r an d middl e within th e weake r intra-crusta l layers . I n a crust i s localize d alon g a numbe r o f norma l simple layere d mode l base d o n verticall y chan - faults, whic h favour s bul k simpl e shear . Th e ging mineralogical composition o f the crust, one geometries o f th e fault s ar e variable , with steep should expec t tha t al l large normal fault s would planar, listri c an d composit e geometrie s bein g detach a t th e weake r interface , fo r example , observed. Th e muc h weake r lowe r crus t i s within th e middl e crust . I n reality , an y crus t i s characterized b y bul k pur e shear , accommo much mor e heterogeneous tha n this, especially a dated b y anastomosin g ductil e shea r zone s continental crus t tha t ha s undergon e repeate d separated b y elongate d roc k bodies . A n east heating an d cooling . On e als o ha s t o tak e int o dipping an d low-angl e ban d o f reflector s i s account th e long-live d zone s o f weaknesses , recorded withi n th e predominantl y transparen t which probabl y hav e a n impac t o n th e faul t upper mantle . Althoug h othe r possibilitie s ar e geometries. I t i s therefor e no t surprisin g tha t evaluated, thes e feature s ar e believe d t o repre some large r listri c faults see m t o detac h within sent fault s o r shea r zones . the middl e crust wherea s other s ente r th e lower crust. I t i s als o notice d (Fig . 14 ) that som e o f This wor k wa s funde d b y th e Commissio n o f th e the intra-mantl e reflection s ar e situate d i n th e European Unio n an d th e Norwegia n Researc h Coun continuation o f large-scal e crustal faults . I f this cil in the framework of Integrated Basi n Studies - Th e o f th e Norwegia n Margin . Nors k Hydr o means tha t thes e ar e crustal-scal e shea r zones , Dynamics a.s., Saga Petroleu m a.s . and Statoi l a.s . provided data this implie s tha t (upper ) crusta l an d mantl e and researc h effort s fo r thi s study . W e ar e especiall y faults ma y b e couple d throug h th e lowe r indebted to J. E. Lie, who reprocessed th e NSDP lines. crust. This can best be explained by a lower crust Reviews an d discussion s b y H . Fossen , a s wel l a s composed o f shea r lense s (Blundel l 1990 ) com- linguistic corrections by H. Brekk e and J. McCrachen , parable wit h lowe r metamorphi c panel s a s dis - are ver y muc h appreciated . W e woul d lik e t o than k played i n the Basin-and-Range province . In thi s H. Hjelmelan d an d a n anonymou s reviewe r for thor model th e lowe r crus t deform s b y bul k pur e ough and very constructive criticism . The authors hav e shear (Allmendinge r e t al. 1987; Hamilton 1987) . honoured thei r suggestions on al l major points . Summary

References

The presen t crusta l architectur e o f the norther n North Se a is the resul t of a long-lasting tectonic history of crustal reconfiguration. Of the several tectonic phase s tha t hav e affecte d th e area , th e Permo-Triassic an d especiall y th e mid-Jurassi c to earl y Cretaceou s extensio n ar e bes t known . The Permo-early Triassic norma l fault s ar e seen to be frequently reactivate d durin g the Jurassicearly Cretaceou s stretching . I t i s also observe d that som e o f th e larg e fault s hav e bee n activ e during the thermal cooling periods that followe d the tw o stretchin g phases. The conceptual model show s that both supra and intra-basemen t detachmen t level s ar e pre sent withi n th e uppe r (Dl) , middl e (D2 ) an d lower (D3) crust, respectively. The Dl leve l may partly b e controlled b y gravity and/or litholog y during th e Jurassic-earl y Cretaceou s phas e (Fossen e t al. 2000). D 2 an d D 3 may b e caused

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types. In : SPENCER , A . M . e t al. (eds ) Habitat o f Hydrocarbons on the Norwegian Continental Shelf. Graham and Trotman, London , 55-60 . 1989. Reactivation of faults on the Norwegian continental shel f and it s implication s for earthquak e occurrence. In : GREGERSEN , S . & BASHAM , P . (eds) Causes and Effects o f Earthquakes a t Passive Margins and in Areas with Post-glacial Rebound on both Sides o f th e North Atlantic. Elsevier, Amsterdam, 69-92 . , F^ERSETH , R . B. , STEEL , R . J. , IDIL . S . & KLOVJAN, O. S . 1990 . Architectural styles of basin fill in the norther n Vikin g Graben. In: BLUNDELL. D. J . & GIBBS , A . D . (eds ) Tectonic Evolution o f the North Se a Rifts. Clarendon , Oxford , 158-179 . , STEEL , R . J . & NOTTVEDT , A . 1995 . Subtl e traps in extensiona l terranes : a mode l wit h referenc e t o the Nort h Sea . Petroleum Geoscience, 1, 223-235 . GIBBS, A . D . 1987 . Deep seismi c profiles in th e north ern Nort h Sea . In : BROOKS , J . & GLENNIE , K . (eds) Petroleum Geology of North West Europe. Graham an d Trotman , London , 1025-1028 . GILTNER, J . P . 1987 . Applicatio n o f extensiona l models t o th e norther n Vikin g Graben . Norsk Geologisk Tidsskrift, 67 , 339-352 . HAMILTON, W . 1987 . Crusta l extensio n i n th e Basi n and Rang e Province , southwester n Unite d States . In: COWARD . M . P. , DEWEY , J . F . & HANCOCK . P. L. (eds ) Continental Extensional Tectonics. Geological Society , London , Specia l Publications . 28, 155-176 . HURICH, C . A . & KRISTOFFERSEN , Y . 1988 . Dee p structure o f th e Caledonid e Oroge n i n souther n Norway: ne w evidenc e fro m marin e seismi c profiling. In : KRISTOFFERSEN , Y . (ed. ) Progress in Studies o f th e Lithosphere i n Norway. Norge s Geologiske Undersokelse . Specia l Publication , 3 . 96-101. JORDT, H. , FALEIDE , J. L , BJORLYKKE , K . & IBRAHIM , M. T. 1995 . Cenozoic sequence stratigraphy of the central an d norther n Nort h Se a Basin : tectoni c development, sedimen t distributio n an d prove nance areas . Marine an d Petroleum Geology. 12 , 845-879. KLEMPERER, S . L . 1988 . Crustal thinnin g and natur e of extension in the northern North Se a from dee p seismic reflection profiling. Tectonics, 7, 803-821. & HURICH , C . A . 1990 . Lithospheri c structur e of the Nort h Se a fro m dee p seismi c profiling . In : BLUNDELL, D . J . & GIBBS , A . D . (eds ) Tectonic Evolution o f th e North Se a Rifts. Clarendon . Oxford, 37-63 . & WHITE , N . J . 1989 . Coaxia l stretchin g o r lithospheric simpl e shea r i n th e Nort h Sea ? Evidence fro m dee p seismi c profilin g an d sub sidence. In : TANKARD , A . J . & BALKWILL . H . R . (eds) Extensional Tectonics an d Stratigraphy o f the North Atlantic Margins. American Associatio n of Petroleu m Geolog y Memoir , 46, 511-522 . KUSZNIR, N . J . & PARK , R . G . 1987 . The extensional strength o f th e continental lithosphere : it s depen dence o n geotherma l gradient , an d crusta l com position an d thickness . In : COWARD , M . P. , DEWEY, J . F . & HANCOCK, P . L . (eds) Continental

NORTH-SEA GEOMETRIE S AN D DEE P STRUCTUR E Extensional Tectonics. Geologica l Society , Lon don, Specia l Publication , 28, 35-52. LEEDER, M . R . 1983 . Lithospheri c stretchin g an d North Se a Jurassi c clasti c sourc e lands . Nature, 305, 510-514 . LISTER, G . S. , ETHERIDGE , M . A . & SYMONDS , P. A . 1986. Detachmen t faultin g an d th e evolutio n o f passive continental margins. Geology, 14, 246-250. MCGEARY, S. , CHEADLE , M . J. , WARNER , M . R . & BLUNDELL, D . J . 1987 . Crusta l structur e o f th e continental shel f aroun d Britai n derive d fro m BIRPS dee p seismi c profiling . In : BROOKS , J . & GLENNIE, K . (eds ) Petroleum Geology o f North West Europe. Graha m an d Trotman , London , 33-41. McKENZiE, D . 1978 . Som e remark s o n th e develop ment o f sedimentar y basins . Earth an d Planetary Science Letters, 40, 25-32. N0TTVEDT, A. , GABRIELSEN , R . H . & STEEL , R . J . 1995. Tectonostratigraph y an d sedimentar y architecture o f rif t basins ; wit h referenc e t o th e northern Nort h Sea . Marine an d Petroleum Geology, 12 , 845-879. ODINSEN, T., REEMST , P. , VANDE R BEEK , P. , FALEIDE , J. I . & GABRIELSEN , R. H . 2000 . Permo-Triassi c and Jurassi c extensio n in the northern Nort h Sea : results fro m tectonostratigraphi c forwar d model ling. This volume. PINET, B . 1989 . Dee p seismi c profilin g an d sedimen tary basins . Bulletin d e l a Societe Geologique d e France, 8, 749-766. PLATT, N . H . 1995 . Structur e an d tectonic s o f th e northern Nort h Sea : ne w insight s fro m deep penetration regiona l seismi c data . In: LAMBIASE , J. J . (ed. ) Hydrocarbon Habitat i n Rift Basins. Geological Society , London, Specia l Publications , 80, 103-113 . RATTEY, R . P . & HAYWARD , A . B . 1993 . Sequenc e stratigraphy o f a faile d rif t system : th e Middl e Jurassic t o Earl y Cretaceou s basi n evolutio n o f the Central an d Norther n Nort h Sea . In : PARKER, J. R . (ed. ) Petroleum Geology o f Northwest Europe: Proceedings o f th e 4t h Conference. Geo logical Society , London , 215-249 . RESTON, T . J . 1990 . Th e lowe r crus t an d th e exten sion o f th e continenta l lithosphere : kinemati c analysis of BIRPS deep seismi c data. Tectonics, 9, 1235-1248. 1993. Evidenc e fo r extensiona l shea r zone s i n th e mantle, offshor e Britain , an d thei r implication s for th e extensio n o f th e continenta l lithosphere . Tectonics, 12 , 492-506. ROBERTS, A. M., YIELDING, G. & BADLEY, M. E . 1990. A kinematic model for th e orthogonal openin g of the Lat e Jurassi c Nort h Se a Rif t System , Den mark-Mid Norway. In: BLUNDELL, D. J. & GIBBS, A. D . (eds ) Tectonic Evolution o f th e North Se a Rifts. Clarendon , Oxford, 180-199 . ,, KUSZNIR , N . J. , WALKER , I . & DORN LOPEZ, D . 1995 . Quantitative analysi s of Triassi c

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extension i n th e norther n Nort h Sea . Journal o f the Geological Society, London, 152 , 15-26 . ROYDEN, L . & Keen, C . E. 1980 . Rifting processes and thermal evolutio n o f th e continenta l margi n o f eastern Canad a determine d fro m subsidenc e curves. Earth an d Planetary Science Letters, 51 , 343-361. SCLATER, J . G. , HELLIGER , S. J. & SHOREY , M. 1986 . An analysis of the importance of extension in accounting for the post-Carboniferous subsidence of th e North Se a basin. Universit y o f Texa s Institute fo r Geophysics, Interna l Report . SCOTT, D . L . & ROSENDAHL, B. R. 1989 . North Viking Graben: a n eas t Africa n perspective . AAPG Bulletin, 73 , 155-165 . SPEKSNJIDER, A. 1987 . The structura l configuration of Cormorant Bloc k I V i n contex t o f th e norther n Viking Graben structural framework . Geologic en Mijnbouw, 65 , 357-379 . STEEL, R . & RYSETH , A . 1990 . Th e Triassic-earl y Jurassic successio n i n th e norther n Nort h Sea : megasequence stratigraph y an d intra-Triassi c tectonics. In : HARDMAN , R . F . P . & BROOKS , J . (eds) Tectonic Events Responsible for Britain's Oi l and Ga s Reserves. Geologica l Society , London , Special Publications , 55, 139-168 . 1993. Triassic-Jurassic megasequence stratigraphy in th e Norther n Nort h Sea : rif t t o post-rif t evolution. In: PARKER, J. R . (ed. ) Petroleum Geology of Northwest Europe: Proceedings of the 4th Conference. Geologica l Society, London , 299-316 . TER VOORDE, M., F^RSETH, R. B. , GABRIELSEN, R. H . & CLOETINGH , S . A. P . L . 2000 . Repeate d litho spheric extensio n i n the norther n Vikin g Graben: a couple d o r a decouple d rheology ? This volume. TORSVIK, T . H. , STURT , B . A. , SWENSSON , E. , ANDERSEN, T . B . & DEWEY , J . F . 1992 . Palaeo magnetic datin g o f faul t rocks : evidenc e fo r Permian an d Mesozoi c movement s an d brittl e deformation alon g th e Dalsfjor d Fault , wes t Norway. Geophysical Journal International, 109 , 565-580. VILCOTTE, J . P. , MELOSH , J., SASSI, W. & RANALLI , G . 1993. Lithospher e rheolog y an d sedimentar y basins. Tectonophysics, 226 , 89-95 . WERNICKE, B . 1985 . Uniform-sens e norma l simple shear o f th e continenta l lithosphere . Canadian Journal o f Earth Sciences, 22 , 108-125 . WHITE, N. J. 1990 . Does the uniform stretchin g mode l work i n th e Nort h Sea ? In : BLUNDELL , D. J . & GIBBS, A. D . (eds ) Tectonic Evolution of th e North Sea Rifts. Clarendon , Oxford , 217-240 . YIELDING, G. , BADLEY , M . E . & FREEMAN , B . 1991 . Seismic reflection s fro m norma l fault s i n th e northern Nort h Sea . In : ROBERTS , A . M. , YIELD ING, G . & FREEMAN , B . (eds ) Th e Geometry o f Normal Faults. Geologica l Society , London , Special Publications , 56 , 79-89. ZIEGLER, P . A . 1982 . Geological Atlas o f Western an d Central Europe. Shel l Internationale , The Hague .

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Repeated lithosphere extension in the northern Vikin g Graben : a couple d o r a decoupled rheology ? M. TE R VOORDE, 1 R . B . F^RSETH, 23 R. H . GABRIELSEN 3 & S . A . P . L . CLOETINGH 1 1

Institute of Earth Sciences, De Boelelaan 1085, 1081 HV Amsterdam, the Netherlands 2 Norsk Hydro, Exploration, P.O. Box 200, N-1321 Stabekk, Norway 3 Department of Geology, University of Bergen, Allegaten 41, N-5007 Bergen, Norway Abstract: Th e Lat e Permian-Earl y Triassi c an d Jurassic-Cretaceou s rif t event s that influ enced th e structur e o f th e norther n Vikin g Grabe n ar e examined . Here , a 2 D forwar d numerical model including faults i s applied o n seismi c line NVGT88-04, whic h crosses th e basin i n th e E— W direction . W e concentrat e o n (1 ) th e amoun t o f Jurassi c basi n subsi dence that can be explained by thermal contracting resulting from th e Permo-Triassic event, (2) the spatia l connection between the tw o rif t event s and (3 ) the apparen t mod e o f flexur e (coupled v . decoupled ) o f th e lithosphere . Modellin g result s indicate tha t post-rif t subsidence a s a resul t o f Permo-Triassi c riftin g ha s practicall y ceased a t th e onse t o f Jurassi c rifting. Th e position of the rif t axi s is shown to migrate with time, from th e Horda Platfor m in Permo-Triassi c tim e t o th e presen t Vikin g Graben centr e i n Jurassic-Cretaceou s time . The majo r differenc e betwee n the consequence s o f a couple d versu s a decouple d mod e o f flexure i s that th e latte r cause s a large r amoun t o f fault-bloc k rotation. O n thi s basis , we suggest a decouple d mod e o f flexur e fo r th e Permo-Triassi c rif t phase , a couple d mod e o f flexure for th e Earl y Jurassic rift phase , and a decoupled mod e o f flexure again for th e Lat e Jurassic period .

The Viking Graben i s part o f the series of linked half-grabens tha t compos e th e Nort h Se a sedimentary basi n (Beac h e t al. 1987 ; Badley e t al. 1988; Faerset h et al . 19950 ; Christiansso n e t al . 2000; Odinsen e t al. 2000; Fig. 1) . This basin was formed durin g severa l extensional events following Caledonian collision. The first event occurred in Devonia n tim e (e.g . Hossac k 1984 ; McCla y et al. 1986 ; Andersen & Jamtveit 1990 ; Fossen & Rykkelid 1992 ) and affecte d a n area that extends far beyon d th e late r margin s o f th e rif t system . This wa s followe d b y th e pronounce d Permo Triassic(?) an d Jurassic-Cretaceou s rif t phases . As wel l contro l i s spars e fo r th e pre-Triassi c sequences, th e ag e of the olde r o f these events is still a matter o f debate (see Table 1) . However, a Late Permian-Early Triassi c ag e is suggested by the occurrenc e o f Permia n sediment s i n th e southern Viking Graben (Lervi k et al. 1989 ) and in the Uns t Basi n (Johns & Andrews 1985 ; Platt 1995), b y palaeomagneti c datin g o f faul t rock s in th e southwester n par t o f Norwa y (Torsvi k et al . 1992) , an d b y th e occurrenc e o f Permia n dykes i n th e southwes t Norwegia n coasta l are a (Faerseth et al. 1976 ; Faerseth 1978; Furnes et al . 1982). Additiona l mino r Lat e Triassi c event s were recognize d b y Morto n e t al . (1987 ) an d

Gabrielsen et al. (1990), and a phase of increased subsidence occurre d i n lat e Cretaceous time . In spit e o f almos t 3 0 years o f researc h i n th e Viking Graben area, severa l problems remai n t o be solved: • Becaus e of overprint of later structuring , the Permo-Triassic rif t stag e wa s identifie d only with difficulty i n seismic sections available for the earl y studies of th e norther n Nort h Sea . Because o f this and th e limite d well control , the relative magnitudes of the Permo-Triassi c and Jurassic-Cretaceous events in the Viking Graben hav e bee n disputed . Som e worker s (e.g. Whit e 1990 ; Lippar d & Li u 1992 ) have argue d tha t Jurassi c riftin g wa s domi nant. However , Whit e (1990 ) cam e t o thi s conclusion b y focusing on th e Eas t Shetland terrace are a i n th e west , wherea s majo r Triassic riftin g seem s t o hav e occurre d i n the east, below the Horda Platform. Lippar d & Li u (1992 ) inferre d non-unifor m thin ning (/3 Crust < Anantie ) i n th e Jurassi c Hord a Platform, t o obtai n th e bes t fi t betwee n observed an d calculate d basi n subsidence . Others (e.g . Giltne r 1987 ; Marsde n e t al . 1990) derived from numerica l modelling that

From: NOTTVEDT , A . e t al . (eds ) Dynamics o f th e Norwegian Margin. Geologica l Society , London , Specia l Publications, 167 , 59-81. l-86239-056-8/00/$15.0 0 © Th e Geologica l Societ y of Londo n 2000 .

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Fig. 1 . Ma p o f Vikin g Grabe n an d it s surroundings , showin g th e locatio n o f th e seismi c line s NVGT88-04. NSDP84-01 an d NSDP84-02 . the majo r par t o f th e extensio n occurre d during th e Permo-Triassi c phase , an d tha t the basi n subsidenc e o f th e secon d phas e i s for a larg e par t du e t o therma l contractio n resulting fro m th e firs t event . Giltner (1987 ) inferred thi s fro m a ID , unifor m extension model, includin g effects o f a finite rif t phase , whereas Marsde n e l al . (1990 ) used a 2 D depth-dependent extensio n model , assumin g instantaneous stretching . Althoug h th e im portance o f th e Permo-Triassi c even t i s now firmly established, b y new , deep-penetratio n regional seismi c line s (Plat t 1995) , improve d interpretations o f dee p seismi c reflectio n data (Christiansso n e t al . 2000 ) an d basi n modelling studie s (Robert s e t al . 1993 , 1995; Odinsen e t al . 2000) , th e amoun t o f exces s heat Inherited ' b y th e Jurassi c even t i s stil l not wel l established . Another matte r o f debat e i s th e positio n of th e Triassi c rif t axis . I t i s ofte n state d (e.g. Lippar d & Liu 1992 ; Yieldin g et al. 1992) that the Triassic rift axi s coincided with that of the Jurassic-Cretaceous basin. Badley et al . (1984, 1988 ) suggested that mos t o f th e major fault s that were active during the latter episode wer e reactivate d basement-involved

faults o f th e firs t rif t episode , an d tha t greatest stretchin g an d subsidenc e occurre d above th e Permia n rif t axis . However , they also observe d occasiona l basemen t fault s i n places unaffecte d b y th e Permia n rif t stage . On th e othe r hand , Faerset h (1996 ) argue d that the thickness distribution of the sediment pile suggest s tha t th e tw o observe d riftin g stages may have little spatial connection. Th e effects o f the Permo-Triassic rif t phas e can b e recognized t o th e eas t belo w th e Hord a Platform (e.g . Stee l & Ryset h 1990 ; Sneide r et al. 1995), whereas no unequivocal evidence for th e rif t i s found below the Jurassic riftin g axis. Therefore, Faerseth (1996) proposed that the Permo-Triassi c extensio n maximum ma y have bee n locate d i n th e easter n par t o f th e basin, with the main rif t situate d beneat h th e present Horda Platform . This is supported b y observations o f othe r dominantl y Triassi c basins in the eastern North Sea , e.g. the Egersund Basin , th e Hor n Grabe n an d th e Ast a Graben (Lervi k e t al . 1989) , which ar e als o found eas t o f th e Jurassi c rift axis . Several numerica l basi n modellin g studies have been carried out o n the northern Viking Graben (e.g . Giltne r 1987 ; Marsde n e t al .

LITHOSPHERE EXTENSIO N I N TH E VIKIN G GRABE N

61

Table 1 . Dating o f th e major rift events, a s proposed o r used by various workers

Reference

First rif t phas e

Second rif t phas e

Giltner (1987 ) Gabrielsen e t al. (1990) Marsden e t al . (1990) (modelled a s instantaneous events: White (1990 ) Lippard & Liu (1992 )

248-2 13 Ma* ?Late Palaeozoic-Scythian 250-2 13 Ma 230 Ma Triassic 260-235 Maf o r 225-200 Maf Early Triassi c 250 Ma Kungurian-Anisian (261-236 MaJ) 261-236 Mat

169-98 Ma* Bathonian-Ryazanian Bathonian-Ryazanian 170 Ma) 160-100 Ma

Roberts et al . (1995) (modelled a s instantaneous events: Odinsen et al . (2000) This study

156-131 Mat Mid- Jurassic-Earliest Cretaceou s 170 Ma ) Callovian-Ryazanian (165-141 Mat) 165-141 Mat

* Time scal e o f Harlan d e t al . (1982). t Time scale of Haq e t al. (1987). tTime scal e o f Harlan d e t al . (1990). 1990; Robert s e t al . 1995 ; Odinse n e t al . 2000). In most o f these studies an estimate of the lithosphere rigidity, often expresse d a s an 'effective elasti c thickness ' (EET) , wa s mad e to calculate th e flexural response o n the mass redistribution as a result of the extension. The EET value s estimate d b y differen t worker s vary considerably, from 1.5 km (Roberts et al. 1995) t o ±4 5 km (Odinse n e t al . 2000). Thi s discrepancy i s ofte n attribute d t o th e possi bility tha t th e crus t i s decouple d fro m th e mantle, becaus e o f vertica l variation s i n the strengt h o f th e lithosphere . However , little attentio n ha s bee n pai d s o fa r t o th e influence o f thi s strengt h distributio n o n the extensio n mechanism. The purpos e o f thi s pape r i s t o examin e th e nature o f th e variou s episode s o f riftin g i n th e Viking Graben , thereb y focusin g on th e follow ing questions: • Ho w much post-rif t subsidenc e was induced by the excess heat, remnant fro m th e Permo Triassic stretchin g event ? An d ho w di d thi s eventually influence the later developmen t o f the basin? • I s there an y spatial connectio n betwee n the rift events ? Wher e wa s th e regio n o f max imum Permo-Triassi c extension , an d di d this influenc e th e positio n o f th e Jurassic Cretaceous basin ? • I s the lithosphere likely to be in a 'decoupled' state, an d wha t doe s thi s mea n fo r th e flexural response o n extension ? Whereas i n earl y numerica l modellin g studie s in th e Vikin g Grabe n (e.g . Badle y e t al . 1988 ;

White 1990 ) statement s abou t th e olde r rif t phase wer e derived fro m studie s on th e amoun t of extensio n an d subsidenc e o f th e latter , Rob erts e t al . (1995 ) focuse d particularl y o n th e pre-Jurassic event , usin g a combine d flexura l backstripping an d forwar d modellin g method . Following thes e workers , w e investigat e th e Permo-Triassic rif t phas e a s wel l a s th e younge r ones. We concentrate o n the northern par t o f the Viking Grabe n (60°20 /-61°N), usin g a depth converted versio n o f seismi c lin e NVGT88-0 4 (Figs 1 an d 2) . A 2D , non-unifor m numerica l model, includin g fault s an d incorporatin g th e effect o f a finit e perio d o f extension , i s used . The mode l allow s fo r bot h couple d an d decoupled behaviou r (Te r Voord e e t al . 1998) . As th e effec t o f compactio n durin g buria l i s not include d in th e model , w e first calculate the decompacted thicknesse s o f th e sedimen t pack ages. Thes e thicknesse s ar e use d t o constrai n the modelling.

Geological structure Seismic lin e NVGT88-04 i s situated betwee n th e deep seismi c reflectio n line s NSDP84-0 1 an d - 2 published previousl y (Beac h e t al . 1987 ; Gibb s 1987) and recentl y reprocessed and reinterprete d (Christiansson e t al . 2000 ; Odinse n e t al . 2000 ; Fig. 3) . Th e thre e line s togethe r provid e a n outline o f th e cross-sectiona l structur e o f th e northern Vikin g Graben. Heavil y faulted Meso zoic synrif t sediment s ar e unconformabl y over lain b y a Cenozoic , almos t unfaulte d post-rif t sequence. Pre-Jurassi c sediment s ca n clearl y

Fie 2 (a ) Seismic lin e NVCJTKK-0 4 Positio n o f th e lin e i s indicated i n Fig . I . Vortica l axis : two-wa y trave l lim e (ms|. (b) Depth-converted interpretation . 0FZ , 0ygarden Faul t Zone- HBS . Hil d Bren t Statfjor d Fault ; ()s , Oseher g Fault ; Br . Brag e Fault ; BrF , Brag e Fas t Fault ; 0y . 0ygarde n Fault .

LITHOSPHERE EXTENSIO N I N TH E VIKIN G GRABE N

63

Fig. 3 . Depth-converte d interpretatio n o f line s NSDP84-0 1 an d NSDP84-02 . Afte r Christiansso n e t al. (2000). Position o f th e line s i s indicated i n Fig . 1 .

be observe d belo w th e Hord a Platform , an d were recently also recognize d furthe r to th e west (Platt 1995 ; Faleid e e t al . 2000) . Th e Permo Triassic even t extende d ove r a wide r are a tha n the Jurassi c rif t (e.g . Gabrielsen e t al. 1990 ; Faerseth e t al . \995a; Robert s e t al . 1995 ; Odinse n et al . 2000) .

intervals (pre-Jurassic , earl y Jurassic , lat e Jur assic an d Cretaceous ) wer e extracted , an d th e decompacted thicknesse s wer e calculate d (Bald win & Butle r 1985 ; Dykstr a 1987 ) (Fig . 4) . A s the Bajocia n to p o f th e Bren t grou p i s clearl y visible o n th e seismi c profiles, we choose thi s t o be th e boundar y betwee n th e earl y Jurassi c an d the lat e Jurassic. Fo r th e compaction correctio n we use d th e porosity-dept h relatio n

Decompaction of th e sediments From lin e NVGT88-04 , th e compacte d thick nesses o f th e sediment s fo r fou r seperate d tim e

where <j) i s porosit y an d z i s depth . W e choos e 0o = 0.56 an d c = 0.39 km"1 , a s suggeste d b y

64

M. TE R VOORD E E T AL .

Fig. 4. Sedimen t thicknes s as measured alon g profile NVGT88-04 , subdivided into fou r dept h intervals . Dashed lines, befor e decompaction; continuou s lines , after decompaction . Sclater & Christie (1980 ) for a n average d sand shale litholog y in th e Nort h Sea . Subsequently , the reconstructe d profile s for eac h tim e interval were obtaine d b y correctin g fo r th e horizonta l fault component s a t eac h leve l (Fig . 5) . I f th e sedimentation rat e wa s hig h enoug h t o avoi d the basi n t o becom e underfilled , th e resultin g profiles represent the basin subsidence during the given time interval. This was not alway s the cas e during th e histor y o f th e stud y area. I n Cretac eous an d earl y Tertiar y time , fo r example , th e water dept h wa s probably mor e tha n 50 0 m an d may eve n hav e increase d t o a maximu m o f 1000m in the central part o f the basin (Nelso n & Lamy 1987 ; Bertram & Milto n 1988) . As wate r depths ar e not include d i n the numerica l model , we regard basin subsidence derived for suc h time intervals as underestimated.

Basin subsidence , subdivide d int o fou r time-intervals

Interval top basement-top Triassic (Fig. 5a) The most strikin g feature s in this profil e are th e large, wedge-shape d basin s a t th e easternmos t part o f th e profile , beneat h th e presen t Hord a Platform. Thes e wedges are thinnin g t o th e west and sho w th e larges t subsidenc e o f th e Permo Triassic basi n syste m in th e stud y area. Decom pacted sedimen t thicknesse s u p t o 6.2k m ar e derived. T o th e west , beneat h th e presen t rif t axis, a sediment pile with an approximatel y even thickness o f 4. 5 km i s found . However , becaus e of th e relativel y poo r qualit y o f th e seismi c image beneat h th e Bren t Grou p a t thi s depth , this estimat e shoul d b e regarde d a s uncertain.

LITHOSPHERE EXTENSIO N I N TH E VIKIN G GRABE N

65

Fig. 5. Decompacte d sedimen t thicknes s along profile NVGT88-04 , subdivide d int o four interval s and correcte d for subsequen t faulting . Fault s A, B , C and D ar e describe d i n the text . We interprete d fou r major , westerl y dippin g faults beneat h th e Hord a Platform , wherea s the faul t patter n wes t o f faul t 'C ' (th e Brag e East Fault ) seem s t o b e dominate d b y easterly dipping faults . Thi s i s emphasize d b y th e sedi mentary wedge s slightl y increasing i n thicknes s towards th e wes t abov e th e Oseber g structure .

Interval top Triassic-top Brent (Fig. 5b) At thi s time , subsidenc e wa s shifted toward s the present rif t axis , reachin g a maximu m basi n

depth o f 3.7k m i n a basi n o f les s tha n 35k m width. Th e positio n o f th e presen t rif t axi s als o corresponds t o th e sit e wher e th e faul t polarit y shifted fro m easterly dipping faults i n the west to westerly dippin g fault s i n th e east . Thi s implie s that thi s sit e has move d fro m th e Brag e are a i n Permo-Triassic tim e t o th e presen t rif t axi s in Earl y Jurassi c tim e i n thi s par t o f th e basin . However, variation s alon g strik e i n th e relatio n between th e Permo-Triassi c an d Jurassi c centr e of extension exist , possibly influence d by (deacti vated?) transfe r faults (Faerset h 1996) .

66

M. TE R VOORD E E T AL .

It should b e noted that considerable faultin g is derived fo r thi s period . Thi s implie s tha t th e early Jurassi c extensiona l episod e ma y b e mor e important tha n might be expected base d o n lines farther t o th e south , wher e thi s sequenc e i s seen as a genera l progressiv e thickenin g t o th e wes t with onl y ver y limited faul t control .

Interval top Brent-base Cretaceous (Fig. 5c) Although th e positio n o f riftin g i n thi s perio d coincided wit h th e Earl y Jurassi c basi n centre , activity alon g east-dippin g fault s i n th e wes t became more pronounced, reflectin g initiatio n of the Shetlan d Platform . Faul t rotatio n bot h eas t and wes t o f th e rif t axi s appear s t o hav e le d t o footwall uplif t an d erosion . Th e maximum basin depth wa s 5.2km , measure d i n th e strongl y westerly thinnin g sedimen t wedg e bounde d b y fault 4 B\ Th e shallowes t poin t o f thi s wedg e lies ±17 km west of the deepest point , at a dept h of 2.1km .

Interval base Cretaceous-base Tertiary (Fig. 5d) Except fo r a slight offset alon g tw o west-dippin g faults a t th e Hord a Platform , th e effec t o f faulting i s hardl y noticeabl e i n th e Cretaceou s sequence. Instead , subsidenc e i n a wide r are a i s observed. Maximu m sedimen t thickness , alon g the rif t axis , is measured t o b e 3.8km . Possibly , thermal subsidenc e an d sedimen t loading , com bined wit h a genera l ris e in se a level, resulted i n gradual buria l o f Jurassi c faul t block s durin g this period . Although we acknowledge the fact that the top of the basement i s difficult t o constrain fro m line NVGT88-04, an d i s therefore no t undisputed , a thick Triassi c sedimen t pil e wa s als o suggeste d by Christiansso n e t al. (2000), who reinterprete d the deep structure of the northern Viking Graben from dee p reflectio n seismi c lines and gravimetric and magneti c data (Fig. 3). Bearing this in mind, some preliminary conclusions can be drawn. The Permo-Triassic rif t phas e affecte d a wide r are a than di d th e subsequen t Jurassic-Cretaceou s event. Thi s i s consistent wit h result s o f Faerseth (1996). Also , Odinse n e t al . (2000) , usin g th e interpretation o f Christiansso n e t al . (2000) , derived a regiona l stretchin g distributio n in th e Permo-Triassic compare d wit h th e mor e loca lized thinnin g i n Jurassi c time . Th e are a o f strongest subsidenc e associate d wit h the Permo Triassic even t i s positione d t o th e eas t o f th e

location o f majo r Jurassi c extension . Thi s i s also consisten t wit h a shif t o f th e rif t axi s to th e west afte r Triassi c times , whic h i s reflecte d i n the reversa l o f faul t polarit y i n tim e wes t o f th e Oseberg area . The faul t polaritie s derive d fro m lin e NVGT88-04 ar e no t necessaril y representativ e for th e whol e Vikin g Graben area , a s th e faul t polarity shift s alon g strik e (Gabrielse n e t al . 1990; Faerseth et al . 19950 ; Faerset h 1996) . However, a shif t o f th e faul t polarit y i n tim e i s supported b y wha t w e regard a s th e mos t likel y interpretations o f line s furthe r t o th e sout h (e.g. NVGT88-02) . I n thi s area , east-dippin g faults withi n the Vikin g Graben see m t o repre sent maste r fault s i n th e Permo-Triassi c perio d as wel l a s i n th e earl y stag e o f Jurassi c rifting , but ar e i n place s cross-cu t b y younge r west dipping fault s durin g th e lat e stag e o f Jurassi c rifting (Faerset h 1996) .

Lithosphere rheology One o f th e critica l factor s i n a flexura l basi n model i s th e presume d lithospher e rheology . In mos t models , th e lithospher e i s assume d t o react o n a loa d i n the sam e wa y a s a thi n elastic plate floating on a viscous fluid. The thickness of this (imaginary ) plate , th e so-calle d 'effectiv e elastic thickness ' (EET ) determine s th e flexural response o f th e lithospher e t o loading : a smal l EET yield s a large-amplitude , short-wavelength response, wherea s a larg e EE T lead s t o small amplitude, long-wavelengt h flexure . Th e EE T can thu s b e estimate d b y considerin g i t a s a factor o f th e 'respons e function 1 describin g th e lithosphere deformation caused b y loading. This approach ha s bee n followe d fo r th e Vikin g Graben b y variou s workers , resultin g in supris ingly larg e variation s i n estimate s o f th e EET . Odinsen e t al. (2000) derived an EE T determined by th e 450 :C isotherm , whic h i n thei r mode l corresponds t o a dept h o f c . 45 km, usin g a kinematic numerica l mode l develope d b y Koo i (1991). O n th e othe r hand , Kuszni r et al. (1991). Roberts e t al . (1995 ) an d Te r Voord e e t al . (1997), usin g flexural models fo r footwal l uplift , arrive at EET values between 1. 5 and 6 km in the same area . W e propos e tha t thi s discrepanc y might b e associate d wit h th e assumptio n o f a 'coupled" versus a 'decoupled ' lithospher e rheology, the latter of which is characterized b y a very weak an d ductil e lower crust (i.e . Kusznir et al . 1991; Te r Voord e e t al . 1998) . The continenta l plate may contai n dippin g o r sub-horizontal weak ductile zones, which cannot support significan t bendin g stresses . I t i s argued

LITHOSPHERE EXTENSIO N I N TH E VIKIN G GRABE N

67

Fig. 6 . Strengt h profile s for a typica l continental lithospher e i n a n extensiona l regime , assumin g a 2 0 km thic k upper crust , a 15k m thic k lowe r crust , an d a tota l lithospher e thicknes s of 125km , fo r variou s lithospher e rheologies. (a) and (b ) assume a 'dry' rheology, the upper crust consistin g of dry quartzite an d th e lower crust of diabase, an d (c ) assumes a 'wet ' rheology , th e uppe r crus t consistin g of wet quartzite an d th e lowe r crus t o f diorite. A zer o por e pressur e wa s assume d i n (a) , an d a hydrostati c por e pressur e i n (b ) and (c) .

(e.g. L e Picho n & Chamot-Rook e 1991 ; Burov & Diamen t 1995 ) tha t th e appearanc e o f suc h zones i n th e lowe r crus t permit s mechanica l decoupling o f th e uppe r crus t fro m th e mantle . A redistributio n o f load s i n th e uppe r crust , for exampl e a s a resul t o f extensiona l faulting , would the n b e compensated fo r b y elastic deformation i n th e uppe r crus t exclusively , whereas the lowe r crus t woul d adjus t b y ductil e flow . This ca n explai n th e ver y lo w (1.5-6 km) effec tive elasti c thicknesse s obtaine d i n modellin g studies where the basi n topograph y wa s used a s the mai n constrainin g factor . Figure 6 show s strengt h profile s fo r th e tensional regim e i n a typica l continenta l crus t before extension . Th e profile s ar e constructe d from th e commonl y used , laboratory-derive d strength equation s (e.g . Goetz e & Evan s 1979 ; Brace & Kohlstedt 1980 ; Carter & Tsenn 1987 ; see Appendi x 1) . Th e depth-strengt h curve s demonstrate clearl y th e effec t o f th e strai n rat e on th e crusta l configuration . I f strai n rate s ar e higher than 1CT 13 s"1 the shallowest weak ductile layer occur s a t th e bas e o f th e lowe r crust , o r perhaps eve n deeper . Fo r lowe r strai n rates , a n additional ductile zone appears a t the base of the upper crust . A full y decoupled mod e o f flexure is likely t o be an exceptional case, becaus e th e rate of lower

crustal flo w i s generall y no t hig h enoug h t o compensate fo r th e entir e upper-crusta l defor mation (Buc k 1988 ; Te r Voord e e t al 1998) . Therefore, par t o f th e compensatio n probabl y mostly occurs b y deep, very low viscosity mantle material causin g th e lithospher e t o b e i n a 'partly decoupled ' mod e (se e Te r Voord e e t al . 1998). However , t o mak e th e effect s o f th e assumption o f couple d v . decouple d behaviou r on th e modellin g result s a s clea r a s possible , we wil l conside r onl y th e end-member s o f th e decoupled-coupled spectru m i n thi s study. The numerica l model We use a forward, 2D difference model , describ ing th e lithospheri c deformatio n resultin g fro m extension, an d includin g the effec t o f faultin g i n the uppe r crust . A finit e syn-rif t phas e i s con sidered (Waltha m 1989 , 1990) . Thi s i s o f majo r importance whe n calculatin g th e post-rif t ther mal subsidence , a s stressed , fo r example , b y Ter Voord e & Cloetingh (1996) . A presentatio n of th e mode l i s give n i n Fig . 7 . Followin g Kusznir e t al . (1991) , th e mode l i s divided int o an uppe r laye r of faultin g wit h vertica l shea r i n the deforming hanging wall, and a lower layer of more distributed deformation. In both layer s the

68

M. TE R VOORD E E T AL .

Fig. 7 . Schemati c representatio n o f the model. Extensio n is achieved b y distributed deformation (i.e. deformation along faults ) i n th e uppe r crust , an d b y distribute d thinnin g in th e lowe r lithosphere.

condition o f volum e conservatio n i s satisfied . Extension rate s ca n b e varie d pe r tim e interva l and pe r fault . Th e fault s ar e assume d t o flatte n into th e detachmen t tha t form s th e boundar y between th e tw o layers . Th e dept h o f thi s detachment i s an importan t factor , a s i t ha s th e same functio n i n th e mode l a s a "neckin g depth' (e.g. Braun & Beaumont 1989 ; Weissel & Karne r 1989; Koo i e t al. 1992) . Th e neckin g dept h i s defined a s the level that remain s horizontal in the absence o f isostati c forces , an d i s a decisiv e factor fo r th e existenc e an d amoun t o f footwal l uplift (e.g . Te r Voord e & Cloetingh 1996) . Temperatures ar e calculate d fro m th e hea t transfer equation :

where T i s temperatur e ( C) , / i s tim e (s) , K i s thermal diffusivit y (nrs" 1 ), v i s velocity , F i s heat productio n (Wm~ 3 ), p i s densit y (kgm~ 3 ) and C i s specific heat (J C C - 1 kg" 1 ). The equatio n is solved usin g a finite difference method o n a rectangula r grid . Therma l proper ties o n eac h gri d poin t ar e obtaine d b y interpo lation from th e second, movin g grid representin g the extending basin. Hea t transfe r during a s well as afte r th e deformatio n i s calculated . Flexure i s calculate d fro m th e thin-plat e approximation (e.g . Bodine e t al . 1981) :

where u - i s deflectio n (m) , D i s rigidit y ( N m) , Arompi i s densit y o f th e materia l underneat h th e plate (kgm~ 3 ), p COmP2 i s densit y o f th e materia l above th e plate (kgm~ 3 ), g is gravitational acceleration (ms~ 2 ) an d q i s vertical loa d (Nm~ 2 ). The vertica l loa d i s calculate d b y vertica l integration o f th e densit y contrast s cause d b y

deformation o f th e crust , temperatur e change s and th e mas s o f deposited sediments , relativ e to that o f th e undeforme d crust :

where a i s thermal expansio n coefficien t ( K ] ). For th e couple d mod e o f flexure , w e assum e the deflectio n u - t o b e constan t i n eac h vertica l column throug h th e whole lithosphere. signifyin g that th e crust an d mantl e are forge d togethe r b y the lowe r crust . Hence , th e integratio n interva l for calculatin g th e vertica l loa d q comprises th e whole lithosphere . p c ompi i n equatio n (3 ) i s the asthenospher e density , an d p ComP2 th e sedi ment density. For th e decoupled mod e of flexure, the deflectio n u - is assumed t o b e different fo r th e crust an d mantle . Fo r th e uppe r crust , q i s calculated b y integrating from th e surfac e t o th e lower crust , p ComPi i s take n a s th e lower-crusta l density, p comp2 a s th e sedimen t density , an d th e rigidity i s assume d t o b e lo w ( E E T < 6 k m ) . This i s differen t fro m earlie r studie s usin g lo w EET value s (e.g . Kusznir e t al . 1991 ; Roberts et al . 1995) , wher e th e influenc e o f decouplin g on th e compensatio n material s i s no t take n into account . The therma l feature s o f th e mode l hav e bee n described i n mor e detai l b y Te r Voord e & Bertotti (1994) . th e structura l feature s b y Te r Voorde & Cloeting h (1996) , an d th e effect s o f decoupling hav e bee n discusse d b y Te r Voord e et al . (1998) .

Model parameters The mode l parameter s w e used i n thi s stud y ar e summarized i n Tabl e 2 . Th e Moh o dept h i s assumed t o b e 35km , base d o n th e crusta l thickness observe d belo w th e platfor m areas . The detachmen t dept h i s assume d t o b e 18km .

LITHOSPHERE EXTENSIO N I N TH E VIKIN G GRABE N Table 2 . Parameters used for th e modelling Parameter

Value

Deviation used fo r sensitivity tes t

EET (coupled ) EET (decoupled ) Density sediment s Density uppe r crus t Density lowe r crus t Density asthenosphere Moho dept h Detachment dept h Young's modulus Poisson rati o

450°C isother m 1.5-6 km 1900 kg m-3 2700 kg m"3 2950 kg m-3 3300 kg m-3

see Fig. 8 see Fig . 8 ilOOkgm- 3 ±50kgm~ 3 ±50 kgm- 3 -

35km 18km 7 x 10 10 Nm 0.25

±2 km ±5 km -

based o n observation s o f Fosse n e t al. (1998). They describe d th e flattening of normal fault s a t a maximu m dept h between 1 8 and 20km , in the Tampen Spu r are a o n th e wester n flan k o f th e Viking Graben . Although Fosse n e t al . (1999 ) suggeste d tha t these flattene d fault s ar e inherite d mechanica l weak zones in the basement, relate d to Devonia n extensional structure s o r Caledonian thrusts , th e detachments coul d also be linked to ductile shear zones in the middle an d lowe r crust an d thu s be explained b y temperature - an d pressure-con trolled change s i n th e lithospher e rheolog y (e.g. Fosse n & Gabrielse n 1996) . Th e differen t detachments ma y have different causes , and ma y also have been initiated at different stage s during rifting. Then , trul y low-angl e structure s migh t also develo p b y flattenin g o f originall y steepe r fault structures , durin g repeate d phase s o f stretching, an d perhap s ac t a s detachments i n a later stag e (Fosse n e t al . 2000) . Althoug h th e flattening o f th e fault s wit h dept h ma y hav e no unequivoca l explanation , i t ma y indicat e the existence and positio n o f ductile zones in the lower crust . I n th e cas e o f th e Vikin g Graben , the maximu m observe d dept h o f flattenin g should the n als o b e interprete d a s th e dept h o f mechanical decoupling . Model sensitivit y an d constraints The overal l stat e o f flexur e depend s o n th e strength o f the layer(s) , the densit y and amoun t of sediments , th e dept h o f th e Moho , an d th e depth o f neckin g durin g lithospheri c extension . To obtai n a n impression of the model sensitivit y to these variable s fo r th e Vikin g Grabe n con figuration, w e teste d th e effect s o f varyin g th e EET, th e densitie s an d th e Moh o depth , usin g

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parameters an d deviation s give n in Table 2 an d the sam e faul t configuratio n a s i n Fig . 5 . Because, in our model , a change in necking level implies a chang e i n faul t shapes , effect s o f varying thi s leve l o n th e resultin g flexura l respons e are difficul t t o tes t independently. However, i t is possible t o calculat e whic h chang e i n sedimen t density would b e equivalent t o a change of 1 km in th e leve l o f necking (see Appendix 2) , and w e used thi s analog y t o estimat e th e mode l sensi tivity fo r th e leve l o f necking. The deviation s i n densitie s an d Moh o dept h cause maximum variations in the resulting basin subsidence o f 60 m fo r th e couple d cas e an d 185m fo r th e decouple d case . A reasonabl e variation o f 5k m i n neckin g dept h ca n b e compared wit h a variatio n i n sedimen t densit y of —14 8 kgm~3 (i n the cas e o f a deepe r neckin g level) o r 262kgm~ 3 (i n th e cas e o f a shallowe r necking level) for th e couple d cas e (Appendi x 2, equation (6)) , leading to a maximum variatio n in basin subsidenc e o f 100m . Fo r th e decouple d case th e leve l o f neckin g ha s n o influenc e (Appendix 2 , equation (6)) . In Fig . 8 the effect s o f changes i n EE T value s are show n fo r th e simulatio n o f th e Permo Triassic tim e interva l (i.e . Fig . 5a) . Th e figur e shows clearl y that th e choic e o f th e EE T i s o f limited influenc e fo r th e couple d scenario , bu t might be important i n the case o f decoupling. In the modelling , w e use d a n EE T define d b y th e 450°C isother m fo r couple d flexure , an d th e EET givin g the bes t fit for decouple d flexure . Finally, effect s o f reasonabl e error s i n th e estimated fault shape s and deformation mechan ism ma y b e o f th e orde r o f 1 km (Whit e e t al . 1986; Whit e & Yieldin g 1991 ; Withjac k & Peterson 1993) . In previou s studie s o n th e Vikin g Grabe n (Marsden e t al. 1990 ; Roberts e t al. 1995 ; Odinsen e t al . 2000) , modellin g result s and (seismic ) data normall y fit within 1000m . Base d o n this , on the estimated sensitivit y of the model, an d o n uncertainties in the seismi c data, fo r exampl e as a resul t o f possibl e error s i n chose n velocit y models, w e regar d a fi t o f th e mode l wit h th e data withi n 1000 m a s a goo d approximation . Modelling result s The numerica l mode l wa s use d t o simulat e th e sequential basin configuration s shown i n Fig . 5 , in order to relate the amount o f basin subsidence to th e amount o f extension along various faults, and t o th e strengt h an d mod e o f flexur e o f the lithosphere . I n th e model , sediment s ar e assumed t o fill the basi n u p t o th e surface . Thi s

70

M. TE R VOORD E E T AL .

Fig. 8. Modellin g result s fo r interva l to p basement-to p Triassic , assumin g differen t value s fo r th e EET. (a) Decouple d rheology ; (b ) coupled rheology .

is a realistic approach, as long as the basin is not underfilled a t th e en d o f th e modelle d episode . The modellin g parameter s w e used ar e summar ized i n Tabl e 2 .

Interval top basement-top Triassic (Fig. 9) Triassic sediment s ar e suppose d t o hav e bee n deposited i n lacustrine environments, dominate d by alluvia l fan s alon g th e grabe n margin s an d with fine r fluvia l o r lacustrin e sediment s i n th e lows' (e.g . Stee l & Ryset h 1990) . A t th e en d o f Triassic tim e th e norther n Vikin g Grabe n wa s approximately a t se a level . The sedimen t thick ness show n i n Fig . 5 a might thu s b e interprete d as th e tota l basi n subsidenc e durin g thi s tim e interval. W e focuse d especiall y o n th e flank s o f the graben , becaus e o f th e poo r seismi c resolu tion in the basin centre. Consequently , th e muc h better fi t o f th e mode l beneat h th e platfor m areas tha n beneat h th e rif t axi s (Fig . 9a ) reflect s the highe r confidenc e o f th e observe d sedimen t thickness i n tha t area . To obtai n th e bes t modellin g result , a faul t configuration wa s require d wit h a chang e i n fault polarit y on th e flanks of the Viking Graben and o f th e Brag e are a (Fig . 9a) . Th e thinnin g of the mappe d sequenc e i n th e Oseber g are a towards th e eas t i s the n explaine d b y extensio n along th e east-dippin g faul t 'A' , boundin g th e Viking Grabe n i n th e west , wherea s th e Viking Graben centr e subside d furthe r alon g th e west dipping faul t W B', 25k m farthe r t o th e east . Th e

major extensio n occurre d o n th e Hord a Plat form, provide d tha t th e to p Triassi c interpreted in th e grabe n centr e i s correct . Two scenario s wer e modelle d fo r thi s tim e interval, on e assumin g a couple d rheolog y wit h an EE T determine d b y th e 45 0 C isother m (corresponding t o a n initia l dept h o f 4 3 km i n the model) , th e othe r assumin g a decouple d rheology wit h a n EE T o f 1.5km . A reasonabl e fit between the modelled basi n configuration and the on e interprete d fro m th e seismi c lin e coul d be obtaine d fro m bot h assumption s (Fig . 9b) . but wit h differen t value s fo r th e amoun t o f extension. I n th e cas e o f a couple d configura tion, th e lithospher e i s to o stron g t o sho w a significant flexura l respons e o n th e extension , and th e basemen t subsidenc e i s almos t entirel y due to crustal movements along the faults. How ever, if we adopt a decoupled mod e o f extension, the mode l show s a syn-rif t uplif t o f th e area , leading t o a smal l amount o f subaeria l footwal l erosion, followe d by thermal subsidence (Fig. 9b and c) . The modelle d uplif t an d subsidenc e patterns for th e decouple d scenari o agre e wit h th e findings o f Robert s e t al. (1995) , an d thos e o f the couple d scenari o wit h finding s o f Odinse n et al . (2000) . Robert s e t al . (1995 ) modelle d th e pre-Jurassic riftin g stag e i n th e Hord a Platfor m region wit h a flexural cantilever model, usin g an EET o f 1.5km , an d foun d tha t a stag e o f foot wall uplif t an d erosio n precede d a stag e o f overall therma l subsidence . O n th e othe r hand , the results of Odinsen et al. (2000), who use d th e

LITHOSPHERE EXTENSIO N I N TH E VIKIN G GRABE N

71

Fig. 9. Modellin g results for interva l to p basement-to p Triassic. (a ) Modelled faul t configuration ; dashe d lin e indicates basi n dept h a s derive d fro m profil e NVGT88-04 . (b ) Long dashes, modellin g resul t fo r couple d rheology; shor t dashes , modellin g resul t fo r decouple d rheology ; continuou s line , basi n dept h a s derive d fro m profile NVGT88-04 . (c ) Modelled basi n dept h fo r decouple d rheology , a t (fro m to p t o bottom ) 0 , 30 , 74 and oo Ma afte r th e en d o f rifting . coupled approach , d o no t indicat e suc h uplift . According t o Robert s e t al. (1995) , th e seismi c expression o f th e bette r image d Triassi c faul t blocks beneat h th e Hord a Platfor m give s th e impression o f bevellin g at th e fault-bloc k crests. This ca n b e notice d i n interpretation s of Beac h et al . (1987 , fig. 2), Robert s e t al . (1993, fig. 6a) and Odinse n e t al . (2000 , fig s 2 an d 5) , an d supports th e decouple d scenari o a s th e mos t likely one . The tota l amount s o f extensio n use d i n ou r model ar e 2 3 km fo r th e couple d scenari o an d 34.5km fo r th e decouple d scenario , signifyin g an averag e (3 valu e o f 1.1 8 fo r th e couple d case , and 1.2 9 for th e decouple d cas e (measure d ove r the entir e profile) . Majo r extensio n i n Permo Triassic tim e took plac e on the Horda Platform, the tota l modelled amoun t o f extension being 14 or 22 % wes t o f th e Brag e fault , an d 2 2 or 38 % east o f it , fo r th e couple d o r decouple d case , respectively.

Roberts e t al . (1995 ) reporte d a f t valu e o f 1.34 in the Horda Platform area , alon g the sam e section a s in the present study , whereas Odinsen et al . (2000 ) calculate d a n averag e /3 facto r o f 1.27 fo r thei r norther n transect , an d 1.1 9 fo r their souther n transect . The y derive d f t factor s for th e Hord a Platfor m o f 1.3 3 an d 1.39 , bu t they di d no t indicat e exactl y over whic h width this was measured . Figure 9c shows the syn-rift uplif t an d thermal subsidence with time for the decoupled scenario . Indicated ar e the top o f the basement a t th e end of rifting, 3 0 Ma late r (i.e. latest Triassic), 74 Ma later (i.e . whe n th e secon d majo r rif t phas e i s assumed t o start ) and afte r tota l therma l relaxation. In contrast t o earlier proposals (e.g. Giltne r 1987; Marsde n e t al. 1990 ; Roberts e t al . 1995), but i n accordance wit h Odinsen et al. (2000), the present mode l suggest s tha t therma l subsidenc e as a resul t o f Permo-Triassi c riftin g canno t have ha d muc h influenc e on th e Jurassi c event .

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(It shoul d b e note d tha t therma l subsidenc e fo r the couple d cas e woul d hav e eve n les s effect. ) Researchers wh o state d tha t a n inherite d ther mal anomaly influence d Jurassic riftin g use d I D thermal calculation s (e.g . Giltne r 1987) , instan taneous Triassi c riftin g (Marsde n e t al. 1990 ; Roberts e t al . 1995) , and/o r derive d thei r statements fro m studie s o f th e Jurassi c even t alone (Giltne r 1987). Interval top Triassic-top Brent Group (Fig. 10) The to p o f th e Bren t Grou p horizo n i s dia chronous o n a regiona l scale , an d i s situated i n the Bathonia n (Middl e Jurassic) sequence i n the northern Vikin g Graben . Th e lowe r Jurassi c sequence consist s predominantl y o f marin e shales, whereas the Middle Jurassic sequence was deposited i n non-marine to paralic environments (Fjellanger e t al . 1996 ; Ravna s e t al . 1999) . Th e Tarbert Formatio n o f th e uppermos t Bren t Group consist s o f shallo w marin e sandstones . As a n approximation , th e profil e show n i n Fig. 5 b might thu s b e regarde d t o represen t th e total basin subsidence during early Jurassic time, until mid-Bathonia n time . This tim e interva l could b e simulate d wit h a much les s complicate d faul t configuratio n than that o f th e Permo-Triassi c phase . Th e positio n of th e rif t axi s was i n th e Vikin g Graben centre ,

coinciding wit h th e sit e o f majo r subsidenc e a s well a s th e positio n o f shif t i n faul t polarit y along th e profil e (Fig . lOa) . However , th e subsidence in the basin centre might be overestimated, because, again , th e seismi c data a t thi s depth in the basi n centr e d o no t allo w fo r a decisiv e solution. Compare d wit h the earlie r rift phase , a change i n faul t di p wa s impose d betwee n th e Oseberg an d Brag e are a (faul t k CT). Figur e 1 0 shows th e modellin g results for th e couple d an d the decouple d case . Fo r th e decouple d case , a n EET o f 6k m wa s used . Th e couple d scenari o gives result s tha t ar e mor e consisten t wit h th e observed basi n configuration , a s hardl y an y tilting o f th e faul t block s i s observed . Fo r thi s reason, therma l subsidenc e relate d t o thi s phase is assume d t o b e negligible . T o mode l th e observed amoun t o f subsidence, modelled offset s along th e fault s ha d t o b e se t u p t o 2 km larger than th e observe d offsets , suggestin g a n over estimation o f th e amoun t o f extension , and th e existence o f a n additiona l source o f basi n subsidence, differen t fro m fault-relate d extension . As show n b y th e modellin g results of th e to p basement-top Triassic interval, only a small part of the additional subsidence (i.e. less then 350m) can b e explained by remnant thermal subsidence caused b y Permo-Triassi c rifting , whic h wa s not include d i n th e modellin g results. Anothe r explanation fo r basi n subsidenc e migh t b e th e existence of a 'proto-rift stage ' (Gabrielsen 1986; Nottvedt e t al . 1995) . Nottved t e t al . (1995 )

Fig. 10 . Modellin g result s fo r interva l to p Triassic-to p Brent , (a ) Modelle d faul t configuration ; dashe d lin e indicates basi n dept h a s derive d fro m profil e NVGT88-04 . (b ) Long dashes , modellin g resul t fo r couple d rheology; shor t dashes , modellin g resul t fo r decouple d rheology ; continuou s line , basi n dept h a s derive d fro m profile NVGT88-04 .

LITHOSPHERE EXTENSIO N I N TH E VIKIN G GRABE N proposed a n idealized , three-stage mode l fo r rif t evolution, involvin g a proto-rift stage, a syn-rif t stage and a post-rift stage . The proto-rift stag e is characterized b y thermally induced domal uplift , or b y depositio n i n a wide , slowl y subsidin g basin with only minor faul t activity . The syn-rif t stage describes the phase of active stretching an d block rotation , an d th e post-rift stage consists of asymptotically decreasing subsidence, caused b y thermal contraction . Interferenc e between post rift subsidenc e relate d t o th e Permo-Triassic rif t phase an d proto-rif t subsidenc e relate d t o th e Jurassic rif t phas e (Stee l 1993 ; Nottved t e t al. 1995) migh t for m a n alternativ e sourc e fo r th e extra subsidenc e observed . The total amount o f extension in the modelled scenario (Fig . lOa ) i s 7.3km , yieldin g an aver age ( 3 value o f 1.05 . Th e modelle d /3 reache s a value o f 1.1 6 i n th e basi n centre . Apparently , a mino r extensiona l even t ha s occurre d i n earliest Jurassic time, as argued earlier by R0e & Steel (1985 ) an d Gabrielse n e t al . (1990) . Th e

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magnitude o f thi s event cannot b e derived fro m the modelling results, because o f the large uncertainty i n the basi n depth . This earl y Jurassi c extension , combine d wit h the additiona l subsidence , cause d th e deposi tional environmen t t o conver t fro m continenta l to marine , implyin g tha t th e creatio n o f ne w accommodation spac e outpace d sedimen t sup ply. Subsequently , a s sedimentatio n continued , the depositiona l environmen t passe d t o non marine again , an d i n lates t Bajocian-earlies t Bathonian time the first rotational movements of the secon d majo r rif t phas e commence d (Faer seth e t al . 19956 ; Ravna s & Bondevik 1999). Interval top Brent Group-base Cretaceous (Fig. 11) The Uppe r Jurassi c sediment s i n th e Vikin g Graben consis t mostl y o f marin e shales . Th e shales o f th e Bathonian-Oxfordia n Heathe r

Fig. 11 . Modellin g result s for interva l top Brent-bas e Cretaceous , (a ) Modelle d faul t configuration ; dashed lin e indicates basi n dept h a s derive d fro m profil e NVGT88-04 . (b ) Long dashes , modellin g resul t fo r couple d rheology; shor t dashes , modellin g resul t fo r decouple d rheology ; continuous line , basi n dept h a s derive d fro m profile NVGT88-04 . (c) Modelled basi n depth fo r decoupled rheology , at (fro m top t o bottom) 0 and o o Ma afte r the en d o f rifting .

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formation ar e overlai n b y th e organic-ric h Draupne Formation (equivalen t t o the Kimmer idge Cla y Formation) , an d bot h ar e interfin gered b y sandston e unit s (e.g . Ravna s e t al. 1999). Base d o n environmenta l interpretations , the wate r dept h a t th e en d o f thi s episod e (Ryazanian) i s estimate d t o b e 200 m o n th e margins t o 300 m i n th e grabe n centr e (Badle y et al . 1988) , althoug h thi s migh t b e a ver y con servative estimat e (e.g . Marsde n e t al . 1990 ; Yielding e t al . 1992) . The modellin g result s ar e show n i n Fig . 11 . The EE T use d fo r th e decouple d cas e wa s 6 km. Broadly th e sam e faul t configuratio n coul d b e used a s fo r th e earl y Jurassi c phase , bu t wit h a few mor e activ e faults . Th e basi n no w becam e asymmetric, deepenin g t o the east in this part of the graben . However , thi s varie s alon g strik e (see, e.g. Odinsen e t al. 2000). The footwal l uplif t in th e wes t i s bes t explaine d b y usin g th e decoupled approach , a s show n i n Fig . l i b . This approac h yield s a n averag e /3 facto r ove r the whole transect o f only 1.0 7 (and a maximum of 1.2) . Together wit h th e earl y Jurassi c phase , this make s a Jurassi c ( 3 of 1.12 . Thi s i s slightl y lower tha n th e values found fo r Jurassic stretch ing b y Odinse n e t al . (2000), wh o obtaine d 1.1 5 for th e norther n transec t an d 1.1 9 fo r th e southern transect , o r b y Robert s e t al . (1993), who reporte d a ( 3 of 1. 3 i n th e basi n centre , of 1.1 5 in th e Eas t Shetlan d Basi n an d west ern Hord a Platfor m an d o f 1.0 5 i n th e easter n Horda Platform . Tha t th e 3 valu e w e derive d is relativel y lo w ca n b e relate d t o th e fac t tha t our mode l assume s a zer o wate r depth , whic h is a n underestimatio n fo r thi s tim e interval .

Figure l i e display s th e basi n configuratio n immediately afte r stretchin g a s wel l a s afte r thermal subsidence . Fro m this figure, it is evident that th e therma l anomal y cause d b y th e extension alon e canno t b e responsibl e fo r th e sub sequent Cretaceou s subsidence . Thi s ca n partl y be explaine d b y th e fac t tha t th e basi n wa s underfilled at the end of this rift phase, which was not include d i n th e modelling . Th e modelle d amount o f extension , an d thu s th e therma l anomaly, i s therefore probabl y a n underestima tion. Durin g Cretaceou s time , muc h o f the sedi mentation wa s accommodated b y th e infillin g o f previous rif t bathymetry , causin g additiona l subsidence becaus e o f sedimen t loading . Thi s effect i s furthe r exaggerate d b y th e compactio n of earlie r deposite d sediment s (Yieldin g e t al . 1992), whic h i s no t take n int o accoun t i n th e present model .

Modelled Moho configuration To compar e th e modelle d Moh o dept h wit h the observations, we superimposed th e Moh o defor mation o f th e thre e tim e inteval s on eac h other , the resul t of whic h is show n in Fig . 12. The 'observed Moho ' i n Fig . 1 2 is derive d b y inter polation betwee n th e Moh o reflector s observe d in seismi c line s NSDP84-01 an d - 2 (e.g . Chris tiansson e t al . 2000 ; Odinse n e t al . 2000) . although w e acknowledge tha t th e Moh o i s no t unequivocally image d east o f the Viking Graben (see Fig. 3) . Starting from a constant initia l value of 3 5 km, th e modelle d Moh o dept h i s too larg e compared wit h th e dept h derive d fro m th e

Fig. 12 . Moh o uplif t a s a consequence o f rifting . Dashe d line , modelled Moh o dept h resultin g from thre e stages of rifting ; continuou s line , Moho dept h alon g lin e NVGT88-04, a s derive d fro m interpolatio n betwee n interpretations o f lin e NSDP84-0 1 an d NSDP84-02 .

LITHOSPHERE EXTENSIO N I N TH E VIKIN G GRABE N observations. Th e maximu m difference , belo w the Vikin g Grabe n axis , i s 6.5km . Thi s i s i n agreement wit h findings o f Odinsen e t al. (2000), who suggeste d tha t th e Permia n pre-rif t Moh o depth migh t hav e varie d fro m ±3 5 km o n th e flanks to les s tha n 30k m i n th e basi n areas . T o obtain th e bes t fi t fo r th e subsidenc e data , Lippard & Li u (1992 ) neede d t o assum e a n initial crusta l thicknes s varyin g fro m 28k m i n the centr e t o 30-3 2 km o n th e flank s o f th e Graben. Thi s pre-Permia n variatio n i n Moh o depth i s likely to b e caused b y Devonian crusta l

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thinning, impose d o n th e thickene d crus t o f the Caledonia n Orogen y (Anderse n & Jamtvei t .1990; Fosse n & Rykkelid 1992) .

Mode o f flexur e i n the Vikin g Grabe n End-members: the coupled versus the decoupled mode An importan t resul t o f thi s modellin g stud y i s that i t i s har d t o discriminat e betwee n th e

Fig. 13 . Mod e diagram s fo r th e lithosphere , indicatin g the possibilit y of a decouple d rheolog y in Moh o depth-Moho temperatur e spac e (Spadin i & Podladchikov 1996) , for thre e differen t lithospher e rheologies . Boundaries between the ductile and brittle lower crust are given for strain rates of (from left t o right) 10~ 17, 10~ 16, 10~15, 10~ 14 an d 10~ 13 s" 1. Blac k dot s indicate th e assume d pre-rif t (r 0) and post-rif t (t\) Moh o dept h (a t th e position o f major Moh o uplift ) i n th e Vikin g Graben, a s derive d from Fig . 11. Th e ligh t grey arrow s indicate the chang e i n th e stat e o f th e lithospher e in cases o f instantaneou s rifting an d ver y slow rifting .

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M. TE R VOORD E E T AL .

coupled an d th e decouple d mod e o f flexur e o n the basis of the fit between modelling result s an d seismic dat a alone (e.g. Fig . 9) . This might b e an explanation fo r th e larg e variation s betwee n EET value s use d fo r previou s modellin g studie s of th e Vikin g Graben. Nevertheless , w e showe d that th e presence or absenc e of footwall uplif t i s a goo d indicato r fo r th e mod e o f flexure : th e suspected presenc e o f erode d faul t block s i n th e Permo-Triassic phas e an d o f footwal l uplif t i n the lat e Jurassic phas e canno t b e explained wit h the couple d approac h (eve n thoug h w e use d a rather dee p neckin g level) , wherea s th e absenc e of footwall uplift i n early Jurassic tim e is impos sible t o reconstruc t b y decouple d flexure . Although caution shoul d b e taken when interpreting th e result s (w e modelle d onl y th e end members o f th e coupled-decouple d spectrum , whereas th e 'partl y decoupled ' mod e i s mos t likely to occur), th e best result s were obtained b y assuming a decoupled mod e o f flexura l respons e in th e Permo-Triassi c phase , a couple d mod e o f flexure in early Jurassi c time , an d a change bac k to decouple d flexur e i n lat e Jurassi c time . Th e choice o f th e mod e o f flexur e i s important , a s it ca n hav e a stron g influenc e o n th e derive d amount o f extension . Fo r th e Permo-Triassi c phase, fo r example , w e derive d a d facto r o f 1.18 fo r th e couple d cas e an d o f 1.2 9 fo r th e decoupled case . A prerequisit e for decouple d behaviou r i s th e existence o f a ductil e lowe r crust . Figur e 1 3 shows whic h Moh o dept h an d temperatur e will caus e th e manifestatio n o f a ductil e lowe r crust accordin g t o th e rheologica l laws . Th e curves wer e derive d b y calculatin g th e dept h a t

which brittl e deformatio n give s wa y t o ductil e deformation (Spadin i & Podladchikov 1996) , i.e. where cr b i s th e brittl e strengt h an d a d i s th e ductile strengt h neede d t o caus e deformatio n a t a give n strai n rat e (see Appendix 1) . The param eters use d are give n in Tabl e 3. The Moh o depths an d temperature s o n th e lef t sid e ar e associated wit h a n entir e brittl e crust , wherea s the condition s o n th e righ t sid e o f th e curve s yield a ductile lowe r crust. I f we assume th e bol d line t o b e th e pre-rif t geother m i n th e norther n Viking Graben , an d th e blac k do t a t t Q t o indicate th e stat e o f th e pre-rif t Moho , w e ca n conclude fro m th e figur e tha t th e lowe r crus t i s at least partly ductile, and that th e first condition for a decouple d mod e o f flexur e i s fulfilled . However, the uplift o f mantle material combined with syn-rif t coolin g wil l eventuall y resul t i n a shift fro m th e decouple d t o th e couple d stat e of the lithospher e (Spadin i & Podladchiko v 1996) . which agree s wit h ou r modellin g result s fo r th e Permo-Triassic an d th e Earl y Jurassi c phases . The chang e bac k t o th e decouple d stat e i n Bathonian tim e cannot b e explained in this way. or b e relate d t o th e chang e i n extensio n (an d thus strain ) rate . I n fact , a t lowe r strai n rates , decoupled behaviou r is more likel y to occur than at hig h strai n rate s (e.g . Bru n & Tro n 1993 ; Ter Voord e e t al. 1998 ; see Figs 6 and 13) . However, a therma l even t coul d explai n th e weaken ing o f th e lithospher e an d a shif t t o decouple d behaviour. A Mid-Jurassi c pre-rif t therma l dom e i n th e so-called 'tripl e junction' o f th e Nort h Se a rif t i s

Table 3 . Parameters used for construction of strength profiles

Upper crus t Lowe

r crus t

Quartzite (w) Quartzite (d) Diorit

e (wet ) Diabas e (dry )

£ p (kJmol-') 172. 6 Ap (Pa~"s-' ) 1.26 x 10~ 13 n 1. 9 £d (kJmor 1 ) -4d(s~ l ) crD (GPa ) R (J(molK)- 1 ) Thermal diffusivit y sediment s Thermal diffusivit y crus t Thickness heat producin g laye r Heat productio n

134 21 6.03 x 1(T 24 1.2 2.72 2.

_ _ -

2 27 6 6 x ID' 16 3.1 6 x 1(T 20 4 3.0 5

Mantle olivine 510 7 x 10~ 14 3.0 535 5.7 x 10 " 8.5

8.314J (molK)- 1 0.75 x lO^nrs- 1 l . O x 10- 6 m 2 s-' 15km 2.3 x 10- 6 Wm- 3

Material constant s fo r quartzite , diorite , diabas e an d olivin e ar e adopte d fro m Tsen n & Carter (1987) .

LITHOSPHERE EXTENSIO N I N TH E VIKIN G GRABE N well established (e.g . Ziegler & Van Hoorn 1989; Hendrie e t al. 1993 ; Underbil l & Partingto n 1993), evidenced b y basin-wide subaerial erosio n (the 'Mid-Cimmerian ' unconformity), th e extrusion o f the Rattra y volcani c rocks, and floodin g events o n basi n margins . Thi s therma l dom e is interprete d t o b e th e effec t o f a short-lived , anomalously ho t asthenospheri c puls e (Hendrie et al . 1993 ) o r transien t mantl e plum e (Underbill & Partington 1993) . However, althoug h th e timing of this event is correct t o explain the shif t to decouple d behaviou r i n Bathonia n time , th e location migh t b e too fa r awa y (i.e. ±250 km t o the south ) fro m th e norther n Vikin g Grabe n to hav e suc h a pronounce d effect .

The partly decoupled mode of flexure As mentione d above , i n thi s stud y w e onl y look a t th e end-member s o f th e spectru m o f decoupled-coupled mode s o f flexure . Never theless, th e mos t probabl e mechanis m t o occu r is th e partl y decouple d mod e o f flexure . I n thi s case, th e viscosit y o f th e lowe r crus t i s to o high t o allo w fo r a n isostati c compensatio n level i n th e lowe r crust . Therefore , a s a firs t approximation, th e lithospher e wil l reac t i n a coupled mode , an d onl y short-wavelengt h lateral variatio n i n th e loa d i s compensate d b y lower-crustal flow. In general , modelle d surfac e deflection s obtained b y assumin g a partl y decoupled lithosphere diffe r fro m th e fully couple d mod e in tha t an extr a short-wavelengt h componen t ca n b e observed. I n contrast , th e differenc e fro m th e fully decouple d cas e i s that th e long-wavelength uplift componen t disappear s (Te r Voord e e t al . 1998). Translatin g thi s t o Fig s 9-1 1 raise s th e suggestion o f partl y decoupled flexure.

Conclusions The evolutio n o f th e Vikin g Grabe n ha s bee n studied usin g a numerical model, constrained by seismic lin e NVGT88-04 . Althoug h th e impor tance o f th e rif t phas e i n Permo-Triassi c tim e has bee n affirme d b y th e modellin g results , th e post-rift subsidenc e i s shown t o hav e practically ceased a t th e onse t o f Jurassi c rifting . Hence , no significan t amoun t o f basi n subsidenc e i n the Jurassi c phas e ca n b e ascribe d t o therma l contraction resultin g fro m th e Permo-Triassi c event. Th e Jurassi c subsidenc e shoul d therefor e be explained entirely by a new extensional event,

77

possibly precede d b y th e 'proto-rift ' stag e a s defined b y Nottved t e t al . (1995). The position o f the rift axi s is shown to change with time . Th e Permo-Triassi c rif t axi s wa s positioned beneat h th e present Horda Platform, whereas th e Jurassi c rif t axi s wa s positione d beneath th e present Vikin g Graben centre . Thi s result migh t be dependent o n th e positio n alon g strike (Faerset h 1996) . Fo r example , Odinse n et al . (2000 ) foun d fo r lin e NDSP84- 1 tha t th e maximum (3 factors fo r bot h th e Permo-Triassi c and th e Jurassi c phase s wer e i n th e presen t Viking Grabe n axis , wherea s fo r th e mor e southern lin e NSDP84-0 2 th e maximu m ( 3 factor i n th e Permo-Triassi c phas e wa s o n th e Horda Platform . To ou r knowledge , this basin modelling study is th e firs t on e tha t attempt s t o discriminat e between th e couple d an d th e decouple d mod e of flexure . Th e majo r differenc e betwee n thes e two mode s i s th e amoun t o f fault-bloc k rota tion, which is larger whe n a decouple d rheolog y is assumed. The coupled and decouple d mode of flexure are sometime s har d t o discriminat e o n the basi s o f basin geometrie s alon e (e.g . Fig . 9) , which ma y explai n th e larg e variatio n i n EE T values foun d i n th e literature . Nevertheless, the modelling result s sugges t a decouple d mod e o f flexure i n th e Permo-Triassi c phase , a couple d mode o f flexur e i n earl y Jurassi c time , an d a change bac k t o decoupled flexure in Lat e Juras sic time . Subsequen t researc h o n thi s topi c should focu s o n furthe r constraint s o n th e mode o f flexure, such as the lower-crustal thickness an d material , an d th e pre-rif t geotherm . We woul d lik e t o than k R . Ravna s fo r bringin g u s together, an d T . Odinse n fo r discussion s abou t th e development o f th e Vikin g Graben . Review s b y R. Boutilier and D. Waltham were greatly appreciated. This researc h wa s supporte d b y th e IB S (Integrate d Basin Studies ) project , par t o f th e Joul e I I researc h programme funde d b y th e Commissio n o f Europea n Communities (Contract JOU2-CT92-0110). This paper is Publicatio n 98100 1 o f th e Netherland s Researc h School of Sedimentary Geology.

Appendix 1 The brittl e yiel d strengt h i s give n b y Byerlee' s law (Byerle e 1978 ; Brace & Kohlstedt 1980) :

where p i s th e density , g i s th e gravitationa l acceleration, z i s the depth , a = 0.75, A =0 for zero por e pressure , an d A = 0.4 fo r hydrostati c pore pressure . Ductil e deformatio n i s assume d

78

M. TE R VOORD E E T AL .

to occu r b y power la w cree p (Kirb y 1983) :

where e is strain rate, A, n and E p ar e empiricall y derived materia l constants , R i s th e ga s con stant an d T i s th e absolut e temperature . Fo r olivine with a strengt h exceeding 200 MPa, duc tile deformatio n i s describe d b y Dor n cree p (Goetze & Evans 1979 ; Tsenn & Carter 1987) :

Parameters ar e give n i n Tabl e 3 .

Appendix 2 The dept h o f necking is the leve l of zero vertical motion i n th e absenc e o f isostati c forces . Thi s level determines the ratio between thinning of the upper crust, where crustal material i s replaced by sediments, and thinnin g of the lower lithosphere, where crusta l materia l i s replace d b y mantl e material. Therefore , th e dept h o f neckin g i s a decisive facto r fo r th e ne w loa d distributio n caused b y crustal extension. For a give n amount of thinning , the chang e i n loa d A P cause d b y a variation i n neckin g dept h A N i s equa l t o th e change i n loa d A P cause d b y a variatio n i n sediment densit y Ap s. Ap s ca n b e calculate d as follows.

Fig. A2 . Thinnin g o f th e crus t aroun d neckin g level s N] o r W 2, assumin g the depressio n o f th e surfac e i s equal (i.e . 3 is different) fo r bot h cases. Reference loa d P I cause d b y thinning around necking leve l A^ , usin g sediment density ps] (see Fig. Al):

where g i s gravitational acceleration (ms~2 ), 3 is a stretchin g factor , p u i s densit y o f th e uppe r crust (kgm~ 3 ), p \ i s densit y o f th e lowe r crus t (kgirT 3 ), p m i s densit y o f th e mantl e (kgrrr 3 ) and M i s the Moh o dept h (km) . AP cause d b y thinnin g around N 2 instea d o f around TV, :

AP cause d b y usin g sedimen t densit y p instead o f p\:

2

This yield s the followin g equivalenc e betwee n &N(=N\ - N 2) an d A p ( =pi - p 2):

Fig. Al . Thinnin g of th e crus t aroun d neckin g level s N] o r A r2, assuming (3 is equal (i.e. the depressio n of th e surface i s different) fo r bot h cases .

If w e assum e th e depressio n o f th e surfac e a s a result of crustal thinning to be constant (as it can

LITHOSPHERE EXTENSIO N I N TH E VIKIN G GRABE N be observed), a change in neckin g depth implies a chang e in th e j3 factor (se e Fig . A2) . The n

and

which is the equatio n to be used for the cas e of the Vikin g Graben .

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LITHOSPHERE EXTENSIO N I N TH E VIKIN G GRABE N HELLAND-HANSEN, W. , L0MO , L. , RY SETH, R . & STEEL , R . J . 1999 . Sedimentatio n history as an indicato r of rif t initiatio n and development: th e late Bajocian-Bathonian evolutio n of the Oseberg-Brag e area , norther n Nort h Sea . Norsk Geologisk Tidsskrift, i n press . ROBERTS, A. , YIELDING, G. , KUSZNIR , N., WALKER , I . & DORN-LOPEZ , D . 1993 . Mesozoi c extensio n i n the Nort h Sea : constraint s fro m flexura l back stripping, forwar d modellin g an d faul t popula tions. In: PARKER, J. R . (ed.) Petroleum Geology of Northwest Europe: Proceedings of the 4th Conference. Geologica l Society , London, 1123-1136 . ,, , & 1995 . Quantitativ e analysis o f Triassi c extensio n i n th e norther n Viking Graben. Journal of th e Geological Society, London, 152 , 15-26 . R0E, S . L . & STEEL , R . J . 1985 . Sedimentation , sealevel ris e and tectonic s at th e Triassic-Jurassi c boundary (Statfjor d Formation) , Tampe n Spur , northern Nort h Sea . Journal o f Petroleum Geology, 8 , 163-186 . SCLATER^ J . G . & CHRISTIE , P . A . F . 1980 . Conti nental stretching: a n explanation o f the post-mid Cretaceous subsidenc e o f th e Nort h Se a basin . Journal o f Geophysical Research, 85, 3711-3739 . SNEIDER, J . S. , D E CLARENS, P . & VAIL , P . R . 1995 . Sequence stratigraph y o f th e Middl e t o Uppe r Jurassic, Vikin g Graben , Nort h Sea . In: STEEL , R. J., FELT , V., JOHANNESSEN , E . & MATHIEU , C . (eds) Sequence Stratigraphy o n th e Northwest European Margin. Norwegia n Petroleu m Society , Special Publication , 5 , 167-197 . SPADINI, G . & PODLADCHIKOV , Y . 1996 . Spacin g o f consecutive norma l faultin g i n th e lithosphere : a dynamical mode l fo r rif t axi s migratio n (Tyr rhenian Sea). Earth and Planetary Science Letters, 144, 21-34 . STEEL, R . 1993 . Triassic-Jurassic megasequence strati graphy i n the Northern Nort h Sea : rif t t o post-rift evolution. In: PARKER, J. R . (ed. ) Petroleum Geology of Northwest Europe. Proceedings of the 4th Conference. Geologica l Society , London, 299-315. & RYSETH , A . 1990 . The Triassic-earl y Jurassi c succession i n th e norther n Nort h Sea : megase quence stratigraph y an d intra-Triassi c tectonics. In: HARDMAN , R . P . F . & BROOKS , J . (eds ) Tectonic Events Responsible for Britain's Oil and Gas Reserves. Geologica l Society , London , Spe cial Publications , 55 , 139-168 . TER VOORDE , M . & BERTOTTI , G . 1994 . Therma l effects o f norma l faultin g durin g rifte d basi n formation. 1 : A finit e differenc e model . Tectonophysics, 240 , 133-144 . & CLOETINGH , S . 1996 . Numerical modelling of extension i n faulted crust : effects o f localized an d regional deformatio n o n basi n stratigraphy . In : BUCHANAN, P . G . & NIEUWLAND , D . A . (eds ) Modern Developments in Structural Interpretation, Validation an d Modelling. Geologica l Society , London, Specia l Publications , 99 , 283-296. , RAVNAS , R. , F^RSETH , R . & CLOETINGH , S . 1997. Tectoni c modellin g o f th e middl e Jurassi c synrift stratigraph y i n th e Oseberg-Brag e area ,

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northern Vikin g Graben . Basin Research, 9 , 133-150. , VAN BALEN, R . T. , BERTOTTI , G . & CLOETINGH , S. A . P . L . 1998 . Th e influenc e o f a stratifie d rheology o n th e flexura l respons e o f th e litho sphere t o (un)loadin g b y extensiona l faulting . Geophysical Journal International, 134 , 721-735 . TORSVIK, f. H. , STURT , B . A., SWENSSON , E. , ANDER SEN, T . B . & DEWEY , J . F . 1992 . Palaeomagneti c dating o f faul t rocks : evidenc e fo r Permia n an d Mesozoic movement s an d brittl e deformatio n along th e extensiona l Dalsfjor d Fault , wester n Norway. Geophysical Journal International, 109 , 565-580. TSENN, M . C . & CARTER , N . L . 1987 . Upper limits of power la w cree p o f rocks . Tectonophysics, 136 , 1-26. UNDERBILL, J . R . & PARTINGTON , M . A . 1993 . Jurassic therma l domin g an d deflatio n i n th e North Sea : implication s o f th e sequenc e strati graphic eveidence . In : PARKER , J . R . (ed. ) Petroleum Geology of Northwest Europe: Proceedings o f th e 4t h Conference. Geologica l Society , London, 337-345 . WALTHAM, D . 1989 . Finit e differenc e modellin g o f hanging wal l deformation . Journal o f Structural Geology, 11 , 433-437. 1990. Finit e differenc e modellin g o f sandbo x analogues, compaction and detachment free deformation. Journal of Structural Geology, 12,375-381. WEISSEL, J . K . & KARNER , G. D . 1989 . Flexural uplif t of rif t flank s du e t o mechanica l unloadin g o f th e lithosphere durin g extension . Journal o f Geophvsical Research, 94 , 1 3 919-13 950. WHITE, N . J. 1990 . Does the uniform stretching model work i n th e Nort h Sea ? In : BLUNDELL , D . J . & GIBBS , A . D . (eds ) Tectonic Evolution o f th e North Se a Rifts. Clarendon , Oxford , 217-240 . & YIELDING , G . 1991 . Calculatin g norma l faul t geometries a t depth : theor y an d examples . In : ROBERTS, A . M. , YIELDING , G . & FREEMAN , B . (eds) Th e Geometry o f Normal Faults. Geolog ical Society , London , Specia l Publications , 56 , 251-260. , JACKSON , J. A . & MCKENZIE , D . P . 1986 . Th e relationship betwee n th e geometr y o f norma l faults an d tha t o f th e sedimentar y layer s in thei r hanging walls . Journal o f Structural Geologv, 9 , 789-795. WITHJACK, M . O . & PETERSON, E . T. 1993 . Predictio n of normal-faul t geometrie s - a sensitivit y study. AAPG Bulletin, 11 , 1860-1873 . YIELDING, G. , BADLEY , M . E . & ROBERTS , A . 1992 . The structural evolution of the Brent Province. In: MORTON, A . C. , HASELDINE , R . S. , GILES , M . R . & BROWN , S . (eds ) Geology o f th e Brent Group. Geological Society, London, Specia l Publications , 61, 27-40. ZIEGLER, P . & VA N HOORN , B . 1989 . Evolutio n o f North Se a rif t systems . In : TANKARD , A . J . & BALKWILL, H . R . (eds ) Extensional Tectonics and Stratigraphy of the North Atlantic Margins. American Associatio n o f Petroleu m Geologists , Memoir, 46, 471-500.

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Permo-Triassic an d Jurassic extension in the norther n Nort h Sea: results fro m tectonostratigraphi c forward modelling TORE ODINSEN,

14

PAU L REEMST,

JAN ING E FALEIDE 1

3

25

PETE R VA N DE R BEEK,

& RO Y H . GABRIELSEN

26

1

Department of Geology, University of Bergen, Allegaten 41, N-5007 Bergen, Norway 2

Faculty of Earth Sciences, Vrije Universiteit, De Boelelaan, 1081 HV Amsterdam, The Netherlands ^Department of Geology, University of Oslo, P.O. Box 1047 Blindern, N-0316 Oslo, Norway 4 5

Present address: S tat oil, N-5020 Bergen, Norway

Present address: Geologica AS, P.O. Box 8034, N-4029 Stavanger, Norway 6

Present address: Universite Joseph Fourier, 15 rue Maurice Gignoux, 38031 Grenoble Cedex, France

Abstract: W e have undertaken 2D forward modelling across th e northern North Sea, based on reprocessed , interprete d an d depth-converte d dee p reflectio n seismi c line s NSDP84- 1 and - 2 (15stwt) an d refractio n data . Tw o separat e stretchin g phases , Permo-Triassi c an d Jurassic, ar e recognized . Th e cumulativ e stretchin g i s consistent wit h th e observe d crusta l structure an d th e overal l basi n configuration , as reproduce d b y forwar d modelling . Good agreement betwee n observe d an d modelle d to p basemen t level , and crusta l thickness below the platfor m area s ar e particularl y emphasized . Crustal-scal e modellin g indicate s tha t crustal thicknes s varied acros s th e norther n Nort h Se a at th e onse t o f th e Permo-Triassi c rifting, fro m c. 35km i n th e platfor m area s t o les s than 30k m i n th e interio r o f th e basin . This ma y b e ascribe d t o Devonian(-Carboniferous? ) crusta l stretching . Thinnin g o f th e crust ha s progressivel y bee n narrowed , fro m post-Caledonia n extensiona l collapse , t o les s regional Permo-Triassi c basins , an d finall y developmen t o f th e Vikin g Graben are a i n th e Jurassic-early Cretaceou s time . Mos t o f th e Permo-Triassi c stretchin g occurre d betwee n the 0ygarden Fault Zon e t o the east and th e Shetland Platfor m (souther n transect) and th e Hutton Faul t alignmen t t o th e west . Th e widt h o f th e Permo-Triassi c basi n wa s c . 120125km, with calculated /3 mean between 1.3 8 an d 1.40 . Permo-Triassi c /3 mean estimates across the present Horda Platform vary between 1.3 3 and 1.39 . The Jurassic /3 mean estimates for the same are a var y betwee n 1.0 8 and 1.13 . Acros s th e Vikin g Graben , Permo-Triassi c /3 mean varies betwee n 1.2 8 (southern transect ) an d 1.4 1 (northern transect) . Thi s i s lowe r tha n estimates fo r th e Jurassi c /3 mean, whic h amount s t o 1.5 3 and 1.42 . Permo-Triassic an d Jurassic /3 mean estimate s acros s th e Eas t Shetlan d Basi n ar e 1.2 9 and 1.11 , respectively . Lithospheric therma l evolutio n reflects th e genera l difference s betwee n Permo-Triassic an d Jurassic stretching , wit h a muc h wide r therma l perturbatio n durin g th e forme r an d a focusing an d latera l migratio n toward s th e eas t o f th e pea k therma l elevatio n durin g th e latter. Ther e ar e stil l uncertaintie s related t o th e degre e o f (de)couplin g between th e uppe r crust an d uppe r mantl e durin g th e Permo-Triassi c an d th e Jurassi c rif t phases . Thes e uncertainties ar e relate d t o th e interpla y between age, strain rate , crusta l rheology , crustal thickness an d long-live d zones o f weaknesses.

The norther n Nort h Se a ha s bee n extensivel y phase . A previous rif t phas e is best recognized in described i n the literature as an extensiona l sedi- reflectio n seismi c line s whic h cros s th e Hord a mentary basin , whic h forme d durin g repeate d Platform , a s deepl y buried , rotate d larg e faul t lithospheric stretching . A s a resul t o f hydro - blocks . Becaus e o f th e lac k o f wel l dat a fro m carbon exploratio n o f th e area , muc h attentio n syn-rif t sediment s associate d wit h thes e struc has focuse d o n rotated faul t block s tha t evolve d tures , th e age of this earlier even t i s still a matte r during th e mid-Jurassi c t o earl y Cretaceou s rif t o f debate , an d bot h Permia n (Eyno n 1981 ; From: N0TTVEDT , A . e t al. (eds ) Dynamics o f th e Norwegian Margin. Geologica l Society , London , Specia l Publications, 167 , 83-103. 1-86239-056-8/00/S15.0 0 © Th e Geologica l Societ y of Londo n 2000 .

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T. ODINSE N E T AL .

Badley e t al. 1988 ; Gabrielsen e t al. 1990 ; Faer seth e t al . 1995 ) and Triassi c (Beac h e t al . 1987 ; Giltner 1987 ; Robert s e t al . 1995 ) age s hav e been proposed . Several studies have also focused on the crustal and lithospheri c configuration across th e north ern North Sea, as interpreted fro m a deep seismic reflection surve y NSDP8 4 an d othe r dee p line s (Beach 1986 ; Beach e t al . 1987 ; Klemperer 1988 ; Klemperer & Huric h 1990 ; Brun & Iron 1993) . Limitations i n th e dat a qualit y hav e permitte d several interpretation s o f th e crusta l an d litho spheric configuration. It ha s bee n suggeste d tha t lithospheric stretchin g wa s governe d b y sym metrical pur e shea r (Giltne r 1987 ; Badle y e t al . 1988), decouple d symmetrica l shea r (Klemp -

erer 1988) , decoupled asymmetrica l shea r (Res ton 1990) , o r asymmetrica l simpl e shea r (Beac h 1986; Beac h e t al . 1987 ; Gabrielsen 1989) . Modelling o f basi n formatio n i n th e Nort h Sea wa s initiate d b y th e wor k o f McKenzi e (1978), wh o use d th e are a a s on e o f hi s typ e examples. Hi s widel y recognize d mode l fo r lithospheric stretchin g ha s bee n followe d b y a number o f relate d studies . Earl y work s wer e concentrated o n th e Central Graben area , where ID subsidenc e analysi s was carrie d ou t (Sclater & Christi e 1980 ; Woo d 1981 ; Wood & Barto n 1983). Mor e recently , acces s t o conventiona l seismic reflectio n dat a an d wel l dat a move d attention toward s th e norther n Nort h Se a (Fig. 1) . Suc h dat a facilitate d calculatio n o f

Fig. 1 . Ke y ma p an d mai n structura l elements, northern North Sea (referenc e leve l i s base Cretaceous). Area s that ar e primaril y affecte d b y Jurassic-early Cretaceou s stretching ar e show n i n grey . Location s of transect s 1 and 2 are shown .

NORTH SE A PERMO-TRIASSI C AN D JURASSI C EXTENSIO N stretching estimate s fro m subsidenc e an d sum mation o f fault heave (Giltner 1987 ; Badley et al. 1988). Som e worker s als o hav e focuse d o n stretching estimate s lookin g a t crusta l thin ning derive d fro m dee p seismi c reflectio n line s (Beach e t al . 1987 ; Klemperer 1988) . The dee p lines have not previousl y been subjected to mor e advanced quantitativ e studie s suc h a s forwar d or revers e modelling. The presen t stud y i s base d o n th e frequentl y used dee p (15stwt ) seismi c reflectio n line s NSDP84-1 and - 2 (transect 1 and 2 of the present study), whic h hav e bee n post-stac k reprocesse d and redisplaye d (Figs 2 a an d 3a ) an d dept h converted (Fig s 2 b an d 3b) . Detaile d strati graphic an d structura l architectur e hav e bee n interpreted fro m high-qualit y conventional seismic line s (7stwt) . T o obtai n optima l velocit y information, velocitie s from well s were used fo r the uppermos t 3- 4 s o f th e section , wherea s refraction velocitie s fro m expande d sprea d pro files (ESPs) were applied to the deeper parts of the transects. Th e dee p basi n geometr y i s furthe r constrained by gravity and magnetic data (Christiansson et al. 2000). We studied the thermal an d tectonic evolution in the northern Viking Graben by forwar d modelling , integrating Permo-Trias sic an d mid-Jurassic-earl y Cretaceou s litho spheric stretching . Thre e mai n topic s wer e th e specific subjec t of our stud y of transects 1 and 2 : to estimat e an d compar e Permo-Triassi c an d Jurassic stretching ; t o evaluat e th e modelle d crustal configuratio n with respec t t o th e strati graphy, an d th e basemen t an d Moh o reliefs ; t o evaluate lithospheric thermal evolution. The stud y extend s previou s studie s i n tha t we full y integrat e commercia l an d dee p seismi c lines t o bette r constrai n th e structura l architec ture an d basemen t topography . Ou r modellin g approach allow s u s t o trac k th e therma l an d mechanical evolutio n o f th e lithosphere , a s wel l as basi n stratigraphy , throug h multipl e finiterate riftin g phases , thu s allowin g meaningfu l predictions o f stratigraph y t o b e made. W e emphasize tha t thi s pape r i s par t o f continuin g work i n th e Nort h Se a Basin , an d th e reade r will therefor e benefi t fro m additiona l readin g of Christiansso n et al. (2000) , Fosse n et al. (1999), Odinse n e t al . (2000 ) an d Te r Voord e et al . (2000) . Th e wor k o f Te r Voord e e t al . (2000) is especially relevant as they take som e of the presen t modellin g result s furthe r int o a new modelling tool . Seismic dat a Transect 1 ha s a lengt h o f 270k m an d run s NW-SE. Fro m SE , th e transec t crosses th e

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0ygarden Faul t Zon e an d Hord a Platform . It continue s acros s th e interio r o f th e Vikin g Graben an d Eas t Shetlan d Basin , onto th e Eas t Shetland Platfor m (Fig . 2) . Transect 2 i s 270k m lon g an d run s E- W (Fig. 3) . A t th e easter n en d o f transec t 2 i t crosses th e 0ygarde n Faul t Zon e an d Hord a Platform, befor e i t passe s acros s th e Vikin g Graben, whic h is both deeper and narrower than in transec t 1 furthe r north . Fro m th e Vikin g Graben, transec t 2 passe s ont o th e Shetlan d Platform. Unfortunately , velocit y dat a fro m ESPs wer e no t availabl e alon g transec t 2 . Thi s forced u s t o calibrat e th e dee p velocit y profil e with data fro m transec t 1 . Furthermore, th e preJurassic stratigraphy and depth t o basemen t ar e not wel l constrained beneath th e western part o f the Hord a Platform . In bot h transects , th e uppe r reflectiv e crus t is dominated b y Cenozoic, relativel y flat-lying and unfaulted post-rif t sequence s (Jord t et al. 1995) , unconformably coverin g heavil y faulte d Meso zoic syn-rif t sediments . Bot h transect s revea l a n asymmetrical pattern , wit h th e deepes t part s o f the basi n situate d t o th e west , i n th e hangin g wall o f th e inne r wester n maste r faul t o f th e Viking Graben . The middl e crus t i s characterize d b y poo r reflectivity. Thi s i s i n stron g contras t t o th e undulating reflective lower crust. Although there are som e difference s i n th e reflectiv e patter n o f the lower crust in the two transects, the principal architecture i s consistent . Als o th e widt h an d area o f Moh o shallowing , an d th e depth s o f Moho, ar e simila r i n transect s 1 and 2 (Odin sen e t al . 2000) . Reflection s belo w Moh o ar e recorded in bot h transects . Thes e feature s dip away fro m th e grabe n axi s an d ar e focuse d a t the transitio n betwee n strongl y thinne d crus t and th e thicke r crus t belo w th e platfor m areas . They follow th e general pattern fo r intra-mantle reflections a s describe d b y Klempere r (1988) . Modelling procedur e an d parameters Early modellin g studie s o f th e norther n Nort h Sea (Giltner 1987 ; White 1990 ) were based on the McKenzie (1978 ) model . Thes e model s predic t the first-order characteristics of the basin, but d o not mode l fault-controlle d syn-rif t stratigraph y because they fail to incorporate th e mechanics of rifting i n a self-consisten t manne r (Koo i e t al . 1992). Alternatively , shorter-wavelength model s have bee n applie d tha t explicitl y tak e motio n along rif t borde r fault s int o accoun t (Marsde n et al. 1990; Roberts e t al. 1993, 1995). The short wave-length model s requir e ver y lo w flexura l

Fig. 2. (a ) Post-stac k reprocesse d transec t 1 in time . (For location , see Fig . 1. ) (b) Depth-converte d an d interprete d transect I . Questio n mark s (? ) refer t o uncertaintie s o f the Permo-Triassi c sequenc e an d basemen t level . N o vertica l exaggeration .

Fig. 3. (a ) Post-stack reprocesse d transec t 2 in time. (For location , se e Fig. 1. ) (b) Depth-converted an d interprete d transect 2. Question mark s (?) refer to uncertaintie s of the westwar d continuation of th e Permo-Triassi c sequenc e an d basemen t level . N o vertica l exaggeration.

88

T. ODINSE N E T AL .

rigidities o f th e lithospher e (effectiv e elasti c thickness (EET ) 1.5- 6 km) an d ar e i n apparen t agreement wit h a gravit y stud y o f th e cen tral North Se a by Barton & Wood (1984) , which suggested loca l isostati c compensation . Th e latter study , however , calculate d th e isostati c response o f an initially unflexed an d undeforme d plate t o post-rif t loadin g b y sediment s an d thermal contraction . A s discusse d b y Koo i e t al. (1992) , suc h a treatmen t o f isostati c com pensation i s internally inconsistent and i s bound to yiel d very low elastic thicknesses . I n addition. Barton & Wood (1984 ) di d no t tak e th e gravity effect o f sedimen t compactio n int o account , another facto r that wil l lead t o an underestimat e of lithospheri c strength (Cowi e & Karner 1990) . A mor e recen t seismi c an d gravit y stud y o f the centra l Nort h Se a b y Hollige r & Klemp erer (1990 ) indicate s significan t departures fro m local isostas y ove r th e basin , suggestin g tha t th e lithosphere ha s retaine d flexura l rigidity . Thi s conclusion i s consisten t wit h inference s fro m dynamic model s o f riftin g (Brau n & Beaumon t 1989; Bass i e t al . 1993 ; Buro v & Diament 1995 ) that unde r moderat e amount s o f stretching , th e lithosphere wil l retai n significan t strength. Ebin ger e t al . (1991 ) an d Va n de r Bee k (1997 ) hav e shown that rifting i n the East African and Baikal rifts, whic h i n man y way s provid e present-da y

analogues t o riftin g i n th e norther n Nort h Sea , is controlle d b y a stron g lithospher e wit h EE T of c . 30 km. We us e a 2 D forwar d mode l (Koo i e t al . 1992). whic h adopt s a finit e strengt h o f th e lithosphere durin g stretching . Th e kinematic s of stretchin g an d th e flexura l isostati c respons e are controlle d b y th e dept h o f neckin g (Z neck ) (Fig. 4 ) (Brau n & Beaumon t 1989 ; Weisse l & Karne r 1989 ; Koo i e t al . 1992) . A surfac e depression wit h a dept h (S) i s then give n by

where 3 i s a variabl e stretchin g factor. In thi s manner, fault-controlle d syn-rif t strati graphy ca n b e simulate d withou t havin g t o revert t o extremel y low flexura l rigiditie s (Koo i et al . 1992 ; Spadin i e t a l 1995) . Th e applie d Zneck is constrained b y seismic data an d tria l and error, an d considere d robust . Th e mode l i s ru n by dividin g th e crusta l an d subcrusta l litho sphere, wit h a constan t pre-rif t latera l thicknes s (35 km), into a number o f 5 km wide boxes. Eac h box i s assigne d a crusta l an d a subcrusta l stretching facto r (Royde n & Kee n 1980) . Th e crustal an d subcrusta l factor s are assume d t o be equal i n ou r model . I t incorporate s latera l hea t flow an d finit e stretchin g (Cochra n 1983) . Th e

Fig. 4 . Schemati c illustratio n o f th e lithospheri c neckin g (afte r Koo i e l al . 1992 ; modifie d fro m Brau n & Beaumont 1989) . The leve l o f necking (Z neck) is defined a s th e leve l o f n o vertica l motio n i n th e absenc e of isostatic forces . A shallo w leve l o f neckin g (bottom ) create s a surfac e depressio n tha t i s shallower tha n th e isostatically compensate d basi n dept h (CD ) an d result s i n a downwar d stat e o f flexure. Conversely, fo r a dee p level o f neckin g (top ) a n upwar d loa d act s o n th e lithosphere , resultin g i n flexura l supporte d rif t flanks .

NORTH SE A PERMO-TRIASSI C AN D JURASSI C EXTENSIO N thermal stat e an d subsequen t subsidenc e o r uplift ar e calculate d usin g a finit e differenc e scheme. Thi s basemen t subsidenc e i s inverte d to th e correspondin g stratigraph y b y infil l o f sediments and calibrate d b y an estimated waterdepth profile . Th e flexura l strengt h o f the litho sphere i s controlle d b y th e 450° C isotherm , which rise s durin g therma l perturbatio n a s a consequence o f stretching . Wit h thi s approac h the flexura l strengt h o f th e crus t an d mantl e are alway s coupled . Thi s i s obviousl y a simpli fication of the elastic response as the model does not allo w a further investigatio n of the interplay between strain rate, age, composition an d thick ness of the crust. As shown by Burov & Diament (1995) an d Cloeting h & Buro v (1996) , thes e factors contro l th e lithospheri c strength , whic h may chang e totall y before , durin g an d afte r rifting. The subject has been further investigated by Te r Voord e e t al (2000) , wh o teste d bot h end-members (couple d o r decoupled) , an d w e will consider that stud y at th e end o f this paper. Compaction o f sediments is calculated using a standard porosity-dept h relationshi p (f(z) = 0.58e~° 41r) fo r al l sediment s (Sclate r & Christi e 1980). Thi s i s a simpl e approach wher e a sand / shale rati o o f 40/6 0 reflect s th e 'averag e lithol ogy' i n th e compactio n scheme . Kyrkjebo e t al . (2000) teste d th e effect s o f varyin g lithology on compaction by backstripping the Cretaceous and Cenozoic sequence s alon g transec t 2 assumin g 100% o f shal e an d 100 % o f sand , usin g th e porosity-depth relatio n o f Sclate r & Christi e (1980). Th e differenc e betwee n shal e an d san d volumes (2 D cross-sectiona l area ) wa s o f th e order o f 0-20%. Hence , th e sensitivit y analysis indicated tha t th e uncertaintie s related t o com paction-decompaction model s d o no t under mine our regiona l conclusions . The mode l als o account s fo r change s i n se a level. We used a long-term sea-leve l curve base d

on Komin z (1984) . For specifie d stages i n basi n evolution, tectoni c subsidenc e an d crusta l struc ture ca n b e examined . I n addition , th e forwar d model ca n predic t a likel y positio n o f seismi c horizons, whic h ar e otherwis e difficul t t o track , for instanc e in the Viking Graben an d belo w the western part of the Horda Platfor m in transect 2. The mode l parameter s applie d i n th e presen t study ar e show n i n Tabl e 1 . The timin g of events responsible fo r th e pres ent structura l framewor k i n the northern Nort h Sea ha s bee n debate d i n th e literature . O n th e basis of published an d ou r ow n data, tw o major stretching phase s wer e use d i n th e modelling , namely Artinskia n t o Ladina n (261-23 6 Ma), referred t o a s Permo-Triassic, an d Bathonia n t o Berriasian (165-14 1 Ma), referre d t o a s Juras sic (tim e scal e o f Harlan d e t al . (1990)) . I t i s acknowledged tha t th e initiatio n o f th e firs t phase i s especiall y controversial , a s n o well s have penetrated th e thick pre-Triassic sediment s observed withi n th e dee p half-graben s i n th e Norwegian par t o f th e norther n Nort h Sea . However, age dating in the southwestern part o f Norway indicat e Permia n age s (Faerset h 1978 ; Torsvik e t al. 1992) , which is in accordance wit h seismic studie s (Badle y e t al . 1988 ; Gabrielse n et al. 1990) . Permo-Triassic riftin g ha s als o been suggested fo r th e adjacen t Uns t Basi n (John s & Andrews 1985) . W e hav e chose n 26 1 Ma a s initiation o f Permian riftin g bu t realiz e that thi s age i s uncertain . Precise datin g of the initiatio n o f the Jurassi c rift phas e ha s als o bee n subjec t t o disagreemen t (Gabrielsen e t al . 1990) . According t o Helland Hansen e t al . (1992), increased subsidenc e rates and relativ e sea-level rise date back as far as late early Bajocia n time , althoug h significan t sub basin formatio n probabl y date s t o Kimmerid gian time in many areas (Steel 1993) . This make it reasonabl e t o assum e tha t mos t o f th e hea t

Table 1 . Parameters used i n th e tectonostratigraphic modelling (after Bodine e t al . 1981; Steckler 1981; Kooi e t al . 1992) Symbol

Definition

Value

•^neck

Depth o f neckin g (regiona l istostasy ) Isotherm describin g th e effectiv e elasti c thicknes s (EET ) Initial effectiv e elasti c thicknes s Initial lithospheri c thicknes s Initial crusta l thicknes s Coefficient o f therma l expansion Temperature o f asthenospher e Thermal diffusivit y Mantle densit y a t surfac e condition s Crustal densit y a t surfac e condition s Water densit y

18km 450°C 44km 130km 35km 3.4 x KT^C- 1 1333°C 7.8xlO-7m2s-' 3.33 gem"3 2.80gcirr3 1.03gcm~ 3

Te EET L C a Tasth

K m c c

89

90

T. ODINSE N E T AL .

was introduced during Bajocian-Callovian time . Accepting th e complexitie s i n precis e datin g o f the norther n Nort h Se a rif t phase s (N0ttved t et al. 1995) , it i s important t o not e tha t reason able change s o f th e age s w e hav e use d her e will no t chang e th e stretchin g result s o r th e crustal model .

Both phase s ar e characterize d b y regional , E-W oriente d extensio n associated wit h normal faulting an d syn-rif t sedimentatio n (e.g . Faerseth et al. 1995), and followe d by thermal cooling and sediment loadin g (e.g . Giltne r 1987 ; Gabrielse n et al . 1990 ; Whit e 1990 ; Robert s e t al . 1995) . Figure 5 summarize s th e stratigraphy , relativ e

Fig. 5. Tectono-sedimentologica l events , norther n Nort h Se a (modified fro m Gabrielse n e t al . 1990 ; Nottvedt et al . 1995) . Tim e scal e fro m Harlan d e t al . (1990).

NORTH SE A PERMO-TRIASSI C AN D JURASSI C EXTENSIO N Table 2 . List o f ages assigned t o th e horizons Stratigraphy Sea floor 0 Base Pliocen e 5 Near to p Oligocen e 2 Inter lowe r Oligocen e 3 Base Oligocen e 3 Base Eocen e 5 Base Tertiar y 6 Inter Cenomania n 9 Inter lowe r Cretaceou s (Ryazanian ) 14 Base Cretaceou s 14 Upper middl e Jurassic 16 Lower Jurassi c 20 Lower middle Triassic 23 Lower Triassi c 244 Upper lowe r Permia n 26

Age (Ma )

5 1 5 6 5 5 1 4 5 0 6 1

sea-level elevation and th e tectonic event s of the northern Nort h Sea . Seismi c horizons (Tabl e 2 ) are calibrate d t o well s along th e transects , an d the age s o f th e horizon s ar e correlate d t o th e time scal e o f Harland e t al (1990) . Stratigraphic modellin g an d crustal thickness estimates A clarificatio n o n ou r us e o f 'Vikin g Graben ' may b e neede d i n th e following , a s previou s workers als o hav e applie d thi s ter m fo r th e whole northern Nort h Se a rift system . We make a distinctio n betwee n th e muc h mor e regiona l northern Nort h Se a an d th e presen t c . 60 km wide Viking Graben, wher e most o f the Jurassi c rifting an d th e post-Jurassic subsidence occurred (Fig. 1) . When th e Permo-Triassi c Vikin g Gra ben and th e Jurassic Viking Graben ar e referre d to, thi s i s fo r compariso n o f stretchin g magni tudes i n th e sam e area , i.e . wit h respec t t o th e present 60k m wid e Vikin g Graben . Thi s als o goes for the Horda Platfor m and Shetlan d Basin and othe r areas . Figures 6 an d 7 sho w th e modelle d an d observed stratigraphy , an d th e basemen t sub sidence of transects 1 and 2 . The Cretaceous an d Tertiary stratigraph y i s identicall y reproduce d during modelling . Also , th e Jurassi c strat a ar e generally in very good agreement, although some discrepancies occu r i n area s tha t hav e bee n up lifted an d eroded . Thi s i s eviden t fo r th e Gull faks are a an d th e wester n ri m o f th e Hord a Platform i n transec t 1 and toward s th e 0ygar den Faul t Zon e i n both transects . Th e modelled and observe d Triassi c strat a ar e mostl y i n concert, althoug h a misfi t o f 300-40 0 m occur s

91

locally belo w th e Gullfak s bloc k i n transec t 1 and below the Horda Platform i n transect 2 . The good agreemen t a t to p basemen t leve l i s em phasized, a s previou s definition s of thi s surface have been affecte d b y considerable uncertaintie s (Klemperer 1988 ; Fichle r & Hosper s 1990 ; Marsden e t al . 1990) . A notio n o f th e apparen t thickness variatio n o f th e Jurassi c an d pre Jurassic sediment distributio n fro m th e Gullfaks block t o th e Vikin g Grabe n (Fig . 6 ) i s neces sary. Thi s portio n o f transec t 1 ha s been * backstripped b y Fosse n e t al . (2000) , an d a t Rogaland Research , an d th e result s hav e bee n consistent whe n decompactio n o f th e sediment s is take n int o account . In area s wher e interpretation o f the stratigra phy is uncertain or no interpretation is available, the model can sugges t a first-order stratigraphy. This i s eviden t i n transec t 2 withi n th e deepes t parts o f th e Vikin g Grabe n an d th e wester n Horda Platform , wher e th e position s o f th e Triassic an d pre-Triassi c horizon s ar e suggeste d by th e forwar d model. Here , to p basemen t level is furthe r constraine d b y gravimetri c an d mag netic data (Christiansso n e t al . 2000). Calculated tectoni c an d basemen t subsidenc e is demonstrated fo r differen t position s along th e modelled transect s (Fig s 6 a an d 7a) . Th e base ment subsidenc e curve s quantif y th e relativ e importance o f th e tw o rif t phases , an d sho w that th e subsidenc e associate d wit h th e Permo Triassic stretchin g was generall y larger tha n fo r Jurassic stretching . I t i s onl y withi n th e Viking Graben tha t Jurassi c subsidenc e outpace d th e Permo-Triassic event . Th e tota l basemen t sub sidence reache s a maximu m o f 12-1 4 km within the deepes t par t o f th e Vikin g Graben . Th e calculated tectonic subsidence shows the flexural and therma l response , wit h it s hanging-wal l subsidence an d footwal l uplif t i n respons e t o fault movement . Exact estimate s o f th e present-da y crusta l thickness ar e importan t durin g forwar d model ling. Thi s constrain s no t onl y th e fina l strati graphic an d crusta l model , bu t als o influence s the parameter s use d durin g modelling. Figure 8 displays th e modelle d crusta l structur e o f transects 1 and 2 . Th e observe d an d modelle d crustal thicknes s ar e i n concer t belo w mos t o f the platform areas. This is important as these are the area s wher e th e crusta l thicknes s i s bes t constrained. Belo w th e basi n area s i n bot h transects, an d belo w th e wester n Hord a Plat form i n transect 2 modelled crust i s thicker than is observe d (locall y u p t o 9 km). According t o Christiansso n e t al . (2000) , refraction dat a (ESP ) clos e t o transec t 1 an d gravity data , sho w tha t th e bas e o f th e crus t

92

T. ODINSE N E T AL .

Fig. 6 . (a ) Modelle d stratigraph y of transec t I . Boxe s a t to p sho w calculate d tectoni c (dotted lines ) an d basement (continuou s lines) subsidenc e alon g th e modelle d transect , (b ) Observed stratigraph y of transec t I . (Note vertica l exaggeration. )

coincides wit h th e bas e o f th e reflectiv e lowe r crust i n bot h transects . A particula r proble m exists belo w th e easter n par t o f th e Vikin g Graben i n transec t 1 an d belo w th e Hord a Platform i n transec t 2 wher e th e Moh o split s into tw o separat e reflectiv e band s (Fig s 2 and 3 ) (Odinsen e l ai 2000) . I n transec t 1 th e up per ban d deepen s fro m 2 0 t o 2 8 km. an d th e lower ban d deepen s fro m 2 4 to 36k m t o th e SE. Here, th e uppe r ban d delineate s velocitie s o f 6+kms" 1 abov e an d 8+kms" 1 belo w (Chris tiansson e t al. 2000) , whic h correspond s t o th e seismic definitio n o f th e Moho . However , thi s

solution wil l plac e the Moh o a t th e sam e or eve n shallower leve l belo w th e norther n Hord a Plat form tha n underneat h th e Vikin g Graben . implying a tota l crusta l stretchin g of abou t th e same orde r o f magnitud e fro m th e Vikin g Gra ben to the 0ygarden Faul t Zone . However , bot h sediment distributio n an d subsidenc e curve s contradict this . Thus , modellin g result s sho w that bul k stretchin g decrease s eastward s fro m the Vikin g Graben. whic h woul d b e compatibl e with a n increas e i n th e crusta l thickness . Hence , when th e bul k modelle d stretchin g is inverted to Moho depth , on e finds that th e modelled Moh o

NORTH SE A PERMO-TRIASSI C AN D JURASSI C EXTENSIO N

93

Fig. 1 . (a ) Modelle d stratigraph y of transec t 2 . Proposed Permo-Triassi c sequenc e emplace d fro m th e Horda Platfor m t o th e Vikin g Graben. Boxe s at to p sho w calculate d tectonic (dotte d lines ) and basemen t (continuous lines ) subsidence alon g th e modelle d transect , (b ) Observed stratigraph y o f transec t 2 . (Note vertical exaggeration. )

corresponds t o th e lowe r o f th e tw o reflectiv e bands (Fig . 8a) . I t i s unclear ho w fa r sout h th e 50km wid e and 5-1 0 km thic k bod y continues , as n o velocit y dat a ar e availabl e sout h o f transect 1 . However , i t i s note d tha t th e sam e lower-crustal spli t ca n b e interprete d fro m th e crustal reflectivit y patter n i n transect 2 (Odinsen et al 2000) , an d tha t th e modelle d Moh o leve l (Fig. 8b ) coincides with that o f transect 1 . Christiansson e t al . (2000 ) obtaine d th e sam e result s •

from gravit y modelling o f th e crusta l transects . They als o addresse d th e natur e o f thi s bod y i n more detail. Stretching distribution and thermal evolution Estimating th e stretchin g distributio n i n th e northern Nort h Se a ha s bee n th e subjec t o f many paper s i n recen t year s (Beac h e t al . 1987;

94

T. ODINSE N E T AL .

Fig. 8. (a ) Modelle d present-da y crusta l structure , an d observe d an d modelle d Moh o depth s fo r transec t 1 . Top o f high-velocit y bod y shoul d b e note d (fro m Christiansso n e t al. 2000). (b ) Modelled present-da y crusta l structure, an d observe d an d modelle d Moh o depth s fo r transec t 2 . It shoul d b e noted tha t th e modelle d Moho leve l coincide s wit h th e bas e o f th e reflectiv e lowe r crus t (simila r t o transec t 1) . Giltner 1987 ; Badley e t al . 1988 ; Marsde n e t al . 1990; White 1990 ; Roberts e t al . 1993 , 1995 ; Ter Voorde e t al. 2000), and hav e yielde d a rang e o f magnitudes fo r th e specifi c areas . Th e larg e variations ar e dependen t o n th e metho d use d (summation o f faul t heaves , forwar d modelling , backstripping, measurement s o f crusta l thin ning), th e cross-sectiona l area , an d th e inter pretation fo r tha t specifi c area . Thi s als o indicates ther e ma y no t b e a uniqu e /3 solution , as ou r understandin g an d simpl e measuremen t of stretching will never b e able t o reflec t th e tru e dynamics o f a n extendin g crust . Recognizing the genera l limitations of stretching estimates , th e presen t forwar d mode l com pares th e magnitud e o f th e Permo-Triassi c an d Jurassic rif t phase s (an d th e coincidental therma l evolution) i n on e mode l run , withou t havin g t o change metho d o r approac h durin g modelling . The 3 value s ar e sample d eac h 5k m alon g th e transects, an d mirro r th e crusta l stretchin g i n a 'spiky' manne r simila r t o Marsde n e t al . (1990 , figs 2.1 5 an d 2.16) . However , unlik e Marsde n et al . (1990 ) an d other s (Robert s e t al . 1993 , 1995; Ter Vord e e t al. 2000), we do no t decoupl e

the lowe r fro m th e uppe r crust , althoug h w e likewise assum e a ductil e deformatio n mechan ism i n th e lowe r crust . Hence , th e presen t 3 profiles ar e no t presente d a s smoot h pur e shea r profiles, a s the y ar e mean t t o resembl e the fault controlled topograph y wher e laterall y adjacen t flexural loads ar e incorporated .

Late early Permian-early mid-Triassic stretching (Artinaskian (261 Ma)-Ladinian (2 36 Ma)) Figure 9 show s th e modelle d latera l stretch ing distributio n alon g th e tw o transects . Th e 3 distribution fo r Permo-Triassi c riftin g indicate s that stretchin g wa s distributed i n a broad basin , delineated t o th e eas t b y th e 0ygarde n Faul t Zone and t o th e west by the Hutto n Faul t align ment i n transec t 1 and th e Shetlan d Platfor m i n transect 2 . Th e basi n i s define d b y a serie s o f large tilte d faul t blocks , whic h ar e bes t image d below the Horda Platform an d the Tampen Spu r area. Th e block-boundin g norma l fault s hav e a

NORTH SE A PERMO-TRIASSI C AN D JURASSI C EXTENSIO N

95

Fig. 9 . (a ) Modelle d latera l stretchin g distribution o f transec t 1 . (b) Modelled latera l stretchin g distribution of transec t 2 . spacing o f 15-2 0 km an d throw s u p t o 4- 5 km (Faerseth e t al, 1995) . Th e orientatio n o f th e faults wa s primaril y th e sam e a s fo r th e Juras sic faults, c . N-S belo w the Horda Platfor m an d the Vikin g Graben , changin g t o NNE-SS W below th e Shetlan d Basin . The modellin g procedur e implie s tha t a n initial Permian crust of lateral uniform thickness existed. W e believ e tha t 35k m i s probabl y a reasonable initia l thicknes s i f w e conside r th e crust i n adjacen t area s tha t ar e littl e affecte d b y the Permo-Triassi c an d Jurassi c stretching . This is also in accordance with other studies (Sellevoll 1977; McCla y e t al . 1986 ; Kinc k e t al . 1991) . Furthermore, test s hav e show n tha t changin g the initia l value fro m 35k m t o 3 0 or 40k m wil l have som e impac t o n th e modelle d stretchin g estimates, bu t wil l no t alte r th e relativ e importance o f th e tw o rif t phases . Tabl e 3 give s an overvie w o f th e modelle d mea n stretchin g values fo r individua l semi-regiona l area s alon g the tw o transects . The modelle d /? mean fo r th e Permo-Triassi c stretching i s 1.2 7 (transec t 1) , an d 1.1 9 (tran sect 2) . I f w e conside r onl y th e interio r o f thi s basin /3 mean become s 1. 4 in transec t 1 (from th e 0ygarden Faul t Zon e t o the Hutton alignment), and 1.3 8 in transect 2 (from the 0ygarden Fault Zone t o th e East Shetlan d Platform) . Th e width of thi s basi n ma y hav e bee n c . 120-125 km

when riftin g ende d i n earl y mid-Triassi c time . Anean fo r th e Hord a Platfor m i s 1.3 3 an d 1.3 9 in transect s 1 an d 2 respectively . Thes e value s are ver y simila r t o thos e obtaine d b y Robert s et al. (1995), who analysed an east-west oriented profile acros s th e Hord a Platform . Thei r pro file wa s locate d sout h o f transec t 1 . Th e /? mean values acros s th e 'Permo-Triassic ' Vikin g Gra ben ar e 1.4 1 i n transec t 1 an d 1.28(? ) i n tran sect 2 (Table 3) . Table 3 . Modelled (3 mean estimates fo r th e PermoTriassic and Jurassic rift phases Area Transect I Horda Platfor m Viking Grabe n Western Eas t Shetland Basi n East Shetlan d Basin Whole transec t Transect 2 Horda Platfor m Viking Graben Shetland Platfor m Whole transec t

Permo-Triassic Jurassic riftin g rifting mea n (3 mean (3 1.33 1.41 1.14

1.08 1.42 1.13

1.29

1.11

1.27

1.15

1.39 1.28(?) 1.0(?) 1.19

1.13 1.53 1.03(?) 1.19

96

T. ODINSE N E T AL .

Permo-Triassic /3 mean fo r th e Eas t Shetlan d Basin i s 1.29 . Here , th e Hutto n alignmen t ha d a simila r statu s t o th e 0ygarde n Faul t Zon e o n the eas t flan k o f th e grabe n system . Model ling result s indicat e a basemen t subsidenc e o f c. 1 km i n the hangin g wal l of th e Hutto n align ment (Fig . 6a) . Th e are a northwes t o f th e Hut ton alignmen t i s considerably les s influence d by Permo-Triassic stretching , wit h a /3 mean reduce d to 1.14 . The modellin g doe s no t recor d an y Permo Triassic stretchin g history fo r th e Shetlan d Plat form. Here , Tertiar y sediment s unconformabl y overlie a thick sequence o f Palaeozoic sediments , which agai n ar e locall y drape d b y a ver y thi n Permo-Triassic sequence . Th e thick pre-Permia n sequence i s supporte d b y gravimetri c an d mag netic dat a (Christiansso n e t aL 2000 ) Accordin g to Plat t (1995) , thes e Palaeozoi c sediment s ar e most likel y o f Devonian(-Carboniferous? ) age . The absenc e o f Mesozoic strat a o n th e Shetlan d Platform suggest s tha t erosio n ha s occurred , initiated b y footwal l uplif t o f th e boundin g fault syste m (Robert s & Yieldin g 1991) . Ou r

modelling analysi s also show s tha t th e Shetlan d Platform becam e elevate d (an d eroded ) durin g rifting. I t i s uncertai n ho w larg e a volum e o f sediment wa s erode d durin g lon g period s o f elevation, an d conversel y ho w muc h sedimen t accumulated durin g th e intervenin g periods o f deposition. Jurassic stretching (Bathonian (165 Ma)Berriasian (141 Ma)) Overall Jurassi c stretchin g amounts t o approxi mately th e sam e i n transec t 1 (/3mean i s 1.15 ) an d transect 2 (/3mean is 1.19) . Although this is similar to th e Permo-Triassi c result s fo r transec t 2 i t is much lowe r fo r transec t 1 (Table 3) . The Jurassic 3 mQan across th e Horda Platfor m amounts t o 1.0 8 along transec t 1 and 1.1 3 alon g transect 2 . I n othe r words , Jurassi c stretchin g for th e Hord a Platfor m are a wa s substantially less tha n th e Permo-Triassi c phas e (1.3 3 an d 1.39) for the same area. Estimate d Jurassi c 3mcan across the Viking Graben i s 1.42 in transect 1 and

Fig. 10 . Modelle d therma l evolution of th e lithospher e along transect 1 from th e Triassi c period t o present . The bas e o f th e lithospher e is described b y th e 130 0 C isother m (contou r interval i s 10 0 C). Latera l EET valu e is displayed.

NORTH SE A PERMO-TRIASSI C AN D JURASSI C EXTENSIO N 1.53 i n transec t 2 . Thi s i s highe r tha n th e les s constrained calculate d Permo-Triassi c values . Anean i n th e Eas t Shetlan d Basi n i s 1.11 , whic h is muc h les s tha n th e Permo-Triassi c stretchin g of 1.29 . Stretchin g acros s the Shetlan d Platfor m (transect 2 ) is merely 1.03 . Modelled thermal evolution of transect 1

Modelled therma l evolutio n an d th e EE T valu e in transec t 1 ar e show n i n Fig . 10 . Maximu m thermal perturbatio n a t th e en d o f th e Permo Triassic rif t phas e affecte d a muc h wide r par t o f the lithosphere tha n Jurassic extensio n (Fig s lO d and f) . Th e latera l exten t o f th e riftin g i s expected t o b e mirrore d i n th e post-rif t sequence s i n Figs 6 and 7 ; these figures demonstrate tha t th e thermal subsidenc e tha t followe d th e firs t phas e produced a broade r an d mor e uniforml y dis tributed post-rif t sequenc e tha n th e latte r phase . It i s als o note d tha t ther e i s a shif t toward s the eas t an d a focusin g o f th e pea k therma l elevation fro m th e Permo-Triassi c t o th e Jur assic rif t phase . The EET , whic h i s controlle d b y th e 450° C isotherm, starte d ou t wit h a n initia l valu e o f 44 km. A t th e en d o f Permo-Triassi c stretching , EET varie d fro m 4 2 km t o 3 0 km belo w th e platform an d basi n areas, respectivel y (Fig. lOf) . At th e onse t o f th e Jurassi c stretching , EE T is increase d t o 38k m belo w th e basi n area s (Fig. lOe) . EE T wa s reduce d afte r th e Juras sic stretching, to 40 km below the platform areas, and 30k m belo w th e Vikin g Graben . Thes e estimates are very close to those derived from th e Baikal an d Eas t Africa n rift s (EE T c. 30km) (Ebinger e t al. 1991 ; Va n de r Bee k 1997) . Several studie s hav e emphasize d tha t pos t Permo-Triassic therma l subsidenc e wa s no t ended whe n Jurassi c riftin g starte d (Giltne r 1987; Gabrielse n e t al . 1990 ; Robert s e t al . 1995), suggestin g tha t th e Jurassi c norther n North Se a wa s a hybri d basi n affecte d simulta neously b y therma l subsidenc e an d renewe d initiation o f stretching . Latera l heat-flo w mod elling o f th e temperatur e structur e for transec t 1 supports thi s conclusion (Fig . lOe) . The residua l Permo-Triassic hea t a t th e onse t o f th e Jurassi c rift phas e indicate s tha t ther e wa s a potentia l for mino r therma l subsidenc e befor e therma l equilibrium. The elevate d therma l perturbation , an d th e reduced EE T (fro m 4 0 km t o 3 8 km), fro m Paleocene t o Eocen e tim e (Fig s lO b an d c) , is a response t o introductio n o f a mino r stretchin g value o f 1.02-1.0 3 durin g Paleocen e time . Thi s elevates th e temperatur e fiel d slightl y an d in -

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creases th e subsidenc e i n Paleocen e an d Eocen e time, a s recorde d i n th e wel l dat a an d seismi c cross-sections. Th e additiona l post-Paleocen e thermal coolin g gav e enoug h accommodatio n space t o reproduc e th e observe d stratigraph y without introducin g unrealisticall y larg e wate r depths. Introducin g thi s (3 valu e wil l obviousl y contribute to the thinning of the crust. Although this crusta l thinnin g i s speculative , i t wil l no t have an y impac t o n th e crustal-scal e modellin g that i s addresse d below . Wit h respec t t o th e modelled Permo-Triassi c an d Jurassi c stretchin g results, th e Paleocen e value s are withi n th e limi t of uncertaintie s o n th e regiona l scal e tha t i s studied here . It i s noted tha t recen t studie s see m t o confirm that a mil d therma l even t occurre d i n th e northern Nort h Se a Basi n i n earl y Tertiar y time (Hall & White 1994 ; Nadin & Kuznir 1995 ; Nadin e t al . 1995) , an d variou s explanation s fo r this departur e fro m McKenzies " (1978 ) post-rif t subsidence mode l hav e bee n offered . Accordin g to Bertra m & Milto n (1989 ) an d Nadi n & Kusznir (1995 ) n o othe r mechanis m i s require d for th e formatio n o f Tertiar y accommodatio n space tha n post-Jurassi c therma l subsidenc e 'buffered' b y a transien t Paleocen e uplif t event . On th e othe r hand , Hal l & White (1994 ) argue d that a rapi d increas e i n water-loaded subsidenc e occurred a t th e beginnin g of Tertiar y time . Thi s increase, u p t o 1 km i n amplitude , i s anomalou s in tha t i t i s no t predicte d b y th e lithospheri c stretching model , unles s a n additiona l phas e of Tertiar y stretchin g i s invoked . Ther e is , however, onl y limite d evidenc e fo r sufficien t Tertiary norma l stretching . Th e present-da y thermal statu s show s tha t th e lithospher e ha s almost reache d therma l equilibrium . Hence , th e EET show s a n almos t unifor m valu e o f 4 2 km (Fig. lOa) . Comments o n erosion , wate r dept h an d tectonic uplif t The proble m o f reproducin g area s tha t hav e been elevate d abov e se a leve l an d erode d i s present i n al l stratigraphi c modelling . On e wa y of dealin g wit h thi s proble m usin g th e 2 D for ward modellin g tool i s to decreas e th e stretching values an d thereb y elevat e th e specifi c areas . This wil l reduc e th e stratigraphi c misfi t tha t appears i n area s tha t ha s bee n uplifte d an d eroded. W e hav e chose n t o d o thi s fo r th e Gullfaks bloc k an d the western rim of the Horda Platform i n transect 1 and clos e t o th e 0ygarden Fault Zon e i n bot h transects . I t i s emphasize d that thi s reductio n o f th e Jurassi c stretchin g

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T. ODINSE N E T AL .

values (<1.1 ) is local, an d th e ne t effec t i s negligible o n th e regiona l an d semi-regiona l scale s relevant i n the discussion o f the present analysis. Trying t o reproduc e th e Triassi c sequenc e poses particula r problems , whic h ar e partl y reflected i n th e large r stratigraphi c misfi t (Fig ures 6 an d 7) . Becaus e o f th e larg e Permo Triassic stretchin g values , whic h lowe r th e dee p stratigraphy significantly , th e proble m her e i s not primaril y elevation and erosio n a s describe d for th e Jurassi c sequenc e above . I t i s relate d t o the model , whic h predict s a thicke r Triassi c post-rift sequenc e tha n i s observed, a s i t simplifies the condition s fo r sedimen t preservatio n b y assuming tha t bas e leve l i s equa l t o se a level . If so, this would probably have produced a much thicker Triassi c sedimentar y coloum n tha n i s observed. Wher e Triassi c unit s ar e encountere d in well s i n th e area , the y ar e deposite d int o terminal floo d basins , ephemera l lake s an d alluvial plain s (Hid e 1989 ; Stee l 1993 ; Nystue n & Fait 1995) . The result is the formation of many bypass surfaces , alon g whic h sediment s wer e transported downslop e int o th e basi n t o th e ultimate bas e level , thu s givin g ris e t o a muc h thinner continenta l Triassi c successio n tha n predicted b y th e model . A s th e mode l assume s marine sedimentatio n u p t o a specifie d wate r

depth, w e chos e t o compensat e fo r thes e pro blems by introducing an enhanced Triassi c wate r depth, which decreases th e thickness of the lower and middl e Triassi c sediment s (Fig . 11) . I t i s important t o not e tha t w e chose t o reduc e th e stratigraphic misfit t o avoi d modellin g problems for th e post-Triassi c sequence . In general , water dept h ma y b e a critica l fac tor i n the subsidence and buria l analysis because it assesse s th e amoun t o f sedimen t underfill . An overestimatio n o f palaeo-waterdept h ma y induce artificia l uplif t throug h tim e in th e subsi dence model , wherea s underestimatio n artifi cially reduce s subsidenc e rate . However , ou r introduction o f enhance d palaeo-waterdepth s actually provide s a bette r stretchin g estimate , because i t allow s u s t o asses s th e subsidenc e relative t o a non-zer o bas e level . When movin g fro m a continenta l Triassi c environment t o a marin e post-Triassi c environ ment, the general water depth shallows to zero at the onset o f Jurassic time, and increase s again in late Jurassi c t o earl y Cretaceou s time . I n th e deepest par t o f th e Vikin g Graben wate r depth s reach a maximum of 2000m i n early Cretaceou s time. I n th e Shetlan d Basi n an d o n th e Hord a Platform th e dept h i s close t o zer o durin g early Cretaceous time . I n previou s studies , Nadi n &

Fig. 11 . Wate r depth estimates for a syntheti c well in the central part o f the Vikin g Graben i n transect 1 . Stars (*) indicate enhance d Triassi c wate r dept h estimate s use d t o matc h th e observe d stratigraphy.

NORTH SE A PERMO-TRIASSI C AN D JURASSI C EXTENSIO N Kuznir (1995 ) predicte d axia l wate r depth s in the Vikin g Grabe n o f u p t o 1300 m fo r earl y Cretaceous tim e an d 900 m fo r th e lat e Cretac eous time , wherea s Bertra m & Milto n (1989 ) and Bar r (1991 ) suggeste d a Cretaceou s wate r depth o f th e orde r o f 1000 m fo r th e norther n North Se a Basin. A t th e onse t o f Tertiar y tim e the wate r depth s varie d betwee n 50 0 and 800 m across th e basi n (Fig . 11), whic h i s i n accor dance wit h faunal trends (Gradstein e t al. 1994) . From thi s tim e th e wate r dept h shallow s to th e present values . Uplift o f th e Norwegia n mainlan d an d adja cent areas in Tertiary tim e (Dore 1992 ) will have some impac t o n th e stretchin g estimates , a s these ar e calculate d a s a functio n of subsidence. This modelling problem i s mostly limite d to th e Horda Platform . I t i s difficul t t o evaluat e pre cisely ho w larg e thi s effec t is , a s th e magnitud e of uplif t i s affiliate d wit h uncertainties . W e ob serve tha t th e wester n limi t o f th e norther n Horda Platfor m (transec t 1 ) i s situate d muc h further eas t tha n see n t o th e sout h (transec t 2) . Therefore uplif t o f wester n Norwa y ha s als o influenced th e norther n par t o f th e Hord a Platform more than the southern part. If we take into accoun t th e ver y fe w Jurassi c fault s see n along the southern Horda Platform, it is reasonable t o expect tha t stretchin g durin g thi s phas e was a t leas t a s great t o th e north . Acknowledg ing this , th e importanc e o f th e Permo-Triassi c stretching with respect t o th e Jurassic phase wil l still b e th e same . Implications o f crusta l thicknes s estimates and th e flexura l strengt h o f th e lithospher e The misfi t betwee n th e observe d an d modelle d Moho leve l below th e basi n area s i n bot h tran sects, an d belo w the wester n Horda Platfor m in transect 2 (Fig . 7) , offer s a t leas t tw o explana tions. First , on e ha s t o tak e int o accoun t tha t Permo-Triassic stretchin g ma y hav e bee n muc h larger i n th e basi n area s tha n ou r result s show . This is not likely , because th e ver y high /3 values needed woul d reduc e th e crusta l thicknes s t o a few kilometres , whic h agai n woul d contradic t the geophysica l dat a (Christiansso n e t al. 2000). The second an d probably th e most reasonabl e explanation deals with uncertainties of the initial crustal configuration . Whe n th e riftin g wa s initiated i n Permia n tim e (se t t o 26 1 Ma), th e crust ha s t o b e define d wit h a unifor m latera l thickness i n the mode l (w e use 35km) . This i s a simple approac h wher e th e influenc e o f th e unconstrained Devonian(-Carboniferous? ) crus tal stretchin g i s no t considered . Accordin g t o

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Fossen & Rykkelid (1992) , th e Devonian exten sion o f th e thickene d Caledonia n Nort h Se a crust was much more important tha n th e PermoTriassic an d Jurassi c extension . W e limi t th e investigation o n th e Permia n crusta l configuration t o changin g th e initia l thicknes s durin g modelling. Thes e test s sho w tha t reducin g the initia l crustal thicknes s fro m 35k m t o 28 30km give s a minimu m misfi t betwee n th e modelled an d observe d Moh o depth s belo w th e basin areas , indicatin g tha t th e Devonian( Carboniferous?) stretchin g di d no t produc e a crust wit h a unifor m latera l thickness . I t ma y have varied fro m c . 35 km i n th e platfor m area s to less than 30k m in the basin areas, if we accept that thi s crusta l configuratio n laste d unti l th e onset o f th e Permo-Triassi c rifting . It i s acknowledge d tha t Devonian(-Carbon iferous?) sediment s probably cove r basemen t o n the platfor m area s o f th e norther n Nort h Se a (McClay e t al 1986 ; Ziegler 1990 ; Faerset h et a l 1995; Platt 1995) , although their presence within the deepe r part s o f th e basi n i s more uncertain . As th e resolutio n o f th e presen t dat a doe s no t allow a furthe r investigation o n thi s subject , w e cannot full y integrat e th e pre-Permia n tectoni c evolution an d sedimentar y distributio n i n th e modelling schem e at present . However , we realize tha t thes e factor s ma y hav e som e impac t on th e discussio n o f th e Permo-Triassi c stretch ing estimate s an d therma l configuration , a s seen abov e fo r th e modelle d Moh o level . Th e Devonian(-Carboniferous?) extensio n may , fo r example, hav e a simila r therma l impac t o n th e Permo-Triassic phas e t o tha t indicate d fo r the Permo-Triassi c phas e o n th e Jurassi c ther mal evolution , dependin g o n whe n th e pre Permian stretchin g ended. This is highly relevant on th e Shetlan d Platfor m (transec t 2) , wher e a considerabl e sedimentar y colum n o f pre Permian age covers the basement in half-grabens (Platt 1995) . The ag e a t whic h thi s topograph y was generate d i s generall y unknown . W e can , however, conside r to p basemen t a s a tim e lin e (261 Ma) i n the model. Thi s i s also a convenient approach fo r th e Shetlan d Platform , wher e 'top basement' i s defined a s the to p o f thi s sequence. We believ e thi s i s a vali d perspective , a s th e Permian an d post-Permia n stretchin g seem s t o be insignifican t i n thi s area . With a n isother m o f 450° C describin g th e EFT, th e flexura l strengt h o f th e crus t an d mantle ar e alway s coupled . Thi s approac h i s simple an d ha s therefor e bee n studie d i n mor e detail b y Te r Voord e e t a l (2000) , wh o teste d both coupled and decoupled EF T along an E- W transect situate d betwee n ou r transec t 1 and 2 . Those worker s suggeste d tha t th e crus t an d

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mantle hav e bee n decouple d (EE T 6 km) durin g the Permo-Triassi c an d Jurassi c rif t phases . Th e resulting stretchin g estimate s ma y yiel d larg e differences, wher e a couple d flexura l strengt h can giv e 50-60 % les s stretchin g tha n a full y decoupled one . However , th e subjec t i s compli cated a s couple d o r decouple d flexura l strength , and thereby stretching estimates, probably represent end-members , wher e a full y decouple d mode o f flexur e i s likel y t o b e a n exceptiona l case (Buc k 1988 ; Te r Voord e e t al. 1998 , 2000) . Hence, on e ha s t o tak e int o accoun t tha t th e degree, o r mod e o f (de)couplin g i s controlle d by a combinatio n o f age , crusta l compositio n and weaknesses , crusta l thickness , an d strai n rate (Englan d 1983 ; Kuzni r & Park 1987 ; Dun bar & Sawye r 1989 ; Bass i e t al . 1993 ; Buro v & Diamen t 1995 ; Cloeting h & Buro v 1996) . Future focu s o n th e norther n Nort h Sea , fol lowing th e presen t pape r an d th e wor k o f Te r Voorde e t al . (2000) , shoul d therefor e includ e studies o n crusta l rheology , an d pre-Jurassi c thickness and structura l configuration, including an investigatio n on th e interactio n betwee n th e Devonian(-Carboniferous?) structura l settin g and th e Permo-Triassi c structure .

the end o f the Permo-Triassic rif t phas e affec ted a much wider part o f the lithosphere than did th e Jurassi c phase . Thi s i s als o reflecte d in th e post-rif t sedimen t distributio n follow ing the two phases, where the first produces a much mor e unifor m sedimentar y sequenc e than th e latter . There i s also a focusin g an d migration toward s th e eas t o f th e pea k ther mal elevation from the Permo-Triassi c t o th e Jurassic event . We assum e tha t thi s focusing relates t o change s o f the crusta l compositio n or rheolog y an d structura l framewor k fro m the Permo-Triassi c t o th e Jurassi c rif t phase . It i s als o note d tha t th e lithospher e di d not full y reac h a therma l equilibriu m at th e initiation o f Jurassi c stretching . Th e poten tial fo r enhance d subsidenc e i n th e Jurassic , however, seem s to hav e bee n small .

Conclusions

We wis h t o than k A . M . Berg e an d P . Christiansso n for co-operatio n an d discussions , an d J . E . Li e fo r reprocessing th e NSD P lines . We ar e als o gratefu l t o W. Sassi, J. P. Nystuen. J. Skogseid an d an anonymous reviewer fo r thei r constructiv e criticis m an d helpfu l comments, an d t o G . Sulliva n fo r linguisti c correc tions. This work wa s funded by the Commission o f the European Unio n an d the Norwegian Researc h Counci l in th e framewor k o f Integrate d Basi n Studies-Th e Dynamics o f th e Norwegia n Margin . Nors k Hydr o a.s., Saga Petroleu m a.s . and Statoi l a.s. provided dat a and researc h effort s fo r thi s study.

Three mai n topic s wer e th e specifi c subjec t o f our forwar d modellin g o f transect s 1 and 2 .

References

• Estimatio n o f th e Permo-Triassi c an d Jur assic stretchin g shows tha t th e firs t wa s generally muc h mor e importan t tha n th e latter. It i s only acros s th e Vikin g Graben tha t th e Jurassic stretchin g exceed s th e olde r event . The 0ygarde n Faul t Zon e an d th e Hutto n alignment appea r t o hav e bee n th e principa l structures o f th e Permo-Triassi c basin . • Th e modelled stratigraph y i s generally similar t o th e observe d one . Mos t o f th e discrepancies ar e limite d t o th e Triassi c sedimentary sequence . Th e modelle d an d observed Mon o relie f show s a minimu m misfit belo w mos t o f th e platfor m areas . Here, th e modelle d Moh o leve l coincide s with th e bas e o f th e reflectiv e lowe r crust . The misfi t increase s below th e basi n area s in both transects , an d belo w the western Hord a Platform i n transec t 2 . Thi s misfi t ma y b e ascribed t o a laterall y varyin g crusta l thick ness a t th e onse t o f th e Permo-Triassi c lithospheric stretching , fro m c . 35 km i n th e platform areas to 28-30 km in the basin areas . • Th e lithospheri c therma l evolutio n show s that th e maximu m therma l perturbatio n a t

BADLEY, M . E. . PRICE , J . D. , RAMBECH-DAHL . C. & AGDESTEIN, T . 1988 . The structura l evolutio n o f the northern Vikin g Graben and it s bearing upo n extensional mode s o f basi n formation . Journal o f the Geological Society, London, 145 , 455-472. BARR, D . 1991 . Subsidence an d sedimentatio n in semistarved half-graben : a mode l base d o n Nort h Se a data. In : ROBERTS , A. M. , YIELDING , G. & FREE MAN, B . (eds ) The Geometry o f Normal Faults. Geological Society , London, Specia l Publications. 56, 17-28 . BARTON, P . & WOOD . R . 1984 . Tectonic evolutio n of the Nort h Se a basin : crusta l stretchin g an d subsidence. Geophysical Journal o f th e Royal Astronomical Society, 79 . 987-1022. BASSI, G., KEEN . C. E . & POTTER, P. 1993 . Contrasting styles o f rifting : model s an d example s fro m th e eastern Canadia n Margin . Tectonics, 12 , 639-655. BEACH. A . 1986 . Som e comment s o n sedimentar y basin developmen t i n th e norther n Nort h Sea . Scottish Journal o f Geology, 22 , 1-20. , BIRD , T. & GIBBS, A. 1987 . Extensional tectonics and crusta l structure : deep seismi c reflection dat a from th e norther n Nort h Se a Vikin g Graben . In : COWARD, M . P. . DEWEY , J. F . & HANCOCK . P. L . (eds) Continental Extensional Tectonics. Geologi cal Society , London . Specia l Publications . 28 . 467-476.

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(ed.) Petroleum Exploration an d Exploitation i n Norway. Norwegia n Petroleu m Societ y Specia l Publication, 4 , 135-179 . NOTTVEDT, A. , GABRIELSEN , R . H . & STEEL . R . J . 1995. Tectonostratigraph y and sedimentar y architecture o f rif t basin s wit h referenc e t o th e north ern Nort h Sea . Marine an d Petroleum Geology. 12. 881-901. ODINSEN. T. . CHRISTIANSSON . P. . GABRIELSEN . R. H. . FALEIDE. J . I . & BERGE , A . M . 2000 . Th e geometries an d dee p structur e o f th e norther n North Se a rif t system . This volume. PLATT. N . H . 1995 . Structur e an d tectonic s o f th e northern Nort h Sea : ne w insight s fro m dee p penetration regiona l seismi c data . In : LAMBIASE . J. J . (ed. ) Hydrocarbon Habitat i n Rift Basins. Geological Society , London. Specia l Publications . 80, 103-113 . RESTON, T . J . 1990 . Th e lowe r crus t an d th e exten sion o f th e continenta l lithosphere : kinemati c analysis o f BIRP S dee p seismi c data . Tectonics. 9, 1235-1248 . ROBERTS, A . M . & YIELDING . G. 1991 . Deformatio n around basin-margi n fault s i n the North Se a midNorway rift . In : ROBERTS, A . M. . YIELDING. G. & FREEMAN, B . (eds ) Th e Geometry o f Normal Faults. Geologica l Society , London , Specia l Pub lications, 56 , 61-78. ,, KUSZNIR . N . J. , WALKER , I . & DORN LOPEZ, D . 1993 . Mesozoic extensio n i n th e Nort h Sea: constraint s fro m flexura l backstripping . forward modellin g an d faul t populations . /// : PARKER, J . R . (ed. ) Petroleum Geology o f Northwest Europe: Proceedings of the 4th Conference. Geological Society . London , 1123-1136 . ,, , & 1995 . Quantitativ e analysis o f Triassi c extensio n i n th e norther n North Sea . Journal o f th e Geological Society, London, 152 , 15-26 . ROYDEN, L . & KEEN. C. E . 1980 . Rifting processe s an d thermal evolution of the continental margin of eastern Canad a determine d fro m subsidenc e curves . Earth an d Planetary Science Letters, 51. 343-361. SCLATER, J . G . & CHRISTIE , P . A . F . 1980 . Conti nental stretching : an explanatio n o f the pos t mid Cretaceous subsidenc e o f th e centra l Nort h Sea Basin . Journal o f Geophysical Research. 85 . 3711-3739. SELLEVOLL, M . A . 1977 . Moho beneat h Fennoscandi a and adjacen t part s o f th e Norwegia n Se a and th e North Sea . Tectonophysics. 20 , 359-366. SPADINI, G. . CLOETINGH . S . & BERTOTTI . G . 1995 . Thermo-mechanical modellin g o f th e Tyrrhenia n Sea: lithospheri c neckin g an d kinematic s o f rifting. Tectonics. 14 . 629-644. STECKLER, M . S . 1981 . Th e thermal an d mechanical evolution of Atlantic-type continental margins. PhD thesis , Columbia University , Palisades . NY . STEEL, R . J . 1993 . Triassic-Jurassi c megasequenc e stratigraphy i n th e Norther n Nort h Sea : rif t t o post-rift evolution . In : PARKER . J . R . (ed. ) Petroleum Geology of Northwest Europe: Proceedings o f th e 4t h Conference. Geologica l Society . London, 299-315 .

NORTH SE A PERMO-TRIASSI C AN D JURASSI C EXTENSIO N TER VOORDE, M. , F^ERSETH , R. B. , GABRIELSEN, R. H . & CLOETINGH , S . A . P . L . 2000 . Repeate d litho sphere extensio n i n th e norther n Vikin g Graben : a coupled o r a decouple d rheology ? This volume. , VAN BALEN, R . T. , BERTOTTI , G . & CLOETINGH , S. A . P . L . 1998 . Th e influenc e o f a stratifie d rheology o n th e flexura l respons e o f th e litho sphere t o (un)loadin g b y extensiona l faulting . Geophysical Journal International, 134 , 721-735 . TORSVIK, T . H. , STURT , B . A., SWENSSON , E., ANDER SEN, T . B . & DEWEY , J . F . 1992 . Palaeomagneti c dating o f faul t rocks : evidenc e fo r Permia n an d Mesozoic movement s an d brittl e deformatio n along the Dalsfjor d Fault , west Norway. Geophysical Journal International, 109 , 565-580 . VAN DER BEEK, P . 1997 . Flank uplif t an d topograph y at the centra l Baika l rif t (S E Siberia) : a tes t of kinematic model s fo r continenta l extension . Tectonics, 16, 122-136 .

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Detachments and low-angle faults in the northern North Sea rif t system HAAKON FOSSEN, 1 TOR E ODINSEN, 1 ROALD B . F^RSETH 2 & RO Y H . GABRIELSEN 1

1

Department of Geology, University of Bergen, Allegaten 41, N-5007 Bergen, Norway 2 Norsk Hydro, 1321 Stabekh, Norway Abstract: Severa l maste r fault s i n th e Nort h Se a basi n ten d t o flatte n t o giv e lo w dip s a t depth, an d i n thi s sens e for m detachment s i n th e rif t system . Suc h lo w angl e fault s ar e identified i n the western flank of the Vikin g Graben (Tampe n Spu r area) , wher e they occu r as bot h intra - an d supra-basemen t detachments . Interferenc e betwee n detachment s an d steeper fault s results in ramp-flat-ramp geometries . I n the eastern part o f the Gullfaks faul t block, a supra-basemen t detachmen t i s probabl y associate d wit h anomalousl y hig h lat e Jurassic extensio n in the Gullfaks Field area . Th e low-angle Gullfaks detachment als o helps explain th e presenc e o f set s o f paralle l east-dippin g fault s (domin o systems) , a commo n feature i n th e collapse d hangin g wal l t o low-angl e detachments . Simila r detachment s probably exis t beneath th e Gullfaks S0r block and S E of the Visund fault block . All of these are interprete d a s lat e Jurassi c collaps e structure s directl y relate d t o activ e lat e Jurassi c extensional tectonics . Stron g indication s o f intra-basemen t detachment s ar e als o foun d i n the Tampe n Spu r area . Thes e detachment s ar e forme d b y major norma l fault s tha t flatte n in the basement , a s seen beneat h th e Visund fault block . This geometr y may t o som e extent be relate d t o faul t rotatio n durin g repeate d phase s o f extensio n i n th e Palaeozoic-Earl y Mesozoic period. However, abrupt flattening of some o f the faults in the basement indicate s that th e master fault s follo w some of the many pre-existing mechanically weak zones i n the basement, primaril y low-angl e Devonia n extensiona l shear zone s o r Caledonia n thrusts .

The mos t commo n hydrocarbo n tra p type s i n the Nort h Se a ar e controlle d o r influence d b y faults, a s delineatin g structure s o f rotate d faul t blocks, a s synsedimentary structure s controllin g the distribution o f source an d reservoi r rocks, o r as barrier s t o flui d flo w (Hardma n & Boot h 1991). Accordingly, faul t geometry i s importan t for understandin g man y oilfield s an d ga s fields in thi s area . Experimental modelling, kinematic considera tions and field examples all indicate that fault s in extended region s ma y b e planar , non-plana r (e.g. listric) , hig h angl e o r lo w angle , o r a combination o f these (e.g. Wernicke & Burchfie l 1982; Gibb s 1984 ; Gabrielse n 1986 ; McCla y & Ellis 1987 ; Fossen & Gabrielsen 1996) . However , there ha s bee n a tendenc y i n th e literatur e t o prefer on e o f thes e type s o f faul t geometrie s when makin g interpretations . Fault s show n a s steep, sub-plana r feature s dominat e th e litera ture, possibl y becaus e downward-flattenin g o r low-angle structure s ar e mor e difficul t t o detec t seismically. However , extensiv e interpretation o f listric faul t geometrie s wa s favoure d b y som e i n the 1980 s (e.g. Beach 1984 ; Gibbs 1984) , and th e discovery o f low-angl e fault s i n extende d terranes (Armstrong 1972 ) also led to consideration

of suc h structure s i n th e sam e tim e perio d (e.g. Wernicke & Burchfiel 1982 ; Lister et al 1986) . In general, a large-scale rif t syste m can b e expected to contai n element s of all of thes e type s of faul t geometries. W e believ e that thi s also hold s tru e for th e norther n Nort h Se a rift system . In thi s paper w e present observation s primar ily base d o n a combinatio n o f commercia l 2 D and 3 D seismi c dat a an d a reprocesse d dee p seismic lin e acros s th e norther n Nort h Se a (NSDP84-1) (Fig . 1) . W e wil l focu s o n th e Gullfaks-Visund-Snorre regio n and on the nonplanar an d low-angl e faul t geometrie s reveale d by th e seismi c data. Downward-flattenin g faults and detachmen t fault s hav e bee n previousl y described sout h o f th e presen t stud y area , where extensiona l detachment s ten d t o for m along th e Zechstei n sal t (Clause n & Korstgar d 1996; Thoma s & Cowar d 1996) . Downward flattening faults ar e also interpreted at deepe r a s well a s highe r level s i n th e Nort h Se a (Beac h 1984; Gibb s 1984 ; Speksnijder 1987 ; Gabrielsen 1988; Graue 1992 ; Platt 1995) . In a complemen tary wor k t o th e presen t contribution , Odinse n et al . (2000 ) sugges t tha t multipl e level s o f detachment, a s wel l a s composit e faul t geome tries, ar e commo n i n th e norther n Nort h Sea.

From'. N0TTVEDT , A . e t al . (eds ) Dynamics o f th e Norwegian Margin. Geologica l Society , London , Specia l Publications, 167 , 105-131. 1-86239-056-8/OO/ S 15.00 © Th e Geologica l Societ y o f Londo n 2000 .

Fig. 1 . Regiona l profil e based o n seismi c lin e NSDP84-1 . Modifie d from Christiansso n ct at. (1999).

NORTH SE A DETACHMENT S AN D LOW-ANGL E FAULT S We define the term detachment as : a low-angle (portion o f a) fault o r shea r zon e tha t separate s rocks tha t sho w differen t amount s o f extensio n and/or deformation styles . By low-angle fault we mean a faul t wit h di p <30° . A listri c faul t i s usually define d a s a spoon-shape d faul t i n thre e dimensions, an d appear s a s a curved , concave upward faul t o n cross-sections . Fault s wit h geometries differen t fro m bot h listri c an d pla nar fault s ar e simpl y referre d t o a s non-plana r faults, wit h additiona l descriptiv e term s suc h as downwar d flattening . Below , w e distinguis h between supra-basemen t detachment s occurrin g in th e sedimentar y sequenc e abov e th e Caledo nian basement, and intra-basement detachments, which partl y o r wholl y ar e confine d t o base ment rocks .

Regional geological development The Nort h Se a rift (Fig . 2 ) is a post-Caledonia n graben syste m wit h a multiphas e extensiona l history tha t starte d wit h Devonian extensio n of the thickened Caledonian crust . This extensional phase affecte d a n are a tha t extend s fa r beyon d the late r (Permo-Triassi c an d Jurassic ) margin s of the North Sea rift syste m (Fossen & Rykkelid

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1992; Fossen 1998) . At least two important post Devonian phase s o f rifting ar e recognize d in th e northern Nort h Sea . The firs t i s a Permo-Triassi c phase , whic h i s not know n in great detail but i s thought t o be of great importance, in terms of both extension and formation o f a structural framework (Gabrielsen et al 1990 ; Faerset h et al. 19950 ; Robert s e t al. 1995). Th e Permo-Triassi c extensio n appear s t o have affecte d a wide r are a tha n th e Jurassi c phase (Gabrielse n e t al . 1990 ; Robert s e t al . 1990, 1995 ; Faerset h et al . 19950 ; Odinse n e t al . 20006). Th e extensiona l fault s definin g th e largest faul t block s i n th e Nort h Se a rif t ar e mostly o f Permo-Triassi c origin , althoug h reactivated i n Jurassi c time . Therma l contractio n and sedimen t loadin g prevaile d throughou t Triassic time , an d aroun d 2- 4 km o f Triassi c sediments ar e deposite d i n th e norther n Viking Graben (e.g . Badle y et al . 1988 ; Steel & Ryseth 1990). A major uplift (erosion ) is recorded i n th e Lower-Middle Jurassi c serie s o f th e centra l North Sea , wher e a majo r rif t dom e ma y hav e been locate d (e.g . Ziegler 1990) . In th e norther n Viking Graben, doming-relate d regression led to the depositio n o f the Bren t Group sandstones . Although loca l faul t activit y i s know n fo r more o r les s the entir e Jurassi c tim e period, i t is

Fig. 2. Regiona l overvie w an d interna l subdivisio n o f th e norther n Nort h Se a rif t system . Interpretation s of lines NSDP84-1 and - 2 are incorporated int o the figure. It shoul d b e noted tha t colour s on section s illustrate sedimentary sequences , whereas map area s ar e colour code d accordin g t o th e topo-tectonic classificatio n suggested b y Gabrielse n (1986) . HSZ , Hardangerfjor d shea r zone .

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well established tha t mos t o f the structural trap s in th e Nort h Se a ar e th e result s o f a lat e Mid to Lat e Jurassi c extensio n phase . Th e rat e o f crustal extension appear s t o hav e increase d sub stantially fro m Mid - t o Lat e Jurassi c time . Th e Shetland Platfor m wa s formed , an d differentia l subsidence wa s associate d wit h th e develop ment an d rotatio n o f larg e faul t blocks . Exten sional structure s relate d t o thi s Jurassi c phas e are clearl y image d o n seismi c sections , an d ar e therefore know n i n more detail tha n th e PermoTriassic ones. Stretching estimates vary from les s than 15 % i n th e platform/sub-platfor m area s t o 40-50% i n th e interio r par t o f th e rif t (Vikin g Graben) (Marsde n e t al . 1990 ; Robert s e t al. 1993; Odinsen e t al. 2000/?). The faul t blocks ex hibit erode d crests , indicatin g that the y were all at o r near se a level during a late (Kimmeridgian) stage o f th e riftin g history . Increase d extensio n rate an d accelerate d faul t movement s resulte d in increased wate r dept h i n lat e Kimmeridgian Volgian time , an d no t unti l thi s tim e di d th e Viking Grabe n becom e th e deepes t structural topographical elemen t o f th e rif t i n th e stud y area (Fig . 1 ) (Gabrielsen e t al . 1990) . The Jurassic Viking Graben di d not develop as a single , straight entity , but a s a syste m o f synrift unit s bounde d b y maste r faults . A syste m of grabe n segment s formed , wher e th e grabe n segments wer e linke d throug h accommodatio n zones as described b y Scott & Rosendahl (1989). The rat e o f extensio n decrease d a t th e Jurassic-Cretaceous transition , an d wa s practi cally terminate d b y Ryazania n time . Therma l and sedimen t loading-relate d subsidenc e influ enced th e entire North Se a until Paleocene time . A genera l ris e i n se a leve l resulte d i n a pro gressive oversteppin g o f the platform an d buria l of Jurassi c faul t block s durin g Cretaceou s tim e (e.g. Ziegle r 1990) , an d th e significan t bathyme try a t th e en d o f th e Jurassi c perio d wa s infille d by marin e Cretaceou s shales .

On the basis of observations i n the North Sea . Gabrielsen (1986 ) propose d a conceptua l topo tectonic mode l fo r extensiona l grabe n system s (Fig. 2) . Thi s mode l allow s fo r non-rotatin g steep plana r faults , rotatin g low-angl e plana r faults an d listri c faults to b e developed a t differ ent stage s o f grabe n formation , an d fo r such faults t o b e simultaneousl y activ e i n differen t parts of the extending graben. Such relationships are als o see n i n analogu e experimenta l model s (Fossen & Gabrielse n 1996) . I n Gabrielsen' s model, the extra-marginal fault syste m separate s the are a tha t ha s undergon e stretchin g from it s undeformed surroundin g terrane . Th e extra marginal faults , whic h ma y constitut e a rela tively small-scal e horst-and-grabe n topography , are stee p structure s tha t wer e activate d a t th e initial stag e of the graben history . The platform , which commonl y i s separate d fro m th e extra marginal faul t syste m b y a mino r horst , repre sents a structura l uni t tha t i s characterize d b y moderate faul t activit y and subsidence . It s width may var y considerabl y alon g th e strik e o f the grabe n margin . On it s graben-ward side , th e platform i s frequentl y bordere d b y a mar ginal platfor m hig h toward s th e heavil y faulted sub-platform, whic h i s characterize d b y a n array o f rotate d faul t blocks . Th e oute r maste r fault syste m separates th e margi n fro m th e interior graben . Following Boswort h (1985 ) an d Rosendah l (1987), w e acknowledg e tha t mos t grabe n sys tems ar e subdivide d int o separat e units , th e geometry o f whic h reflect s along-strik e spatia l interaction o f the maste r faul t system s and thei r interconnections a t depth . Suc h relationship s include transfe r zones o f differen t types , reflect ing th e degree o f overlap an d eventuall y shifting polarities betwee n th e faul t segments .

General structure and low-angle faults of the northern North Se a

The Nort h Se a basi n display s mos t o f th e fea tures describe d above , a s illustrate d i n Fig . 2 . Also, th e interio r grabe n o f th e lat e Jurassic Cretaceous Vikin g Grabe n display s severa l centres o f subsidenc e (Faerset h 1983) , indicatin g that separat e grabe n unit s exist . A simila r pat tern o f grabe n unit s wit h shiftin g polaritie s i s recognized i n th e Permo-Triassi c basi n i n th e present Hord a Platfor m (Scot t & Rosendah l 1989; Gabrielse n e t al . 1990 ; Faerset h e t al . \995a). Finally , comple x an d composit e faul t geometries i n the Nort h Se a have bee n reporte d by severa l worker s (Badle y e t al . 1984 , 1988 ; Spekschnijder 1987) . This ha s led to speculation s

General structure of graben systems Several structura l feature s ar e commo n i n extensional grabe n system s tha t hav e reache d a certain stag e o f developmen t (Johnso n 1930 ; Robson 1971 ; Hardin g 1984) . Severa l studie s suggest tha t suc h feature s reflect th e bul k geom etry o f th e extendin g crus t o n a regiona l scal e (Wernicke 1985 ; Cowar d 1986) , an d comple x fault interactio n a t dept h o n th e semi-regiona l and loca l scale s (Wernick e & Burchfie l 1982) .

Structure of the northern North Sea rift system

NORTH SE A DETACHMENT S AN D LOW-ANGL E FAULT S about whethe r additiona l level s o f detachmen t may hav e bee n activ e durin g it s developmen t (Gabrielsen 1988) . Utilizing reprocesse d dee p reflectio n data , high-quality commercia l seismi c lines and gravi metric an d magneti c data , th e deep structur e o f the norther n Vikin g Grabe n ha s recentl y bee n reinterpreted b y Christiansso n e t ai (2000 ) and Odinsen e t al (2000) . These studie s support th e view tha t th e are a i s truncate d b y a principa l east-dipping crustal-scal e fault , whic h subcrop s along th e easter n margi n o f th e Eas t Shetlan d Basin an d flatten s i n th e highl y reflectiv e lowe r crust beneat h th e wester n borde r o f th e (Jur assic) Vikin g Graben . Intra-mantl e eastward dipping reflection s mappe d beneat h th e Hord a Platform (als o reporte d b y Klempere r (1988) ) may represen t th e continuatio n o f thi s maste r fault. A shallowe r (intra-Triassic ) detachmen t was postulate d fo r th e Gullfak s Fiel d (Fosse n

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1989; Koestle r e t al 1992 ) and fo r th e Cormor ant Fiel d (Speksnijde r 1987) , an d ha s bee n supported i n th e late r wor k b y Odinse n e t al . (2000). A simila r detachmen t wa s observe d below th e Hord a Platform , whic h i s dominate d by antithetica l an d mor e plana r faults . Th e present stud y concentrate s o n analysi s of thes e detachments, and elaborates upon their complex geometries. A transec t acros s th e norther n Nort h Se a (Fig. 1 ) reveals the asymmetrical geometry of the rift. Thi s par t o f th e Nort h Se a basi n i s a t present boun d b y downwar d flattening , mar ginal faults, called the 0ygarden Fault Zone and Hutton alignment , respectively. These faults , a s well a s most maste r fault s withi n the rift , ar e o f Permo-Triassic origin , bu t wer e reactivate d i n Jurassic time . Between thes e margina l faults , th e strat a ar e rotated an d generall y dip towar d th e margins ,

Fig. 3. Faul t map of the Gullfaks-Visund-Snorre area, with profile locations. Faults a t top Brent Group or base Cretaceous leve l (wher e the Bren t Group i s eroded).

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away fro m th e Permo-Triassi c rif t axi s eas t o f the Jurassi c Vikin g Grabe n (Fig s 1 an d 2) . Similarly, the main fault s dip toward th e PermoTriassic rif t axis , but somewha t mor e steepl y o n the easter n (Hord a Platform ) side . Thi s differ ence i n di p an d th e asymmetr y o f th e syste m reflect tha t a large r portio n o f th e extensio n accumulated o n th e wester n sid e o f th e Permo Triassic rif t axi s than o n th e easter n side . Low-angle fault s ar e particularl y relate d t o the western margin o f the Viking Graben proper (e.g. Swallo w 1986 ; Harri s & Fowle r 1987 ; Cherry 1993 ; Faerset h e t al. 19956 ; Plat t 1995) . Thus, th e are a wes t o f th e Hutto n alignmen t represents th e wester n footwal l o f th e entir e asymmetrical Jurassi c grabe n system . However , low-angle faults of less regional significance also occur withi n the basin , an d i t ma y b e usefu l t o distinguish betwee n margina l an d intra-basi n low-angle fault s because o f their differen t spatia l and kinemati c significanc e in rif t systems . The marginal faults to the northern North Sea rift exhibi t stee p uppe r part s (c . 55-60° fo r th e 0ygarden fault) , an d flatte n downward s int o the basemen t t o locall y for m low-angl e faults . It i s possible tha t thes e fault s ar e linke d t o low angle o r sub-horizonta l ductil e shea r zone s i n the ductil e middl e t o lowe r par t o f th e crust , and i n thi s sens e ar e mechanicall y linked . Th e downward flattenin g o f thes e margina l fault s i s reminiscent of that of simple extensiona l model s where rigid footwall s require the marginal faults to b e listri c to develo p set s o f rotate d (domino ) fault block s in their hanging walls (e.g. Wernicke & Burchfiel 1982 , fig. 7). Rotation o f the domin o fault block s i s made possibl e b y th e non-plana r geometry of the relate d margina l fault, an d con sequently, a n abrup t chang e i n di p i s seen fro m the relativel y horizontal bed s i n th e footwal l t o rotated bed s i n th e hangin g wall . However , models tha t includ e (flexural-isostatic ) footwal l deformation (e.g . Kuszni r e t al . 1991 ) ar e no t restricted b y the same boundar y conditions , an d non-planar margina l fault s ar e therefor e no t a necessity i n rif t systems . The Tampe n Spu r are a Intra-basin low-angl e faults or detachment s ar e most commo n o n th e western sid e o f th e Vi king Graben, particularly th e Gullfaks-Visund Snorre par t o f th e Tampe n Spu r are a (Fig . 3) . The hig h densit y o f oilfield s and prospect s ha s made th e concentratio n o f geologica l an d geo physical dat a in this area remarkabl y high . Low angle fault s an d detachment s i n that regio n will be th e mai n focu s of th e res t o f this paper .

Fig. 4 . Mai n faul t block s i n th e Gullfaks-Visund Snorre are a an d th e name s applie d i n the text . A. th e Statfjord fault ; B, the Snorre fault ; C. the Visund fault : D, th e Gullfaks fault; E . the Gullfaks S0r fault: F. th e Viking Grabe n boundar y fault .

The most significan t fault s i n th e are a ar e referred t o a s follow s (Fig . 4) : A , th e Stat fjord fault ; B , th e Snorr e fault ; C , th e Visun d fault; D , th e Gullfak s fault ; E , th e Gullfak s S0r fault; F, th e Viking Graben boundary fault . Similarly, th e mai n faul t block s o f th e are a (Fig. 4 ) are calle d th e Statfjord , Gullfaks , Gull faks S0r , Visun d and Visun d s0r0st fault blocks . All thes e faul t block s ar e tilte d s o tha t th e bedding i s dipping gentl y to th e wes t o r north west betwee n th e generall y east - o r southeast dipping faults . W e hav e foun d indication s tha t several o f these fault s (A , B , C , D ) hav e non-planar geometries , and tha t the y form local basement-involved detachment s i n th e are a (Fig. 5) . I n addition , w e se e indication s o f a supra-basement detachmen t beneat h th e Gull faks Fiel d (Gullfak s detachment) and sout h an d northeast o f Gullfaks (Gullfaks S0r an d Visund s0r0st detachments ; se e below). Supra-basement detachment s An exampl e o f supra-basemen t detachmen t i s found beneat h th e Gullfak s oilfiel d (Petterso n et al . 1990 ; Fossen & Hesthammer 1998) , which occupies th e easter n par t o f th e Gullfak s faul t

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Fig. 5 . Simplifie d 3 D illustratio n of the Gullfaks-Visund-Snorr e area . Layerin g at middl e Jurassic level is indicated.

Fig. 6 . Profil e acros s th e Gullfak s fault block , showin g the presenc e of a low-angl e detachment. Carefull y interpreted 3 D seismi c data, wel l informatio n and th e dee p seismi c line shown in Fig . 7 form th e basi s for th e interpretation. A standard 'lin-vel ' depth conversio n metho d wit h well data from the Gullfaks Field wa s applie d and extende d down to basement, which is assigned a constant velocity of 6.1 kms"1 . Parameters used for interval sea floor-base Cretaceous: V 01 = 1790ms-' , £ =0.2653; base Cretaceous-to p Statfjor d Fm : K 02 = 800ms-1 , /c = 0.8523; base Cretaceous-to p Triassic: K 0 2= 1000ms- 1 , £ =0.8523. (Se e Fig . 3 for location. )

block (Fig . 4) . Maste r fault s wit h kilometre scale displacement s separat e th e Gullfak s faul t block fro m th e Statfjor d faul t bloc k t o th e west (fault A ) an d th e Visund-Gullfak s S0 r are a t o the eas t (fault s D an d B Fig . 4) . Amon g these , faults D an d B for m a highl y non-plana r structure tha t wrap s aroun d th e Gullfak s Field

in a very characteristic manner . Thi s faul t struc ture splay s bot h t o th e sout h an d nort h (fault s D, E , B and C) , and delineate s th e Gullfaks S0r and Visun d faul t blocks , respectivel y (Fig. 4). Detailed knowledg e abou t th e structura l geology o f th e Gullfak s Fiel d ha s bee n gaine d through systemati c analysi s o f exploratio n an d

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production dat a (Fosse n & Hesthamme r 1998) . In genera l terms , th e wester n par t o f th e Gull faks faul t bloc k i s dominate d b y a classica l domino faul t system . The domino fault s ar e dipping abou t 30 ° t o th e east , wherea s th e beddin g within th e block s exhibit s mor e gentl e dip s (generally 10-18° ) t o th e west. Thi s distinc t an d geometrically unifor m domin o syste m extend s

from th e Tordi s faul t i n th e wes t (Fig . 3 ) t o a horst comple x i n th e easternmos t par t o f th e Gullfaks faul t bloc k (Fig . 6) , spannin g abou t 10-15 km i n th e E- W directio n an d slightl y more i n the N- S direction . The extensio n i n th e domin o are a i s considerably highe r than i n the res t o f the Gullfak s faul t block. A recen t map-vie w restoratio n o f th e

Fig. 7. (a ) Part o f the reprocessed dee p seismi c line NSDP-1 that traverses the Gullfaks fault block . The low-angle detachment i s indicated togethe r wit h som e o f th e majo r basement-involvin g faults. Th e interpretatio n of faults withi n the basemen t i s speculative an d no t wel l constrained . (Se e Fig. 3 for location. ) (b) Close-u p o f th e reflection interprete d a s th e Gullfak s detachment (arrows).

NORTH SE A DETACHMENT S AN D LOW-ANGL E FAULT S Gullfaks Fiel d (Roub y e t al. 1996 ) show s tha t the seismicall y resolvabl e Jurassi c E- W exten sion acros s th e field is of th e orde r o f 40-50% (0 = 1.4-1.5) . A simila r estimat e o f th e west ern par t o f th e Gullfak s faul t bloc k give s onl y 10-15% extensio n (/3 = 1.1-1.15) . Simila r results ca n b e obtaine d b y summin g th e heave s of faults acros s E- W profil e lines . There i s good evidence tha t additiona l sub-seismi c deforma tion is present i n the domino syste m (Koestler et al. 1992 ; Fosse n & Hesthammer 1998) , possibly increasing the total extension to as much as 80% (0=1.8).

The Gullfaks detachment The presenc e o f a low-angle , shallo w detach ment beneat h th e Gullfak s Fiel d ha s bee n suggested b y previou s worker s (Fosse n 1989 ; Koestler e t al . 1992) , althoug h n o compellin g evidence fo r it s existenc e ha s bee n presented . Reprocessing o f lin e NSDP84- 1 has , however , revealed a gentl y E-dippin g seismi c even t at about 3- 5 s (Fig . 7) . I n th e wes t th e reflectio n appears to merge wit h a steeper fault . Ther e ar e no simila r reflector s shallowe r i n th e section , where the reflections are either sub-horizontal or westerly dipping . Th e reflectio n is therefore no t a multipl e o f shallowe r reflections , an d ca n b e traced fo r a horizonta l distanc e o f mor e tha n 10km. Considerin g th e geometry , positio n an d extent o f th e reflection , w e interpret it a s a lowangle fault (detachment) underlyin g the Gullfaks Field. A proble m wit h thi s interpretatio n i s that th e possible detachment reflecto r cross-cut s a stron g an d continuou s W-dippin g signa l i n its western par t (Fig . 8a) . However , a simpl e restoration exercis e (Fig . 8a-c ) reveal s tha t thi s continuous seismi c signa l ca n b e explained a s a coincidental alignmen t o f tw o reflection s fro m different stratigraphi c level s on eac h sid e o f th e low-angle E-dippin g reflectio n (detachment) , a case that is familiar to mos t experience d seismic interpreters. An interprete d geologica l sectio n alon g thi s seismic line, based o n both 2 D and Gullfak s 3D seismic data an d wel l information (Fig . 6) , indicates tha t th e detachmen t ha s a di p o f 5-30° , and reaches a depth of about 7- 8 km in the eastern part of the Gullfaks fault block . Gravity and magnetic dat a ar e consisten t wit h th e seismi c interpretation o f top basement beneat h the Gullfaks Fiel d a s shown in Fig. 6 although a precise definition o f to p basemen t i s difficul t fro m an y available dataset . Accordin g t o ou r interpreta tion, th e detachmen t sole s ou t jus t abov e th e basement-cover interface.

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The contact relationship with the steepe r an d older (Permo-Triassic ) maste r faul t t o th e eas t (fault B-D ) i s no t clea r fro m th e reflectio n seismic data . A s n o evidenc e i s foun d fo r th e detachment i n the hanging wall to this fault, i t is assumed tha t th e detachmen t merge s wit h faul t B-D t o for m a ram p structure , simila r t o th e Visund S0r0s t detachmen t (se e below) . I n th e west, the detachment appears to be connected t o the Tordis fault , which therefore may have acted as the breakaway fault t o th e detached Gullfaks domino system.

The Gullfaks S0r detachment Fault D (Fig . 4) , whic h separate s Gullfak s S0r from th e Gullfaks fault bloc k t o th e west, has a flattening-downward appearanc e o n 3 D seismic data (Fig . 9) . Although the geometr y is not wel l constrained i n it s easter n part , i t appear s tha t fault D flatten s a t depth s o f abou t 5.5- 7 km to for m a detachmen t a t simila r dept h t o th e Gullfaks detachment . Th e entir e Gullfak s S0 r fault bloc k ride s o n thi s detachment , whic h appears t o joi n faul t E eas t o f Gullfak s S0r . Faults abov e th e detachmen t ar e characterize d by domino-style , sub-paralle l E-dippin g faults , similar t o th e styl e see n i n th e Gullfak s Field . The general interpretation o f depth t o basemen t in thi s are a implie s tha t th e detachmen t i s located withi n th e Permo-Triassi c sedimentar y sequence. We therefore classify th e Gullfaks S0r detachment a s a supra-basemen t detachment .

The Visund sorost detachment Faults B an d F forme d th e boundarie s o f th e initial Visun d mega faul t bloc k (combine d are a of th e Visun d an d Visun d s0r0s t faul t block s in Fig . 4) . Th e uppe r par t o f faul t F (Fig . 3 ) appears t o b e straigh t i n cross-section . I t i s o f Permo-Triassic origi n (Faerset h e t al . 19956) , although th e mai n thro w i s relate d t o Jurassi c extension, reaching a maximum of about 5 km at Middle Jurassic level. The are a a t presen t bounde d b y faults C an d F (Visun d s0r0st faul t bloc k in Fig . 4 ) was par t of th e initia l Visund meg a faul t block , bu t wa s separated a s a resul t o f Volgia n footwal l col lapse of the southeastern par t of the mega-bloc k (Faerseth et al. 19956) . Fault C and som e minor, east-dipping fault s forme d a t thi s stage . Th e faults appea r t o converg e an d merg e a t dept h (Fig. 10) . Depth conversio n o f fault C reveals a rampflat-ramp geometr y wit h a presen t maximu m

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Fig. 8 . Simpl e restoration o f the portio n o f line NSDP84-1 wher e the detachmen t reflectio n merges with the steeper norma l faul t (see Fig . 7 a for location) . Fro m (a ) the detachment reflectio n appears t o transec t a continuous reflectio n (circled) . The sectio n was restored by rigid rotation o f th e hanging wall t o th e detachment (b) and subsequentl y b y translatio n alon g th e stee p faul t (c) . In (c ) the sectio n i s restored, afte r remova l o f a n offset o f a fe w hundred metre s along th e detachment . The rotatio n that wa s require d i s consistent with anticlockwise rotation o f the hangin g wall of the detachment durin g deformation . Th e successfu l restoratio n show s that th e continuou s reflectio n may b e th e resul t of coincidental alignmen t of two differen t reflectin g surface s o n each

Fig. 9 . Seismi c profil e throug h th e Gullfak s S0r faul t block , showin g th e downward-flattenin g geometr y o f faul t D withi n th e (Permo-?)Triassi c sediments . Depth converted versio n (n o vertica l exaggeration) i s shown i n th e uppe r righ t corner . Th e low-angl e par t o f faul t D define s th e Gullfak s S0 r detachment .

NORTH SE A DETACHMENT S AN D LOW-ANGL E FAULT S dip o f 45-50°, decreasing toward s 15 ° below th e Visund s0r0s t faul t bloc k befor e mergin g wit h the steepe r faul t F (Fig . 10) . The exac t positio n of to p basemen t i s ambiguous fro m th e seismi c data, bu t fro m regiona l stratigraphi c informa tion i t is likely that th e faul t detache s just abov e top basement . W e therefor e classif y th e faul t a s a supra-basemen t detachmen t simila r t o th e Gullfaks detachment . Detailed tectono-stratigraphi c wor k b y Faer seth e t al. (19956) gives strong evidenc e tha t th e Visund s0r0s t detachmen t (faul t C ) formed dur ing latest Jurassic footwall collapse of the Visund fault block , whic h before thi s collaps e wa s con tinuous wit h th e Visund s0r0st faul t block .

Intra-basement detachment s The Visund detachment The Snorr e faul t (B) , which separates th e Visund fault bloc k fro m the Snorr e Field , i s well defined from fault-plan e reflection s and/o r terminatin g reflectors i n the hanging wall and footwal l down to abou t 4 s o n 2 D line s (Fig s 1 1 an d 12 ) (see also Nelso n & Lam y 1987) . Belo w thi s dept h the faul t enter s a zon e o f reduce d reflectivit y where th e faul t plane is not s o clearly displayed . However, a t lowe r depth s (5-6.5s ) anothe r reflector appear s o n th e seismi c line . Thi s sig nal i s muc h stronge r tha n th e surroundin g reflections, an d i s gently east dippin g (Fig . 12b ) or locall y sub-horizonta l (Fig . lib) . I t i s clea r from thei r strengt h an d geometr y tha t thes e reflections canno t b e multiple s fro m overlyin g events (they are overlain by a low-reflective zon e below W-dippin g reflections) , and th e fac t tha t they generat e unequivoca l sea-botto m multiples (Figs l i b and c , and 12 b and c ) shows that the y represent rea l physica l structure s i n th e crust . It i s possible t o correlat e this signal from lin e t o line, an d a contoure d ma p o f th e reflectio n i s shown i n Fig . 13 . W e sugges t tha t th e lower , strong reflectio n represent s a low-angl e faul t within th e basement , and tha t i t i s connected t o the Snorr e faul t t o defin e th e Visun d detach ment. O n som e o f th e lines , a sub-horizonta l reflection als o appear s t o th e wes t of the Snorr e fault (e.g . Fig . 11) , suggestin g tha t th e Visun d detachment i s par t o f a mor e extensiv e detachment syste m a t abou t 6s . Dept h conversio n o f this detachment (se e Fig. 15a , below) shows tha t

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the detachmen t remain s sub-horizonta l afte r depth conversion , an d exist s a t a dept h o f about 14km .

Gullfaks area Interpretation o f th e dee p seismi c dat a acros s the Gullfaks Field, together wit h commercial 2 D seismic lines, has resulte d i n the faul t geometrie s shown i n Fig . 1 4 (Odinse n e t a l 20000) . Th e Statfjord faul t (A ) i s here interprete d a s a non planar faul t tha t become s a low-angl e detach ment structur e in th e basement . A domino-styl e fault-block arrangemen t i s tentativel y interpre ted t o b e situate d abov e thi s detachment , simi lar t o th e Jurassi c domin o syste m abov e th e overlying Gullfak s detachment . Hence , ther e are indication s tha t tw o level s o f detachment s (a supra - an d intra-basemen t detachment ) ma y exist beneat h th e Gullfak s Field .

Evolution of the intra-basement detachments Low-angle detachment s o f lat e t o post-Caledo nian ag e ar e know n fro m bot h side s o f th e North Se a rift . Th e sub-horizonta l decollemen t zone betwee n Caledonian nappe s an d basemen t in souther n Norway , th e NW-dippin g Hard angerfjord shea r zone , an d th e Nordfjord-Sog n detachment ar e al l well-expose d example s o f low-angle structures accommodating substantia l Devonian extensio n o n th e easter n sid e o f th e North Se a (Norto n 1987 ; Serann e & Segure t 1987; Fosse n 1992) . Similarly , reactivatio n o f Caledonian thrust s a s low-angl e extensiona l features i s recognized in Scotlan d (McCla y et al. 1986; White & Glasser 1987 ; Powell & Glendinning 1990) . Some o f these detachments continue under th e Nort h Se a an d for m zone s o f weak ness i n th e crust . Fo r mechanica l reasons , thes e and simila r low-angl e structures o f Caledonia n or Devonia n ag e wer e easil y reactivate d a s de tachments durin g th e Permo-Triassi c an d Jur assic extension phases, especially in areas of high post-Devonian extension , and cause d ramp-fla t geometries o f th e typ e see n i n Fig . 11 . Reacti vation o f orogeni c structure s a s low-angl e extensional fault s o r shea r zone s i s a commo n phenomenon i n extende d areas , suc h a s i n th e western US A (Allmendinge r e t aL 1983 ; Coney & Harm s 1984 ; Wernick e e t al . 1987 ; Parris h et al. 1988; Constenius 1996) . Assuming that thi s

Fig. 10 . Seismi c profil e (fro m 3 D data ) acros s faul t C (Visun d s0r0s t detachment) , whic h clearl y appear s a s a downward-flattenin g faul t o n th e seismi c data. Depth-converte d versio n (n o vertica l exaggeration) is shown in the uppe r righ t corner . (See Fig. 3 for location. )

Fig. 11 . (a ) Seismic line NVGT-88-08 across the Visun d faul t block, showin g ho w th e Snorr e faul t flatten s beneath th e Visun d faul t bloc k t o for m a n intra-basemen t detachment (Visun d detachment) a t about 6 s. (b) Uninterpreted versio n of (a). The upper par t of the Snorre faul t is clearly visible as a downward-flattening reflection, an d a sub-horizonta l reflectio n i s interpreted a s a low-angl e detachment, (c ) Close-up o f the stron g sub-horizonta l reflectio n and it s multiple. (Se e Fig . 3 for location. )

Fig. 12 . (a ) Seismic section NVGT-88-0 9 show s ho w th e low-angl e Visund detachmen t i s somewhat steepe r tha n i n Fig . 11 . (b ) Uninterprete d versio n o f (a) . The Visun d detachment i s clearly visibl e a s a multiple-generatin g reflectio n i n th e basemen t (c) . (See Fig . 3 fo r location) .

Th

H . F O S S E N EJ A L

Fig. 13 . Time-contoure d ma p o f the Visund detachmen t (Snorr e fault ) base d o n 2 D seismi c lines (stippled lines) . Grey are a indicate s are a wher e th e fault-plan e reflectio n i s visible, o r wher e terminatin g reflector s defin e th e fault plane . development als o hold s fo r th e Nort h Sea , the easterly di p o f low-angl e intra-basi n fault s o r detachments i n th e Tampe n are a ma y indicat e that th e Caledonia n sutur e zon e i s locate d eas t of the Tampen area , beneat h th e Viking Graben or clos e t o th e Permo-Triassi c rif t axis . On th e othe r hand , i t i s possible tha t th e detachments wer e initiall y steepe r fault s tha t rotated t o becom e low-angl e structure s dur ing th e Permo-Triassi c an d Jurassi c extensiona l phases. A maximum estimat e o f the original (pre basin-fill) di p o f th e detachment s i s obtaine d from th e di p o f th e top-basemen t surfac e i n th e footwall o f th e detachmen t faults . This estimat e is valid only for rigi d block rotatio n withou t significant internal , small-scale deformation . I n th e case o f the Visund detachment, th e top basemen t surface i n th e footwal l i s dippin g abou t 25 ° t o the west (Fig. 15a). To p basemen t restoratio n b y predominantly rigi d bloc k rotation s an d transla tions woul d resul t i n origina l di p o f detachmen t

of th e sam e orde r (Fig . 15c) . If additional smallscale deformatio n i s allowe d fo r (e.g . inclined shear, no t show n i n Fig . 15), origina l faul t dip s may b e found t o hav e been lower than 2 5 . Low ering o f faul t dip s throughou t th e extensiona l history o f th e Nort h Se a Rif t i s thu s possible , but onl y t o a limite d extent (maximu m 20-3 0 ) . Hence, fault s tha t ar e currentl y dipping 3 0 t o the east ma y have exhibite d dip s u p to 50-6 0 i n Permian time . I t i s possibl e tha t som e o f th e faults develope d int o low-angl e structure s afte r repeated phase s o f extension , and acte d a s lowangle fault s or detachments onl y during the Lat e Jurassic extensio n phase . The abov e discussio n relate s to th e mode l en visaged b y Yieldin g e t al. (1991) . Thi s mode l suggests that the downward shallowin g of master faults i n th e Nort h Se a i s a resul t o f multiple episode s o f deformation . A n early , Permo Triassic phas e o f extension caused th e formation and subsequen t rotatio n o f th e fault s i n th e

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Fig. 14 . Dept h profile acros s the Gullfak s faul t block , showin g the interpretatio n o f fault A a s downward flattening, formin g a n intra-basemen t detachmen t underneat h th e Gullfak s supra-basemen t detachment. The interpretation on this line is based on a combination of line NSDP84-1 (Fig. 7), the Gullfaks 3D seismic data and well information , an d conventiona l 2D seismi c data fro m th e area .

basement, wherea s Jurassi c extensio n cause d the faults to grow into the Jurassic sequence with steeper dips . Faults with a lower (Permo-Triassi c level), low-angle segment an d a n upper (Jurassi c level), steepe r segmen t woul d resul t fro m thi s model. However , example s suc h a s Fig . 11 , where th e ramp-fla t transitio n is located withi n the basement , demonstrat e tha t th e mode l o f Yielding e t al. (1991 ) canno t accoun t fo r al l o f the fault geometries relate d t o low-angle faults in the northern Nort h Sea . We think tha t bot h th e multiple-deformation mode l an d th e contro l o f older (Devonia n o r Caledonian ) shea r zone s in the anisotropi c basemen t influence d th e occur rence o f low-angl e faul t movement s i n Meso zoic time .

Evolution of the supra-basement detachment From th e discussio n above , w e conclud e tha t most o f the intra-basemen t detachment s hav e a history tha t goe s bac k a t leas t t o th e Permo Triassic extensio n phase , an d probabl y t o th e Devonian extensiona l and/o r Caledonia n con tractional events . Th e supra-basemen t detach ments ar e clearl y o f younge r age , becaus e the y occur i n rock s o f (late ) Triassi c age . The y ar e therefore likel y to be partl y or wholl y relate d to th e lat e Jurassi c extensio n phase , a s dis cussed below.

This is particularly clear for th e Visund S0r0st detachment, whic h ha s recentl y bee n show n t o be a collaps e structur e tha t forme d durin g lat e Jurassic rotatio n of the Visund fault block. Thi s collapse occurre d i n th e southeaster n an d high est portio n o f th e origina l Visun d mega-faul t block. The initiation of the collapse can be dated to mid-lat e Kimmeridgia n time , bu t th e mai n activity i s believe d t o hav e bee n i n th e Volgia n period (Faerset h e t al . \995b). Thes e constraint s show tha t supra-basemen t detachment s forme d at a relativel y lat e stag e i n th e lat e Jurassi c extension phas e an d afte r th e mai n fault s wer e established. The nearb y Gullfak s Fiel d i s locate d o n a bend i n th e Gullfak s faul t block , whic h occur s between th e souther n ti p o f faul t B an d th e northern terminatio n of faul t D . Th e character istic 'Gullfak s bend ' ma y b e interprete d a s a n accommodation zon e (family 3 case G') betwee n faults B an d D accordin g t o th e schem e o f Rosendahl (1987 ) (Fig . 16) . Thro w variatio n data (Fig . 17 ) fo r fault s B an d D indicat e a significant decreas e i n thro w fo r bot h fault s towards th e ape x of th e 'Gullfak s bend' . This is consistent wit h a mode l i n whic h thes e fault s grew toward s th e Gullfak s field , wher e the y finally intersecte d t o for m th e Gullfak s ben d accommodation zone . Durin g thi s process , th e Gullfaks Fiel d are a becam e a n elevate d part o f the rotatin g Gullfak s faul t block . An y elevate d

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H. FOSSE N E T AL .

Fig. 15 . (a ) Depth-converte d versio n o f Fig . 11, showin g that th e detachmen t occur s a t abou t 12-14k m depth . It shoul d b e noted tha t th e chang e i n di p o f th e maste r fault s occur s wel l withi n th e basement . Th e lin-ve l depth conversio n metho d base d o n wel l data fro m th e Gullfak s Field i s applied dow n t o basement , where a constant velocit y of 6.1 kms-1 i s applied. (See text for discussion.) (b) Simple restoration to the Triassic top Teist Fm level , (c ) Restoration o f th e to p basemen t surface . (Note chang e i n di p o f th e detachmen t fault. ) Th e balancing performed her e restore s th e sectio n b y making the marke r horizon approximatel y horizontal. Rigid-body rotatio n an d mino r amount s o f vertical shear ar e applied . Compaction-relate d effect s ar e no t considered. Vertica l shea r wit h n o rigi d bod y rotatio n give s a simila r result. part o f a larg e faul t bloc k i n a rif t syste m i s the potential objec t o f extensiona l collapse , an d i t appears tha t th e hig h portio n o f th e Gullfak s fault bloc k collapse d abov e th e Gullfaks detach ment. This process was probably dynami c rather than strictly sequential, so that th e gradual uplif t of th e Gullfak s are a le d t o formatio n o f th e Gullfaks detachmen t and repeate d sli p along the detachment a s fault s B and D grew. The presenc e o f a detachmen t beneat h th e Gullfaks Fiel d explain s severa l o f the character istics o f th e area . First , i t provide s a soun d explanation fo r th e presenc e o f th e domin o

system. All the main faults in the Gullfak s Field (i.e. i n th e uppe r plat e o f th e Gullfak s detach ment) di p t o th e east , antithetica l t o bedding , except fo r th e margina l hors t comple x a t it s eastern edg e (Fig . 6 , sp.6800) . Th e formatio n of suc h uniforml y dippin g set s o f fault s i s favoured i f a dippin g basemen t o r detachmen t exists underneat h th e paralle l faul t system . In tha t case , th e fault s i n th e uppe r bloc k ma y easily develo p domino-styl e faul t block s wit h faults dippin g syntheticall y t o th e detachmen t or basement . Thi s has bee n demonstrate d by extensional experiment s performe d wit h tilte d

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A Gullfak s detachmen t als o help s explai n the unusuall y larg e extension s i n th e Gullfak s Field (domin o system ) a s compare d wit h th e rest o f th e Gullfak s faul t block. Roub y e t al . (1996) use d a numerica l map-vie w restoratio n method t o sho w tha t th e Lat e Jurassi c E- W extension i s muc h highe r i n th e Gullfak s Fiel d (c. 40%) tha n i n th e Gullfak s bloc k i n genera l (c. 15%). Thi s observatio n fit s wel l wit h a detachment model , wher e th e uppe r plat e ca n collapse an d exten d independentl y fro m th e lower and western part of the block. I n a similar way, th e Gullfak s S0 r bloc k showe d hig h extension (33% ) compare d wit h th e are a t o th e west, whic h ca n b e explaine d b y a n underlyin g Gullfaks S0 r detachment .

Fig. 16 . Mode l for th e developmen t of th e Gullfak s spur. The Snorre and Gullfak s S0r faults (fault s D and B in Fig . 4) are though t t o hav e defined half-graben s with a geometry similar to tha t described by Rosendahl (1987 ) (a). As the y grew (b), the fault s approached on e another, and a s the y crossed (c), the Gullfaks spu r was defined. I t i s possible that part of the complex fault patter n mapped t o th e east of th e Gullfaks Fiel d (see Fig. 3) is a resul t o f simultaneou s movement o n fault s D an d B .

Insights fro m plaste r experiment s

sandboxes (McCla y & Elli s 1987 ; Vendevill e et al. 1987 ) (Fig . 18) , an d i s als o characteristi c for th e upper-plat e deformatio n abov e detach ments i n th e wester n US A (e.g . Wernick e 1985 ; Lister & Davi s 1989) . Domin o system s ma y also develo p in imtilted san d models , bu t rarel y with th e consistency and numbe r o f fault block s as see n i n th e Gullfak s Fiel d (Fig . 6) . A simila r argument ca n b e mad e fo r th e existenc e o f a gentl y east-dippin g detachmen t underneat h Gullfaks S0 r durin g developmen t o f th e upper block faults .

The observation s fro m th e Gullfak s Fiel d an d physical modellin g sugges t tha t th e geometrie s and interaction between master faults are crucial for th e fina l resul t o f th e extensiona l deforma tion. Thre e basi c relationship s exis t betwee n supra-basement detachment s an d th e steepe r master faul t associate d wit h the basemen t scar p (ramp). I n th e firs t cas e (Fig . 19a) , th e detach ment is older than the steep fault and is therefore offset b y th e latter . I n th e secon d exampl e (Fig. 19b ) th e detachmen t i s the younge r struc ture, an d accordingl y cross-cut s an d offset s the stee p fault . Finally , i f th e tw o ar e activ e at th e sam e tim e (Fig . 19c) , the y ma y join an d merge into a single structure i n the ramp region .

Fig. 17 . Throw-variatio n diagrams for fault s D an d B (Fig. 4) at to p Bren t Grou p level . Faul t D show s a maximum at about 6 POO', and a decrease in throw is recorded towards the Gullfaks Field. Similarly, fault B has maximum throw values north of the Gullfaks Field (north of 61°30'), and show s a decrease in throw to the south towards Gullfaks. These data indicate that faults D and B grew towards the Gullfaks Field, where they eventually met, a s suggested in Fig . 16.

126

H. FOSSE N E T AL .

Fig. 18 . Result s o f sandbo x experiment s above a n incline d base . Redraw n fro m McCla y & Elli s (1987 ) an d Vendeville el al. (1987). In both cases a prominent domino-styl e faul t patter n wa s produced, simila r t o th e style seen abov e th e Gullfak s supra-basemen t detachment .

However, i f the displacemen t alon g th e detach ment i s large, the upper par t o f the stee p maste r fault is likely to become deactivated as it is transported toward s the graben center . In this case, a new stee p faul t typicall y forms abov e th e ram p in th e footwal l to th e forme r faul t (Fig . 19c) . It is possible to model supra-basement detach ments an d ramp-flat-ram p geometrie s experi mentally wit h sand , cla y o r plaster . I n suc h models, th e tempora l developmen t o f faultin g and faul t interactio n ca n b e studie d i n detail . Plaster experiment s o f th e typ e describe d b y Fossen & Gabrielsen (1996 ) have been foun d t o be useful , particularl y if wet plaste r i s extended together wit h a stronger , ductil e basemen t o f wet barit e powder . I n on e o f thei r experiments (run 1) , a supra-basement detachmen t an d fault block geometries similar to the Gullfaks example were produced . Durin g thi s ru n (Fig . 20) , th e plaster in the right-han d en d o f the deformatio n box move d relativ e t o a barit e basemen t ramp , indicating the presence of a faul t o r detachmen t between th e two . Subsequently , a straigh t faul t (fault 3 i n Fig . 20 ) develope d t o th e lef t an d merged wit h th e detachment , an d a large , non planar detachmen t structur e formed (Fig . 20d). The resultin g detachment doe s no t cu t th e stee p faults forme d a t a n earlie r point , bu t togethe r these structure s for m a kinematicall y coheren t system simila r t o Fig . 19c . Geometrically , this ramp bears similarities to the master fault eas t of the Gullfak s detachment (Fig. 6 except that th e Gullfaks maste r faul t continue s downward). In a differen t experiment , tw o stee p maste r faults ( 1 an d 2 in Fig . 21b ) forme d i n a volume

of we t plaste r unde r plan e strai n extension . A new , listric fault ( 5 in Fig . 21c ) grew from th e free surfac e of th e previousl y undeformed foot wall, flattened out at about one-third o f the total height o f th e model , an d eventuall y cut acros s fault 1 to joi n faul t 3 (Fig . 2Id) . Faul t 1 was consequently deactivate d an d offset , an d th e resulting geometry and deformation history is an example o f th e cas e show n i n Fig . 19b . The tw o plaste r experiment s describe d her e illustrate how tw o differen t situation s can occu r during a singl e plan e strai n extensio n history . For natura l examples of supra-basement detach ments, suc h a s th e Gullfak s an d Gullfak s S0 r detachments discusse d above , th e geometri c relationships i n th e ram p zon e (a s illustrate d in Fig. 19 ) shoul d b e examined . Th e apparen t continuity o f th e relativel y stee p maste r faul t that transect s the basement an d th e sedimentary cover o n th e seismi c dat a exclude s th e mode l shown i n Fig . 19b . Althoug h th e availabl e seismic dat a ar e no t goo d enoug h t o safel y distinguish betwee n Fig . 19 a and c ther e i s n o indication o f a continuatio n o f th e Gullfak s o r Gullfaks S0 r detachmen t o n th e downthrow n side o f th e steepe r maste r fault . Togethe r wit h the similaritie s betwee n th e mode l show n i n Fig. 2 0 and th e Gullfak s Field (basemen t ramp, marginal horst) , thi s support s th e mode l shown in Fig. 19c, although the model shown in Fig. 19a was suggeste d b y Koestle r e t aL (1992) . A s fa r as th e Visun d are a i s concerned , th e ag e con straints discusse d abov e (Faerset h e l al . \995b) show tha t th e Visun d s0r0s t detachmen t i s clearly younger than the steep master fault t o the

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Fig. 19 . Th e thre e possibl e relationship s between detachments an d steepe r master fault s abov e a basement escarpment . Number s indicat e relativ e timing. I n (a ) the stee p faul t post-date s th e detachment, wherea s in (b ) the timin g is opposite . In (c ) the detachmen t an d th e stee p faul t are activ e simultaneously. I f the displacemen t alon g th e detachment is significant, th e stee p fault i s transported across th e ramp , an d a new fault form s i n its footwall (stippled). Th e resul t ca n b e a complex faul t zon e above th e ramp . east, bu t tha t the y move d simultaneousl y in latest Jurassi c time . Hence , th e mode l show n i n Fig. 19 c also applie s t o thi s detachmen t system .

Conclusions Some o f th e maste r fault s on th e wester n flank of th e Vikin g Grabe n exhibi t lo w dip s i n th e basement, an d acte d a s intra-basemen t detach ments durin g th e lat e Jurassi c stretchin g phase . Some of these detachment fault s might have ha d higher initia l dip s an d rotate d int o les s stee p orientations throug h bloc k rotation s an d inter nal deformatio n durin g pre-Jurassi c riftin g phases. I n thi s model, fault s tha t wer e originally intermediate o r hig h angl e rotated t o for m low-

Fig. 20. Tempora l developmen t o f plaster model , where th e barit e basement i s stifTe r tha n th e overlying plaster, an d therefor e forms a ram p a s the right-hand wall is pulled t o th e right . Faul t 3 develops int o a downward-flattening fault , which form s a detachmen t towards th e en d o f the run . Modified fro m Fosse n & Gabrielsen (1996). angle faults , whic h acte d a s detachment s i n Jurassic time . However , som e fault s ar e see n t o flatten within th e basement , indicatin g that the y follow pre-existin g E-dippin g wea k zone s tha t

128

H. FOSSE N E T AL . Gullfaks S0r. Thes e detachment s ar e therefor e not reactivate d Devonia n extensiona l detach ments o r Caledonia n thrusts . Th e developmen t of a classica l domin o syste m o n Gullfak s (con sistent faul t di p polarity ) indicate s tha t th e detachment existe d a s a n active , E-dippin g sli p surface a t earl y stage s o f the lat e Jurassi c devel opment o f th e domin o system . Suppor t fo r thi s interpretation i s foun d i n experimenta l work , where suc h domin o system s ar e foun d t o develop mor e easil y i f th e basa l plat e o f th e experiment i s tilted . Th e Gullfak s detachmen t model als o provide s a soun d explanatio n fo r th e unusually high extension recorded i n its hanging wall. Contemporaneou s movemen t alon g th e detachment an d th e maste r faul t support s a model wher e th e detachmen t merge s wit h th e master faul t t o for m a singl e faul t structur e a t depth (Fig . 19c) , a mode l tha t certainl y applie s to th e Visun d S0r0s t detachmen t an d probabl y also t o th e Gullfak s S0 r detachment . All o f th e thre e supra-basemen t detachment s occur i n th e hig h part s o f rotated , first-orde r fault block s i n th e Triassic-Jurassi c sequenc e o f the Tampe n area . Thi s suggest s tha t the y ar e gravity-controlled collaps e structure s forme d during extensiona l rotatio n o f th e first-orde r fault blocks . Both supra - an d intra-basemen t detachment s may hav e a significan t influenc e o n th e develop ment o f upper-crusta l structures . A bette r mapping an d understandin g o f thes e low-angl e structures i s importan t fo r improve d under standing o f th e upper-crusta l developmen t o f the Nort h Se a rif t system . This paper has benefited from comments by J. Akselsen. N. Platt , M. Seranne, J. Walsh and D . Couturier.

References

Fig. 21 . Sequentia l developmen t o f plane strai n extensional plaste r mode l carried ou t a t th e 199 2 TSGS meeting, Bergen . No basemen t i s introduced, and a low-angl e detachment develop s and cut s earlie r steep faults . (Se e Fossen & Gabrielsen (1996 ) fo r description o f th e experimenta l method. ) most likel y ar e Devonia n extensiona l shea r zones and/o r reactivate d Caledonia n thrusts . Supra-basement detachmen t i s identifie d i n Triassic sediment s beneat h th e Gullfak s Field , SE o f th e Visun d faul t bloc k an d underneat h

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Syn-rift sedimentar y architecture s in the Northern North Se a R. RAVNAS, 1 A . N0TTVEDT, 2 R . J . STEEL 3 & J . WINDELSTAD 4 1

Norske Conoco A.S, P.O. Box 488, 4001 Stavanger, Norway Norsk Hydro Research Centre, Sandsliveien 90, Postboks 646, 5001 Bergen, Norway (Present address: Norsk Hydro Canada, 1190 lll-5th Avenue SW, Calgary, Alberta T2P 3YZ, Canada) 3 University of Wyoming, Department of Geology & Geophysics, Laramie, Wyoming 82071-3355, USA 4 Statoil, LTEK, P.O. Box 300, 4001 Stavanger, Norway Abstract: Fro m Permia n t o Jurassi c time s the norther n Vikin g Graben an d adjacen t plat form area s experienced multiple rifting, th e Permian-early Triassi c and middle-late Jurassi c rift episodes , separate d b y a n intervenin g middle Triassic-middle Jurassi c inter-rif t perio d dominated b y relative tectonic quiescence. The associated syn - and inter-rift strat a show large variations i n sedimentar y architectur e a s a resul t o f tempora l an d spatia l variation s i n tectonic deformation and subsidence, sediment supply, climate and accommodation creation . The Permian-early Triassi c syn-rift successio n is believed to consist predominantly of nonmarine, arid to semiarid, aeolian, sabkha, alluvial and lacustrine strata, probably interbedde d with marine strata o n the Horda Platfor m an d i n the Viking Graben . The middle Triassic-middle Jurassi c experienced several subsidence stages which, together with climatic variations, exerte d a majo r control o n the periodic outbuildin g and retrea t o f rift marginal , alluvia l and shallo w marin e clasti c wedges . Evidence fo r faul t bloc k rotatio n suggests tha t th e subsidenc e wa s cause d partl y b y mino r extensiona l stages . A s such , th e middle Triassic-middl e Jurassic doe s no t fit the type of development assume d t o b e typica l for eithe r post- o r pre-rift basins. Hence, th e notation inter-rift i s assigned t o this period an d the associated succession . The middle-late Jurassic rift episode was characterized b y multiple rift phases separate d b y intervening stage s o f relative tectonic quiescence . Th e syn-rift infil l is mixed non-marine an d marine and consists o f fluvial through shallo w marine and shelfa l deposit s t o deeper marin e sediment gravit y flo w an d (hemi-)pelagi c strata . A t th e large r scale , relate d t o th e entir e middle-late Jurassic rif t episode , th e syn-rif t infil l i n genera l show s a two-fol d sandstonemudstone litholog y motif, typical of underfilled rift basins . At the intermediate scale , relate d to singl e rif t phases , threefol d sandstone-mudstone-sandstone , twofol d sandstone mudstone an d singl e mudstone litholog y motif s ar e present , typica l o f sedimen t overfilled / sediment balanced , sedimen t underfille d an d sedimen t starve d rif t basins , respectively . Th e spatial an d tempora l variation s i n the syn-rif t infil l reflec t relativ e distanc e t o th e rif t basi n hinterland area s (whic h had a large sedimen t yiel d potential) an d overal l increased tectoni c subsidence and enhance d rif t topograph y a s the rif t basi n evolved. This sugges t tha t th e tectonostratigraphi c evolution of the norther n North Se a rift basi n can b e viewed at several scales: at the largest scale the rif t basi n evolved through multiple rift episodes, whic h commonl y ha d a duratio n o f severa l ten s o f Ma . Th e rif t episode s ar e separated b y inter-rif t periods . Rif t episode s ar e subdivide d int o interval s representin g distinct rif t phases . Thes e rif t phase s wer e separeted b y tectonic relatively quieter intervals, here referred to as tectonic quiescence stages. Inter-rift periods are subdivided into prolonged tectonic quiescence intervals separated b y short-lived rift stage s or minor rift phases . Distinct rift phases and inter-rif t tectoni c quiescence intervals commonly represent periods of a few to 10+ Ma, and correspond t o second-order sequences or 'megasequences'. At the smaller scale, syn-rift succession s can be subdivided into packages related to distinct rotational tilt event or faulting event s (deformatio n spans) , representin g hundred s o f k a t o fe w M a an d corresponding t o third-orde r sequences . Solitary , large-magnitud e faultin g event s (deformation clines ) ar e likel y t o exer t a majo r contro l o n hig h frequenc y base - o r sea level fluctuation s an d thu s on th e development o f higher-order sequences . However, suc h a control is difficult t o prove and can probably only be recognized in sub-basins with abundant wells and a dense well spacing .

From: N0TTVEDT , A . e t al. (eds ) Dynamics o f th e Norwegian Margin. Geologica l Society , London , Specia l Publications, 167 , 133-177 . 1-86239-056-8/OO/ S 15.00 © Th e Geologica l Societ y o f London 2000 .

Fig. 1 . Ma p o f norther n Nort h Se a wit h location o f seismi c section s i n Fig . 2 and stud y areas . Not e differen t area l exten t of th e Permia n earl y Triassi c an d th e middl e Jurassic earl y Cretaceou s rif t basins .

NORTH SE A SYN-RIF T SEDIMENTAR Y ARCHITECTURE S During recent years , considerable effor t ha s bee n put int o recognizin g th e architectura l signa ture diagnosti c o f syn-rif t basin-fill s i n orde r t o develop predictiv e model s fo r syn-rif t plays . Distinct stratigraphica l signature s o f contrast ing type s o f rif t basins , non-marin e a s wel l as marine , hav e bee n documente d o r concep tualized (e.g . Surly k 1978 , 1989 ; Alexande r & Leeder 1986 ; Frostick & Reid 1987 ; Leeder & Gawthorpe 1987 ; Hamblin & Rust 1989 ; Morle y 1989; Mac k & Seage r 1990 ; Schlisch e & Olse n 1990; Lambias e 1990 ; Prosse r 1993) . Althoug h attempts hav e bee n mad e t o erec t unifyin g syn rift infil l models , ther e is growing consensus tha t a variet y o f syn-rif t infil l pattern s ma y b e present. Thei r signatur e depend s o n th e inter play betwee n creatio n o f syn-rif t accommoda tion, climate , typ e an d volum e o f sedimen t supply, an d relativ e base - o r sea-leve l stan d and thei r change s (Leede r & Gawthorp e 1987 ; Surlyk 1989 ; Prosser 1993 ; Leeder 1995 ; Ravnas & Stee l 1998) . The presen t contributio n describe s a serie s of syn-rift infil l pattern s observe d i n Permian Triassic an d Jurassi c sub-basin s o f th e Nort h Sea (Fig . 1) . W e hav e focuse d o n half-grabe n type sub-basin s i n area s underlyin g and nex t t o three norther n Nort h Se a regional seismi c lines (Fig. 2) . Th e structura l interpretation s o f thes e lines ar e discusse d i n detai l b y Odinse n e t al. (in press) and by ter Voorde et al. (in press). They illustrate differen t style s o f rif t basi n architec ture: a southern segmen t (Fig. 2a), which forms a solitary, wid e half-grabe n wit h flanking , broa d platforms; a centra l segmen t (Fig . 2b) , whic h is represented b y a full-grabe n wit h flankin g faul t blocks an d terraces , an d borderin g platfor m areas; an d a norther n segmen t (Fig . 2c) , which consists o f a serie s of half-graben sub-basins . The intentio n o f th e presen t contributio n i s threefold: first, the significan t variabilit y present in non-marine and , i n particular, marin e syn-rif t infills o f regionall y subsidin g rif t terrain s i s described. Secondly , th e mai n controllin g fac tors o n syn-rif t sedimentar y architectur e fo r specific faul t block s ar e discussed . Emphasi s i s put o n th e effect s tha t Permian-Triassi c an d Jurassic extensiona l tectonic s ha d upo n th e distribution o f lithofacie s an d stratigraphi c surfaces suc h a s hiatuses , unconformitie s an d condensed section s o n a half-grabe n scale . Thi s discussion illustrat e ho w specifi c syn-rif t signa tures, a s recognized fro m well s and seismi c data, can be used to date activit y on extensional fault s and t o separat e distinc t tectoni c event s an d discrete rif t phases . Lastly , the detailed tempora l and spatia l evolutio n o f the syn-rif t sedimentar y architecture a s wel l a s o f th e larger-scal e infil l

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trends i n th e basin-histor y ar e discussed . Thi s discussion ai m t o demonstrat e wh o a detaile d interpretation o f th e syn-rif t sedimentar y archi tecture ca n b e use d t o furthe r deciphe r th e structural evolutio n o f th e Nort h Se a rift basin . The databas e fro m th e norther n Nort h Se a is wide ranging, an d consis t o f 2D an d 3 D seismi c coverage, an d abundan t wel l data . Th e studie d sections have been core d extensivel y in all areas , which, i n additio n t o detaile d biostratigraphy , have allowed interpretatio n o f sedimentary pro cesses an d depositiona l environments , a s wel l as th e recognitio n an d datin g o f ke y stratigra phical surfaces . Thi s has , i n turn , allowe d th e detailed datin g o f th e sequentia l evolutio n o f the rif t basin' s depositona l an d structura l his tory, an d resulte d i n recognitio n o f a serie s of tectoni c event s fro m th e Triassi c through out th e Jurassic . Mor e dat a exis t o n syn-rif t strata an d thei r sedimentar y architecture s i n the upper Triassic-Jurassic , an d especiall y fro m the middle-uppe r Jurassic , whic h thu s ha s received a mor e thoroug h attention . A general ized litho-stratigraphical column of the TriassicJurassic sectio n i n th e northernmos t Nort h Sea and associate d tectoni c event s are presente d in Fig . 3 .

Geological settin g The northernmos t Nort h Se a rif t basin , i.e . th e Viking Grabe n an d adjacen t platfor m areas , were moulde d b y a t leas t tw o episode s o f litho spheric stretching , date d a s ?lat e Permian-early Triassic and middl e Jurassic-earliest Cretaceou s (e.g. Eyno n 1981 ; Badle y e t al . 1984 , 1988 ; Beach e t al . 1987 ; Giltne r 1987 ; Lervi k e t al . 1989; Blundel l & Gibb s 1990 ; Gabrielse n e t al . 1990; Ziegler 19900 , b\ Yielding et al 1992) . The Permian-early Triassi c an d middl e Jurassic early Cretaceous rif t episode s were characterize d by distinct , bu t differen t structura l configura tions (Fig . 1 ; Lervi k e t al. 1989 ; Faerseth 1996) . The tectoni c extensio n i n th e norther n Vikin g Graben durin g thes e stretchin g episode s i s known t o hav e varie d bot h spatiall y an d tem porarily (e.g. Roberts e t al. 1993<2 , 1995 ; Odinsen et al . i n press) , leadin g t o considerabl e vari ability i n architectur e an d compositio n o f th e syn-rift infill . Moreover , a s a result o f the multiphase extensio n an d crusta l thinning , isostati c rebound cause d a gradua l lowerin g o f th e surface relie f fro m mainl y continenta l durin g Permian-early Triassi c riftin g t o essentiall y marine durin g lat e Jurassic-earl y Cretaceou s rifting (N0ttved t e t al . 1995) .

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Fig. 2. Depth-converte d interpretation s of seismic lines (a) NSDP84-2, (b ) NVGT88-04 an d (c ) NSDP84-1. Note difference i n structura l geometry fro m solitar y wide half-graben i n (a ) throug h a full-grabe n i n (b ) t o a series of half-graben sub-basin s in (c). See Fig. 1 for locatio n o f seismic sections. Profile s 1 and 3 (a an d c , respectively ) are fro m Odinse n e l al. (in press) whereas profil e 2(b ) is from Te r Voord e (i n press).

Fig. 3. Triassi c -Jurassic lithostratigraphic nomenclature and tectoni c subdivisions, i.e. rif t episodes , inter-rif t period s and rif t phases , fo r the northernmost Nort h sea . Th e lithostratigraphy i s modified fro m Vollse t & Dor e (1984 ) an d Richard s e t al. (1993) .

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The Permian-earl y Triassi c rif t basi n extend s westwards fro m th e Norwegia n coas t an d includes th e majo r par t o f th e Eas t Shetlan d Platform an d Eas t Shetlan d Basi n (John s & Andrews 1985 ; Lervik e t al 1989 ; Roberts et al 1995). It consists o f a series o f half-grabens wit h master-fault spacin g o f c . 15-35 km . (Badle y et al . 1984 , 1988 ; Lervi k e t al . 1989 ; Gabriel sen e t al . 1990 ; Marsde n e t al . 1990) . Wedge shaped package s tha t ar e believe d t o contai n mainly non-marine strat a fill in the half-graben s (Lervik e t al . 1989 ; Stee l & Ryset h 1990 ; Stee l 1993). The to p o f the syn-rif t sequenc e ha s bee n argued t o b e withi n th e Teis t Formatio n an d assigned a Scythia n ag e (Fig . 3 ; Steel & Ryset h 1990). B y analog y wit h Eas t Greenland , th e West Shetlan d basins , th e souther n Nort h Se a and th e Osl o Grabe n (e.g . Surly k e t al . 1984 ; Glennie 1990#,6 ; Boot h e t al . 1993 ; Olausse n et al . 1994) , i t seem s reasonabl e t o assum e tha t the lower part o f the syn-rif t infil l ma y date back to th e lat e an d possibl y earl y Permia n (se e als o Johns & Andrew s 1985 ; Lervi k e t al . 1989 ; Gabrielsen e t al . 1990 ; Steel 1993 ; Faerseth et al . 1995a). This notio n i s supported b y the presenc e of Permia n dyke s (Faerset h et al . 1976 ; Faerseth 1978; Furne s e t al . 1982 ) an d Permia n exten sional reactivatio n o f fault s o f th e Nordfjord Sogn detachmen t onshor e wester n Norwa y (Torsvik e t al . 1992) . Recen t wor k b y Chris tiansson e t al. (2000) suggest s tha t th e Permianearly Triasi c syn-rif t successio n i s underlai n b y yet anothe r sedimentar y package, possibl y early Permian and/o r Carboniferou s i n age , an d similar t o wha t ha s bee n describe d belo w th e East Shetlan d Platfor m (Plat t 1995) , an d i n th e Oslo grabe n (e.g . Olausse n e t al. 1994) . The overlying middle Triassic-middle Jurassic succession has generally been assigned a post-rif t status, followin g Permian-earl y Triassi c stretching. However , th e middl e Triassic-middl e Jur assic basi n developmen t wa s no t accompanie d by simple , exponentiall y decreasin g subsidence rate a s a resul t o f therma l relaxatio n (e.g . se e McKenzie 1978) . Instead, th e interva l was characterized b y distinc t phases o f increased basina l subsidence, probabl y take n u p b y differentia l movement acros s rif t margina l an d certai n intrabasinal lineament s (Stee l & Ryset h 1990 ; Steel 1993) . Thes e variation s i n basina l subsi dence rate s an d accompanyin g variation s i n sediment suppl y controlle d th e outbuildin g an d retreat o f fiv e large-scale , regressive-transgres sive alluvia l an d shallo w marin e clasti c wedge s from th e Norwegia n an d Scottis h hinterlands . It i s argue d below , base d o n th e sedimentar y architecture of some of these clastic wedges, that gentle faul t bloc k tiltin g accompanie d a t leas t

some o f th e intervenin g interval s wit h hig h basinal subsidenc e rates . Thus , mino r exten sional stage s apparentl y occurre d repeatedl y also throughout th e middle Triassic-middle Jurassic (Faerset h & Ravnas 1998) . The Jurassi c rif t episod e affecte d th e Vikin g Graben fro m th e Bathonia n t o th e Ryazanian . Onset o f bloc k tiltin g wa s no t synchronou s throughout th e basin , an d i s argue d t o hav e commenced a s lat e a s Kimmeridgia n o n th e bordering Hord a Platfor m (Badle y e t al . 1988) . A diagnosti c featur e o f th e middl e Jurassic earliest Cretaceou s evolutio n o f th e norther n North Sea , i s tha t th e tilte d faul t bloc k sub basins appear t o have developed throug h a series of distinc t tectoni c phase s (Underbil l 199la. b: Stewart e t al . 1992 ; Partingto n e t al . 1993 ; Rattey & Ha y ward 1993 ; Faerset h e t al . 19956 ; Faerseth & Ravna s 1998) . Datin g an d chronos tratigraphical resolutio n i s stil l to o poo r t o determine whethe r ther e wa s a synchroneit y or diachronism o f tectoni c maxim a an d minima , though ther e ar e indication s o f a slightl y dia chronous riftin g acme . I t i s als o wort h notin g that th e intensit y o f th e differen t rif t phase s i s observed t o var y spatiall y acros s th e rif t basi n (see below) . Th e middl e Jurassic-earl y Cretac eous stretchin g resulte d i n characteristi c wedge shaped package s whic h onl y partiall y infil l th e fault bloc k topography/bathymetry . I n contras t to th e Permian-earl y Triassi c syn-rif t basin fill, th e sedimen t infil l o f th e middl e Jurassic earliest Cretaceou s rif t basin s wa s essentiall y of marine nature . Deltai c an d parali c sediment s were deposite d i n rif t margina l areas , periph eral t o majo r intrabasina l high s an d i n som e platform areas , wherea s th e mor e axia l an d deeper part s o f th e rif t syste m receive d mainl y (hemi-)pelagic clay s an d interbedde d sedimen t gravity flo w deposit s (e.g . Faerset h & Pedersta d 1988; Stee l 1993) . Permian-early Triassi c syn-rif t infill s In th e norther n Nort h Sea , th e Permian-earl y Triassic syn-rif t successio n ha s s o fa r onl y bee n drilled o n the Margareta' s Spur an d i n the Uns t Basin, wher e i t consist s o f alluvia l re d bed s (alluvial fa n an d fluvial sandstones, an d alluvial plain mudstones ) an d lacustrin e mudstone s (Johns & Andrew s 1985 ; Lervi k e t al . 1989) . Basal Zechstein equivalent carbonates an d evaporites ar e presen t o n th e Margareta' s Spur , where the y res t unconformabl y upo n erode d Devonian strat a (Lervi k e t al . 1989) . O n th e Horda Platform , underneat h th e Viking Grabe n and i n the East Shetlan d Basi n syn-rift strat a are

NORTH SE A SYN-RIF T SEDIMENTAR Y ARCHITECTURE S postulated base d o n thei r wedge-shape d strata l geometries (e.g . Lervi k e t al. 1989 ; Gabrielse n et al . 1990 ; Faerset h 1996) . I n thes e area s th e Permian-early Triassi c syn-rif t successio n i s u p to 2k m thick , infill s N- S oriente d half-graben s 15-35 km wide , an d i s believe d t o b e mainl y alluvial i n nature . Th e uppermos t par t o f the syn-rif t successio n ha s bee n drille d i n wel l N31/2-4 (Stee l & Ryset h 1990 , thei r fig . 9) , where i t consist s o f a sand-pron e alluvia l package. The overlyin g mud-rich package i s dated a s Scythian-Anisian i n th e nearb y N31/6- 1 wel l (Lervik e t al . 1989) . Alluvial to lacustrin e environments hav e als o bee n inferre d fo r th e earl y Triassic syn-rif t Teis t Formatio n i n th e Bery l Basin, souther n Vikin g Graben (Frostic k e t al . 1992), a s wel l a s fo r time-equivalen t strat a further sout h i n th e Norwegian-Danis h Basi n (e.g. Stee l & Ryset h 1990) . I n th e latte r are a there i s a predominanc e o f flood-basi n an d lacustrine deposit s a t increasin g distanc e fro m the rif t margina l hinterlands. Extending th e rif t episod e bac k int o th e lat e and possibl y earl y Permian , an d acceptin g th e notion tha t th e Permian-earl y Triassi c rift-axi s was centre d underneat h th e Hord a Platfor m (Faerseth 1996 ; Odinse n e t al . i n press) , impl y that potentia l marin e strat a (Zechstei n an d Kupferschiefer equivalents ) an d possibl e ari d or deser t sediment s (Rotliegende s equivalent ) may for m par t o f th e Permian-earl y Triassi c syn-rift succession . Possibl e deser t deposit s ma y include inlan d sabkh a t o aeolia n an d interdun e deposits, simila r t o th e Rotliegende s o f th e Moray Firt h an d Norther n Permia n Basi n (e.g. S0rensen & Martinse n 1987) . B y analogue with East Greenlan d an d th e souther n Nort h Sea , potential marin e deposit s ma y includ e marin e transgressive sandstone s t o offshor e claystone s and possibl e carbonate s an d evaporites . Earl y syn-rift deser t sediment s may b e presen t acros s the entir e Permian-earl y Triassi c rif t basin , whereas marin e strat a ma y hav e bee n deposite d in th e deepe r half-graben s only , i.e . o n th e Horda Platfor m an d possibl y th e norther n Viking Graben . However , th e norther n Nort h Sea i s underlain b y Caledonia n crus t tha t ma y have bee n thickene d compare d t o th e souther n North Sea during Permian-early Triassic rifting , leaving a possibilit y tha t isostas y cause d con tinental facies and environments to prevail in the northern Nort h Se a relative to th e encroachin g marine water s fro m th e north an d south . Hence, a wid e variet y o f Permian-earl y Triassic non-marin e syn-rif t deposit s ma y b e present, rangin g fro m ari d t o fluvial and lacus trine, possibl y includin g marin e strat a i n sub basins clos e t o an d alon g th e rift-axis . Spatia l

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and tempora l variation s i n infil l architectur e across th e rif t basi n ar e likel y to b e present a s a result o f th e interpla y betwee n sedimen t suppl y and extensiona l faulting , i.e . th e creatio n o f a tilted faul t bloc k topograph y an d th e half graben's dependenc e o n sedimen t suppl y fro m local or through flowing (sub-) regional drainag e (Leeder & Gawthorpe 1987 ; Steel 1993). The rate and styl e o f riftin g durin g th e Permian-earl y Triassic is uncertain. It is not know n whether the Permian-early Triassi c rif t episod e experience d pulsatory rifting , simila r to the middle Jurassicearliest Cretaceou s rif t episod e (se e below) , o r evolved throug h a singl e stretchin g interval , characterized b y repeate d faul t event s an d associated topograph y rejuvenatio n without any prolonged intervening tectonic quiescence stages. Climatic variations, from a dry, desert-like environment i n the Permian t o more humi d bu t stil l semiarid, fluvia l an d lacustrin e condition s i n the earl y Triassic , ar e als o likel y t o hav e influ enced o n th e sedimentar y architecture o f the rif t basin infill . Underlining th e likel y complexit y o f th e Permian-early Triassi c syn-rif t infill , i t i s diffi cult t o speculat e furthe r upo n th e detaile d sedimentary architectur e o f th e syn-rif t infil l without bette r wel l control . Middle Triassic-middle Jurassi c succession s Based o n th e documentatio n o f a serie s o f large-scale clastic wedges which built fro m bot h the eas t an d wes t int o th e norther n Nort h Se a basin, Stee l (1993 ) suggeste d tha t th e middl e Triassic-middle Jurassi c wa s characterize d b y pronounced tempora l variation s in basina l subsidence an d sedimen t supply . Relativel y hig h basinal subsidenc e rate s ar e postulate d fo r th e Late Carnian-Earl y Norian , Lat e Rhaetian , Middle Sinemurian , Early Toarcian, an d possi bly als o th e Lat e Pliensbachia n (Stee l & Ryseth 1990; Stee l 1993) . Moreover , loca l subsidenc e rates varie d acros s a numbe r o f majo r faul t zones, an d i t shoul d b e note d tha t som e o f th e successions deposite d durin g period s wit h hig h basinal subsidenc e rate s ar e locall y represented by wedge-shape d strata l package s (e.g . Gabriel sen e t al . 1990 ; Steel & Ryseth 1990 ; Faerseth & Ravnas 1998) . Late Carnian-earl y Noria n an d lat e Rhae tian clay-prone successions (lower-middle Lunde Formation an d uppermos t Lund e Formation Raude Membe r o f th e Statfjor d Formation) , are characteristicall y thic k downflan k o n larg e Permian-early Triassi c faul t blocks , wherea s they thi n progressivel y i n a n updi p directio n

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(e.g. Stee l & Ryset h 1990 , thei r fig . 19) . O n a basin-wide scale, the intervening sandprone unit s of th e Lomvi , uppe r Lund e an d Statfjor d For mations hav e mor e tabula r geometrie s (Stee l & Ryseth 1990 ; Stee l 1993) , indicatin g basinwid e more unifor m subsidenc e rates . Locall y als o these sand-pron e unit s ma y sho w distinc t thick ness variations acros s certai n faults, e.g. as in the Gullfaks are a (A . Dalland , pers . comm . 1997) . Increased basina l subsidenc e rate s durin g th e middle Sinemuria n resulte d i n drownin g o f the Statfjor d Formatio n alluvia l plai n an d th e establishment o f marine conditions , th e 'Dunli n sea\ i n the northern Nort h Sea . I t is noteworthy that th e sediment s representin g th e transitio n from alluvial-deltai c t o marin e conditions , th e Nansen Member , ar e present mainl y in downdi p areas o f Permian-earl y Triassi c faul t blocks . In updip reache s o n th e same faul t blocks , e. g in the Tampen Spu r area , marin e mudstone s o f the Amundsen Formatio n i s separate d b y a trans gressive la g from , o r res t directl y upo n alluvia l strata o f th e Eiriksso n Membe r (e.g . Nystue n el al 1989 , their fig . 14) . During th e middl e Triassic-earl y Jurassic , subsidence seem s t o hav e bee n accentuate d along certain masterfaults, especially those forming th e wester n boundar y fault s t o th e Vikin g Graben, th e easter n 0ygarde n Faul t Zone , an d some intraplatforma l fault s o n th e Hord a Plat form (Gabrielse n e t al . 1990 ; Stee l & Ryset h 1990; Faerset h & Ravna s 1998) . Othe r faults , which gre w t o for m masterfault s durin g th e middle-late Jurassi c rifting , probabl y wer e initiated durin g th e lat e Triassic-earl y Jurassi c (Faerseth & Ravna s 1998) . Intermittent period s o f uplif t hav e bee n documented i n th e middl e Triassi c (Stee l & Ryseth 1990 ) and i n th e lat e Triassic-earl y Jur assic (R0 e & Steel 1985) , althoug h thei r precis e dating i s difficul t du e t o th e biostratigraphica l barren natur e o f th e successions . A s mos t well s are situate d o n structura l highs , th e uplift s ma y possibly represen t intermitten t period s o f foot wall uplif t relate d to rotational faulting (see als o Hollander 1987 ; Roberts e t al . 1987). Periods o f mino r footwal l uplif t ar e recorde d for th e lates t Pliensbachian-earl y Toarcia n i n the footwal l o f th e Oseber g Faul t (Livbjer g & Mj0s 1989) , th e Alwyn-Ninian-Hutto n align ment (Johnso n & Essautie r 1987 ; Sawye r & Keegan 1996) , an d i n th e Statfjord-Gullfak s area alon g th e wester n flan k o f th e norther n Viking Grabe n (Robert s e t al . 1987 ; T. Dreyer , pers. comm . 1995) . Th e uplif t an d ?submarin e erosion resulte d i n the formatio n o f partly foot wall derive d shallo w marin e san d ridge s o f the uppe r Coo k Formatio n o n th e Oseber g

Fault Bloc k (Livbjer g & Mj0 s 1989) . Thi s occurred concomitantl y wit h th e retrea t o f th e rift margina l Coo k Formatio n shorelin e (Stee l 1993) an d drownin g o f downfaulte d basina l areas t o th e wes t (Faerset h & Ravna s 1998) , and thu s mirror s th e Callovian-Kimmeridgia n formation o f shallo w marin e syn-rif t wedge s in th e sam e are a (se e below ; Ravna s & Bon devik 1997) . Part s o f th e Coo k Formatio n i n the Gullfak s are a ma y hav e a simila r origi n (T. Dreyer, pers . comm. 1995) , although th e unit in thi s are a probabl y represent s tidall y influ enced shoreline s o r tida l estuarie s (Dreye r & Wiig 1995 ; Marjana c & Stee l 1997) . I t i s als o noteworthy tha t som e rotationa l faultin g ha s been invoke d t o explai n th e sedimentar y archi tecture o f Sinemurian-Earl y Toarcia n strat a i n the Bery l Embaymen t (Richard s 1991) . Minor rotationa l faultin g ha s furthe r bee n postulated t o occu r locall y i n th e Aalenian , e.g. i n th e Eas t Shetlan d Basi n (Robert s e t al . 1987; Ratte y & Ha y ward 1993 ; N0ttved t e t al . 1995) an d o n th e Hord a Platform , wher e i t exerted a majo r contro l o n th e presenc e an d distribution o f the Aalenian Oseber g Formatio n (Graue e t al . 1987 ; Mut o & Stee l 1997) . Thi s faulting even t was possibly related t o th e middle Jurassic domal uplif t centre d beneat h th e tripple junction o f th e Nort h se a rif t basi n (Ziegle r 19900,6; Underhil l & Partingto n 1993) , o r t o uplift an d rejuvenatio n o f th e rif t margina l hinterlands (Stee l 1993) . The abov e revie w indicate s tha t th e middl e Triassic-middle Jurassi c perio d i n th e norther n North Se a wa s characterize d b y a serie s o f repeated subsidenc e events, which where at leas t in som e area s associate d wit h gentl e rotationa l faulting. Deformatio n i s generally located abov e the Permian-earl y Triassi c rif t structure , bu t apparently focuse d t o area s beneat h an d adja cent t o th e lat e Jurassic-earl y Cretaceou s rift axis (Faerseth & Ravnas 1998) . On th e bordering platformal areas , i.e . the Hord a Platfor m an d i n the East Shetlan d Basin , the late Triassic-middl e Jurassic successio n is , i n general , characterize d by tabula r t o subtabula r sedimen t geometries , resembling pre - o r post-rif t strata . The mechanis m behin d th e middl e Triassic middle Jurassic subsidence events is presently not clear. Two non-extensiona l modes o f subsidenc e can b e envisaged ; differentia l compactio n o f the strongl y wedge shaped Permian-earl y Trias sic syn-rif t half-grabe n infill s mus t hav e le d t o differential subsidenc e an d sli p o n man y o f th e tilted bloc k boundar y fault s durin g subsequen t thermal relaxation , an d th e uneve n natur e o f the therma l subsidenc e itsel f fro m a minimu m at th e basin margin s to a maximum towards th e

NORTH SE A SYN-RIF T SEDIMENTAR Y ARCHITECTURE S basin axi s i s als o likel y t o hav e bee n accom modated b y severa l o f th e olde r faul t systems . Compaction drive n subsidence was most impor tant immediatel y subsequen t t o th e stretchin g phase. Later , an d a s a n alternativ e mechanism , basin development was accompanied b y renewed crustal stretching , albei t no t o f significan t mag nitude. The indications of differential subsidenc e and rotationa l faultin g continuin g int o th e early and middle Jurassic may support a component o f repetitive, wea k lithospheric stretching . This suggest s tha t th e typica l threefol d sub division int o proto- , syn- an d post-rif t stages , commonly use d i n th e analysi s o f rif t basi n evolution an d infil l history , need s furthe r modifications whe n dealin g wit h rif t basin s that experience d multipl e rif t episodes . Pro longed interval s with a duratio n o f ten s o f M a occurring betwee n rif t episodes , an d character ized b y mor e diffus e extension , ma y b e bette r referred t o a s inter-rift periods. Middle Jurassic-earliest Cretaceou s marin e syn-rift architecture s

Viking Graben at 60°N (Profile 1) The Vikin g Grabe n a t 60° N a s show n i n Pro file 1 (Fig. 2a) , consists of a solitary half-graben , some 50k m wide . The half-grabe n i s separate d from borderin g platfor m area s b y a masterfault to th e wes t an d a flexure d an d faulted , gentl y sloping eastern margin . Som e intrabasinal high s are present . Alon g th e profil e ther e ar e a fe w scattered 'basinal ' well s (e.g. N30/7-7, N30/10-5, N30/10-6, N30/11-3 , N30/11-4) , whic h giv e some insigh t int o th e sedimentar y architectur e of the early, Bajocian-Bathonian syn-rif t basina l infill. A t presen t ther e ar e to o littl e dat a t o speculate on the sedimentary architecture o f the remaining part o f the syn-rif t succession , i.e. the late Bathonian-Volgia n Vikin g Group , alon g this profile . The uppermost Nes s and Tarbert Formation s are late Bajocian-Bathonian i n age and togethe r they constitut e thre e transgressive-regressiv e sequences, eac h som e 150-20 0 m thic k (Fig . 4) . The uppermos t par t o f th e Tarber t Forma tion form s th e transgressiv e segmen t o f a n additional sequence . Th e transgressive-regres sive sequence s constitut e threefold , sandstone mudstone-sandstone packages , an d are believed to b e a resul t o f a pulsator y styl e of rotationa l faulting (Fig . 5) . Th e sedimentar y architectur e of eac h transgressive-regressiv e sequenc e sug gests tha t progradatio n occurre d durin g tecto nically dorman t stage s wherea s backsteppin g o f

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the shorelin e too k plac e durin g intervenin g stages wit h highe r rate s o f rotationa l fault ing. Fou r distinc t rotationa l faultin g event s ar e recognized fro m th e earl y Bajocia n t o th e middle/late Bathonian , an d ar e groupe d to gether a s th e Bajocian-Bathonia n rif t maxim a (Figs 3, 5). A similar relationship is found in th e uppermost Nes s an d Tarber t Formation s i n the Hil d an d Oseber g area s t o th e nort h (se e below; Ravnas e t al 1997) . Two trend s o f sedimen t partitionin g withi n the sequence s ar e wort h noting . First , th e transgressive segment s ar e thic k i n downdi p areas whereas they thin in an updip direction on the flexure d half-grabe n margin . Toward s th e bordering platfor m the y ar e represente d b y a n unconformity. I n contrast , th e regressiv e seg ments appea r t o b e o f mor e eve n thicknes s i n downdip an d media l reache s o f th e flexured , hangingwall margin , wherea s the y als o thi n i n an updi p directio n o n th e hangingwall . Sec ondly, the transgressive segment s ar e commonly the thicker member of the lower sequences. Thi s is gradually reversed upward s i n th e succession , where the regressive segments become thicke r a t the expens e o f th e transgressive . I t shoul d als o be noted tha t th e transgressive segment s usually are dominate d b y coarse r sedimen t tha n th e regressive. Althoug h tida l influenc e i s commo n throughout, it tend s to be more prevalent i n the transgressive segments . In downdip areas, th e regressive-transgressiv e turnaround i s characterized b y an aggradationa l shoreline an d thick , heterolithic an d ofte n coal bearing strat a i n mor e landwar d position s (e.g . well N30/10-6 , Fig . 4) . Thi s i s interprete d a s a result o f increased basina l subsidenc e rate s an d relative sea-level rise, which, in turn, ar e relate d to increase d rate s o f rotationa l faulting . Initi ally, the high sediment supply kept pace with the increasing rates o f relative sea-level rise, leading to a temporary stationary position of the palaeocoastline an d a n aggradationa l stackin g patter n of parali c an d shorelin e lithosomes . Coal-bear ing strat a ar e therefor e particularly commo n i n the lat e progradationa l t o earl y backsteppin g stage, before marine conditions were (re-) established alon g th e axi s of the rif t basin . Sediment gravit y flo w deposit s ar e presen t in downdi p area s i n th e uppe r sequences , im plying tha t a t thi s stag e deepe r marin e condi tions had bee n established alon g the axi s of the early Viking Graben (e.g . wel l N30/7-7 , Ravna s et a l 1997 , thei r fig . 16) . Thus , th e retrea t an d eventual drownin g of the Brent-delt a was due to sedimen t suppl y no t bein g abl e t o kee p pace wit h increasin g basina l subsidenc e rates . Because o f th e assume d ver y lo w gradien t o f

Fig. 4. E W wel l correlation o f th e late Bajocia n Bathonia n Uppe r Bren t Group, centra l Vikin g Grahen segment . Fou r interval s of higher tilt-rates are inferre d fro m th e sedimentary architecture . Syn-rotational strat a for m aggradational-to-hacksteppin g packages , includin g the lowe r par t o f th e progradationa l packages . Intervenin g tectonic quiescenc e strat a ar e represente d b y th e mai n bod y o f th e progradationa l packages .

Fig. 5 . Schemati c lat e Bajocian-Bathonia n infillin g o f th e centra l Viking Graben segmen t (southern half of Norwegian Quadran t 30) . The faulte d righ t hal f o f the profiles represent th e hanging-wal l dip slop e i n Fig . 4 whereas th e deepes t basi n an d terrac e t o th e lef t represen t position s o f well 30/10-6 an d 30/10-5 , respectively . The inferre d footwal l further t o th e lef t i s not show n i n Fig . 4 .

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the Bren t del plain alon g a basinal N- S strik e profile, singl e larg e earthquakes wit h associate d slip o n masterfault s (i n th e rang e o f a fe w metres), potentiall y coul d hav e le d t o floodin g events reachin g severa l kilometre s landwar d along th e axi s o f th e earl y rif t basin . Repeate d earthquake event s i n combinatio n wit h eustas y may therefor e hav e resulte d i n repeated , south wards encroachin g floodin g event s acros s th e Brent-delta and the formation of stacked higherorder sequence s or parasequences . Eac h higher order transgressive-regressiv e sequenc e ma y

accordingly represen t distinc t faultin g events , although suc h a linkage is difficult t o confirm . Ahvyn-Hild area (Profile 2) Profile 2 (Fig. 2b) crosses north o f the Hild area, which form s a horst-and-graben faulte d terrace , located a t th e termina l junctio n betwee n tw o crescent shape d masterfaults . Thes e fault s con stitute th e wester n boundar y faul t syste m o f the northern Viking Graben (Fig . 1) . During the

Fig. 6. E— W wel l correlatio n of the lat e Bajocian-Bathonian Tarber t Formation and th e Bathonian-Ryazanian Viking Group in the Hil d area . Interval s of higher tilt-rate s in the Tarbert Formatio n typicall y are represente d by a threefol d aggradational-to-backstepping-to-intial progradational package, whereas tectonic quiescence strata constitut e the reminde r of the progradationa l package, i.e. similar t o wha t is inferred fo r th e Tarbert Formation i n Fig. 4. (For a detailed correlatio n o f the Tarbert Formatio n stackin g pattern, th e reader is referred to Ravna s & Steel (1998) , their Fig . 9) . Syn-rotational strata i n the mud-pron e Heather Formatio n ar e overall backstepping , wherea s intervening tectonic quiescenc e strat a ar e overall forestepping-to-aggradational . Syn-rotational strat a i n the Draupne Formation for m a forestepping-to-backstepping wedge , underlain an d overlain b y tectonic quiescence, condensed claystones.

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Fig. 7 . Schemati c structura l developement and infillin g during one of the Bathonian-Kimmeridgian (Heather Formation) rif t phases in the Hild area. The fault bloc k to the left represen t the footwall to the Hild terraces, i.e. the southern continuation of the Alwyn structure. The central graben and horst topograph y represent s th e Hild terraces, whereas the deeper basin to th e right represents the Viking Graben proper .

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early (Bajocian-Bathonian) rif t maxima the Hild area formed the western, downfaulted margin of the asymmetrica l earl y Vikin g Graben (Faerset h & Ravnas 1998) , and the area experienced fault related subsidenc e whic h cause d sligh t tiltin g o f several sub-basins . Alread y a t thi s earl y stage , differential subsidenc e acros s segment s o f th e initially broade r Vikin g Grabe n cause d west dipping palaeoslope s wit h ope n marin e connec tions i n downdip area s an d coal-bearing coasta l plain environment s i n a n updi p an d landward , i.e. southward , direction . The Tarber t Formatio n (lat e Bajocian-Bath onian) is unusually thick in this area (Fig. 6), and four transgressive-regressiv e sequences , possibly reflecting th e sam e numbe r o f differentia l sub sidence events , ca n be mapped (se e also Ravnas & Steel 1997 , their fig. 9). There were significant focus and sediment volume partitioning between the coastal plain and the marine shorefac e area s at thi s stage , thoug h th e tiltin g wa s never suffi cient t o produc e unconformitie s ove r cresta l reaches. Successive sequences show an irregular, but overal l landward-steppin g tren d prio r t o the Heathe r transgression . Sedimen t volume s were partitione d i n th e sequence s suc h tha t th e transgressive segments , whic h includ e mos t o f the coal-bearin g back-barrie r an d coasta l plai n deposits, thicke n landwards , whereas the shore face-dominated regressiv e segments thicken seawards. An y sequenc e therefor e consist s o f tw o oppositely thickenin g components . From th e Lat e Bathonian , majo r faultin g shifted t o th e masterfault s t o th e eas t o f th e Hild area , i.e . th e presen t wester n boundar y faults of the Viking Graben. Thi s caused the Hild area t o becom e a terrac e betwee n th e Vikin g Graben prope r an d th e structurall y shallowe r half-grabens o f th e Eas t Shetlan d Basi n t o th e west (Fig. 6) . As such, the Hild are a experienced a dramati c chang e i n structura l settin g fro m being par t o f a wide r basina l are a durin g th e Bajocian-Bathonian rif t maxim a t o a positio n in th e footwal l o f th e rif t basin' s wester n boundary faul t syste m durin g th e subsequen t

phases. Accordingly , th e Callovian-?Volgia n rift phases ar e recorde d b y a series of unconformity strands , eac h formin g a separat e syn-rif t unconformity (Faerset h & Ravna s 1998) , an d overlying clay-prone or condensed sections . The unconformities ma y a t leas t locall y b e o f sub marine natur e (Fig . 7). During th e Callovian-?Volgian , mino r tilting of th e terrac e are a i n additio n t o th e forma tion o f a n intraterrac e horst-and-grabe n topo graphy, resulte d i n continue d depositio n i n intraterrace structura l lows . However , th e syn rotational deposit s wer e derive d b y erosio n o f older, dominantl y fine-grained middle-late Jurassic syn-rif t sediments . Th e lat e Bathonian Kimmeridgian Heathe r Formatio n i s therefor e dominated b y a relativel y homogenou s mud prone succession (Figs 6 , 7). Possible candidate s for coarse-graine d syn-rotationa l deposit s ar e located clos e t o an d abuttin g agains t th e western an d easter n terrace-boundin g masterfaults . Despite th e mudprone natur e o f the succession , three syn-rif t wedge s ar e recognize d whic h appear t o mirro r thei r time-equivalen t strat a on th e Oseber g Faul t Bloc k o n th e opposin g eastern flan k o f the Vikin g Graben (se e below). The well data, however, are scant an d giv e little detailed informatio n o n th e three-dimensiona l sedimentary architectur e of the syn-rif t wedges. The Kimmeridgian-early Ryazanian Draupne Formation is too thin to be subdivided on seismic data. Again , the availabl e seismic an d wel l dat a suggest that the tectono-sedimentary regime prevalent durin g depositio n o f th e Draupn e For mation i n th e Hil d are a mirrore d tha t o f th e Draupne Formation o n the Oseberg Fault Block. Seismic dat a sugges t tha t coarse r materia l i s present i n th e hangingwall o f th e large r faults , forming ?isolate d fault-scar p aprons . Introduc tion o f coarse r materia l int o th e deepe r marine realm i s correlate d wit h period s o f increase d tectonic activity . This wa s accompanied b y pronounced faul t bloc k rotation , uplift , sedimen t instability an d gravitationa l resedimentation . The latte r wa s particularl y commo n dow n th e

Fig. 8. E- W wel l correlation o f the late Bajocian-Bathonian Tarber t Formatio n an d th e Bathonian-Ryazanian Viking Group i n the Oseberg-Brage area. Interval s of higher tilt-rate s i n the Tarbert Formatio n typicall y are represented b y wedge-shaped, threefol d aggradational-backstepping-intia l progradationa l package , whereas tectonic quiescenc e strat a constitute the main body o f the progradational package , i.e. similar to what is inferred for th e Tarbert Formatio n in Fig. 4 . Syn-rotational strat a i n the Heather Formatio n ar e represented by forestepping-to-aggradational-to-backstepping sand-prone , fault-scar p derive d successio n downdip on th e hangingwall. mud-prone basinal successions , and forestepping-to-backsteppin g sand-prone footwall successions. Intervening tectonic quiescence strata show a progradational-to-aggradational stacking pattern. Facies variations are observed between sub-basins, and between the syn-rotational and tectonic quiescence strata, reflecting spatia l and tempora l variation s in basin physiograph y durin g both Tarbert an d Heathe r Formation s depositio n (Ravnas & Bondevik 1997 ; Ravnas el al. 1997) . Syn-rotationa l strata i n the Draupn e Formatio n for m a forestepping-to-backstepping wedge, underlain and overlain by tectonic quiescence, condensed claystones. similar to wha t is inferred fo r th e Hil d are a (Fig . 6).

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sub-basin-bounding masterfaults . I n contrast , intervening period s o f relativ e tectoni c quies cence wer e dominate d b y drapin g o f hemipela gic an d pelagi c claystones , commonl y wit h a high source-roc k potential , acros s th e entir e terrace area .

Oseberg-Brage area (Profile 2) In th e Oseber g an d Brag e area s (Fig. 2b) , which form th e easter n flan k o f th e norther n Vikin g Graben an d th e wester n margi n o f th e Hord a Platform respectively , middle Jurassi c stretchin g commenced i n th e Bajocian-Bathonia n (Hel land-Hansen e t al 1992 ; Ravna s e t al. 1997) . During th e earl y (Bajocian-Bathonian ) rifting , the area forme d the faulted eastern margin of the asymmetrical (westwar d sloping ) earl y Vikin g Graben. Fro m th e lat e Bathonia n t o th e Kim meridgian, th e Oseber g Faul t Bloc k togethe r with th e smalle r faul t block s alon g it s western flank, forme d a n u p t o 2 5 km wide , N- S elon gate, eastward-tilte d mega faul t block . Collaps e and segmentatio n o f th e Oseber g meg a faul t block occurre d i n th e Kimmeridgian-Volgian , and resulte d i n th e presen t structura l configu ration o f th e are a (Faerset h & Ravna s 1998) . This interva l wa s characterize d b y increase d subsidence alon g th e NE-S W trendin g struc tural grain . The overall , large-scal e tren d o f th e middle late Jurassi c syn-rif t infil l i s on e o f fining upwards, fro m th e parali c an d shallow-marin e sandstones o f th e Tarber t Formation , throug h the shallow-marin e t o shelfa l san d an d mud stones o f the Heathe r Formation , t o th e deepe r marine claystone s o f th e Draupn e Formatio n (Fig. 8) . Thi s large-scal e fining-upwar d succes sion record s th e overal l progressiv e deepenin g of th e marin e rif t basi n a s th e rif t evolved , from a mixe d non-marin e an d shallo w marin e basin i n th e Bajocian-Bathonia n int o a shallo w marine basi n i n th e Callovian-Kimmeridgian , and the n int o a deepe r marin e basi n i n th e Kimmeridgian-Volgian. Successiv e rift phase s in the Bajocian-Oxfordian/Kimmeridgia n resulte d in th e depositio n o f a serie s o f superimpose d shallow marine , wedge-shape d packages . Eac h wedge-shaped packag e correlate s wit h a distinct rift phas e (Ravna s & Bondevi k 1997 ; Ravna s et al . 1997 ; Faerset h & Ravna s 1998) . Sedimen t was supplie d fro m rif t margina l sources , firs t from th e axia l Brent-delt a an d late r fro m th e transverse Troll-delta , durin g period s o f relative tectonic quiescence . Durin g th e intervenin g periods wit h hig h rate s o f rotationa l faulting ,

i.e. th e syn-rotationa l stages , sedimen t wa s sup plied predominantl y fro m th e uplifte d updi p areas o f th e faul t block s o r rif t interio r sources . The intervals corresponding t o hig h extension rates i n th e Bajocian-Bathonia n ar e character ized b y a threefol d lithosom e model , consistin g of basal an d capping sandstone s an d intervening mudstones (Ravna s e t al . 1997) . Three intervals of hig h extensio n rate s ar e recognized , th e first in th e lates t Earl y Bajocia n (uppermos t Nes s and lowermos t Tarber t Formations) , th e nex t two i n th e lat e Bajocia n (intr a Tarber t Forma tion) an d th e early-middle Bathonia n (Tarbert Heather transition) , respectively . Although thes e units ar e relativel y thi n o n th e Oseber g Faul t Block, thicke r succession s (<200m ) ar e recog nized i n th e borderin g terrac e are a an d i n th e basinal Vikin g Grabe n well s (Fig . 8 ; se e also Ravnas e t al . 1997 , their fig . 15) . The sedimen tary architectur e and inferre d tectono-stratigra phical developmen t o f th e uppermos t Nes s and Tarber t Formatio n transgressive-regressiv e sequences mirro r tha t inferre d fo r th e time equivalent strat a i n th e centra l segmen t o f th e Viking Grabe n a t 60 CN (se e above) . Further more, Ravna s e t al . (1997 ) argued , base d o n the stackin g pattern o f th e transgressive-regres sive sequences , tha t th e lat e Bajocian-earl y Bathonian rift-maxim a wa s characterize d b y a series o f smaller-scal e rotationa l til t events . These shorter-ter m rotationa l til t event s ar e here referre d t o a s deformatio n spans . Ravna s et al . (1997 ) als o suggeste d tha t th e threefol d lithosome mode l establishe d fo r mixe d non marine an d shallo w marin e syn-rif t succession s applies a t severa l scales ; bot h th e longe r ter m late Bajocian-earl y Bathonia n rif t maxim a a s well a s th e shorte r ter m rif t event s o r deforma tion spans . The intervals corresponding t o hig h extension rates i n th e middle/?lat e Bathonian-Kimmerid gian shallo w marin e syn-rif t succession s ar e characterized b y a twofold , sandstone-claystone lithosome mode l (Fig s 8 , 9 ; Ravna s & Bonde vik 1997) . Again, thre e interval s o f hig h exten sion rat e ar e recognized . Thes e occurre d i n th e middle-late Bathonia n (lowe r Heather member , upper part) , lat e Callovian-earl y Oxfordia n (middle Heathe r member , uppe r part ) an d th e late Oxfordian-Kimmeridgia n (uppe r Heathe r member, uppe r part) . Th e lac k o f sandston e cappings to th e syn-rotational intervals is attributed t o th e fac t tha t th e sediment s deposite d during hig h extension rate s wer e mainly source d from adjacen t footwal l highs. The peneplanatio n of th e rif t interio r source s durin g th e lat e syn rotational an d subsequen t tectoni c quiescenc e stages, togethe r wit h th e larg e distanc e t o th e

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Fig. 9 . Schemati c structural developement and infillin g durin g one of the Callovian-Kimmeridgian (Heathe r Formation) rif t phase s i n the Oseberg-Brage area an d o n th e Horda Platform . The half-graben t o th e right represent th e Oseberg Faul t Block , whereas the larger, meg a faul t bloc k t o th e right represen t th e Hord a Platform.

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Fig. 10 . Schemati c structura l development an d infillin g o f during Kimmeridgian-Volgian (Draupn e Formation ) rifting i n the Oseberg-Brag e area. Th e Oseber g Fault-Bloc k is presented i n simplified, schematic manner as a medium-scale half-grabe n wit h a faulte d hanging wall dip slope.

NORTH SE A SYN-RIF T SEDIMENTAR Y ARCHITECTURE S (then progradational ) rif t margina l shoreline , rendered th e Oseberg are a sediment starve d well into th e subsequen t tectoni c quiescenc e stage . During th e Bathonian-Oxfordia n rif t phases , complete infillin g o f th e half-grabe n coul d no t occur befor e th e rif t margina l shorelin e onc e again ha d approache d th e are a (Fig . 9 ; Ravnas & Bondevi k 1997) . The syn-rotationa l strat a i n the Brag e area an d o n th e Oseberg Faul t Bloc k show a simila r sediment architechture, onl y differing i n sedimen t calibr e o f thei r infil l (Fig . 8) . Intervals of higher extension rate in the deepe r marine Kimmeridgian-Volgia n syn-rif t succes sion ar e als o characterize d b y a twofold , ?conglomerate-sandstone-claystone lithosom e model (Fig s 8 , 10) . Durin g thi s interva l th e sandprone strat a an d coarse r lithologie s ar e interpreted t o have been derived by erosion fro m adjacent highs , and redeposite d b y gravitationa l resedimentation processe s int o th e adjacen t sub-basins (Ravna s & Stee l 1997) . Th e coarse r syn-rotational materia l constitut e stacke d ?con glomerate-sandstone apron s alon g th e half graben-bounding masterfaults , and sand-pron e turbidite succession s i n hangingwal l structura l lows. The latter units probably represent a series of ponded , sheet-lik e turbidites . Th e hanging wall fault-scar p apron s ar e assume d t o b e overall mud-prone , consistin g o f debri s flo w deposits an d slumps/slides . The intervenin g tectonic quiescenc e interval s are characterize d b y suspension fall-ou t wit h a lo w influ x o f clasti c sediment. Th e sediment s deposite d durin g th e tectonic quiescenc e stage s for m half-graben wide clayston e drapes , whic h appea r t o hav e blanketed als o mos t o f th e footwall s o f th e half-graben-bounding masterfault s (Ravna s & Steel 1997) . Claystones wit h a hig h source-roc k poten tial predominantl y forme d durin g intervenin g tectonic quiescence intervals from th e late Bathonian t o th e Volgian , especiall y durin g th e Kimmeridgian-Volgian. I n th e latte r period , displacement alon g N- S an d NE-S W trendin g faults probabl y restricte d water circulation considerably, resultin g i n th e formatio n o f rhomb shaped sub-basin s with a stratified water column and anaerobic bottom water conditions (Faerseth & Ravnas 1998) . Horda Platform (Profile 2) Prior to the late Oxfordian , th e Horda Platfor m can b e considere d a s a gentl y eastwar d tilte d mega faul t block , bounde d b y th e 0ygarde n Fault Zon e i n th e eas t an d th e Vikin g Graben and it s flanking fault blocks an d terrace s t o th e

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west (Fig. 2b ; Faerseth 1996) . Figure 1 1 shows an E-W wel l cross-sectio n fro m th e Trol l fiel d i n the eas t t o th e Brag e fiel d i n th e west . Thi s section illustrate s th e comple x outbuildin g an d retreat o f transvers e rif t margina l shallo w marine clasti c wedge s durin g interval s wit h gentle syn-rif t extension an d faul t block tilting . Following a perio d o f wea k tiltin g an d drowning i n th e Bathonian , th e Fensfjor d Formation shorelin e advance d toward s th e Brage are a durin g th e earl y an d middl e Callo vian (Helle m e t al 1986 ; Stee l 1993 ; Stewar t et al 1995 ; Ravnas & Bondevik 1997) . Renewed extension an d increasin g rate s o f basina l sub sidence resulte d i n eastwar d retrea t o f th e rif t marginal shoreline , acros s th e Hord a Platfor m in th e lat e Callovia n (Fig . 9 ; Stee l 1993) . Concomitantly, ther e wa s uplif t o f th e Brag e area a s a resul t o f th e combine d effect s o f footwall uplif t alon g th e Brag e faul t an d gentl e clockwise rotation o f the Horda Platfor m meg a fault bloc k (Faerset h & Ravna s 1998) . Thi s resulted i n th e formatio n o f locall y derive d shallow marin e lithosomes , bot h o n th e Brag e area an d i n th e Oseber g sub-basi n t o th e wes t (Ravnas & Bondevi k 1997) . Pea k eastwar d progradation of the Brage hangingwall shoreline took plac e close to the rotational til t climax, and was nearl y coeva l wit h maximu m eastwar d retreat o f th e rif t margina l shorelin e o n th e Horda Platform . Syn-rotationa l sand-pron e strata aroun d th e expose d footwal l thu s corre late wit h a clay-pron e interva l i n a downdi p and landwar d direction , i.e . toward s th e eas t (Figs 9 , 11) . Durin g th e followin g period wit h waning an d lo w rate s o f extensiona l faulting , sediment suppl y fro m th e latera l hinterland s again wa s sufficien t t o outpac e basina l subsi dence, an d ther e wa s outbuildin g o f th e rif t marginal Sognefjor d Formation shorelin e (Steel 1993). Th e earl y stage s o f progradatio n o f th e Sognefjord Formatio n shorelin e wa s contem poraneous wit h sourceward retrea t o f the locally derived shoreline s i n th e Brag e are a an d th e eventual drownin g and drapin g o f th e expose d footwall islands by marine condensed claystones . These tempora l an d spatia l variation s i n subsidence/uplift rate s an d sedimen t suppl y resulted i n a marke d diachroneit y i n th e devel opment o f clay - an d sand-pron e intervals . In shallo w marin e half-graben s dominate d b y infilling fro m th e transvers e rift margina l shoreline suc h a s th e Callovian-Oxfordia n Hord a Platform, th e rotationa l til t clima x was accom panied b y maximu m subsidence rate s an d pea k landward floodin g i n downdi p areas , wherea s updip area s experience d uplif t an d commonl y subaerial erosion. Althoug h th e basin as a whole

NORTH SE A SYN-RIF T SEDIMENTAR Y ARCHITECTURE S is dominate d b y a threefol d sandstone-mud stone-sandstone syn-rift successio n (Fig s 9 , 11) , there ar e unique syn-rif t sedimentar y signature s to eac h half-graben-specifi c setting : half-grabe n downdip or in basinal areas are characterized by a syn-rotationa l transgressive-regressiv e devel opment (e.g . a s see n i n th e Trol l are a wells) , whereas faul t bloc k updi p area s sho w pro gradation-aggradation-backstepping o f locall y derived sedimentar y lobe s (e.g . a s documente d for th e evolutio n i n th e footwal l o f th e Brag e Fault, i.e . th e Brag e field , se e als o Ravna s & Bondevik 1997) . A simila r scenari o i s als o envisage d fo r th e late Oxfordian-Kimmeridgia n rif t maxima . Higher rate s o f rotationa l faultin g an d basi n floor subsidenc e resulte d i n th e retrea t o f th e rift margina l Sognefjor d Formatio n shorelin e (Fig. 11) . Th e formatio n o f late Oxfordia n an d younger unconformities in the footwall o f intraplatformal fault s (e.g . wel l N31/2-5 , Fig . 11) , suggest tha t th e forme r Hord a Platfor m meg a fault bloc k the n ha d becom e disrupte d int o a series o f smalle r faul t blocks . Accordingly , during th e lat e Oxfordian-Kimmeridgia n an d subsequent rift phases there were several smaller intraplatformal sourc e area s whic h complicated the mor e simpl e pictur e describe d fo r th e lat e Callovian-early Oxfordian rif t phase . However , it i s anticipated tha t th e easternmost , i.e . land ward, half-graben s wil l sho w a simila r syn rotational infil l tren d t o tha t outline d above . Increased basina l subsidenc e rate s toward s th e late Jurassi c rift-axis , sugges t tha t half-graben s located successivel y toward s th e wes t wer e dominated b y progressively deepe r depositiona l environments. Thus, a chang e fro m mixe d nonmarine and shallow marine environments on the Horda Platform through predominantly shallow marine i n th e Brag e are a an d o n th e Oseber g Fault Block , t o deepe r marin e environment s in the Vikin g Graben prope r i s expected.

Penguin half-graben (Profile 3) The Penguin half-graben (Fig. 2c), located i n the footwall o f th e End-of-the-Worl d Fault , i s a

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late Jurassic-earl y Cretaceou s half-grabe n deli neated b y N-S strikin g middle-late Jurassic an d NE-SW strikin g lat e Jurassic-earl y Cretaceou s normal fault s (Fig . 1) . This sub-basi n contain s an exampl e o f deepe r marin e syn-rif t infill , th e late Oxfordian-Ryazania n Kimmeridg e Cla y Formation an d it s intercalate d Ptarmiga n an d Magnus Sandston e Member s (Fig . 12) . Tw o intervals o f hig h extensio n rate s followe d b y periods of relative tectonic quiescence are recognized i n th e lat e Oxfordian-Ryazania n interva l (Ravnas & Stee l 1997) . Eac h o f these tw o rift phases spa n som e 5-6 Ma. The Bathonian Oxfordian shallo w marine syn-rift strata , i.e. the Tarbert an d Heathe r Formations , hav e no t been studie d i n detai l an d wil l therefor e no t be described . The older rif t phase , date d a s late Oxfordian Kimmeridgian, wa s accommodate d b y move ments mainl y o n N- S trendin g structures . The interva l correspondin g t o hig h rate s o f extensional faultin g i s represente d b y a three fold, sandstone-mudstone-sandston e lithosome model, sometimes with a thin basal, coarsening upward part (Fig . 13 ; Ravnas & Steel 1997) . The basal coarsening-upwar d par t reflect s sheddin g of detritus from adjacen t highs during the initial tilting stage s (Fig. 13a , b). Depositio n wa s fro m a wid e range o f sedimen t gravit y flows, including slides/slumps , cohesiv e debri s flows , non cohesive sand y debri s flows , an d high - an d low-density turbidit y currents. Th e backgroun d mudstones represen t hemipelagi c an d pelagi c sedimentation. Th e overlyin g fining-upward s interval is a highly disorganized succession. The fining-upward signatur e is defined b y an upwar d increase i n mu d content , wherea s sandie r inter vals commonl y ar e thinne r an d volumetricall y less abundant . Sand y sedimen t gravit y flo w lobes (Ptarmiga n Sandston e Member ) stac k i n a random, bu t overall downdip shiftin g manner. These ar e particularl y commo n durin g th e interval o f inferre d maximu m subsidenc e rate , i.e. th e til t clima x (Fig . 13c) . Sedimen t wa s delivered bot h fro m longitudinall y disperse d flows a s wel l a s latera l sources . Fault-scar p successions, which represent submarin e fans an d

Fig. 11 . E- W wel l correlation o f the Bathonian-Ryazania n Vikin g Group on th e Horda Platform , i.e. from th e Troll t o th e Brag e fields. In th e Bathonian-Kimmeridgian Heather , Krossfjord , Fensfjor d an d Sognefjor d Formations syn-rotationa l strata for m threefol d backstepping-to-forestepping packages in downdip positions, i.e. t o th e east, wherea s forestepping-to-backsteppin g infil l trend s are develope d i n updip positions, i.e . to th e west. In th e upper syn-rotationa l sandstone packag e (Sognefjord-Draupn e Formations ) th e upper foresteppin g part i s less developed. In th e Kimmeridgian-Volgia n Draupn e Formatio n potentia l syn-rotational strata show similar trend s in updip an d downdi p positions a s described fo r underlyin g syn-rotational packages, bu t ar e overall clay-prone. Syn-rotationa l forestepping-to-backstepping sandstone s ar e onl y present locall y on som e fault-blocks, e.g . well s 31/2- 5 (thin transgressive lag) and 31/2- 1 (uppermost thin sandstones stringers in overall silty succession) .

Fig. 12 . E W wel l correlation o f th e Kimmeridg e Clay Formatio n an d it s intercalated Ptarmigan an d Magnu s Sandston e Member s i n the Pengui n half-graben (modified from Ravna s & Steel 1997) . Two syn-rotational packages are present: th e lower form a threefold sandston e mudston e sandston e packag e with abundant sedimen t gravity flow sandstones interbedde d wit h th e middle mudstones. The uppe r syn-rotationa l packages show s a twofol d sandston e mudston e motif . Forestepping-to-aggradational to-backstepping fault-scar p succession s ar e presen t i n bot h th e lowe r and th e uppe r syn-rotationa l package. Tectoni c quiescen e strat a separatin g th e tw o syn-rotationa l m<'k:i»™ mmnrises a nronradational-to-atitzradationa l sandstone nackaccc, whereas the upper tectonic quiescence/early post-rift interval consists o f condensed claystones .

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Fig. 13 . Schemati c structura l developmen t an d infillin g during Late Oxfordian-Kimmeridgia n riftin g an d th e subsequent Kimmeridgian-Earl y Volgia n tectonic quiescenc e stage i n the Penguin half-graben . Th e Pengui n half-graben i s presented in simplified, schemati c manner as a medium-scale fault block with a faulted hangin g wall dip slope .

Fig. 14 . (a ) NW-SE an d (b ) SW-NE wel l correlations o f th e Draupn e Formatio n an d intercalate d Muni n Sandston e Membe r i n th e Statfjor d an d Statfjor d Nort h areas . Syn-rotational deeper-marine sandstones are present in half-graben sink s to th e west and northwes t where they are overlain by offshore mudstones , which, in turn, are overlain by shallow-marine sandstones. In more proximal positions the deeper-marine sandstone s ar e lacking; only offshore mudstone s and/or shallow-marin e sandstone capping s ar e present . In th e mos t proxima l wells , i.e. towards th e faul t bloc k crest , th e interva l is represented b y a n unconformity , locally with overlying transgressive lag deposits .

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aprons, commonl y hav e a coarsening-to-finin g upwards signatur e (Fig . 12) . Larger, thicke r an d commonly mor e sand-pron e sedimentar y lobe s occur basinwar d o f transfe r zones . Th e fining upward successio n i s capped b y thick claystones (Fig. 13d ) which ar e partl y coeva l wit h sandie r sediment gravit y flo w deposit s t o th e north , i.e. proximall y wit h respec t t o th e sub-basin' s hinterland area. Th e sandstone cappings , i.e . the lower par t o f th e Magnu s Sandston e Member , have a shar p base , an d recor d th e incom e o f confined, high-continuity , foresteppin g o r pro gradational, sand y sedimen t gravit y flow lobes. These ha d a northerl y source , entere d th e Penguin half-grabe n axially , and defin e a n over all progradationa l o r foresteppin g stacking pat tern; th e olde r flow s initiall y infille d basina l lows, wherea s subsequen t flow s reache d pro gressively highe r u p o n th e hangingwal l an d fault-scarp slope s an d furthe r ou t i n th e basi n (Fig. 13e) . The younge r interva l characterize d b y hig h extension rates , tentativel y dated a s lates t Early Volgian-middle/?late Volgian , wa s dominate d by significan t displacemen t o n NE-S W trend ing faults , i n additio n t o reactivatio n o f th e N-S trendin g structura l grai n (Ravna s & Stee l 1997). Thi s interva l i s represente d b y a two fold, sandstone-clayston e lithosom e model. Th e basal fining-upward part i s related t o the sourceward retreat o f the Magnu s submarin e fan/sheet system, an d depositio n o f a backsteppin g turbi dite package . Increase d faultin g alon g th e NE-SW trendin g End-of-the-Worl d Faul t cu t off th e northerl y sedimen t source , resultin g i n temporal basina l sedimen t starvation . Renewed input o f coarse r silicilasti c material occurre d a s the half-grabe n updi p are a an d Pengui n Horst / Makrel Hors t becam e expose d durin g th e sub sequent Middl e Volgia n til t climax . Thi s uppe r syn-rotational successio n i s terminate d b y lat e Volgian-Ryazanian claystone s (Fig . 12) . Th e claystones, formin g a condense d section , wer e deposited durin g th e lat e syn-rif t an d earl y post-rift stage s (se e developmen t interprete d for th e Draupn e Formatio n i n th e Oseber g area. Fig . 10) . The tectoni c quiescenc e strat a diffe r con siderably. A deep-marin e sedimen t gravit y flow sandstone package , th e mai n bod y o f th e Mag nus sandston e member , separate s th e tw o syn-rotational packlages , whera s th e uppe r syn rotational packag e i s cappe d b y a condense d claystone (compar e Fig . 1 3 with Fig . 10) . Thi s reflect a close r distanc e t o a sandie r hinterlan d source wit h hig h sedimen t yiel d durin g th e for mer tectoni c quiescenc e stage . Changin g struc tural regim e durin g th e las t rif t phas e lef t th e

Penguin half-grabe n a s a starved , intrabasina l sub-basin, source d fro m smal l borderin g high s (Ravnas & Steel 1997) .

S tat fjord and S tat fjord North areas (Profile 3) The Statfjor d an d Statfjor d Nort h area s con stitute westward to northwestward tilte d N-S t o NE-SW trendin g half-graben s situate d betwee n the norther n Vikin g Grabe n an d th e Marul k Basin (Fig . 1) . The sub-basin s are delineate d by N-S an d NE-S W trendin g masterfault s which downthrow t o th e eas t an d southeast . Th e Statfjord an d Statfjor d Nort h sub-basin s con tain example s o f sand-pron e Kimmeridgian Volgian deeper-to-shallo w marin e syn-rif t infill , the Draupn e Formatio n an d it s intercalate d Munin Sandston e Member . Th e earl y syn-rif t Tarbert Formatio n ha s bee n discusse d i n som e detail b y Johannesse n e t al. (1995) . The mud prone Heathe r Formatio n i s to o thi n t o allo w any detaile d subdivision . Downdip o n th e faul t blocks , th e Kimmer idgian-Volgian syn-rif t strat a for m a threefol d sandstone-mudstone-sandstone successio n with basal sand y sedimen t gravit y flo w deposits , overlying offshore mudstones , and cappin g shallow marin e sandstones . I n th e media l reache s of th e hangingwal l dip-slope , th e successio n i s represented b y a heterolithi c interva l wit h interbedded offshor e mudstones an d sediment gravity flow sandstones , overlai n b y shallo w marin e sandstones. I n updi p reache s onl y th e shallo w marine sandstones are present, commonly resting upon o r showin g a n intra-sandston e unconformity, wherea s th e entir e interva l i s represented by a n unconformit y an d overlyin g reworke d sandstones toward s cresta l area s o f th e faul t blocks (Fig . 14) . The Kimmeridgian-Volgia n successio n ha s been discusse d t o som e exten t b y Gradija n & Wiik (1987 ) and Soll i (1995) . Solli (1995) related each sandston e depositiona l even t t o distinc t uplift phase s o f th e footwall s alon g th e half graben masterfaults , especiall y th e Inne r an d Outer Snorr e Faults . A n alternative explanation, which i s favoure d here , relate s th e basa l sedi ment gravity flow deposits to the initial and early uplift stages . Thi s earl y uplif t resulte d i n sedi ment instabilit y on th e tilted slope, gravitational collapse an d resedimentatio n o f sedimen t for merly stored alon g the fault block updip areas by ensuing sediment gravit y flows, leading to deposition o f downdi p o r basina l sedimen t gravit y flow lobe s (Fig . 15a,b) . Continue d extensional

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Fig. 15 . Schemati c structura l development an d infillin g durin g Volgian rifting i n th e Statfjord-Statfjor d Nort h area. Onl y th e uppe r par t o f th e faul t bloc k dip-slop e i s shown. Sub-basin s represen t depositiona l sink s on the fault-bloc k dip-slop e itself , e.g . a s o n th e Stattfjor d Nort h structure , or smalle r sub-basin s betwee n two separate fault blocks , e.g . betwee n th e Statfjor d faul t block an d th e Statfjor d Nort h structure .

NORTH SE A SYN-RIF T SEDIMENTAR Y ARCHITECTURE S faulting resulte d i n furthe r tiltin g o f th e faul t blocks an d continued uplif t o f their updip areas . The uplif t resulte d i n widenin g o f th e footwal l islands an d create d shallowin g acros s th e faul t block i n updi p reaches . Th e widenin g o f th e footwall island s increased th e faul t blocks ' sedi ment yiel d potential , wherea s th e progressiv e steepening ma y hav e increase d erosio n rates . These tw o processe s i n combinatio n wit h th e shoaling i n updip reaches promote d th e progra dation o f locally derived hangingwall fan-deltas and shoreline s (Fig . 15c) . Thi s progradatio n i s argued t o have continued wel l into th e late synrotational stage , resultin g i n depositio n o f th e shallow marin e sandston e cappin g o f a t leas t the updi p reache s o f th e half-grabens . Source ward retrea t o f th e hangingwal l shoreline s wa s due either to diminishing sediment supply as the footwall island s becam e progressivel y deepe r eroded, o r becaus e th e shoreline s ha d entere d deeper water s downdi p o n th e hangingwal l an d delivered thei r sedimen t directl y int o deepe r waters. In either case, the capping o f most of the fault bloc k updi p area s b y thi n Ryazania n offshore mudrock s sugges t tha t b y thi s tim e there wa s littl e subaeria l relie f an d tha t th e fault block s b y the n ha d bee n peneplane d an d drowned (Fig . 15d,e ; cf . Robert s e t al. 19936) . Accordingly, i t i s postulated tha t th e threefol d sandstone-mudstone-sandstone syn-rif t succes sion represent s on e singl e rif t phase . The most diagnosti c feature s o f th e syn rotational hangingwal l deposit s are illustrate d in th e wel l cross-section s i n Fig . 14 . Not e especially tha t th e lowe r par t o f th e successio n defines a finin g upwar d tren d fro m deeper marine sedimen t gravit y flo w deposit s t o off shore mudstones, which at least in updip reaches represents a perio d o f upwar d shallowing . Seismic resolutio n i s poor, ye t ther e appear s t o be rapi d latera l an d basina l pinch-ou t o f th e turbidite lobe s (Ghaz i & Schmidt 1995) , as als o suggested b y wel l data . Moreover , thes e earl y turbidite lobe s wer e apparentl y ponde d i n structurally define d bathymetri c lows . Th e coarsening upwar d par t o f th e successio n t o the sandston e cappin g ma y recor d a variet y of relative sea-leve l change s dependen t o n th e

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structural positio n withi n th e half-graben . Downdip o f the fault block fulcru m the coarsening upwar d par t o f th e successio n reflect s a normal regressio n (sens u Posamentie r e t al . 1992). Th e mos t basina l o r downdi p well s suggest tha t ther e wer e at leas t tw o punctuate d progradational events . Althoug h th e floodin g across th e Muni n shorelin e sandstone s ma y a t least partly reflec t eustati c sea-leve l fluctuations , in view of the tectonic setting they are tentatively correlated wit h distinc t faultin g events . Suc h faulting event s wil l hav e resulte d i n punctuate d tilting and gentle rejuvenation of the half-graben relief. Renewe d progradatio n o f th e Muni n hangingwall shoreline would then have occurred as a respons e althoug h slightl y delaye d relativ e to th e postulate d faulting . Updi p o f th e faul t block fulcrum , th e coarsenin g upwar d par t o f the successio n reflect s a force d regression , an d no evidence of the punctuated shorelin e progra dations i s evident. Furthe r updip , transgressiv e shoreline sandstone s o r ravinemen t lag s res t upon a deepl y truncate d substrate . I n cresta l settings i t i s onl y th e lat e syn-rotationa l an d subsequent post-rif t transgressiv e stag e tha t i s recorded b y deposits ; mos t o f th e rotationa l stage i s represented b y a n unconformity .

Snorre half-graben (Profile 3) The westwar d tilte d Snorr e half-grabe n (Soll i 1995) is situated at the southern extension of the southerly plungin g Snorre Faul t Block, between the Statfjor d Eas t an d Visun d Faul t Block s (Fig. 2c) . Rotation along these major fault zones caused a rotation o f the Snorre- H pre-rift strat a to abou t 9 degree s b y Ryazania n time s (N0U vedt et al. this volume). To the north, th e Snorre and Statfjor d Faul t Block s merg e t o for m a complex hors t structur e tha t ha s suffere d c. 1.5km of composit e uplift . The Snorr e hal f graben i s delineate d an d dissecte d b y N- S and NE-S W trendin g faults , commonl y wit h an E-SE dip, an d contains a Bathonian t o Ryazanian shallow-to-deepe r marin e syn-rif t in fill. The Heathe r Formatio n (Bathonian-earl y

Fig. 16. E— W well correlation of the Viking Group and Tarbert Formatio n fro m th e Statfjord Faul t Block acros s the Snorr e an d Visun d Fault Block s to th e terraces southeas t o f the Visund Fault. Several syn-rift wedge s are recognized i n downdip basina l sink s and correlate s wit h unconformity strands acroos footwall highs. The syn-rotational wedges are separate d b y tectonic quiescenc e packages . Sy n rotational strat a ar e sand y in the Tarbert Formation , mudpron e in the Heathe r Formation , an d become s more sandy in the Draupn e Formation. Syn-rotationa l shallow-marin e sandstones ar e present i n the Tarbert Formation . Syn-rotationa l shallow-marine sandstones are also present updip on th e hanging-wall of individual fault block s in the Draupn e Formation wedges , whereas deep-marin e gravity flow sandstones occup y th e halfgraben depositiona l sinks .

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Oxfordian) consist s of silt y mudstone s wit h few sandstone interbeds . Th e Draupn e Formatio n (late Oxfordian-Ryazanian ) include s a lower , dark an d pyriti c clayston e uni t coarsenin g up wards int o a micaceou s sandston e uni t i n th e upper par t i n medial hangingwall reaches an d in the norther n axia l area . N o well s hav e bee n drilled of f the Statfjor d Eas t footwall , but fro m seismic interpretation an d structura l reconstruc tion i t i s suggeste d tha t th e lowe r Draupn e Formation ma y contai n sedimen t gravit y flow sandstones passin g upward s int o mor e fine grained siltstone s an d claystone s i n th e uppe r Draupne Formation . The mos t diagnosti c feature s of th e hanging wall syn-rif t infil l o f th e Snorr e Faul t Bloc k ar e illustrated i n th e wel l cross-sectio n pane l i n Fig. 16 . An earl y Bathonia n til t initiatio n ha s been advocate d fo r th e neighbourin g Statfjor d and Visun d Fault Block s (Robert s e t al. 19936 ; Fasrseth e t al . 1995#) , wherea s a lat e Bajocia n initiation o f tiltin g ha s bee n suggeste d fo r th e Gullfaks are a t o th e sout h (Johannesse n e t al. 1995). Thi s suggest s tha t th e sandstone s o f th e Tarbert Formatio n ma y be part of the early synrift infil l an d tha t initia l weak extensio n i n th e late Bajocian-Bathonia n cause d som e thicken ing o f th e Tarber t Formatio n acros s th e majo r fault blocks . Accepting th e interpretatio n tha t th e lat e Bajocian-Bathonian Tarber t Formatio n repre sents th e earlies t rotationa l movement s o f th e fault blocks , the syn-rif t infil l a s a whole forms a threefold sandstone-mudstone-sandston e pack age i n media l hangingwal l an d norther n axia l reaches o f the-plungin g Snorr e Faul t Block . Locally, o n th e hangingwal l dip-slope , off shore marine mudstones of the Draupne Forma tion envelo p sedimen t gravit y flow sandstones. On a lithological basis, the data suppor t a simple model o f extensio n wit h a prolonge d perio d o f moderate rotationa l rate s extendin g fro m th e Callovian-Volgian. Th e mode l relate s depo sition of the upper Draupne san d unit to reduced rates o f rotationa l faultin g durin g th e wanin g rift stage . A n unconformabl e relationshi p be tween th e Heathe r an d Draupn e Formation s (see below) , o n th e othe r hand , ma y favou r a n interpretation o f a t leas t tw o rift-maximas , one i n th e Callovian-earl y Oxfordia n an d on e in th e lat e Kimmeridgian-Volgian , separate d by a perio d o f tectoni c quiescenc e i n th e lat e Oxfordian-early Kimmeridgian . T o th e south , in th e mor e rapidl y subsidin g part s o f th e half-graben, a twofol d lithosom e model , a sandstone-mudstone succession , i s developed . Sediment gravity flow sandstones are assumed t o have bee n she d of f th e Statfjor d Eas t Faul t

Block i n respons e t o earl y uplift , correspond ing t o erosio n o f th e Heathe r Formatio n an d deposition o f claystone s i n hangingwal l areas . Towards th e cresta l area s o f th e Snorr e Faul t Block, th e entir e interva l i s represente d b y a n unconformity. Th e unconformit y cut s int o th e Brent Group in the south and deep into the lower Jurassic an d uppe r Triassic , i.e . th e Statfjor d Formation, i n th e north . O n th e neighbourin g Statfjord Eas t Faul t Block , th e uppermos t Draupne Formatio n i s restin g unconformabl y on eroded middl e Jurassic Bren t Group strata . Slightly increased rates of extension in the late Bathonian-Callovian le d t o drownin g o f th e Brent Grou p an d depositio n o f th e Heathe r Formation acros s a gently tilted, but submerged fault bloc k topograph y (Fig . 17a) . At thi s point, mudstones wer e widel y deposited, an d wit h th e exception o f th e norther n Tampe n Spur , sedi ments wer e source d fro m th e rif t margina l shoreline t o th e wes t an d fro m th e retreatin g Brent-delta t o th e south . Accelerate d rate s o f rotation i n th e lat e Oxfordia n le d t o exposur e of footwall crest s and segmentatio n of the Tam pen Spu r int o distinc t half-grabe n sub-basin s (Fig. 17b) . The relie f across th e emergin g footwall highs was stil l to o smal l to caus e significan t local erosio n an d redeposition , bu t becaus e o f the segmentation , the axia l an d transvers e sediment supply no longer fed directly into the outer half-grabens, leading to starved basin conditions. The Heather Formatio n appear s t o be unconformably overlai n and slightl y truncate d by th e overlying Draupn e Formatio n i n upflan k posi tions o n th e Snorr e hangingwall , as a resul t o f renewed rotatio n and footwal l uplif t i n the early Volgian (Fig. 17c) . This uplift increase d bot h th e extent an d relie f o f th e emergen t footwal l islands, an d thereb y their sedimen t yield poten tial. I n updi p hangingwal l areas, thi s le d t o a forced regressio n an d significan t loca l erosio n of th e Heathe r Formation . I n th e Statfjor d East footwall , th e Bren t Group wa s extensivel y eroded. Because of the steep gradient of the foot wall slope, sediments were unable to b e stored in significant quantitie s along th e footwal l shore line, an d mixe d san d an d mu d lithologie s were therefore ponde d a s gravitationa l flow and tur bidite deposit s i n structurall y define d bathy metric lows. Maximum uplif t o f th e footwal l islands were obtained i n th e lat e Volgian-Ryazania n (Fig . 17d), resultin g in dee p erosio n int o th e Bren t Group o n th e Snorr e hangingwal l dip-slope . Most o f th e erode d san d wer e store d alon g th e hangingwall shorelines , bu t sedimen t instability and gravitationa l collaps e cause d som e o f th e sand to be resedimented by ensuing gravity flow s

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Fig. 17 . Schemati c tectonostratigraphica l developmen t o f the Statfjor d (left) , Snorr e (central) and Visun d (right) Fault Block s during th e middl e Jurassic-earliest Cretaceous , i.e . throughout th e entir e rif t episod e an d th e immediate post-rift stage. At th e larger scale the evolution can be viewed as representing the early syn-rotational (a) through initial rotation climax (b) and late rotation climax (c), to the late syn-rotational (d) and early post-rift/ tectonic quiescenc e stage s (e) . However, th e longer-term evolutio n i s likely to have been interrupte d b y a series of shorter-term rif t phases , e.g . (a) trough (c) , characterized b y higher tectonic activit y and increasin g rates of faulting (se e text fo r furthe r discussion) .

down th e hangingwal l dip-slope . Progradatio n Visund half-grabens (Profile 3 ) and accumulation o f the upper Draupne shallo w marine sand s acros s th e Snorr e H-area , i s Th e Visun d faul t bloc k constitute s th e eastern interpreted a s a resul t o f wanin g faul t activit y mos t structur e i n a serie s o f N- S oriented , and decreasin g rate s o f rotatio n (Fig . 17d,e) . westerl y tilted fault block s betwee n th e norther n

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Viking Grabe n an d th e Eas t Shetlan d Platfor m in th e northernmos t Nort h Sea . Th e bloc k i s 23-26 km wide , i s bounde d i n th e wes t b y th e Inner Snorr e Faul t an d i n th e eas t (agains t the Vikin g Graben proper ) b y th e Visund Faul t (Fig.l). Extensio n alon g thes e majo r basement involved fault s cause d th e pre-rif t surfac e o f the Visun d half-grabe n t o rotat e b y abou t te n degrees durin g th e middle-lat e Jurassi c syn-rif t interval. Thi s i s steeper tha n norma l i n th e are a (Faerseth e l al. 19956 ) Syn-rift tiltin g o f th e Visun d faul t bloc k occurred at a lower rate during the first c. 10 Ma interval (earl y Bathonian-middl e Oxfordian ) and mor e intensel y fo r anothe r 1 7 Ma (mid dle Oxfordian-earl y Ryazanian) , causin g bot h Heather an d Draupn e Formation s t o hav e clear wedge-shaped geometrie s (Fig . 17) . The syn-rif t infill i s more than 1. 5 km thic k (decompacted ) i n the deepest par t o f the half-graben, but i s absent over th e crestal areas . The pre-rif t surfac e i s estimate d t o hav e undergone a 2 " til t durin g depositio n o f th e Bathonian-middle Oxfordia n Heathe r Forma tion (Faerset h et al . 1995b) , durin g whic h tim e calcareous siltstone s an d mudstone s forme d a n early syn-rif t wedg e o f significan t dimension s (Fig. 16) . Ther e appear s t o hav e bee n littl e sand depositio n a t thi s early stag e an d i t i s un likely tha t footwal l uplif t create d an y substan tial subaeria l landscape . Althoug h mos t o f th e present-day tilt on the fault bloc k occurred later , this early syn-rift rotatio n wa s sufficient t o create a modest syn-rif t unconformit y across the crestal areas o f th e block . The middl e Oxfordian-Ryazania n Draupn e Formation, als o develope d a s a clea r wedge shaped uni t withi n th e Visun d half-grabe n (Fig. 16) , accumulate d durin g accelerate d rota tion o f th e faul t block , resultan t fro m a serie s of rif t phase s fro m middl e Oxfordia n (fro m a 2 tilt ) t o lat e Volgian-Ryazania n (t o a 10 ° tilt; Faerseth e t al . 19956) . Structura l complication s on th e bloc k occurre d b y middle Kimmeridgia n times whe n a c . 6° til t ha d bee n attained . A sizable part of the crestal area collapsed at this time, formin g th e smaller , tilte d faul t bloc k which now occurs o n the SE-flan k of the Visund structure (Fig s 17a,b) . Th e present-da y cresta l area therefor e develope d a s a n emergen t foot wall islan d durin g th e fina l 4 ° o f til t o n th e structure (Faerset h et al . 19956) . Both th e mai n Visun d half-grabe n an d th e smaller flankin g half-grabe n were infilled mainl y from middl e Kimmeridgia n times . Th e mai n sub-basin wa s sediment-starved an d filled largely by claystones , thoug h ther e wer e a t leas t tw o episodes o f sand y turbidit e infill . Th e smalle r

half-graben receive d muc h coarse r sedimen t from accelerate d (Volgian ) footwal l uplif t an d drainage fro m th e ne w cresta l area . A wedge shaped uni t wit h som e 150 m o f slumpe d material, an d conglomerati c debri s flow s an d turbiditic sandstones has been documente d here, at a distanc e o f som e 2 km ou t fro m th e mai n fault scar p (wel l 34/8-7 , Fig. 16) . As wit h othe r syn-rift infills , upward-fining (sometimes upward coarsening-to-fining) signatures , o n a scal e o f 20-30 m ar e commo n i n thi s conglomerati c succession, an d ar e possibl y symptomati c o f sediment suppl y pulses , o r o f sedimen t suppl y passing certai n threshol d condition s i n th e drainage basin . The overal l successio n i n th e smalle r half graben, a s i n th e mai n half-grabe n succession , shows a genera l finin g upwar d trend . Earl y sandy o r gravell y mass flo w depositio n evolved through tim e t o giv e a mudstone-pron e succes sion, probabl y symptomati c o f underfillin g and sediment starvatio n i n half-graben s a t larg e distance fro m th e rif t hinterlan d region s (Ravnas & Stee l 1998) .

Lomre Terrace—Uer Terrace-Horda Platform (Profile 3) This section describes deeper marin e successions deposited durin g the late Oxfordian-Volgian rif t on the Uer and Lomr e terrace s alon g the easter n flank o f th e norther n Vikin g Grabe n (Fig . 1) . The Ue r an d Lomr e Terrace s for m tw o wide , segmented, gentl y eastward an d westwar d tilte d terraces, respectively . These structure s ar e situ ated betwee n the norther n Viking Graben, Sog n Graben an d th e Hord a Platfor m (Fig . 2c) . Th e two terrace s ca n b e considere d a s formin g a large-scale rela y zon e o r transfe r zon e betwee n the Hord a Platfor m an d th e dee p Sog n Grabe n to th e nort h (Fig . 1) . Two syn-rif t wedge s ar e recognized an d thes e correspon d t o a lat e Oxfordian-Kimmeridgian an d a Volgia n rif t phase, respectivel y (correspondin g t o th e lowe r wedge betwee n timelin e 18/19-2 2 and th e over lying wedge , respectively , in Fig . 18) . In basina l area s o f th e Lomr e an d Ue r Terraces, severa l kilometre s from the sub-basin' s boundary faults , th e syn-rotationa l successio n of th e lowe r wedg e form s a finin g upwar d suc cession o f mudstones and intercalated , generally thin (commonl y les s tha n fe w ten s o f metres ) sediment gravit y flo w sandstone s (e. g 35/ 8 wells, Fig . 18) . Th e basa l par t o f th e wedg e i s generally sand-prone , an d consist s o f a fining upwards successio n o f stacked , smaller-scal e

Fig. 18 . NW-S E well correlation o f the upper part s of the Viking Group (middle and uppe r Heathe r members , an d th e Fensfjord, Sognefjor d an d Draupn e Formations ) from th e Sog n Graben/Lomr e Terrac e acros s th e Ue r Terrac e t o th e Hord a Platform . Tw o syn-rotationa l wedge s ar e recognize d i n th e uppe r Heathe r member , an d the Sognefjor d an d Draupn e Formations . Th e lowe r wedg e i s represented b y a n unconformit y i n th e footwall , i.e. updip o n th e Hord a Platform ; forestepping-to backstepping shallow-marin e sandstones o n th e platform-boundin g terraces ; an d backsteppin g o r a fining-upwar d offshore/deeper-marin e sedimen t gravit y flow sandstone-hemipelagic mudstone package i n basinal positions , i.e. to th e northwest. Th e uppe r wedg e is represented b y an unconformit y i n footwall positions an d o n th e terrace area; an d a forestepping-to-aggradational-to-backstepping hemipelagite packag e i n the sub-basins to northwest. Intervening and cappin g tectoni c quiescence/earl y post-rift strat a consis t o f condensed claystones , whic h locall y in th e basina l sink s envelop sedimen t gravit y flow sandstones (e.g . wel l 35/8-2) .

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Fig. 19 . Schemati c structura l developmen t an d infillin g durin g Lat e Oxfordian-Volgia n riftin g alon g th e northwestern edg e o f the Hord a Platfor m an d Lomr e Terrace . The faul t bloc k t o th e righ t represent s the updip edge o f th e Hord a Platform , wherea s th e next , small fault bloc k t o th e lef t represen t th e terrace s boundin g the Horda Platfor m t o the northwest. The central horst an d grabe n topograph y represent s the larger Uer and Lomr e Terraces, wherea s the deepe r basi n t o th e lef t represent s th e norther n Vikin g Grabe n o r Sog n Graben .

NORTH SE A SYN-RIF T SEDIMENTAR Y ARCHITECTURE S fining-upwards sedimen t gravit y flow-hemipelagite packages . Thes e ar e interprete d t o hav e been fed by a shallow marine system temporaril y stationed a t th e seawar d margi n o f th e Hord a Platform (Fig . 19a , b). Th e stationar y positio n of th e shallo w marin e feede r syste m wa s du e t o increasing subsidenc e rate s alon g th e Hord a Platform's wester n boundar y fault , which , i n turn, positione d th e palaeo-shelfbreak alon g th e surface trac e o f this deep-seated basemen t fault . In hangingwal l position s clos e t o th e sub basin-bounding masterfault s (e.g. well N35/11-5, Fig. 18) , a coarsening-to-fining upwar d wedg e is present overlyin g th e basa l sedimen t gravit y flow-hemipelagite unit. Thi s reflects th e advanc e and retrea t o f a coars e submarin e apron , an d consists o f interbedde d sandstone , sometime s several ten s o f metre s thick , an d finer-graine d mud- an d siltstones . Thes e wedge s wer e pre dominantly derive d b y erosio n o f th e adjacen t footwall, thoug h temporaril y an d locall y th e overall retreating rift marginal Sognefjord shore line ma y hav e delivere d sedimen t directl y to th e deeper marin e wedge s (Fig . 19c , d). Th e latte r may b e particularl y commo n o n th e smaller , downfaulted faul t blocks or terraces immediately basinward o f th e Hord a Platform . I n foot walls o f th e majo r basin-boundin g fault s th e syn-rotational stag e i s represente d b y a foot wall unconformity and overlying , thin transgressive deposit s o r ravinemen t lag s onl y (e.g . well N31/2-15, Fig . 18) . The uppe r wedg e i s fine-graine d throughout , although wit h a distinc t coarsening-to-finin g upward signatur e (e.g . wel l N35/11-5 , Fig . 18) . This i s relate d t o th e renewe d influ x fro m adjacent high s durin g th e rotationa l til t climax. The syn-rotationa l stag e i s represente d b y a n unconformity acros s th e most pronounce d foot wall highs, similar to that observe d i n the underlying wedge . Thi s unconformit y may , i n som e places, hav e forme d b y submarine wav e erosio n and/or b e related t o gravitationa l collapse along the fault-scarps . Across smalle r and les s distinct highs the syn-rotational stag e is represented b y a condensed section . This suggests that these highs were submerge d throughou t th e rotationa l til t stage. Thei r starve d syn-rotationa l sedimentar y signature indicate s tha t coarse r sedimen t deliv ered b y sedimen t gravit y flows were ponde d o r bypassed i n th e adjacen t lows . As i n othe r deepe r marin e sub-basins , th e intervening tectoni c quiescenc e stage s ar e repre sented b y clay-pron e deposit s wit h a hig h source-rock potentia l whic h blanke t th e entir e sub-basins (Fig . 19e) . The claystone blankets ar e commonly thi n acros s th e intrabasina l highs , whereas the y thicke n toward s th e intrabasina l

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structural lows . Thi s i s probabl y du e t o th e addition o f som e fine-graine d turbidite s i n suc h settings. Thicker , coarse , tectoni c quiescenc e deposits ar e observe d onl y locall y (e.g . wel l N35/8-2, Fig . 18) . In th e western , updi p reache s o f th e Hord a Platform, Cretaceou s an d locall y Tertiar y deposits res t unconformabl y upo n truncate d Late Jurassi c strata . Th e updi p area s therefor e appear t o hav e bee n subaeriall y exposed durin g the lat e syn-rif t stag e (possibl y fro m th e Lat e Oxfordian o r Kimmeridgian ) an d wel l int o th e post-rift stage . The limite d erosio n o f the updi p reaches suggest s tha t ther e wa s limite d relie f along th e fault-scarp .

Northern North Seas rif t basi n fills: discussion an d conclusions Controls on syn-rift sedimentation Cyclic developmen t o f syn-rif t deposit s ma y result fro m variation s i n either the rate o f extensional faulting , sedimen t suppl y and/o r eustati c fluctations. I n th e basina l situation s discusse d above, varyin g sedimen t suppl y superimpose d on a varyin g rate o f rotationa l faultin g i s advo cated a s the majo r contro l o n th e resultan t syn rift sedimentar y architecture . Syn-depositiona l temporal variation s i n th e rat e o f rotationa l faulting i s evidenced by a combination o f two o r more o f the followin g factors: (1 ) wedge-shape d stratal packages showing internal stratal fanning alternating wit h package s wit h interna l tabula r geometries; (2 ) differentia l growt h o f time equivalent strata l package s and/o r evidenc e o f differential chang e i n palaeobathymetr y acros s extensional faults ; (3) change i n drainag e devel opment an d sedimen t suppl y direction s fro m a predominantly regiona l rif t margina l drainag e or starve d basina l condition s t o loca l fault block drainag e system s an d sedimen t supply ; (4) evidenc e o f downdi p drowning/deepenin g concomittant wit h updi p shallowing/emergenc e in individua l half-grabens ; (5 ) chang e i n basi n physiography fro m a large r ope n basi n wit h similar o r gradua l change s i n depositiona l environments acros s th e entir e rift-basi n t o larg e variations i n basina l condition s betwee n individual sub-basins . I n additon , th e rat e o f exten sional faultin g appea r t o hav e varie d spatially . Spatial variation s i n extensio n rat e resulte d i n variations i n th e onse t o f syn-rif t condition s across the basin, as well as in variations in structural topograph y an d accommodatio n creatio n between individua l sub-basins . Als o th e back ground, basinwid e subsidence/uplif t rat e ma y

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have varie d temporaril y an d spatially , thoug h this i s mor e difficul t t o addres s an d quantify . The background , basinwid e uplift/subsidenc e nevertheless wer e likel y highly influential on th e resultant tectonoeustati c evolutio n o f th e rif t basin (e.g . see Surlyk et al. 1993), and need s to be taken int o accoun t whe n discussin g basinwid e tectonic evolution , depositiona l histor y and/o r correlation o f events an d sedimentar y strata . Sediment suppl y t o individua l sub-basin s i s likely to hav e changed a s a response t o changing structural style s an d th e evolvin g structurall y formed topography/bathymetr y (e.g . see Ravnas & Stee l 1998) . Also th e regiona l or rif t margina l drainage an d sedimen t suppl y i s likel y t o hav e changed durin g th e rif t cycl e a s a respons e t o hinterlands tectonics . Change s i n rif t margina l drainage were , however , mos t likel y mor e instrumental o n th e sedimen t suppl y t o th e rif t marginal sub-basins , i.e . on th e Hord a Platfor m and Mal0 y Faul t Blocks , an d t o th e rif t mar ginal sub-basin s o n th e opposin g U K margin . With respec t t o th e middle-lat e Jurassi c rif t episode, th e syn-rif t infil l wer e derive d b y ero sion o f poorly consolidate d o r semi-consolidate d Triassic an d early-middl e Jurassi c sedimentar y successions. Drainag e developmen t an d sedi mentary respons e t o tiltin g an d uplif t event s are therefor e likel y t o hav e bee n fairl y rapid , at leas t fo r th e Nort h Se a rif t interio r sub basins. Wit h respec t t o th e rif t margina l deposi tional systems , thes e probabl y draine d a base ment terrai n wit h overlying Paleozoic-Mesozoic strata. Hence , th e sedimentar y respons e o f rif t marginal drainag e system s t o hinterland s tec tonics may als o hav e bee n fairl y rapid . Potentia l lags i n respons e tim e du e t o erosiona l threshol d needing t o b e exceeded , ma y hav e forme d i n areas wer e the hinterlan d consisted o f crystalline basement terrain s o r wa s underlai n b y a meta morphosed substrate . The mai n contro l o n th e sedimentar y archi tecture o f rif t basi n infil l wa s thu s sedimen t supply i n relatio n t o rat e an d styl e o f accom modation creation . Creatio n o f accommodation space wa s controlle d b y eustasy , technicall y induced subsidence/uplift , pre-rif t waterdepth / elevation, basi n geometr y an d basinfloo r til t rate. Durin g th e middle-late Jurassic , which wa s a non-glacia l period , th e rat e an d magnitud e o f eustatic sea-leve l fluctuatio n ar e likel y t o hav e been suppresse d b y technicall y induce d base and sea-leve l variation s (Stee l 1993 ; Ravna s & Stee l 1998) . Sedimen t flu x wa s controlle d by a wide range o f interrelated factors , th e most important bein g climate , tectonism , hinterlan d physiography, hinterlan d lithologies , hinter land an d basina l drainag e patterns , depositiona l

patterns, especiall y whethe r poin t o r linea r sourced, basina l sedimen t dispersa l patterns , and relativ e sea-leve l fluctuations . I n rif t basin s consisting o f a serie s o f sub-basin s suc h a s th e northernmost Nort h Sea , extensiona l faultin g causes considerabl e topograph y withi n wha t used t o b e a single , broader, low-relie f basin, o r enhances th e previou s topography . Sediment s delivered fro m rif t margina l source s wa s thu s trapped i n sub-basins adjacent t o th e hinterland. These sub-basin s becam e sediment-overfille d o r sediment-balanced (Nottved t et al. 1995 ; Ravnas & Stee l 1998) . Sub-basins a t increasin g distance from th e hinterland s relied progressively on sup ply fro m rif t interio r sources . A s thes e source s commonly wer e smal l compared t o th e adjacen t half-graben sinks , sub-basin s a t increasin g distance fro m th e hinterland s receive d progres sively les s sediment , an d thu s becam e eithe r sediment-underfilled o r sediment-starved . Th e detailed syn-rif t sedimentar y architectur e o f these fou r basi n infil l types , i.e . th e sediment overfilled, sediment-balanced , sediment-under filled an d sediment-starved , i s discusse d b y Ravnas & Stee l (1998). Milton (1993) demonstrated how the dip of the depositional foundatio n influenc e th e resultan t aspect ratio , i.e . th e di p length : average thick ness, o f th e syn-rif t sedimentar y geometries . Narrower facie s belt s with more rapi d transtion s between individua l facie s tract s shoul d b e expected o n steepe r dippin g slope s compare d t o gentlier dippin g slope s (Ravna s & Stee l 1998) . Such change s i n depositiona l geometrie s ar e generally observe d whe n comparin g syn-rif t infill relate d t o earlie r phase s o f th e middle-lat e Jurassic rif t episode s t o thos e relate d t o late r phases. Syn-rif t infil l relate d t o late r phase s generally hav e narrowe r an d sometime s thicke r facies belt , mainl y du e t o th e stee p dip s o f the. depositiona l foundatio n (u p t o 10" ) du e to th e cummulativ e til t resultin g from recurren t rift phases . Climate als o exerte d a majo r contro l o n th e resultant rif t basi n infill , a s demonstrate d b y a change fro m ari d t o semiari d condition s i n th e late Permian-earl y Triassic , fro m a semi-ari d t o a mor e humi d environmen t i n th e lat e Triassic early Jurassic , an d a possibl e retur n t o a drye r climate in the Kimmeridgian-Volgian (e. g Clemmensen e t al . 1980 ; R0 e & Stee l 1985 ; Wignall & Ruffel l 1990 ; Wignal l & Pickerin g 1993) . Although suc h climati c change s ma y b e dia chronous acros s a larg e area , an d significan t in large , elongat e rift s (suc h a s th e Eas t Africa n Rift-Red Sea-Gul f of Suez, and the Rio Grande Rift), the y ar e likel y t o hav e occure d relativel y rapidly an d cause d les s lateral variation s within

NORTH SE A SYN-RIF T SEDIMENTAR Y ARCHITECTURES a restricte d are a suc h a s th e norther n Vikin g Graben an d adjacen t platfor m areas . Hierarchy of rift basin successions

Rift basi n succession s commonl y contai n in fill segment s relate d t o th e variou s stage s o f active riftin g an d ar e envelope d b y sediment s deposited prio r to and subsequent t o the rift epi sode proper , leadin g to a subdivisio n into pre- , proto-, syn - an d post-rif t strat a (e.g . Nottved t et at. 1995) . I n case s wher e th e rif t basi n wa s subjected t o repeate d stretchin g episodes , th e sediments deposite d durin g th e intervenin g periods wit h relativ e tectoni c quiescence , e.g . th e middle Triassic-middle Jurassic o f the Norther n North Sea , diffe r i n man y respec t fro m typica l post- o r pre-rif t strata . W e sugges t therefor e that suc h succession s shoul d b e referre d t o a s inter-rift. As demonstrate d b y th e middle-lat e Jurassi c syn-rift succession s o f th e norther n Nort h Sea , rift episode s ca n consis t o f higher-frequenc y tectonic event s o r rif t phases . I n suc h case s th e syn-rift successio n consist s o f a serie s o f syn rotational an d intervenin g tectoni c quiescenc e stratal packages . Whethe r suc h rif t phase s reflect tru e temporal variation s i n stretching rat e recorded i n basinwid e near-synchronou s o r slightly diachronou s riftin g acmes , o r ca n b e attributed t o initiation , propagation, migration , lateral shif t o r cessatio n o f faultin g alon g th e sub-basin-bounding masterfaults , i s ye t no t clear. A t present , i t i s also difficul t t o sugges t whether a pulsator y or a single-phas e rift devel opment i s the mor e common . In an y case , th e presenc e o f repetitive rif t epi sodes an d th e potentia l o f repetitiv e rift phase s occuring within a single rift episode , suggest that the rif t basi n succession s ca n b e ordere d hierarchically (Fig . 20a , b). A t th e highes t ran k are th e syn-rif t strat a representin g th e mai n rif t episodes an d th e intervenin g inter-rif t strata , and possibl e pre - an d post-rif t strata . Thes e commonly represen t a tim e interva l o f a fe w tens o f millio n years, an d ar e give n a lowes t o r first order . Nex t ar e th e syn-rotationa l strat a and th e intervenin g tectoni c quiescenc e pack ages whic h relat e t o a singl e rif t phase . I n th e middle-upper Jurassi c o f th e norther n Nort h Sea, a combine d syn-rotationa l an d tectoni c quiescence uni t commonl y represent s som e 4-6 Ma. I n som e case s i t i s possibl e t o furthe r subdivide th e syn-rotationa l strat a int o higher order package s relate d t o distinc t rotationa l tilt events , e.g . a s establishe d fo r th e lat e Bajocian-early Bathonia n Tarber t Formatio n

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transgressive-regressive sequences . Thes e unit s commonly spa n som e 1- 2 Ma an d possibl y represent a cluste r o f faultin g events . Suc h rotational til t event s correlat e wit h discret e deformation spans . Theoretically , i t shoul d b e possible t o furthe r subdivid e syn-rotationa l successions int o eve n higher-orde r unit s repre senting discrete , short-live d faultin g event s o r deformation clines . A t present , thi s i s rathe r difficult, i f no t impossible , i n th e subsurfac e without goo d seismi c databas e and/o r abun dant wells. As wit h th e syn-rif t strata , th e intervenin g inter-rift strat a ca n sometime s b e subdivide d into packages which represents prolonged inter vals o f tectoni c quiescenc e separate d b y unit s which correlate with discrete rift stage s or minor rift phase s (Fig . 20c) . In suc h cases, syn-tectonic strata ar e expected to represent quit e short timeintervals compared t o th e time span represente d by th e entir e inter-rif t package . Althoug h th e inter-rift tectoni c stage s ma y b e insignifican t with respec t t o th e basinwid e stretching factor, they appea r t o exer t a majo r influenc e o n th e sedimentary architectur e o f the inter-rif t succession (Stee l 1993) . As such , thes e event s nee d t o be take n int o consideratio n i f a prope r under standing of the development and distributio n of potential inter-rif t reservoi r an d sourc e roc k intervals ar e t o b e achieved .

Sedimentary architecture of syn-rift packages related to the entire rift episode At a larg e scale , a rif t episod e ma y b e repre sented b y a threefold , sandstone-mudstone sandstone package , an d thi s i s typicall y developed i n half-grabens wher e there is an abundan t sediment supply , for example non-marine, lacus trine o r fluviall y dominate d rift s (Fig . 2la ; se e also Hamblin & Rust 1989 ; Lambiase 1990) , but has bee n advocated fo r marine rift basin s as well (Fig. 21b,c ; Prosse r 1993 ; Leinfelde r & Wilso n in press) . I n th e norther n Nort h Sea , marin e syn-rift succession s relate d t o a n entir e rif t episode, e.g . th e middle-uppe r Jurassic , appea r to b e dominate d b y twofold , fining-upward , sandstone-mudstone package s (Ravna s e t al. 1994; Nottved t e t al . 1995) . I n thi s rif t basi n individual sub-basin s see m t o b e completel y infilled eithe r wel l int o th e post-rif t perio d or subsequen t to continent-continent separation (in th e cas e o f abortiv e o r successfu l rifting , respectively). These large-scal e threefol d o r twofol d archi tectural lithosom e model s ar e i n som e case s

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Fig. 20 . Hierarch y of tectonic events during rif t basi n formatio n and development , and propose d nomenclature for rif t basi n sedimen t infill , (a ) Relate s to th e entir e rif t basi n evolutio n wherea s (b ) relate s t o a single, multiphase rif t episode , an d (c ) relate s t o th e inter-rif t period .

NORTH SE A SYN-RIF T SEDIMENTAR Y ARCHITECTURE S paralleled b y a vertica l stackin g o f repetitiv e syn-rotational and tectonic quiescenc e packages . In suc h cases, th e entire (or parts of the) syn-rif t succession may displa y a cycli c development , the cyclicit y reflectin g th e comple x interpla y between tempora l variatio n i n rate s o f exten sional faulting , sedimen t suppl y and , fo r marin e syn-rift successions , eustati c sea-leve l fluctua tions. S o far , syn-rif t cyclicit y du e t o tempora l variation i n th e rat e o f extensiona l faulting , whether basin-wid e o r pertainin g t o individua l half-grabens, ha s bee n demonstrate d onl y fo r marine, mixe d non-marine/marine , an d lacus trine syn-rif t succession s (e.g . Sellwoo d & Netherwood 1984 ; Purser e t al 1990 ; Lambiase 1990; Underbill 19910 , b\ Partington e t al 1993 ; Rattey & Hayward 1993 ; Faerset h e t a l 19950) . This documentation ma y be achieved i f there is a reference leve l suc h a s relativ e sea-leve l o r loca l lake level, which is sensitive to an d immediately responds t o temporally changing tectonics in the basin. Non-marine rif t basin s may similarly have a multiphas e rift-histor y (P . Theriault , pers . comm. 1995) , thoug h ofte n ther e i s a lac k o f stratigraphic evidenc e to suggest suc h a comple x basin development .

Sedimentary architecture of syn-rotational packages related to a single rift phase In th e middle-uppe r Jurassi c o f th e norther n North Sea , syn-rotationa l strat a relate d t o a single rif t phas e sho w a wid e rang e o f pos sible sedimentar y architecture s a s well . Th e subdivision o f rif t basin s int o a continuu m o f basin-types fro m sediment-overfille d throug h sediment-balanced an d sediment-underfille d t o sediment-starved, allow s a distinc t sequentia l development an d henc e specifi c sedimentar y architecture t o b e erecte d fo r eac h basi n type . The tectoni c significanc e o f stratigraphi c sur faces an d th e variabl e sedimentar y infil l trends , from th e earl y syn-rif t throug h th e rif t clima x and t o th e lat e syn-rif t stage , relate d t o a single rift phase fo r eac h o f these fou r rif t basi n types, ar e discusse d i n mor e detai l b y Ravna s & Stee l (1998) . Notably , lithosom e model s for marin e syn-rif t succession s relate d t o a single rift phase , whethe r shallow - o r deep-marine , may sho w threefol d sandstone-mudstone sandstone, two-fol d sandstone-mudstone , o r single mudstone) signatures . Th e main differenc e between th e threefol d an d twofol d marin e syn rift signature s relate s t o th e exten t t o whic h th e (changing) sedimen t suppl y is able t o kee p pac e with varyin g basinal subsidenc e rates .

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Shallow marine , threefol d syn-rif t signature s formed i n sub-basin s wit h lo w accommodatio n generation (lo w basinal subsidenc e rates) and/o r high sediment supply . I n the northern Nort h Sea rift basin , suc h infil l appears t o b e restricte d t o the earl y phas e o f th e middle-lat e Jurassi c stretching episode , an d t o rif t margina l sub basins durin g late r phases . Example s o f suc h infill type s ar e th e Bajocian-Bathonia n Tarber t Formation i n th e Vikin g Graben , an d th e Bathonian-Kimmeridigian Heather , Krossfjord , Fensfjord an d Sognefjor d Formation s o n th e Horda Platfor m (Fig s 4 & 11). Shallow marin e twofol d syn-rif t signature s formed i n sub-basin s adjacen t t o sourc e area s with limite d sediment yield potential. Suc h subbasins wer e commonly locate d a t som e distanc e from th e regiona l hinterland s o f th e rif t basin , or represen t th e transitio n fro m a mixe d non marine an d shallo w marin e t o a deepe r marin e rift basi n a s th e rif t evolve d an d subsided . Th e Bathonian-Kimmerdigian Heathe r Formatio n in th e Vikin g Grabe n prope r an d i n flankin g half-grabens suc h a s th e Oseber g Faul t Bloc k (Fig. 8) , are goo d example s o f thi s infil l type . Deeper marin e sub-basin s wit h a threefol d syn-rift stratigraph y ha d acces s t o larg e hinter land areas , eithe r th e rif t margina l hinterlan d or larg e rif t interio r sources . In suc h infil l types , shallow marin e lithosome s ma y o r ma y no t be presen t updi p o n large r faul t blocks . Thi s variant i s rathe r unusua l i n th e middle-uppe r Jurassic o f th e norther n Nort h Sea . Th e lat e Oxfordian-early Volgia n o f th e Pengui n half graben (Fig . 12 ) constitutes th e onl y presentl y recognized examples . Th e deeper-marine varian t of the threefol d lithosom e mode l contrast s wit h its shallo w marin e counterpar t b y containin g abundant sandstone s in the strata correspondin g to the rotational til t climax, and b y the commo n sharp basa l contac t o f it s cappin g o f forest epping turbidites . I n th e combine d shallow to-'deeper'-marine varian t exemplifie d b y th e Kimmeridgian-Volgian o f th e Stafjor d an d Snorre area s (Fig s 1 4 & 16 ) there i s a gradua l transition fro m th e offshor e mudrock s t o th e shallow marin e sandston e capping . Deeper marin e sub-basins with a twofold synrift signature , ow e thei r lithosom e characte r t o the limite d sedimen t yiel d potentia l o f th e adjacent sourc e areas . Thi s infil l typ e i s b y fa r the mos t commo n o f th e deepe r marin e vari ants, an d example s includ e th e Earl y Volgian Ryazanian i n the Pengui n half-grabe n (Fig . 12) , and th e Kimmeridgian-Ryazanian i n the Viking Graben (Fig . 6) , Oseber g Faul t Bloc k (Fig . 8) , Visund Faul t Bloc k (Fig . 16 ) and Sog n Grabe n (Fig. 18) .

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Fig. 21 . Rif t basi n infill models: (a) Non-marine rift basins ; (b) mixed non-marine-marine and shallo w marine rift basins; an d (c ) deep-marine rift basins .

Concluding remarks The syn-rif t infil l o f the northernmost Nort h Sea rift basi n var y fro m predominantl y non-marin e

(Fig. 2la) during the late Permian-early Triassi c rift episod e t o dominantl y marin e (Fig . 21b , c) during th e middle-lat e Jurassi c rif t episode . The middle Triassic-middle Jurassic non-marin e

NORTH SE A SYN-RIF T SEDIMENTAR Y ARCHITECTURES to marin e inter-rif t successio n sho w evidenc e that discret e rif t stage s o r mino r rif t phase s occurred intermittentl y als o throughou t th e inter-rift period . The Permian-early Triassic rif t basi n an d th e broader middle-lat e Triassi c non-marin e basi n probably coul d no t easil y access marin e waters . Data fro m th e Permian-earl y Triassi c syn-rif t infill ar e to o scan t t o sugges t an y evolutionary trends i n time . Probabl e subsidenc e maxim a i n the middle-uppe r Triassi c non-marin e inter-rif t succession correspon d wit h part s o f Teist , Middle Lund e an d uppermos t Lund e Forma tions lacustrin e interval s wit h som e brackis h incursions (Stee l 1993) . Thi s ha s resulte d i n a n infill relate d t o each inter-rif t rift stage/rif t phas e that sho w a threefol d architectura l signature s typical fo r sedimen t overfille d an d sedimen t balanced basins . Otherwis e ther e appear s to be no particula r evolutionar y trend s i n tim e asso ciated wit h th e syn-rif t infil l relate d t o middle late Triassi c tectoni c maxima . Potentia l sourc e rocks includ e lacustrin e claystone s an d coal bearing successions . The mixe d non-marin e an d marin e syn-rif t infill relate d t o early-middl e Jurassi c rif t stage s show a sedimentar y architectur e mimicin g th e infill relate d middle-lat e Jurassi c rif t phases , although wit h a subdued asymmetr y of the synrift infil l an d onl y weakl y develope d footwal l unconformities. The middle-lat e Jurassic marin e syn-rif t suc cessions o f th e norther n Nort h Se a rif t basi n reflect a n evolutionar y path fro m a mixe d nonmarine/marine t o a full y marin e environment . Characteristically th e middle-lat e Jurassi c rif t episode developed throug h a series of rift phase s separated b y interval s o f relativ e tectoni c quiescence. Furthermore , th e natur e o f the synrift infil l varie s considerably between the distinct rift phases . Th e transitio n fro m a threefol d mixed non-marin e an d shallo w marin e syn-rif t infill throug h a twofol d shallo w marin e syn rift infil l t o a twofol d deep-marin e syn-rif t infill , reflects th e evolutio n fro m a sediment-overfilled or sediment-balance d statu s t o a n increasingl y sediment-underfilled an d sediment-starve d sta tus. This evolution tend s to have been parallele d by a n increas e i n th e pre-rotationa l elevation / waterdepth an d th e distance t o the rift margina l hinterland areas , a s well as t o a decreas e i n th e rift interio r sedimen t yiel d potential . In th e lowe r par t o f th e middle-lat e Jurassi c syn-rift infill , bot h syn-rotationa l an d tectoni c quiescence sand-pron e strat a for m reservoi r intervals. A s the rif t evolve d an d subsided , shal low marin e syn-rif t reservoi r interval s becam e confined t o narro w belt s alon g th e rif t basin' s

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margin, i.e . t o th e eas t o n th e Hord a Platfor m and wes t in the East Shetlan d Basin . Within th e rift itself , significan t shallo w marine strat a wer e deposited onl y alon g large r footwal l islands . Deeper marin e reservoi r interval s ar e eithe r syn-rotational i n nature , typicall y developed a s fault-scarp o r rela y ram p apron s an d fans ; or sub-basinwide , sheet-lik e sedimen t gravit y flow packages, deposite d durin g tectoni c quiescence stages . Influ x o f syn-rif t sandstone s fro m rift interio r source s usuall y represen t distinc t tectonic phase s accompanie d b y pronounce d footwall uplif t o r majo r reorganizatio n o f th e tectonic regim e whic h lea d t o th e initiatio n o f new masterfault s (e.g . Faerset h e t al. 19950 ; Faerseth & Ravna s 1998) . Rif t interio r sand prone tectoni c quiescenc e package s shoul d b e expected only in sub-basins bordering large sediment source s (Fig . 21b,c ; Ravnas & Steel 1997) . Syn-rift sourc e rock s are more common i n the latest Jurassi c deepe r marin e successions . Although interval s wit h hig h sourc e roc k potential ar e presen t i n th e shallo w marin e t o shelfal Heather Formation , i t is the claystones of the Kimmeridge Cla y and Draupn e Formation s that ar e th e prim e sourc e roc k i n th e norther n North Se a rif t basin . Th e sourc e roc k potentia l varies considerabl y als o withi n thes e units : claystones deposite d durin g tectoni c quiescenc e intervals generall y appear t o hav e bette r sourc e rock potentia l tha n thei r syn-rotationa l equivalents (Ravnas & Steel 1997) . Secondly, clayston e drapes acros s intrabasinal highs and faul t bloc k updip area s als o seem s t o b e riche r tha n thei r downdip correlative . Thi s ma y b e attribute d t o an extrem e starvatio n o f siliciclasti c materia l across suc h areas , an d a highe r expor t rat e o f organic matte r t o th e botto m du e t o th e rela tively shallo w palaeowate r depths . Th e sourc e rock potentia l o f claystones deposited i n downdip areas may be diluted b y the higher frequency of intercalate d silt y an d sand y turbidites . Th e formation o f th e Lat e Jurassi c organi c ric h claystones i n th e norther n Nort h Se a wa s a t least partly du e t o th e formatio n o f tectonically silled, rhomb-shaped sub-basins , the basin shap e reflecting th e interplay betwee n the existing N- S and th e subsequentl y forme d NE-S W trendin g structural grain s (Faerseth & Ravna s 1998) . We woul d lik e t o than k ou r fello w colleague s an d research student s i n th e Joul e II : Integrate d Basi n Studies-Dynamics of the Norwegia n Margi n (IBS DNM) project , and colleagues at respective companies and researc h institution s fo r stimulatin g co-operation and discussions. Reviews by Arne Dalland, Bj0r n Tore Larsen, Snorre Olausse n an d Sara h Prosse r sharpene d the fina l produc t an d ar e appreciated . Th e draftin g office a t Statoil , Nors k Hydr o an d Norsk e Conoc o

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prepared th e illustrations . Th e wor k ha s bee n funde d by th e Researc h Counci l o f Norwa y a s par t o f th e Joule I I Researc h programm e (CE C contrac t No . JOU2-CT92-0110).

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Syn-rift evolutio n an d resulting pla y models in the Snorre-H area , northern North Se a ARVID N0TTVEDT, 1 ANKE R M . BERGE, 2 NANCY E H . DAWERS, 3 ROALD B . F^RSETH, 4 KJEL L O . HAGER, 2 GUNN MANGERUD 2 & CA I PUIGDEFABREGAS 5 1

Norsk Hydro Research Centre, N-5020 Bergen, Norway Present address: Norsk Hydro Canada, lll-5th Avenue SW, Calgary, AB T2P 3Y6, Canada 2 Norsk Hydro Research Centre, N-5020 Bergen, Norway ^Department of Geology and Geophysics, University of Edinburgh, Edinburgh EH9 3JW, UK 4 Norsk Hydro Exploration, 1321 Stabekk, Norway 5 Norsk Hydro Research Centre, N-5020 Bergen, Norway Present address: Centre for Advanced Studies, Cami de Santa Barbara, E-17300 Blanes, Girona, Spain Abstract: A n uppe r Jurassic , wedge-shaped syn-rif t succession , comprising the Heather an d Draupne Formations , i s present i n the hangingwal l troug h o f the Snorr e Faul t Block . Th e succession is bounded t o th e west by the Statfjor d Eas t Fault , wherea s it onlaps the Snorr e Fault Block t o the east. I t consists o f a two-fold coarsening-upward sequenc e fro m shal e to sandstone o f shallow marine/shorelin e origin . Active fault bloc k rotatio n and subsidenc e in th e Snorre-H are a commence d i n th e Mid Bathonian an d lasted throug h th e Ryazanian. Th e Heather Formatio n was deposited durin g the earl y rif t stag e (Mid-Bathonian-Earl y Oxfordian ; 3 ° cumulativ e tilt) , wherea s th e Draupne Formatio n (Lat e Oxfordian-Ryazanian ; 9 ° cumulative tilt) accumulate d durin g the main and lat e rif t stages . The lower part o f the Heather Formatio n wa s likely deposite d across a submerge d tilte d faul t bloc k terrain , wit h a predominan t extra-basina l sedimen t supply. Depositio n o f the uppe r Heathe r Formation , however , wa s governed b y graduall y emerging footwal l islands , albei t ye t withou t significan t loca l erosion . A s a resul t o f increased fault/bloc k rotatio n durin g depositio n o f th e Draupn e Formatio n Shale , Sequences I-I I (lat e Earl y Oxfordian-earl y Mid-Volgian ) footwal l island s becam e firml y established, providing a predominant local sediment source to the Snorre-H sub-basin . Clay and sil t were supplied fro m erosio n of the Heather Formation o n the Snorre-H hangingwall, with subordinat e inpu t o f sand fro m th e Statfjor d Eas t footwall . Subsequent depositio n o f the Draupne Formation , Sequence s III-V (late Mid-Volgian-Ryazanian), was governed by significant relie f on th e footwal l islands, causing dee p erosio n int o th e Bren t Grou p on th e Snorre-H hangingwal l dip-slop e an d leadin g t o progradatio n o f th e Uppe r Draupn e Sandstone shorelin e complex acros s th e Snorre- H area . Depositio n o f th e Draupn e For mation an d tempora l shorelin e position were likely partly controlled by northwards fault-ti p propagation o f the Statfjor d Eas t Fault . Various syn-rif t pla y model s an d depositiona l reservoi r facie s ar e presen t withi n th e Snorre-H hangingwal l basin. They includ e dip-slop e shallo w marine/shorelin e sands , basi n floor gravit y transporte d sand s an d likel y footwal l talu s sand s envelope d i n organi c ric h shales o f the Draupne Formation . Th e distribution of reservoir facie s i s intimately linked t o exposure an d erosio n o f the middle Jurassic Bren t Group below th e syn-rif t unconformity .

In recen t years , considerable effor t ha s bee n put an d sedimentation , leadin g t o localize d an d into developin g predictiv e model s fo r syn-rif t subtl e occurrenc e o f reservoir facies . The signa plays, recognizin g tha t such play s ma y hol d tur e o f syn-rif t play s depend s o n th e interpla y significant hydrocarbo n volumes . Syn-rif t play s betwee n creatio n o f accommodatio n space , stem from a complex interplay between tectonics climate , sedimen t supply , an d relativ e base- o r From: N0TTVEDT , A . e t al. (eds ) Dynamics o f th e Norwegian Margin. Geologica l Society , London , Specia l Publications, 167 , 179-218 . l-86239-056-8/00/ $ 15.00 © Th e Geologica l Societ y o f London 2000 .

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sea-level stan d an d thei r change s (Leede r & Gawthorpe 1987 ; Surly k 1989 ; Prosse r 1993 ; Gabrielsen e t al 1995 ; N0ttved t e l al. 1995 ; Ravnas & Steel 1997 , 1998 ; Ravna s et al . 2000) . In the North Sea, oil has been proven i n uppe r Jurassic syn-rif t sandston e reservoir s o f variou s kinds. The Pipe r (Mahe r 1981) , Troll (Ronnevi k & Johnsen 1984) , Fulma r (Johnso n e t al . 1986) , Claymore (Maher & Harker 1987) , Clyde (Smit h 1987) an d Ul a (Hom e 1987 ) discoverie s com prise margina l t o shallo w marin e sandston e

reservoirs partl y fringin g emergen t footwal l islands, wherea s submarin e fa n sandstone s deposited i n topographi c low s constitut e th e reservoir i n th e Magnu s (De'At h & Scuylema n 1971), Bra e (Turne r e t al . 1987 ) an d Mille r (McLure & Brown 1992 ) discoveries. The presen t contributio n describe s th e hang ingwall dip-slop e evolutio n o f th e Snorr e Faul t Block, th e so-calle d Snorre- H area , i n th e Norwegian norther n Nort h Se a Bloc k 34/ 7 (Fig. 1 ) (Dah l & Soll i 1993 ; Soll i 1995) . W e

Fig 1. (a ) Location map, showing the study area and arra y of westerly tilted faul t block s on the Tampen Spur to the west o f the norther n Viking Graben. (b) Geoseismic crustal sectio n acros s the norther n Vikin g Graben. See Fig. l a fo r location.

SYN-RIFT EVOLUTIO N I N TH E SNORRE- H ARE A attempt t o present a dynamic model fo r the synrift evolutio n and to outline the various reservoi r facies an d pla y type s associate d wit h extension and rotatio n o f th e Snorr e Faul t Bloc k an d deposition o f th e uppe r Jurassi c Draupn e For mation i n particular (Fig . 2) . The focus is on th e erosion an d redistributio n o f san d relate d t o footwall uplif t an d exposur e o f bot h th e Snorr e and neighbourin g Statfjord Eas t Faul t Blocks . In addition , w e sho w ho w a systemati c an d integrated approac h ma y significantl y increas e our understandin g of syn-rif t pla y potential an d predictive abilit y o f syn-rif t facie s an d reservoi r distribution. Th e stud y involve s detailed struc tural an d seismi c stratigraphi c interpretatio n o f multiple 3D-seismi c data , combine d wit h sedi -

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mentological an d biostratigraphica l i n situ an d reworking analyses in a number o f wells. Spatial and tempora l linkin g of processes an d mappin g of mass budget between local drainage areas and associated depositiona l basin s hav e bee n ke y elements i n the process .

Database The seismi c interpretatio n o f th e souther n par t of th e Snorr e Faul t Bloc k i s base d o n 3 D surveys E86 , SG9201 , SG843 1 an d GE8 3 (Fig. 3) . Regional information from merge d an d reprocessed surve y CTM9 4 i s als o used . Thi s survey combines E86 , SG920 1 an d SG843 1 with

Fig. 2. Jurassi c lithostratigraphy of the Tampen Spu r area, following th e nomenclature of Vollset & Dore (1984). Timescale i s after Gradstei n e t al. (1994).

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Fig. 3. Bas e Cretaceous structura l map o f the Snorre-Visund-Gullfaks area , showin g location of the 3 D seismic surveys E86, SG9201 and SG843 1 and well s used in this study. Note annotated well symbols, showing which part of th e syn-rif t successio n i s present an d i f the Bren t Grou p is truncated. SF , Statfjor d Fault ; SNF , Statfjord North Fault ; SEF, Statfjord Eas t Fault; OSF, Outer Snorre Fault; CSF, Central Snorre Fault; ISF, Inner Snorre Fault; TF , Tordi s Fault ; ST9101, ST851 1 and ST9207 . Th e quality of the surveys varie s an d th e qualit y withi n a singl e survey ca n var y over th e area . Th e seismi c data are assume d t o b e zero-phase, bu t difference s i n phase betwee n survey s mak e i t difficul t t o interpret th e sam e reflector s i n detai l ove r th e

whole area . I n th e area s adjoinin g th e structu rally comple x Gullfak s field the horizo n defini tion i s particularl y poor . Th e densit y o f interpretation o f th e uppe r Jurassic-lowe r Cre taceous syn-rif t sequenc e varie s betwee n ever y second an d ever y tent h inline . A detaile d

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Table 1 . Well database an d main us e o f wells Well

Main us e

Well

Main us e

33/09-03 33/09-04 33/09-05 33/09-07 33/09-09 33/09-12 33/09-15 33/09-16 34/07-01 34/07-02 34/07-03 34/07-04 34/07-05 34/07-06 34/07-07 34/07-08 34/07-09 34/07-10

S S S S S S C, B C, B 0 B 0 0 S S S

34/07-12 34/07-13 34/07-14 34/07-15S 34/07-16 34/07-17 34/07- 17A 34/07-18 34/07-19 34/07-20 34/07-21 34/07-21A 34/07-22 34/07-23A 34/07-23S 34/07-24S 34/07-25 34/10-18

0 O 0 B O B, S O B 0 B C. B . S C. B . S B C. B . S CBS B. S . B 0

o o

0

C, ne w cor e descriptions ; B , biostratigraphica l revie w o r ne w analysis ; S . use d i n seismic interpretation ; O, use d fo r genera l information. interpretation o f th e Snorre- H basi n ha s bee n done usin g th e E8 6 survey, whic h i s considere d to b e o f goo d qualit y with reflector s and fault s being well imaged. Besides th e discover y wells , the wel l databas e consists of well s on the hig h area s surroundin g the Snorr e half-graben , bot h o n th e Snorr e footwall itsel f an d o n th e Statfjor d Eas t Faul t

Block (Fig . 3) . The well s use d i n th e stud y ar e listed in Table 1 together with notes o n how they have bee n used . A goo d ti e betwee n well s an d seismic dat a ha s bee n obtaine d (Fig . 4) . Core description s hav e bee n mad e o f th e Snorre-H well s (Tabl e 1) , a s wel l a s biostrati graphical in situ zonation and reworkin g studies. As ca n b e see n fro m th e table , several o f th e

Fig 4. Seismi c rando m line , oriente d approximatel y E-W , showin g tie s t o well s 34/7-23 A and 23S.

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analyses are quantitative and considered o f good quality. The zonations are mainly based o n new analyses carried ou t b y Norsk Hydro/IKU . Geological setting The Snorre Fault Block (34/7) belongs to an array of westerl y tilted, majo r faul t block s constitut ing the Tampen Spu r area (Fig. 1) . The Tampe n Spur is bordere d to the eas t by the norther n North Se a Viking Graben an d t o the west by the Marulk an d Magnu s Basins . The Vikin g Grabe n forme d i n respons e t o Jurassic rifting , commencin g i n th e lat e Mid dle Jurassi c an d i n place s continuin g int o th e earliest Cretaceous . However , neithe r th e onse t nor th e cessatio n o f bloc k tiltin g wa s synchro nous throughou t th e basin . Ther e i s evidenc e for variabl e rate s o f extensio n an d faul t bloc k rotation, wit h th e characteristic s o f a pulse d syn-rift sedimentatio n (Partingto n e t al. 1993 ; Rattey & Ha y ward 1993 ; Faerset h e t al . 1995 , 1997; Ravna s & Bondevi k 1997 ; Faerset h & Ravnas 1998) . The earl y rift-stag e coincide s approximatel y with drownin g o f th e axia l Bren t delt a syste m and depositio n o f th e lowe r par t o f th e mud prone, marin e Heathe r Formatio n (Fig . 2) . It ma y in places encompass deltai c sediments of the uppermos t Bren t Group (Grau e e t al . 1987; Badley e t al . 1988 ; Helland-Hansen e t al . 1992 ; Johannesen e t al . 1995 ; Ravna s e t al . 1997) . A characteristi c feature of th e earl y rift-stage is that major faul t block s exhibi t modest rotation , without developin g major footwall islands. The mai n rif t stag e i s characterize d b y in creased extensio n an d accelerate d faul t bloc k rotation, leadin g t o th e developmen t o f deep , hangingwall half-graben s an d footwal l island s with significan t H - f an d t o furthe r segmenta tion o f th e rif t structure . Th e clay-pron e uppe r Heather and Draupn e Formation s (Fig. 2 ) were deposited locall y in thes e sediment-starved subbasins, a s wel l a s withi n th e axia l part s o f th e northern Nort h Sea , where there was significan t bathymetry. Erosio n o f the Heathe r Formatio n shales o n th e emergen t footwal l island s con tributed t o mu d deposition , bu t wher e erosio n was dee p an d san d pron e strat a lik e th e Bren t and Statfjor d Formation s became exposed, local deposition o f sand too k place . The lat e rif t stag e i s characterized b y waning rates o f faul t bloc k rotation , allowin g fillin g and, i n som e places , progradatio n o f uppe r Draupne Formatio n sandstone s ove r th e pre viously deposite d shale s o f th e Heathe r an d lower Draupn e Formations .

Some author s hav e invoke d a significan t strike-slip componen t withi n th e Nort h Se a Jurassic rif t syste m (Gower s & Saebo e 1985 ; Cartwright 1987 ; Beac h e t al . 1987 ; Larse n 1987). Elevate d structure s i n th e Tampe n Spu r area (Gullfaks , Visund Fault Blocks ) have been interpreted a s result s o f transpressio n (Beac h 1985; Speksnijder 1987 ; Fossen 1989) . However, it i s no w generall y accepte d tha t structurin g results from extensio n and that most of the main Jurassic fault s ar e clos e t o dip-sli p (Faerset h 1996; Roub y et al . 1996 ; Faerseth et al . 1997). The Marul k an d Magnu s Basin s for m essen tially a n extensio n of th e M0r e Basin , which is part o f th e Nort h Atlanti c rif t structure . Th e timing of this structure differ s slightl y from tha t of the northern North Sea , with rifting continuing into the early Cretaceous, an d th e structura l evolution o f Bloc k 34/ 7 mus t therefor e b e viewed i n thi s context. Major NE-S W strikin g fault s i n th e More , Marulk an d Magnu s basin s cu t acros s N- S trending Jurassi c structura l element s o f th e Viking Graben . Th e northernmos t par t o f th e Tampen Spu r (Block 34/4) has suffere d c . 1.5km of composit e uplif t (Yieldin g 1990 ) i n th e footwall o f such crosscuttin g master faults . Th e block-bounding (first-order) master fault s pene trate th e brittle , uppe r crus t (12-1 4 km) and presumabl y sol e ou t a t th e brittle-ductil e boundary (Fig . 1) . Fault s tha t defin e detach ments above or within basement ar e particularly related t o th e easter n Tampe n Spu r (Faerset h et al . 1995 ; Odinse n e t al . 2000 ; Fosse n e t al . 2000). The Snorre-H Structur e The Snorre- H structur e ( = Snorre-H Faul t Block) is defined herei n a s th e souther n par t o f the Snorr e Faul t Block , comprising the uplifte d Snorre-H footwal l an d neighbourin g Snorre- H hangingwall sub-basin to th e west. It i s confined to th e west by the NNE-SSW striking Statfjor d and Statfjor d East Fault s and t o the east b y the Inner Snorr e Faul t (Fig . 3) . T o th e north , th e main Snorr e Faul t Bloc k merges to som e extent with the Statfjord Faul t Block . The main Snorre Field i s segmente d b y th e Oute r an d Centra l Snorre Faults . The Statfjor d Eas t Faul t syste m is composed of a serie s o f NE-striking , left-stepping , en echelon faul t segment s that ar e linked by northerly-striking segment s (Fig . 3 ) (Dawers e t al . in press). The NE-striking en echelon segments are particularly evident at the top Bren t level, where they defin e a serie s o f sub-basin s alon g th e Statfjord Eas t footwall .

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Fig. 5. (a ) Time map o f the bas e Cretaceous horizon. The horizo n wa s interpreted o n ever y secon d inlin e and every tenth crossline using surveys E86, SG9201, SG8431 and GE83. (b) Dip-azimuth map of the autotracked base Cretaceou s map, with th e gre y scal e simulatin g illumination fro m th e northwest . Th e ma p wa s generated with ligh t fro m th e N W a t a n elevatio n o f 30° . As can be seen from the Base Cretaceous tim e map (Fig . 5a) , th e large r Gullfaks-Snorr e foot wall structur e plunge s North . Th e Snorre- H hangingwall sub-basin , o n th e othe r hand , plunges S W an d becomes increasingl y wide r along strike . The Snorre-H structur e can be seen

to wide n fro m les s tha n 10k m nort h o f well s 34/7-23 an d -23 S t o mor e tha n 15k m sout h o f wells 34/7-2 1 an d -21 A (Fig . 3) . Clos e t o wel l 34/7-24S i t i s about 12k m i n width . The dept h varies betwee n 172 2 and 3558ms , wit h a maxi mum i n th e down-throw n sid e o f th e Statfjor d

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Fig. 6 . Snorr e Fault Bloc k cross-section panel . See map inser t for location o f cross-sections. Fault abbreviations are th e sam e as i n Fig. 3.

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Fault. Th e Snorre- H hangingwal l sub-basi n exhibits a curren t di p o f 9-13 ° a t pre-rif t (to p Brent Group , o r close ) leve l withi n th e are a under consideration . Maximu m dip s ar e recorded i n th e SS W of the area . The internal faul t patter n o f the Snorr e Faul t Block i s illustrated o n th e bas e Cretaceou s dip azimuth ma p (Fig . 5b) . Th e smalle r intra faul t block (second-order ) fault s tha t resul t i n th e compartmentalization o f th e mega-block s ar e generally basement-detached. Withi n the Snorr e Fault Bloc k suc h fault s strik e predominantl y N-S, NE-S W and E-W . A mor e detaile d descriptio n o f th e structural characteristics an d evolutio n o f th e Statfjor d East Faul t an d associate d Snorre- H sub-basi n is given i n Dawer s e t al. (1999). The Snorre-H syn-rif t successio n A wedge-shape d syn-rif t succession , comprisin g the Heather an d Draupne Formations , i s present in the hangingwall trough o f the Snorre-H Faul t Block. Th e wedg e range s fro m Mid-Bathonia n to Ryazania n in age. It is not clea r if the syn-rif t succession als o include s th e uppe r par t o f the Bren t Group , a s ha s bee n suggeste d o n th e Statfjord Faul t Bloc k t o th e sout h (Johannese n et al. 1995) . A recent stud y of the Bren t Group/ Tarbert Formation in the Snorre area , however, suggests that intra-basina l faults controlle d very small depocenter s i n Bajocia n times , signifyin g the earlies t patter n o f faul t initiatio n (Davie s et al . 1998) . The Heathe r Formatio n (Mid-Bathonian Early Oxfordian ) consist s predominantl y o f silty shale s of offshor e marin e origin . The Draupne Formatio n (lat e Earl y Oxfordian Ryazanian) consist s o f anoxic , blac k shale s o f offshore marin e to slope origin (Draupne Shale) , coarsening upwards into fine-grained sandstones of a progradin g shorelin e comple x (Uppe r Draupne Sandstone) , occasionall y includin g an intercalated, assume d incise d slop e channe l sandstone uni t (Intr a Draupn e Sandstone) . Th e Draupne Formatio n rests , partl y unconform ably, on the Heather Formation . Syn-rift deposits are generally absent in crestal areas o f the Snorre- H an d Statfjor d Eas t Faul t Blocks, bu t thi n Draupn e Formatio n shale s o f Late Volgia n ag e ca p part s o f th e presen t da y crestal areas of the Snorre Faul t Block . In cross sectional view , th e uppe r Jurassi c syn-rif t infil l can be seen t o but t agains t the NE-SW striking Statfjord Eas t Faul t t o th e north , wherea s i t blankets part s o f th e degrade d footwal l t o th e south (Fig . 6) . Th e syn-rif t successio n i s pre sently deepes t burie d i n th e south-wes t o n th e

Snorre-H hangingwal l an d toward s th e eas t o n the Visun d Fiel d (Fig s 5 & 6) . I t thicken s southward, reaching som e 200m s TW T sout h of wel l 34/7-24S . Significant thicknes s variation s occu r i n th e upper Jurassic syn-rift infil l succession across the en echelon segment s o f the Statfjor d Eas t Faul t (Dawers et al. 1999) . A series of fault-controlled depocenters develope d durin g th e Oxfordia n through Kimmeridgian , involvin g a progressiv e younging o f th e syn-rif t infil l fro m sout h t o north i n the Snorre- H sub-basin . A thi n limeston e marke r be d (Crome r Knol l Group) o f Hauterivian/Valanginia n ag e drape s the Snorr e Faul t Block . It overlie s the Draupn e Formation i n th e dista l hangingwal l area s an d rests directl y o n th e syn-rif t unconformit y i n updip footwal l crestal areas .

Stratigraphic zonation A fairl y complet e Stratigraphi c successio n i s recorded i n th e Snorre- H well s fro m th e to p Bajocian throug h Kimmeridgian , althoug h sev eral mino r Stratigraphi c break s see m t o b e present (Fig . 7) . Th e type s an d numbe r o f events observe d var y a s a resul t o f variabl e sample densit y withi n an d betwee n th e wells , however, givin g a variabl e Stratigraphi c resolu tion. Th e samplin g density , o n average , i s 1-2 m in wells 34/7-21, 21 A, 23A, 23S and 2-3 m in wel l 34/7-24S . In wel l 34/7-21 , zones PJ5A2 and PJ6C 1 are not observed , suggestin g mino r Stratigraphi c breaks ma y b e presen t i n th e Mid-Callovia n and Mid-Oxfordian . Missin g zone s suggest ing minor breaks ar e speculate d als o i n th e late Early-early Mid-Volgian . Th e lat e Mid-Vol gian-Late Volgian and Late Ryazanian intervals are al l considere d presen t i n th e well , bu t th e transition fro m Mid - to Lat e Volgian is not wel l defined. Fro m comparison s t o wel l 34/7-2 1 A, deposition o f th e Uppe r Draupn e Sandston e is interprete d t o begi n a t thi s boundary . Th e sandstone i s mainly Lat e Volgia n i n age , whil e the uppermos t fe w meter s ar e Ryazanian . A n acme o f Botryococcus is recorde d i n th e uppe r part of the Upper Draupn e Sandstone . Well 34/7-21 A i s locate d onl y 600 m wes t o f 34/7-21, an d th e Stratigraphi c successio n i s very similar. Som e difference s ar e see n i n th e lat e Early-Mid Volgia n interval , however , wher e some zone s ar e identifie d tha t wer e not see n i n 34/7-21, an d vic e versa. No zone s hav e bee n identifie d i n th e Mid Oxfordian i n wel l 34/7-23A . Similarly , th e Mid-Volgian i s no t conclusivel y identified . Therefore, a Stratigraphic break ma y be presen t

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Fig. 7. Biostratigraphi c zonation and wel l correlation, includin g chronostratigraphic and lithostratigraphi c correlations. Palynozonations for the Snorre-H wells are indicated. Depths to formation top s are in mRKBMD. in th e succession , although , a s the Mid-Volgia n interval i s no t ver y thick , thi s ma y reflec t a sampling density problem. Lat e Ryazania n tax a have been identified lower in the Upper Draupn e Sandstone tha n i n th e 34/7-2 1 wells, suggestin g that the main part of the sandstone i n this well is Ryazanian i n age . Th e lowermos t par t i s thought t o b e o f Lat e Volgia n age . A n acm e o f Botryococcus ha s bee n identifie d in the ver y to p of th e Draupn e Shale , just befor e onset o f san d deposition an d anothe r acm e i s recorde d jus t beneath th e top Kimmeridgian . Well 34/7-23 S is located abou t 250 m wes t of 34/7-23A an d th e successio n i s ver y similar . However, th e Mid-Volgia n i s les s wel l define d compared t o wel l 34/7-23A . A Botryococcus acme i s evident a t th e sam e relativ e position a s in wel l 34/7-23A.

The 34/7-24 S wel l ha s a successio n ver y similar t o well s 34/7-2 1 and -2 1 A, bu t wit h n o Upper Draupn e Sandston e present . Th e upper most par t o f th e shal e i s Ryazania n i n ag e an d is underlai n b y Lat e Volgia n strata . Th e Mid-Oxfordian an d earl y Mid-Volgia n are als o here absent , o r a t leas t n o indicativ e tax a hav e been recorded . The acme s o f Botryococcus describe d abov e have n o ag e significance , bu t ma y indicat e environmental stress , notabl y influ x o f fres h water int o th e basi n o r reworkin g o f previously deposited fresh/brackis h wate r sediments . Th e acmes observe d nea r th e bas e o f th e sand s i n 34/7-23S and -23 A coincide with a cored sectio n showing frequent coal fragment s in the sedimen t (see below) . Thi s even t i s though t t o b e cor relatable betwee n th e wells.

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Biostratigraphical reworking A considerabl e amoun t o f biostratigraphica l reworked materia l i s presen t i n th e syn-rif t succession, originatin g fro m erosio n o f loca l source uplands . Reworke d tax a i n th e well s 34/7-21, -23&an d -24 S have bee n quantitatively represented an d relate d t o th e formatio n wher e they normall y occu r i n situ (Fig . 8) , providin g information o n the history and timing of erosion and unroofin g of upland sourc e area s (Hager & Smelror 1997) . Typical Brent/Dunli n dinocys t flagellate s (Nannoceratopsis gradlis, Susadinium scrofoides) first appear i n zon e PJ6A2/6 B (Lat e Callovian / Early Oxfordian) , bu t i n ver y limite d num bers. 'Dunlin/Bren t taxa ' thereafte r ar e presen t throughout th e section. The distribution o f taxa, according t o th e lithostratigraphica l group ing, i s relativel y eve n unti l th e Ryazanian , where 'Draupne taxa' dominate. N o indication s of reworkin g o f strat a olde r tha n Toarcian Aalenian hav e bee n seen , i. e correspondin g t o Statfjord Formatio n or Triassi c rocks , eve n though i t ha s bee n documente d tha t suc h dee p erosion occur s furthe r nort h an d t o th e eas t o n the Snorr e Faul t Bloc k (Soll i 1995) . A marke d increas e in reworking occurs in th e Kimmeridgian an d continue s int o Earl y Vol gian. In the Middle Volgian there is an apparen t

decrease an d lac k o f reworking . However , thi s probably relate s mostl y t o th e fac t tha t th e Middle Volgia n interva l ha s no t bee n con clusively identifie d i n an y o f th e Snorre- H wells. I t i s considered unlikel y tha t ther e wa s a significant reductio n i n reworkin g acros s thi s interval. I n Lat e Volgian-Ryazania n time s re working wa s ver y intens e an d encompasse s taxa fro m Draupne , Heathe r a s wel l a s Bren t lithologies. The timin g an d sequenc e o f reworkin g var y between th e wells , wit h th e event s comin g i n slightly earlie r i n well 34/7-21 compared t o wel l 34/7-23S. Th e reworke d tax a foun d i n th e Ryazanian sandstone s i n well s 34/7-23 S an d -23A ar e mor e numerou s tha n th e i n situ form s and man y of the species are characteristic of the Oxfordian-Volgian, whil e the y d o no t occu r (or ar e les s common ) i n th e Ryazanian . Thi s indicates activ e erosio n o f previousl y deposite d upper Jurassi c sediment s an d redepositio n within th e progradin g Uppe r Draupn e Sand stone system . Sinc e th e mai n sedimentar y en d product i n thi s cas e i s sand , i t i s believe d tha t erosion (cannibalization ) o f earlie r deposite d syn-rift sand s i n structurall y highe r position s may hav e occurred . In som e case s i t ca n b e show n tha t rework ing o f a depositiona l uni t occurre d a 'fairl y short time ' afte r deposition . I n wel l 34/7-23S ,

Fig. 8 . Reworke d tax a i n selected Snorre- H well s related t o formatio n o f origin an d structura l position. Th e amount o f reworked tax a is shown relative to th e ag e of th e strat a where they were found and fro m whic h formation/group the y originate (Draupne , Heathe r an d Brent/Dunlin) .

SYN-RIFT EVOLUTIO N I N TH E SNORRE- H AREA reworking of Late Kimmeridgian-Early Volgian sediments started alread y in the Mid-Volgian, as shown b y th e presenc e o f Perisseiasphaeridium pannosum. I n wel l 34/7-24S , reworkin g o f Mid-Oxfordian-Early Kimmeridgia n sediment s occurred i n th e Earl y Volgia n (Ctenidodinium chondrum, Gonyaulacysta jurassica longicornis) and o f Kimmeridgian-Earl y Volgia n sediment s in lat e Mid-Volgia n time s (Cribroperidinium longicorne). Thes e reworke d tax a occu r withi n Draupne Formatio n shale s o f unifor m appear ance, representin g low-energy , offshor e marin e conditions. Shortl y after , however , i n Lat e Volgian times , san d progradatio n reached th e area o f wells 34/7-21, -21 A, -23S, and -23A . I t is worth noting , however , tha t a tim e lag between deposition an d erosion/reworkin g o f som e fe w Ma, a s note d above , i s not a particularl y 'shor t time' i n term s o f faul t activity . Because erosio n o n footwal l scarp s tend s t o cut int o th e entire elevated stratigraph y simulta neously, erosio n i s likely to hav e cu t int o Bren t Group strata i n the Statfjord Eas t footwal l at an early stage. On the adjoining Snorre-H hangingwall, erosio n likel y cu t sequentiall y int o th e stratigraphy incorporate d i n faul t bloc k rota tion. Hence , Heathe r Formatio n strat a expect edly woul d b e erode d fo r som e tim e o n th e Snorre hangingwal l prio r t o exhumatio n an d erosion o f th e Bren t Group . Ther e i s a wea k indication of an inverted succession o f reworked sediments, suggestin g progressiv e unroofing . Overall, however , erosio n int o th e Heathe r Formation an d Bren t Grou p probabl y wa s more o r les s simultaneou s withi n th e availabl e time resolution . Thi s i s likel y a resul t o f th e semi-enclosed basi n physiograph y o f th e plun ging Snorr e Faul t Block , providin g both simpl e updip derive d an d longshor e transporte d sedi ments an d leadin g t o a mixin g o f reworke d taxa fro m sourc e area s tha t wer e a t differen t levels o f erosion .

Description of th e Draupne Formation Fades characteristics Nineteen core s fro m th e well s 34/7-21, 34/-21A , 34/7-23A and 34/7-23 S have been described an d interpreted (Fig . 9) . The uppermost Heathe r Formatio n i s cored in well 34/7-21 , wher e i t i s represente d b y 4 m o f strongly bioturbate d silt y shal e wit h abundan t glauconite an d pyrit e nodules . Paleophycus an d Skolithos ar e th e commo n ichnofacies . Th e to p surface i s well defined , intensel y burrowed, an d includes Belemnit e remnants .

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The lowe r par t o f th e overlyin g Draupn e Formation consist s o f thinl y laminated , littl e bioturbated blac k shale , th e Draupn e Shale , with thin interbeds of carbonate shell-debri s an d pyrite lamina e an d nodules . Th e blac k shale s pass graduall y upward s int o laminate d dar k grey shal e wit h thin , grade d o r rippl e cross bedded sand y lense s o r sand y laminae , ofte n appearing a s doubl e layering . Outsiz e sand y beds o f 1 to 2c m ar e presen t i n th e lowe r par t and become more frequent upsection . Contorted bedding structure s an d small-scal e syn-sedimen tary faultin g have bee n observed . Coa l debris is abundant i n the upper part , whereas Belemnites are presen t throughout . Muensteria an d Asterosoma ar e th e commo n ichnofacies , bu t som e Planolites burrowin g ha s als o bee n observed . The Draupn e Shal e i s interprete d t o reflec t deposition o n a tilte d hangingwal l dip-slop e with genera l stagnan t wate r circulatio n an d anoxic wate r botto m conditions . Th e rippl e bedded sand y lenses , togethe r wit h small-scal e slumping an d resedimente d cla y an d coa l frag ments, ma y indicat e overall gradient steepening accompanied b y coastlin e progradatio n an d sand reachin g th e uppe r slop e environmen t b y storm processes . A s show n b y th e reworkin g analysis, local erosion o f the Heather Formatio n contributed to deposition o f the Draupne Shale . The presenc e o f san d layer s an d coa l frag ments i n th e uppe r par t indicat e erosio n o f th e Brent Group . A 3 m thic k interbedde d sandston e unit , th e Intra Draupn e Sandstone , ha s bee n observe d i n well 34/7-21 . I t i s als o recognize d o n well-log s in 34/7-21A , bu t i n thi s wel l i t i s no t cored . It consist s o f fine- to medium-grained , flat-bedded t o cross-bedded , micaceou s sandston e wit h abundant mu d clasts and coal debris. Contorte d bedding an d frequen t interna l erosion surfaces testify t o frequen t failure . Th e Intr a Draupn e Sandstone i s interpreted as a submarin e channel deposit wit h ensuin g gravit y flo w an d tractio n currents, indicativ e of hig h rate s o f faul t bloc k rotation an d gradien t steepening . Th e blac k shale clast s wer e probabl y derive d fro m th e underlying Draupn e Shale , possibl y throug h channel wal l failure . Th e micaceou s san d an d coal fragments , o n th e othe r hand , wer e likel y sourced fro m erosio n o f th e Bren t Group . I t i s suggested that the sand was supplied fro m failur e and remobilizatio n o f a n incipien t updi p sand y shoreline o r mout h ba r system , likel y du e t o oversteepening an d instabilit y resultin g fro m progressive faul t bloc k rotation . Thickness an d frequenc y of sand y lamina e i n the Draupn e Shal e increas e upward s int o 1 0 to 40cm thic k grade d units , whic h hav e a lowe r

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Fig. 9 . Wel l correlation, includin g core descriptions , an d sequenc e architecture of the Snorre- H Draupn e Formation.

burrowed sand y par t an d a n uppe r laminate d shaly part . Thi s heterolithi c fade s passe s tran sitionally upwards into a tabular sandstone unit, the Uppe r Draupn e Sandstone , whic h i s u p to 35m thick. It consists of cm to dm thick units of micaceous , fine - t o medium-grained , ripple and cross-laminated sandstone . Bioturbation decreases upward s an d horizonta l laminatio n and cross-bedding, occasionall y low-angle , charac -

terise the uppermost part . Paleophycus an d Skolithos burrowin g occasionall y overprin t th e stratification. A lowe r t o uppe r shoreface , prograding shorelin e environmen t is suggested. The lack of top se t deposits inhibits a conclusive interpretation o f th e shorelin e type , bu t th e general characte r suggest s a mixe d fluvial (protuberated) an d wave-dominate d (linear ) shoreline progradation .

SYN-RIFT EVOLUTIO N I N TH E SNORRE- H AREA The uppermos t 6 m o f th e Uppe r Draupn e Sandstone i n well s 34/7-23 A an d -23 S consis t of mediu m grained , paralle l bedde d sandston e with intercalate d cla y laminae. Fain t cross bedding ha s been observe d an d th e sandstone i s moderately bioturbated , mostl y b y Skolithos burrows. It is interpreted as a transgressive sandstone cappin g th e Uppe r Draupn e Sandston e in thes e wells. In wells 34/7-21 A, -23A and -23S, the top par t of the Upper Draupn e Sandston e i s impregnated by diffus e carbonat e growt h passin g graduall y into carbonat e nodules , relate d t o th e subse quent Crome r Knol l transgression . Th e lower most part o f the Cromer Knol l Group was cored in wel l 34/7-2 1 A, wher e it consist s o f limeston e nodules interstratifie d wit h pyrit e ric h greenis h marls passin g int o interbedde d nodula r carbo nates an d re d shales . Shel l debri s i s ver y abundant an d ammonite s ar e occasionall y pre sent. Th e Crome r Knol l Grou p i s interprete d as a deepenin g carbonat e succession , terminat ing the overall shallowin g upward natur e o f the Upper Draupn e Sandstone . Th e red-coloure d interbeds indicat e oxygenate d wate r condi tions, possibl y belo w th e carbonat e compensa tion depth . Key bounding surfaces Five mai n surface s (S- l t o S-5 ) ca n b e distin guished betwee n the top o f the Heather Forma tion an d bas e o f the Cromer Knol l Group . The four lowe r surfaces are tie d t o well 34/7-21, as it

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is the onl y well being cored al l the wa y down t o the to p Heathe r Formation . Th e fifth surface is only represente d i n well s 34/7-23A and -23S . A correlation pane l ha s bee n constructe d betwee n the well s 34/7-21 , -21A , -23A , -23 S an d -24 S (Table 2 ; Fig. 9) . Surface 1 (S-l ) occur s a t 2580.0 m i n wel l 34/7-21, withi n th e Draupn e Shale , an d corre sponds t o th e chang e fro m a mudd y sectio n with frequen t mino r syn-sedimentar y slump s and fault s t o a laminate d mudd y sectio n tha t gradually intercalate s millimetr e thi n sand y laminae. Th e surfac e give s a distinc t gamma ray signal . I n wel l 34/7-21 A, i t is well below th e deepest core d interva l (cor e no . 5) . The gamm a ray signa l seem s t o b e correctabl e t o well s 34/23-A an d -23S , althoug h n o lithologica l changes hav e bee n noticed . S- l correspond s approximately to the top Kimmeridgian timeline and intra Draupne seismi c event. It is characterized by an acm e of Botryococcus (fresh-brackish water algae ) i n th e 34/7-23 S an d -23 A wells , suggesting increase d freshwate r run-of f o r reworking of older fresh-brackish water deposits at thi s time . S- l i s interpreted a s a floodin g o r major basin deepening surface. The combination of increased run-off/reworkin g an d accompany ing basi n deepenin g i s believe d t o indicat e a major til t even t o r puls e o f increase d rate s o f fault bloc k rotation . Surface 2 (S-2 ) occur s a t 2573.0 m i n wel l 34/7-21. I t correspond s t o th e bas e o f the Intr a Draupne Sandstone and appears to be related to

Table 2 . Ke y surface calibration Surface

Type

Biostrat calibration

S-5 S-4

FS FS

S-3

FS

S-2

Base ID S

NP FO S . cf. palmul a FO O . diluculum , B, SED FO B . pomum, ACME Pterospermella & Tasmanites LO G. mutabilis , B, L, SED LO S . jurassica, ACME Botryococcu s L, SED

Near Bas e TL IDS S-l FS/TL

Top T Oxfordian

L

LO O. patulum, FO D . spinosu m LO G. jurassica, F O C. panneum , L O P. pannosum LO S . crystallinum

34/7-21 34/7-21

A 34/7-24

S 34/7-23

NP NP B, L , SE D B

A 34/7-23

S

L?, SE D L, SE D

B?. L? , SED L

L

B, L

B?, L

B?. L

L

Merges with S 3? B

Merges with S 3? L

Merges with S 3? B, L

B, L, SED L

B, L

B

B

B, L

B, L

B, L

L

B, L

L

,L

(FS, floodin g surface ; TL , timeline ; NP , no t present ; B, correlated o n palynologica l event ; L , correlated o n log response; SED, correlate d o n sedimentar y features ; ? , calibration uncertain .

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the channe l morpholog y o f thi s sand body . S-2 is characterize d b y a stron g gamma-ra y signal . It i s still present in wel l 34/7-2 1 A, bu t probabl y merges wit h surface S- 3 in th e othe r wells. Surface 3 (S-3 ) occur s a t 2566.0 m i n wel l 34/7-21, immediatel y abov e th e Intr a Draupn e Sandstone. I t mark s th e bas e o f th e Uppe r Draupne Sandston e progradin g shorelin e sequence. S- 3 has a distinc t gamma-ra y signal . It ca n b e correlate d wit h the bas e o f th e uppe r Draupne coarsening-upwar d sequenc e i n wel l 34/7-23A and with the strongest gamma signal in the non-cored sectio n of well 34/7-23S. As for S-l an acm e of Botryococcus is observed at this level in th e 34/7-23 S and " -23A wells , suggestin g increased freshwate r run-of f o r reworkin g o f older fresh-brackis h wate r deposits a t thi s time . S-3 is believed to represen t a floodin g o r majo r basin deepenin g surface . Th e combinatio n o f increased run-off/reworkin g an d accompanyin g basin deepening i s believed t o indicate a pulse of increased rate s of fault bloc k rotation . Surface 4 (S-4 ) occur s a t 2521.0 m i n wel l 34/7-21 an d correspond s t o th e to p o f th e progradational Uppe r Draupn e Sandston e shoreline sequence . Th e surfac e correlate s t o 2878.0 i n wel l 34/7-2 1 A, t o 3244.0 m i n wel l 34/7-23A an d fall s i n th e no n core d sectio n i n well 34/7-23S , where " it i s indicate d b y a stron g gamma peak . The correlation o f S-4 is based o n biostratigraphical dating , however , an d i s no t well constraine d i n well s 34/7-23 S an d -23A . It coincide s approximatel y wit h th e boundar y between Lat e Volgia n an d Ryazanian . S- 4 i s interpreted a s a flooding or major basi n deepening surface. Surface 5 (S-5 ) occur s a t 3211.8 m i n wel l 34/7 23A, a s a conspicuou s shel l be d o f Belemnites. The facie s descriptio n o f core no.l o f well 34/7 23 S suggest s tha t th e surfac e correlate s t o th e lithology chang e a t 3086.5m . S- 5 delineate s the bas e o f th e uppermos t par t o f th e Uppe r Draupne Sandston e sequence in these wells, an d is truncated b y the unconformity underlying the Cromer Knol l Group . S- 5 is believe d t o repre sent th e fina l floodin g o r basi n deepenin g sur face i n th e syn-rif t successio n prio r t o lowe r Cretaceous transgressio n an d drownin g o f th e rift terrain . Sequence architecture The key surfaces described abov e can be used to define depositiona l sequence s (Seq- I t o Seq-V )

within th e Draupn e Formation , wit h th e to p Heather Formation , S-l , S-3 , S- 4 an d bas e Cromer Knol l Grou p a s th e mai n sequence bounding surfaces . A sequenc e correlatio n be tween th e well s i s show n i n Fig . 9 wher e th e more obviousl y correctabl e floodin g surface s have bee n used a s sequenc e bounding surfaces. Sequence I (Seq-I ) i s lat e Earl y Oxfordia n t o Kimmeridgian i n age . I t i s bounde d b y th e unconformity/hiatus a t to p Heathe r Formatio n and th e floodin g surfac e S-l , whic h coincide s approximately wit h th e intr a Draupn e seismi c event. Seq-I has bee n cored onl y in well 34/7-21, where it consists of anoxic and laminate d clays, typically characterized b y frequent indications of minor slumping and syn-sedimentary faulting. It is onl y a fe w metre s thic k i n wel l 34/7-21 , bu t from wel l log correlation it appears t o thicken in wells 34/7-2 3 an d -24 . Th e cor e description s point t o starve d conditions , likel y relate d t o basin compartmentalizatio n an d steepenin g depositional gradients . Sequence I I (Seq-II ) i s bounded b y th e S- l an d S-3 surfaces. It comprises th e Early Volgian an d the lower part o f Mid-Volgian. Seq-II is cored in wells 34/7-21 , -23A and -23 S and coincide s with the appearanc e o f th e firs t san d influx , firs t a s millimetre thick laminae , and gradually thicken ing up t o mor e frequent , centimetre-thic k intercalations o f rippl e cross-laminate d sandstones . The 3 m thic k Intr a Draupn e Sandston e slop e channel cored in well 34/7-21 is not see n in wells 34/7-23A an d -23S , bu t a n associate d stron g gamma pea k seem s to b e recognizable through out th e area . Thi s suggest s that , wherea s th e sandstone uni t itself is essentially discontinuous and localized , the time surface at which it occurs is o f mor e latera l extent . Th e correlatio n i s supported b y th e positio n o f th e acm e Botryococcus i n th e las t tw o wells . Th e bas e an d to p surfaces o f th e Intr a Draupn e Sandston e likel y merge into on e single surface locally, depending on th e presenc e o r absenc e o f channe l sand bodies. Seq-I I i s interprete d a s a shallowin g upward slop e succession , reflectin g increase d local sedimen t supply. Sequence HI (Seq-III) is Mid- to Late Volgian in age. I t i s bounde d b y th e S- 3 and S- 4 floodin g surfaces an d include s most o f th e Uppe r Draupne Sandston e i n well s 34/7-2 1 an d -2 1 A. It consist s o f a 4 5 m thic k shorelin e succession, coarsening upwar d fro m laminate d mu d t o coarse-grained an d bioturbate d shorefac e sand stones. I n well s 34/7-23 A an d -23S , i t i s repre sented only by a 3 to 4 m thick, more fine-grained

SYN-RIFT EVOLUTIO N IN TH E SNORRE- H AREA coarsening upwar d sequence . Seq-II I reflect s a gradual shallowing of the basin, likely as a result of reduced rate s o f fault block tilting. Sequence IV (Seq-IV ) is bounded by S-4 and S- 5 in well s 23/7-23A and -23S , and b y S- 4 and th e base of Cromer Knoll in wells 34/7-21 and -2 1 A. It i s Ryazania n i n age . I n well s 34/7-23 A an d -23S i t i s represented b y a n upwar d coarsenin g shoreline succession that is similar, in all aspects, to Seq-II I i n well s 34/7-2 1 an d -2 1 A. I t i s interpreted t o onla p S-4 , an d i n well s 34/7-21 and -21 A it appears a s a thin, transgressive san d capping th e Uppe r Draupn e Sandstone . Seq-I V is interprete d a s a progradationa l incremen t i n an overal l back steppin g shorelin e system. Sequence V (Seq-V) is bounded b y S- 5 and th e base of the Cromer Knoll Group and is assumed latest Ryazania n i n age . I t i s represente d b y about 6 m of transgressive sand constitutin g the uppermost par t o f th e Uppe r Draupn e Sand stone i n well s 34/7-23 A an d -23S . Seq- V ma y represent th e fina l retrea t an d drownin g o f th e coastal progradation, bu t could as well be seen as the lateral equivalent to a thicker sandy sequence deposited even further updip i n the basin . Draupne Formatio n seismic characte r Two well-define d seismi c reflectors , th e bas e Cretaceous an d to p Heathe r Formation , mar k the top and base of the Draupne Formatio n an d have been mapped ou t regionall y (Figs 4 & 10) . In addition , severa l intra-Draupn e Formatio n reflectors have been mapped locall y across part s of th e Snorre- H area . Here , th e Draupn e Formation i s seismicall y defined a s consistin g of a lowe r uni t and a n uppe r unit, separated by an intra-Draupn e seismi c event . Th e intr a Draupne seismi c even t underlie s th e bas e o f the Uppe r an d Intr a Draupn e Sandstone s (Figs 6 & 10) and correspond s approximately to the to p Kimmeridgia n timelin e (bounding sur face S-l) . Th e intr a Draupn e seismi c event, i s observed t o onla p th e to p Heathe r alon g th e hangingwall dip-slope in the southern part of the Snorre-H basin . However , alon g th e north eastern edg e o f th e sub-basin , nea r well s -23 S and 23A , it appears t o b e truncated by the Base Cretaceous Even t (BCE) , possibl y relate d t o footwall uplif t alon g a small fault nea r the wells. Alternatively, it may just be too thi n to resolve. The lowe r uni t i s relativel y homogeneous i n seismic characte r an d correspond s t o deposi tional sequence Seq-I, the shale dominated lower part o f th e Draupn e Formatio n ( = Draupne

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Shale). However , a s th e uni t i s no t mor e tha n one cycl e thick , it s interna l characte r an d geo metry ar e no t wel l resolved. I t thicken s slightly westward, wit h severa l loca l depocenter s devel oped alon g segments of the Statfjord Eas t Fault . The upper seismi c unit is more heterogeneous in character an d contain s a number o f spatially discontinuous interna l reflectors . It correspond s to depositional sequence s Seq-I I through Seq-V, the layere d mudstones , siltstone s an d sand stones o f th e uppe r Draupn e Formatio n (incl . the Uppe r Draupn e Sandstone) , an d i t ha s a distinct wedg e shape d geometry . T o th e sout h in th e Snorre- H are a (Fig . 6) , th e uppe r uni t onlaps onto the lower unit. The lower part of the upper uni t ha s a sub-paralle l internal reflectio n pattern, wherea s the uppe r par t ha s a multiple lobe shape d character . I t i s describe d i n mor e detail below. In th e souther n par t o f th e area , relatively continuous reflection s overste p th e degrade d Statfjord Eas t footwal l (Fig . lOc) . Th e comple tion lo g fo r wel l 33/9-12 , locate d jus t nort h of th e mos t intensivel y degrade d footwal l area, record s a thi n drap e o f Lat e Volgia n Draupne shale .

Sandstones in the upper Draupne Formation Sandstones i n th e uppe r Draupn e Formatio n have bee n observe d i n man y well s in th e large r Statfjord an d Snorr e are a (Fig . 3) . It i s difficul t to ma p ou t thes e sand s seismically , a s the y are thin and do not commonly set up reflections that can b e followe d regionally . However , a se t o f reflections i n th e uppe r par t o f th e Draupn e Formation, correspondin g t o th e leve l of sandstone development , hav e bee n mappe d locall y across th e Snorre- H are a (Dawers el al. 1999). Figure 1 0 shows several traverses through the Snorre-H wells. The base of the Upper Draupn e Sandstone i s seismicall y resolvabl e i n well s 34/7-23S an d -23 A (Fig . lOa) . Th e sandston e i s interpreted a s a lobe-like unit lying immediately below th e BCE . I n well s 34/7-2 1 and -2 1 A, a reflector correlate s with the position of the Intra Draupne Sandstone (Fig. lOb) . This horizon can be mappe d northwar d an d southwar d fro m wells 34/7-2 1 and -2 1 A, formin g the bas e o f a separate lobat e unit , whic h i s interprete d t o contain bot h th e Intr a an d Uppe r Draupn e Sandstone units in this area. The seismic lobe in wells 34/7-23S and -23A overlies the seismic lobe in wells 34/7-21 and -2 1 A, supporting the note d biostratigraphical youngin g o f th e Uppe r Draupne Sandston e fro m lat e Volgia n i n well s 34/7-21 and -21 A to Ryazanian in wells 34/7-23S and -23A .

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Fig. 10. Seismi c sections through the Snorre-H are a wells. See map insert for location of lines, (a) Section through wells 34/7-23 A and -23S , showing the bas e of the Uppe r Draupn e Sandston e in pink, (b ) Traverse through wells 34/7-21 A an d -21 , showing the bas e o f th e sand-pron e sectio n i n yellow , (c ) Inline 17 0 from surve y E86 showing the Statfjor d Faul t (left ) an d Statfjor d Eas t Faul t (right) . Severa l internal reflectors overlie the to p Kimmeridgian horizon , shown in red; the lower ones onlap top Kimmeridgian , whereas the uppermos t one s appear t o drap e a degraded portio n o f the Statfjor d Eas t footwal l whic h has bee n displaced b y the Statfjor d Fault, (d ) Detail o f part c showing the chaotic wedge above the top Kimmeridgia n horizon. Note th e subtle downlap o f internal reflectors.

Both seismi c lobe s wer e mappe d i n seismi c section o n every second inlin e in the E86 survey. In area s wher e a bas e sandston e reflecto r was observed i n seismi c section , a weakenin g in th e amplitude o f th e bas e Cretaceou s reflecto r was also observe d (Fig . 1 1 a). Thi s reductio n o f th e otherwise stron g impedanc e contras t associate d with the passage fro m Cromer Knol l limestone s to Draupn e shale s i s though t t o aris e whe n sandstones are present i n the uppermos t par t o f the Draupn e Formation , a s a resul t o f les s velocity contras t betwee n limestone s an d sand -

stones tha n betwee n limestone s an d shales . Thus, ther e i s a potentia l t o discriminat e th e Draupne sandstone s using the seismic attributes of th e BCE . I n orde r t o investigat e thi s possibility a numbe r o f differen t seismi c attri bute maps , using variou s measure s o f ampli tude and variou s window widths, were made for the autotracke d interpretatio n o f th e BC E reflector i n th e E8 6 dat a set . Th e result s sho w that amplitud e anomalie s alon g th e BC E correlate wit h th e patter n o f bas e sandston e reflectors mappe d in the uppermos t par t of the

SYN-RIFT EVOLUTIO N I N TH E SNORRE- H AREA

Draupne Formation . Fo r example , Fig . li b shows th e spatia l variatio n o f th e rm s ampli tude withi n a windo w fro m 4m s abov e th e BCE t o 24m s below . Th e yello w throug h re d areas correspon d t o lo w amplitude s an d th e blue area s correspon d t o highe r amplitudes . Note tha t th e sandston e occurrenc e i n th e 34/7-21 well s and 34/7-2 3 wells corresponds t o low amplitudes . Th e aeria l exten t o f th e inter pretation o f th e bas e o f th e sandstone , whic h is resolvabl e i n seismi c section s (Fig s lO a an d lOb), i s als o show n o n th e attribut e ma p (Fig. lib) . Th e interpretatio n o f th e mappe d sandstone exten t correlate s ver y wel l wit h the areas of low amplitude observed i n this attribute map .

197

Other significant Draupne Formation seismic anomalies In th e deepes t par t o f th e Snorre- H basin , several stacke d reflector s are observe d t o onla p the top Kimmeridgia n reflector describe d abov e (Fig. lOc). Th e lowermos t o f thes e reflector s form th e top o f a seismically chaotic wedge that occasionally show s interna l reflection s tha t downlap ont o th e to p Kimmeridgia n horizon (Fig. lOd) . The exten t of th e wedg e is shown in Fig. 12 . The chaotic wedge appears as an area of low amplitude s relativ e t o th e surroundin g region, suggestin g that i t ma y b e lithologicall y distinct. Th e Statfjor d Eas t footwal l immedi ately adjacen t t o thi s are a ha s bee n degrade d

Fig. 11. (a ) Seismic traverse across th e Snorre-H basi n showing the interpretation o f several amplitude anomalies. See part b for location, (b) Spatial variation in the GeoQuest IES X seismic attribute calculation of rms amplitude for th e base Cretaceou s reflector . Thi s ma p wa s calculated usin g a window extending from 4 ms above th e bas e Cretaceous t o 2 4 ms below. The exten t of the Draupn e Formatio n subcro p i s outlined an d th e uppe r Draupn e seismic lobe s mappe d i n th e vicinity of the 34/7-2 3 and -2 1 wells. Not e tha t ther e ar e a numbe r o f localize d amplitude anomalies that ma y represent subtle , thin sands tha t are not readil y interpreted fro m seismi c sections.

SYN-RIFT EVOLUTIO N I N TH E SNORRE- H ARE A down through th e Brent/Dunlin levels (Figs lOc , 13). Thi s suggest s tha t th e chaoti c wedg e an d deeper interna l reflector s ma y b e relate d t o processes an d product s o f footwall degradation , and ma y i n fac t b e sandy . Simila r lo w ampli tudes ar e observe d als o i n th e vicinit y o f wel l 34/7-24S, o n th e othe r hand , whic h i s entirely shaly. However , footwal l degradatio n adjacen t to thi s well does no t involv e substantial erosio n of sand pron e formations lik e the Bren t Group , as seen t o th e south , nex t t o th e chaotic wedge . East o f th e wedge , som e elongate d N- S trending anomalie s ca n b e see n (indicate d b y arrow i n Fig . 12) . When flattenin g o n th e bas e Draupne/top Kimmeridgia n reflector , they seem to li e i n th e deepes t par t o f th e half-grabe n during thi s time . Th e interpretatio n o f thes e features i s no t clear , bu t a s the y ar e oriente d parallel t o th e colum n directio n o f th e 3 D survey, the y ma y jus t represen t geophysica l noise. Alternatively , bu t les s likely , the y ma y represent axiall y transported gravit y deposits . Figure 1 3 summarizes th e ke y elements o f th e seismic interpretation, includin g the distributio n of th e subtl e Uppe r Draupn e Sandstone , base d

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on th e seismi c attribute maps an d i n relation t o the mai n structura l elements . Discussion In th e following , th e tectonostratigraphic evolution o f th e Snorre- H are a i s summarize d an d erosion an d redistributio n of sand i n the syn-rif t succession i s discusse d i n relatio n t o potentia l hydrocarbon pla y models. Tilting of the S norre Fault Block In orde r t o understan d th e tectonostratigraphi c evolution of rotational fault blocks it is necessary to determin e th e til t history . A s a firs t approx imation, til t ca n b e measure d a s th e angl e be tween a pre-rotational , assume d horizontal , datum an d a give n syn-rif t surfac e in a seismi c section (Fig . 14a) . A s a resul t o f thinnin g an d pinch-out o f syn-rif t package s ont o th e foot walls, however , post-depositiona l differentia l compaction wil l caus e thicke r basina l deposit s

Fig. 12 . Spatia l variatio n i n th e rm s amplitud e fo r th e intr a Draupne to p Kimmeridgia n reflector. Thi s ma p was calculated usin g a windo w extending fro m 8m s below th e reflecto r to 24m s above . Th e outline d are a represents th e extent of the wedge-shaped uni t lying on top of the Kimmeridgian horizon. Note th e N-S trendin g amplitude anomalie s to th e sout h i n the ma p view . The y mos t likel y represen t geophysical noise, bu t could , alternatively, represen t axiall y transported gravit y deposits .

200

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Fig. 13 . Ma p showin g important subcro p trajectories, reflector terminations and interprete d aeria l extent of the Upper Draupn e Sandston e lobes.

SYN-RIFT EVOLUTIO N I N TH E SNORRE- H ARE A to compac t mor e tha n thinne r flan k deposits , reducing th e angl e betwee n to p an d bas e o f th e packages. Assumin g a typical compaction coeffi cient betwee n som e 30-40 ° (Sclate r & Christi e 1980), th e measure d angl e therefor e need s t o be correcte d b y a simila r numbe r (Fig . 14b) . In addition , th e assume d depositiona l gradien t in th e basi n a t th e tim e o f depositio n need s t o be added . The widesprea d occurrenc e o f coal s i n th e Brent Group , especiall y i n the Nes s Formation , provides a n excellen t definitio n o f palaeo-delta top (i.e . sea-level) in th e Snorre- H area . Appar ent parallel reflector s from th e top Bren t Group downwards sugges t i t wa s deposite d largel y before onse t o f stretchin g an d faul t bloc k rotation. T o determin e th e til t histor y o f th e Snorre Faul t Block , th e to p Bren t Grou p horizon i s therefor e use d a s datu m an d se t horizontal. The amoun t o f tilt varies from nort h to sout h alon g the Snorre-H structure , however, the til t histor y discusse d belo w ha s bee n estimated i n th e centra l Snorre- H area . Th e til t measurements hav e bee n average d ove r th e width o f th e hangingwal l sub-basin . Locall y along th e steepes t part s o f th e hangingwal l di p slope, muc h highe r dip s are attained (Fig . 6c-d) . The Bren t Grou p i n th e centra l Snorre- H area i s tilte d relativ e t o th e bas e Draupne/to p Heather (Early/Mid-Oxfordian ) surface by som e 2° toward s th e WSW . Decompaction o f th e Heather Formatio n b y 30-40 % (i.e . Sclate r & Christie 1980 ) increase s th e til t t o abou t 2.5 C . In addition , th e to p Heathe r depositiona l sur face ha d a primary gradien t tha t i s set to 0.5-1 . If higher , th e depositiona l surfac e woul d mos t likely becom e unstabl e an d trigge r significan t gravity collapse, o f which ther e i s little evidence . This implie s tha t th e faul t bloc k til t b y end Heather Formatio n (Mid-Oxfordian ) time s i s likely t o hav e bee n c . 3° (Fig. 14b , 15). The to p lowe r Draupn e seismi c surfac e (to p Kimmeridgian), whic h ca n onl y b e mappe d i n the southern par t o f the Snorre-H sub-basin , ha s less tha n l c o f additiona l til t compare d t o th e base Draupn e surface . However , depositio n o f the laminated blac k shale s of the lower Draupn e seismic unit (= Draupne Shale ) suggests a bathymetric deepenin g an d increas e i n depositiona l gradient durin g depositio n o f thi s unit , tenta tively set to T. Consequently, th e fault bloc k til t by Lat e Kimmeridgia n i s estimated t o c . 4-5c. The til t o f th e Bren t Grou p increase s t o 5 relative t o th e to p Draupn e (Lat e Ryazanian ) surface. Correctin g fo r a 30-40% decompactio n of th e syn-rif t infil l increase s th e til t t o 6.5-7° . In wel l 34/7-24S , th e gamma-ra y lo g suggest s a weakly coarsenin g upwar d sequenc e fro m blac k

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shales t o silt y shales , bu t wit h n o evidenc e o f sand depositio n (Fig . 9) , resultin g fro m under filling o f th e basin . A s core s o f th e Draupn e Formation shale s i n well s 34/7-2 1 and -21 A t o the south-east contai n slum p deposits, indicative of gravitationa l instability , the primar y deposi tional gradien t i n the 34/7-24 S area i s set higher than fo r th e Heathe r Formation , tentativel y to 2° . Thi s give s a to p Draupn e (Lat e Ryaza nian) faul t bloc k til t o f c. 9: , as measured i n th e central Snorre- H area . Thi s represent s a n incre mental increas e o f 4 : fro m Lat e Kimmeridgian. The cumulative late Jurassic-early Cretaceou s tilt amount s t o 1 3 i n th e souther n par t o f th e Snorre-H sub-basin , suggestin g tha t thi s part o f the Snorre- H Faul t Bloc k experience d highe r rates of rotation tha n di d th e central and north ern area . A to p Draupn e depositiona l gradien t in exces s o f 4 i s needed i n orde r t o explai n th e present til t i n thi s area . Suc h a hig h gradien t

Fig. 14. (a ) Methodology fo r estimating tilt angles in a syn-rift succession , (b ) Measured an d estimate d (corrected) til t angle s fo r th e Snorre- H hangingwall basin.

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Fig. 15 . Plo t showin g the Snorre- H Faul t Bloc k cumulative tectonic tilt an d sedimen t thicknes s vs. geological time. Th e timescal e i s after Gradstei n e t al. (1994).

would normall y trigge r significan t instabilit y and gravit y collapse , o f whic h ther e i s littl e evidence. Th e drapin g o f th e Draupn e Forma tion b y the Cromer Knol l limestones and lac k of differential loading , however , ma y hav e pre vented suc h collapse . Alternatively , i t raise s a question i f there has bee n som e additional, post rift tiltin g of th e Snorre- H Faul t Block .

Snorre-H football uplift model To determin e th e natur e an d histor y o f th e Snorre-H Faul t Bloc k footwal l uplift , th e so called 'shorelin e trajectory ' durin g syn-rif t evo lution ha s bee n investigate d (Fig . 16) . Th e shoreline trajector y refer s t o th e cross-sectiona l path of the shoreline a s it migrates in response t o changing relativ e sea-leve l (Helland-Hanse n & Gjelberg 1994) . Th e shorelin e trajector y i s synonymous wit h th e tempora l migratio n o f the 'basement emergence point ' of Roberts e t at. (1993) in cross-sectional view . It gives a measure of footwal l islan d widt h throug h time . Th e basement emergenc e poin t shoul d no t b e con fused wit h th e structura l fulcru m or leve r point , which separate s are a o f ne t uplif t fro m are a o f net subsidenc e acros s a rotatin g faul t block .

Depending o n th e positio n o f th e fulcru m relative t o th e basemen t emergenc e point , th e shoreline trajectory will vary and rotationa l faul t blocks ma y experienc e footwal l islan d expan sion, stillstan d o r contractio n subjec t t o a n identical til t histor y (Ravna s & Steel 1998) . In th e centra l Snorre- H are a th e BC E trun cates th e to p Bren t Grou p abou t halfwa y be tween the Statfjord East Faul t and the Snorre- H footwall crest (6/ 6 km). The top Heather Formation i s truncate d b y th e BC E som e 7k m westward o f th e Snorre- H footwal l cres t an d 5km eastwar d o f th e Statfjor d Eas t Faul t (Fig. 13) . As argue d i n th e discussio n o n tecto nostratigraphic evolutio n below , th e lowe r par t of th e Heathe r Formatio n i s believe d pri marily t o hav e extended fa r t o th e east, possibl y covering th e Snorre- H footwal l cres t (Fig . 16) . The shorelin e positio n (Abasemen t emergenc e point) b y th e Mid-Oxfordia n (en d Heathe r Formation) i s no t wel l understood , bu t w e suggest i t wa s locate d som e distanc e (2- 4 km) towards th e west. The to p Kimmeridgia n onlap s th e Heathe r Formation approximatel y 1- 2 km wes t o f th e BCE/top Heathe r truncatio n lin e (Fig . 13) . However, th e Kimmeridgia n successio n i s clearly offshore/dee p marin e i n origi n an d th e

Fig. 16 . Composit e tcctonostratigraphi c mode l o f th e Snorre- H area , linkin g Snorre- H hangingwal l depositional unit s to uplif t an d erosio n o f th e Snorre- H an d Statfjord Eas t footwalls . Not e tha t th e model represent s a combination o f th e tectonostratigraphi c evolutio n in th e souther n par t o f th e sub-basin , i n th e 34/7-2 1 wells area (Scqs- I II) , and th e norther n part o f th e sub-basin , nea r th e 34/7-2 3 wells (Seq-III-V).

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observed onla p must therefor e b e of submarin e origin, possibl y resultin g fro m axia l an d trans verse sedimen t suppl y t o th e Snorre- H half graben fro m th e Statfjor d Eas t footwall . Th e biostratigraphical dat a show s a n overal l com pression o f th e Lat e Oxfordian-Kimmeridgia n interval i n th e 34/7-2 1 well , indicating tha t th e shoreline b y th e en d o f Kimmeridgia n pro bably wa s locate d eas t o f th e presen t BCE/to p Heather truncatio n lin e (Fig. 16) . Volgian strat a furthe r onla p an d drap e Kim meridgian deposits , an d Lat e Volgia n prograd ing shorelin e sandstone s (Seq-III ; Fig s 9 , 13 ) fringe th e Snorre- H hangingwal l roughl y alon g the BCE/to p Heathe r Formatio n truncatio n line, c . 1-2 km eastwar d o f th e to p Kimmerid gian onla p line . The shoreline unit is overlain by a thi n Ryazania n transgressiv e sandstone . Thi s suggests initia l rapi d expansio n o f th e footwall island durin g lat e Earl y Oxfordian-Earl y Vol gian, followe d b y reduce d expansio n an d stag nation/contraction durin g th e Lat e Volgian Ryazanian (Fig . 16) . To th e north , toward s th e 34/7-2 3 wells , the stratigraphic developmen t i s slightl y different . Here, the top Kimmeridgia n is truncated b y the BCE immediately wes t of the BCE/to p Heathe r truncation lin e (Fig . 13) . However , th e to p Kimmeridgian i s als o truncate d b y th e BC E some further 1- 2 km to the west, along an intrabasinal faul t tha t wa s likel y activate d i n th e Volgian. Volgian-Ryazanian strat a (Seqs-II-V ; Fig. 9) are banked against the northern segments of th e Statfjor d Eas t Faul t an d ar e largel y bounded t o th e eas t b y th e intra-basina l fault . Patches o f assume d Lat e Volgian/Ryazania n sandstones (Seqs-III-V , Fig . 9 ) are present als o eastward o f th e intra-basina l faul t (Fig . 6a-c) , but thes e strat a see m t o hav e bee n late r partl y uplifted and eroded . In summary , the obser vations indicat e a continuousl y expandin g footwall islan d sinc e th e Oxfordian . Footwal l contraction in thi s are a was likel y limite d to the Lat e Ryazanian , i n respons e t o termina tion o f riftin g an d therma l subsidenc e o f th e rift terrain . The shorelin e trajector y o n th e Snorre- H hangingwall shifte d basinwar d a t a lo w angl e through mos t o f th e syn-rif t stag e (Fig . 16) , representing a n 'accretionar y force d regression ' in th e classificatio n b y Helland-Hansen & Mar tinsen (1996) . Th e accretionar y force d regres sion typicall y result s fro m abundan t sedimen t supply (i.e . footwal l erosion ) superimpose d on a fallin g relativ e sea-leve l (i.e . footwal l expansion). Th e turnaroun d an d revers e move ment o f th e shorelin e trajector y i n th e ver y late stage s o f rif t evolutio n represent s a 'non -

accretionary transgression' , resultin g fro m a high rat e o f subsidenc e relativ e t o rat e o f sedi ment supply . The modifie d domino/variable rate s o f fault ing mode l discusse d b y Robert s e t al. (1993 ) predicts a n increasin g widt h o f footwall island s as a result of successive fault bloc k rotation , bu t not o f th e magnitud e see n here . I t i s believed, therefore, tha t uplif t an d expansio n o f th e Snorre-H footwal l islan d becam e amplifie d b y sub-regional uplif t cause d b y th e buoyantl y rising mai n Snorr e Faul t Bloc k an d Tampe n Spur t o th e north . Footwal l islan d expansion also wa s amplifie d b y a n overal l eustati c sea level fal l fro m th e Lat e Kimmeridgia n throug h Ryazanian (Halla m 1969 ; Ha q e t al . 1988 ; Surlyk 1989) . A genera l sea-leve l ris e fro m Bathonian t o Earl y Kimmeridgia n probabl y reduced th e effec t o f footwall uplift an d genera tion o f subaeria l relie f durin g th e earl y stage s of rifting . In the domino/variable rates of faulting model progressive increas e i n rate s o f faultin g pro duces continuous offlap , a s seen in th e norther n Snorre-H area . An offlap/onla p pattern , a s seen in th e centra l an d souther n Snorre- H area , results fro m increased , followe d b y reduced , rates o f faulting . Thi s variabilit y i n styl e o f faulting i s coheren t wit h th e note d northwar d propagation b y segment linkage of the Statfjor d East Faul t (Dawer s e t al . 1999 ; Dawer s & Underhill in press), leaving the southern portio n of th e Statfjor d Eas t Fault , thoug h apparentl y quite activ e fro m th e Bathonia n throug h Kim meridgian, tectonicall y inactive and drowne d by Late Volgia n times . A t thi s time , faul t move ments in the Snorre-H are a localize d around th e main Statfjor d Faul t furthe r to th e west. The tota l thro w o n th e Statfjor d Eas t Faul t next to well 34/724S is roughly 1 km correspond ing t o a faul t displacemen t rat e o f 0.0 5 mm/y over th e Mid-Bathonian-Mid-Volgia n interval . By comparison , measurement s o f recen t tiltin g of fault blocks in Greece (Gawthorpe e t al. 1994) show displacemen t rate s u p t o 5 mm/y, o r tw o orders o f magnitude higher . Thi s indicate s tha t North Se a stretching wa s generally slow . Alter natively, riftin g ma y hav e bee n considerabl y faster i n shorte r periods , correspondin g t o a n overall pulsed mode o f rifting . Cumulative uplift and erosional yield In clastic syn-rift terraines , usually local erosio n follows a s a n immediat e respons e t o footwal l uplift. Du e t o th e unlithifie d natur e o f th e sediments involve d in rotatio n of the Snorr e

SYN-RIFT EVOLUTIO N I N TH E SNORRE- H AREA and Statfjor d Eas t Faul t Blocks , suc h erosio n expectedly woul d be very fast an d efficient . A s a simplification, th e erosiona l yiel d ca n b e considered proportiona l t o the vertical relief create d and th e total are a o f uplift . Based o n th e estimate d cumulativ e til t an d temporal an d spatia l positio n o f th e shoreline / basement emergenc e point , theoretica l footwal l uplift an d hangingwal l subsidenc e acros s th e central Snorre- H Faul t Bloc k hav e bee n calcu lated (Tabl e 3) . A s th e Snorre- H Faul t Bloc k dips becom e significantl y reduce d toward s th e footwall cres t (Soll i 1995 ; his Fig . 2), however, footwall uplif t estimate s base d o n averag e til t angles i n th e Snorre- H sub-basi n wil l b e to o high. Estimate s o f hangingwal l subsidence , o n the other hand , ar e considered mor e representa tive. I n orde r t o correc t th e uplif t model , estimated uplif t ha s bee n compare d wit h th e current dept h o f Late Jurassic erosion along the Snorre-H footwal l crest. A cumulativ e til t o f 9 C by th e en d o f Ryazania n give s approximatel y 1100m o f cumulativ e uplift . Estimate s b y Soll i (1995) suggest between 500-1000m of erosion of this part o f the Snorr e Faul t Block , likel y close to 700 m du e east o f well 34/7-24S. This i s some 400m o r 40% , les s tha n th e estimate d uplif t using 9 ° o f tilt . A s th e Snorre- H footwal l scar p erosional bas e wa s a t sea-leve l th e observe d depth of erosion likely corresponds closely to the total uplift , suggestin g tha t al l estimate d uplif t numbers b e correcte d accordingly . Correcte d uplift estimate s sugges t som e 6 0 m o f uplif t b y Mid-Oxfordian increasin g to about 70 0 m by the end-Ryazanian (Tabl e 3) . Hangingwal l subsi dence increase d less , fro m 400 m b y Mid Oxfordian t o some 800 m by the end-Ryazanian, reflecting th e rapidl y expandin g footwal l islan d width. Th e Snorre- H di p variation s hav e no t been measure d an d integrate d ove r th e entir e fault bloc k width , but, as a sensitivity , a cumu lative uplif t o f abou t 700 m correspond s t o a Late Ryazania n tilt angl e averaging 6°.

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To tes t th e uplif t model , th e sedimentatio n rate i n the Snorre-H hangingwal l basin ha s been investigated (Fig . 15). The amoun t o f reworke d bioflora i n th e syn-rif t successio n (Fig . 8 ) ha s been use d a s a measure o f actua l footwal l relief generated an d loca l erosiona l yield . Th e well s 34/7-23 an d -24 S i n th e Snorre- H sub-basi n clearly hav e a n exponentia l reductio n i n sedi mentation rat e through time , although th e effec t of variations in sediment suppl y cannot be easily separated fro m variation s i n generatio n o f ac commodation space . This observation may seem contradictory t o th e til t evolutio n and footwal l uplift model , whic h suggest s overal l increasing footwall uplif t an d generatio n o f relief throughout th e sam e tim e period. I t i s suggested, however, that the data be explained in terms of initial high rates of sedimentation from a predominant regional, extra-basina l sedimen t suppl y i n th e Bathonian-Oxfordian, durin g deposition o f th e Heather Formation . Th e scarcit y o f reworke d taxa befor e th e Lat e Oxfordia n support s th e notion of a pre-dominant extra-basinal sediment supply. A reduced rate of sedimentation fro m th e Kimmeridgian onward s likel y reflect s a shutof f of the extra-basinal sedimen t sourc e an d transi tion to erosion and supply of sediment from loca l footwall highs . Thi s interpretatio n i s supporte d by th e observe d increasin g number of reworked taxa i n this part o f the succession. It should be noted tha t the observed reduction in reworke d tax a i n th e Mid-Volgia n (Fig . 9; transition Seqs-II/III ) doe s no t fi t wit h a continuously expandin g footwal l island o r uplif t o f the magnitud e estimate d above . Thi s obser vation, however , i s regarde d erroneou s an d not representativ e fo r th e Mid-Volgia n uplif t and erosiona l yield . Th e Lat e Volgia n an d Ryazanian Uppe r Draupn e Sandston e lobe s (Seqs-III-V) contai n abundan t reworke d taxa , indicating significan t loca l erosio n i n th e Ryazanian. Th e reductio n i n sedimentatio n rat e observed fo r th e uppe r Draupn e Formatio n i s

Table 3 . Estimates o f footwall uplift an d hangingwall subsidence o f th e Snorre-H Fault Block, along a n E-W transect near well 34J7-24S Time

Measured Footwal l Cumulativ tilt-angle islan d hangingwal width subsidenc

Mid-Oxfordian End Kimmeridgia n End Ryazania n

3C 5C 9

2-4 km 7km 7 km

e Cumulativ l footwal e uncorrecte

400-600 m 450m 800 m

e Footwal l uplif t l uplif t correcte d d (40 % reduction)

1 00-200 m 600m 1100m

60- 120m 350m 700m

In orde r t o accommodat e a decreasin g fault-bloc k til t fro m th e hangingwall acros s th e footwall , initia l footwall uplift estimate s have been corrected by applying a fixed 40% reduction.Not e that a fixed tilt angle of 5 C has bee n applied fo r to p Kimmeridgian , whereas th e estimated til t angl e i s between 4-5 c.

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believed to reflec t bevellin g of the footwall highs and reduce d erosiona l yiel d durin g th e wanin g rift stage . Tectonostratigraphic evolution In orde r t o summariz e th e tectonostratigraphi c evolution of the Snorre-H area , each of the main sequences described abov e has bee n discussed in terms o f seismi c character , structura l context , depositional facie s and palaeogeographic settin g (Fig. 17) . Pre-Mid-Bathonian The pre-Mid-Bathonian successio n has not been included i n this study. Nevertheless, durin g Lat e Bajocian-Early Bathonia n (= Tarber t Forma tion) th e Snorre- H are a wa s likely governed b y differential subsidenc e acros s th e mai n bloc k bounding fault s (Fig . 17a) . Ou r understandin g of th e stratigraphi c developmen t o f the Tarber t Formation i s hampered b y the lack of well data , but emergin g evidenc e fro m th e Tampe n Spu r area (Johannese n e t al 1995 ; Davies et al. 1998) suggest tha t i t ha s a wea k wedg e shape d geo metry an d wa s deposite d i n respons e t o inci pient rotationa l movement s alon g th e synrif t master faults . Mid-Bathonian-Early Oxfordian (Heather Fm) Mid-Bathonian-Early Oxfordia n subsidenc e (= Heather Formation ) wa s concentrated alon g NNW-SSE strikin g maste r fault s (Fig . 3) . Th e seismic dat a doe s no t allo w a detaile d inter pretation o f th e Heathe r Formatio n interna l reflectors an d i t i s no t possibl e t o determin e i f rates o f faul t bloc k rotatio n di d var y ove r thi s interval. Overal l rate s o f rotatio n wer e signifi cant, however , reachin g a til t o f c . 3° b y th e Early Oxfordian (Fig . 15) . This was sufficient t o create a wedge shaped geometr y of the Heathe r Formation an d a pronounced depocente r developed alon g th e souther n segmen t o f th e Statfjord Eas t Faul t (Dawer s e t al 1999) . It i s presentl y unclea r t o wha t exten t th e Heather Formatio n wa s deposite d acros s th e southern Snorr e Faul t Bloc k durin g Mid Bathonian-Early Oxfordia n tim e (fig . 17b) , a s demonstrated o n th e Statfjor d an d Visun d Fault Block s (Robert s e t a l 1993 ; Fasrset h et a l 1995) . Th e scarcit y o f reworke d bioflor a before the Late Oxfordian suggests that footwal l crestal island s wer e poorl y develope d a t thi s stage. I t shoul d b e noted als o tha t n o san d ha s been detecte d withi n the Heathe r Formatio n i n the Snorre- H area , whic h ma y indicat e tha t a t least th e lower part th e Heather Formatio n wa s

deposited acros s th e footwall crestal area s o f the southern par t o f the Snorre Faul t Block . By Mid-Oxfordia n time s th e Snorre- H hang ingwall shorelin e an d footwal l morphologica l apex (= palaeodrainage divide ) wer e located fa r to the east on the fault bloc k (Fig . 16) . The main sediment sourc e for the Heather Formation was likely extra-basinal. It is suggested, however, that forced regressio n o n th e Snorre- H hangingwal l dip-slope during deposition o f the upper Heathe r Formation le d t o erosio n an d cannibalisatio n of th e lowe r Heathe r Formation . Reflectin g the infan t stag e o f uplift , erosio n o f th e Bren t Group was confined to th e footwall sides o f the Snorre an d Statfjor d Eas t Faul t Blocks , a s suggested b y th e presenc e o f reworke d Bren t Group tax a i n th e Callovia n successio n i n well 34/7-24S (Fig. 8) . Figure 18 a show s a schemati c palaeogeo graphic reconstructio n o f th e large r Snorr e area durin g Mid-Callovia n times . A s note d above, th e Snorre- H an d Statfjor d Eas t foot walls wer e likel y submerged , wherea s a sourc e upland i s believed t o hav e existe d t o th e north , as a result of progressive southward tilting of the Tampen Spur/Snorr e Faul t Block . Late Early Oxfordian-Kimmeridgian (Draupne Fm; Seq I) Continued rotatio n durin g the late Early Oxfordian-Kimmeridgian (Blowe r Draupn e Forma tion; Seq-I) increased th e tilt to some 5° (Fig. 15) . Subsidence stil l wa s concentrate d alon g th e NNW-SSE strikin g maste r faults , leadin g t o a gradual thickenin g of th e lowe r Draupn e For mation, Seq- I ( = lower Draupn e seismi c unit ; Draupne shale) , toward s th e Statfjor d Eas t Fault, wherea s i t thin s mor e abruptl y ont o th e Snorre-H hangingwall . Small basinward dippin g faults occu r alon g the top o f the lower Draupn e Formation o n th e hangingwal l dip-slope , bu t these ar e tentativel y suggeste d t o b e post depositional and relate d t o subsequen t tilting. During depositio n o f th e lowe r Draupn e Formation, Se q I , footwal l cresta l island s ha d become firml y established , leadin g t o increase d rift-basin segmentatio n (Fig. 17c) . Consequently, axial an d transvers e sedimen t suppl y di d n o longer fee d directl y into th e oute r half-grabens . This i s supported b y the apparen t condensatio n or hiatus in the Mid-Oxfordian (Fig. 7) , suggesting a perio d o f widesprea d starvatio n a t thi s time. The high gamma-ray lo g responses seen in the Kimmeridgia n interva l (Fig . 9 ) furthe r indicate basi n deepenin g an d restricte d circulation. As the entire interval is less than on e cycle thick ther e i s n o seismi c resolutio n withi n th e time interval , bu t th e seismi c dat a indicate s a

Fig. 17 . Schemati c illustratio n of th e tectonostratigraphi c evolutio n of th e Snorre- H Faul t Block .

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Fig. 18 . Schemati c illustratio n o f th e large r Snorr e are a palaeogeographi c evolution , (a ) Callovian . (b) Composite Lat e Volgian-Ryazanian .

weak onlap of the lower Draupne Formatio n an d top Kimmeridgia n surfac e ont o th e Heathe r Formation o n th e Snorre- H hangingwal l (Fig . 13). Ther e i s also a pronounce d updi p thinnin g from well s 34/7-24S to 34/7-21A and -21 (Figs 6e, 9), an d a slightl y angular , unconformabl e rela tionship betwee n th e lower Draupn e Formation and th e underlying Heather Formatio n locall y is observed i n upflan k position s o n th e Snorr e hangingwall. Thi s ma y indicate a pulse o f rapi d uplift followin g deposition o f th e Heathe r For mation, wit h subsequen t intensificatio n of foot wall erosion an d loca l sedimen t yield. By th e en d o f Kimmeridgia n th e Snorre- H hangingwall shoreline , a s wel l a s th e foot wall erosional apex, had bee n shifte d westwards, in respons e t o footwal l uplif t an d expan sion an d accompanyin g fault-scar p degradatio n and retrea t (Fig . 16) . Depositio n o f th e lowe r Draupne Formation , Seq-I , wa s source d b y significant loca l erosio n o f th e Heathe r Forma tion an d t o som e degre e th e Bren t Group ,

particularly fro m th e Statfjor d Eas t footwall . The presenc e o f blu e cla y fragment s (probabl y Heather clay ) in the dark laminated mu d is likely derived fro m neighbourin g footwal l islands , although a n efficien t drainag e networ k ha d no t yet bee n established . Alternatively , th e cla y particles ma y ste m fro m failur e an d resedimen tation o n th e destabilize d submarin e slope . B y end-Kimmeridgian erosio n o n th e hangingwal l had cu t int o th e Bren t Group , suggestin g tha t incipient shorelin e san d system s may hav e bee n fringing th e Snorr e hangingwal l a t thi s tim e (Fig. 16) . As a result of continued footwall uplift and expansio n int o th e Volgian , however , the y had zer o preservatio n potentia l an d late r became cannibalized. Early-early Mid-Volgian (Draupne Fm; Seq-H) In th e Earl y Volgian-earl y Mid-Volgia n (= upper Draupn e Formation ; Seq-II) , subsi dence alon g th e NNW-SS E strikin g maste r faults becam e accompanie d b y second-orde r

SYN-RIFT EVOLUTIO N I N TH E SNORRE- H AREA N-S, NE-S W an d E- W faulting , leadin g t o segmentation o f th e large r faul t block s (Fig s 3 & 6) . Th e seismi c dat a doe s no t provid e suffi cient resolutio n t o properl y ma p ou t th e uppe r Draupne Formation , but , generally, thicker synrift package s appea r t o b e presen t i n th e bathymetric low s (Fig. 6) . During depositio n o f th e uppe r Draupn e Formation, Seq-II , continue d hig h rate s o f fault bloc k rotatio n an d associate d uplif t le d t o increased footwal l islan d relie f an d erosio n (Fig. 17c) . At thi s time, rotational faultin g alon g the Statfjor d Eas t Faul t ha d propagate d t o the nort h o f th e centra l Snorre- H area . Th e Snorre-H sub-basi n evolve d int o a northwar d closed an d southwar d deepenin g embayment , separated fro m th e Statfjor d Eas t sub-basi n t o the west . Eas t o f th e Snorr e Faul t Block , th e Visund sub-basi n wa s close d t o th e sout h an d not connecte d t o th e Snorre- H sub-basin . Thi s suggests that eac h o f the sub-basin s by now wa s sourced predominantl y fro m erosio n o f th e adjacent footwal l crestal islands . As th e rat e o f sediment inpu t wa s no t sufficien t t o kee p pac e with th e fault-controlle d subsidence, thi s le d t o starved, under-fille d condition s an d increase d bathymetry i n th e half-graben . A pronounce d onlappin g relationshi p o f th e upper Draupn e seismi c uni t ( = upper Draupn e Formation) onto the lower Draupne seismic unit and sub-paralle l passiv e infil l characte r o f th e upper Draupn e seismi c uni t (Fig s 6d , f, lOc , d) may indicat e tha t ther e wa s a puls e o f rapi d rotation an d uplif t followin g depositio n o f th e lower Draupn e Formation . A s a resul t o f th e onlapping relationshi p (Fig . 6d , f) thicke r syn rift package s are present i n the bathymetric lows and a depositiona l hiatu s ma y b e speculate d a t the Kimmeridgia n boundar y betwee n Seq- I an d Seq-II i n updip area s (Fig . lOc , d). The biostrati graphical dat a show s an overal l compression o f the Lat e Kimmeridgian-Earl y Volgia n interva l in the 34/7-21 well, but no apparent stratigraphi c break ha s bee n observe d (Fig s 7 & 9). This ma y be a result of lack of stratigraphic resolution an d deposition i n the bathymetric lo w occurring ove r a relativel y shor t tim e perio d accompanyin g the increase d rotatio n an d uplift . Th e seismi c onlap o f th e uppe r Draupn e Formatio n i s observed i n a positio n immediatel y adjacent t o the Statfjor d Eas t Fault , suggestin g tha t th e deeper interna l reflector s relate t o processe s o f fault-scarp degradation . By earl y mid-Volgia n th e Snorre- H hanging wall shorelin e an d footwal l erosiona l ape x had shifte d furthe r westwar d relativ e t o th e end Kimmeridgian , indicating continued force d regression a s wel l a s increasingl y deepe r ero -

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sion o n th e Snorre- H hangingwal l durin g th e Early Volgian (Fig. 16) . Deposition o f the uppe r Draupne Formation , Seq-II , wa s source d pre dominantly by erosion o f the Heather an d Bren t Formations and , t o som e degree , th e lowe r Draupne Formation , a s evidence d b y th e re worked taxa present (Fig. 8). The Intra Draupn e Sandstone is believed to hav e been source d fro m erosion o f Kimmeridgia n an d Earl y Volgia n basin steppin g shoreline deposits, in response t o the note d force d regression. The seismi c character , sedimentar y facie s and sequenc e architectur e o f th e lower/uppe r Draupne Formatio n Seqs-I-I I ma y sugges t a pulsed mod e o f riftin g an d genera l rif t cli max durin g th e Lat e Oxfordian-Mid-Volgian . It shoul d b e emphasized , however , tha t varia tions i n faul t displacemen t rate s alon g a faul t system may promote 'rift climax ' stratal pattern s near activ e faul t segmen t centre s contempora neously wit h 'early/lat e rift ' strata l pattern s a t fault tip s and i n overlap zones between unlinke d faults (Dawer s & Underbil l i n press) . Th e neighbouring Visun d Faul t Bloc k t o th e eas t also ha d maximu m stretchin g an d footwal l uplift i n th e Volgia n (Faerset h e l al. 1995) , whereas th e Brent/Statfjor d Faul t Bloc k t o th e SW ha d maximu m rate s o f rotatio n i n th e Oxfordian (Robert s e l al . 1993 ; Nottvedt e l a l 1995). I n a regiona l perspective , thi s ma y poin t to a spatia l an d tempora l partitionin g of faul t displacement i n the Snorre- H area . Late Mid- Volgian-Ryazanian (Draupne Fm; Seqs III—IV) During lat e Mid-Volgian-Ryazania n ( = uppe r Draupne Formation , Seqs-III-V ) th e Snorre- H area experience d wanin g rate s o f faul t bloc k rotation an d increasingl y mor e unifor m sub sidence, leadin g to a n estimate d til t of 9-13 b y Late Ryazania n (Fig . 15) . I n respons e t o continued intra-bloc k deformation , however , the Snorre- H are a becam e compartmentalize d and subjec t t o differentia l subsidenc e b y N- S as wel l a s E- W faults , an d secondar y half graben structure s develope d (Fig . 6a~c) . Foot wall cresta l uplif t alon g som e o f th e fault s le d to th e emergenc e o f minor , intra-basina l foot wall islands. Deposition o f the upper Draupn e Formation , Seq-III-V, wa s governe d b y reduce d rate s o f fault bloc k rotatio n an d uplift , leadin g t o sig nificant reductio n i n ne w topographi c relie f generated. A s a resul t o f continue d westwar d shift o f th e Snorre- H hangingwal l shoreline and footwall erosiona l ape x (Fig . 16) , erosio n no w cut dee p int o th e Bren t Grou p o n th e hangingwall dip-slope , releasin g substantial volumes o f

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sand int o th e Snorre- H sub-basin . I n combina tion with reduced rate s of subsidence, this led to progradation an d accumulatio n o f th e Uppe r Draupne Sandston e acros s th e easter n an d northern Snorre- H are a (Fig . 13) . Th e overal l shallowing-upward natur e of the upper Draupn e Formation, Seqs-III-IV , a s see n i n th e well s in th e centra l Snorre- H sub-basi n (Fig . 9) , and th e fac t tha t i t fills across th e entir e basin in th e well s 34/7-2 3 an d 23 S area , suggest s reduced faul t relate d subsidenc e and generatio n of ne w accommodatio n spac e afte r depositio n of Seq-I L A majo r depocente r develope d alon g th e Statfjord Eas t Faul t t o th e sout h o f wel l 34/724S. Th e depocente r represent s largel y (axial) infilling o f pre-existing bathymetry, however, as the drapin g o f th e souther n portio n o f th e Statfjord Eas t Faul t b y th e uppe r Draupn e seismic unit indicate s i t had becom e tectonicall y inactive b y Lat e Volgia n times . Th e uppe r Draupne seismi c uni t i s oversteppin g th e lowe r Draupne seismi c uni t (to p Kimmeridgia n sur face) an d onlappin g als o ont o th e souther n hangingwall o f th e Snorr e Faul t Block , aroun d wells 34/7-2 1 an d 21 A (Fig s 6e , 13) , suggest ing overal l transgressio n i n thi s area . A s th e Draupne Formatio n oversteppe d th e Statfjor d East Faul t Block , however, th e suppl y o f san d from th e S W was significantl y reduced . The seismi c as wel l a s biostratigraphica l dat a suggest tha t ther e i s a youngin g o f sandston e depositional units towards the north (Figs 9, 13). The sandstone s are believe d to consis t o f minor regressive increment s tha t prograde d int o th e basin durin g regional , bu t punctuated , trans gression. Th e clos e associatio n betwee n sand stone occurrenc e an d structura l positio n le d Dawers e t al (1999 ) t o sugges t tha t th e dis tribution o f thicker shoreline sandstones may be partly controlle d b y th e northward s fault-ti p propagation o f th e Statfjor d Eas t Fault . I t i s important t o emphasise , however , tha t faul t displacement durin g Lat e Volgian-Ryazania n were less compared t o those obtained durin g the Late Oxfordian-Earl y Volgian . A composit e palaeogeographi c reconstruc tion o f the Lat e Volgian-Ryazanian i s shown in Fig. 18b . Shorefac e an d beac h clif f erosio n together wit h fault-scar p retrea t fro m gravit y collapse probabl y contribute d mos t t o th e denudation o f uplifte d faul t bloc k crests . As loca l uplan d area s wer e to o smal l t o allo w larger rive r system s t o develo p fluvia l incisio n most likel y wa s les s important . Sand s wer e deposited i n a progradin g shorelin e comple x along th e ri m o f th e emergen t faul t block s an d filled across the entire embaymen t t o th e north .

The lack of evidences for significant rotational faulting durin g th e Ryazania n suggest s tha t subsidence a t thi s tim e ma y hav e bee n (partly ) related t o therma l relaxatio n an d th e rif t structure enterin g int o a post-rif t stat e o f development. A s a corollary , th e neighbourin g Visund Faul t Bloc k ha s th e syn - t o post-rif t transition documente d a s lates t Volgia n (Faer seth e t al . 1995). Latest Ryazanian-post-Ryazanian From th e lates t Ryazania n onward s th e Snorre-H sub-basi n subside d i n respons e t o post-rift therma l relaxatio n an d post-deposi tional compaction . Followin g depositio n o f th e Draupne Formation , Hauterivian/Valanginia n Cromer Knol l Grou p limestone s drapin g th e syn-rift successio n an d rif t terrai n sugges t a relative sea-leve l rise , leadin g t o progressiv e drowning o f th e inheren t footwal l relie f an d shutofT o f th e loca l a s wel l a s extra-basina l clastic suppl y t o th e half-grabens . Th e relativ e sea-level ris e i s believe d t o hav e bee n accentu ated b y a n earl y Cretaceou s eustati c sea-level rise (Ha q e t al . 1988 , Surly k 1989) . It shoul d b e note d tha t regiona l seismic lines indicate a n onlappin g relationshi p o f lowe r Cretaceous strat a ont o th e top Draupne surface. The lowe r Cretaceou s successio n i s believe d t o represent primaril y passiv e infil l o f pre-exist ing bathymetry . B y th e lates t Ryazania n th e Statfjord-Statfjord Eas t an d Snorr e footwal l islands wer e largely levelled by erosion, and th e subaerial relie f remainin g was therefor e insuffi cient t o provid e for an y significan t clasti c input to th e Snorre- H area . Erosion and redistribution of sand The bas e Cretaceou s subcro p ma p outline s roughly area s o f Lat e Jurassi c erosio n vs . deposition. Th e boundar y betwee n the tw o was very dynamic , however, an d shifte d laterall y in response t o progressiv e faul t bloc k rotatio n (cf. discussio n above) . B y superimposin g th e apex lines of the base Cretaceous structura l time map t o delineate the palaeodrainage divides, the area ca n b e subdivide d int o separat e sediment source basin s an d correspondin g depositiona l basins (Fig. 19a) . By further combinin g the base Cretaceous an d base Draupne subcro p maps , an estimate o f erosio n o f previousl y deposite d sandy formation s tha t provide d a sourc e o f sand t o th e Snorre- H sub-basi n ca n b e mad e (Fig. 19b) . Thi s approac h i s base d o n th e assumption tha t ther e ha s bee n n o tiltin g o f the are a sinc e Lat e Ryazanian , however, which is probably untrue . Post-rif t therma l subsidenc e

SYN-RIFT EVOLUTIO N I N TH E SNORRE- H AREA commonly lead s t o rotatio n o f th e entir e rif t stratigraphy toward s the rift axi s (N0ttvedt e t al. 1995) an d th e presen t N plung e o f th e Snorr e footwall ma y yield velocity effects tha t may als o shift th e faul t bloc k apex . An y o f thes e effect s are considere d minor , however , relativ e t o th e scale o f the estimate s presente d here . As show n i n Fig . 19b , th e Bren t Grou p i s partly t o completel y erode d bot h t o th e west , north an d eas t o f the Snorre-H area. A s a result of th e southerl y structura l plunge , th e Tampe n Spur has the largest uplift and deepest erosion to the north , exposin g Middl e Jurassic-Triassi c lithologies i n thi s area . Uppe r Jurassi c syn-rif t sands i n th e Snorre- H sub-basi n wer e source d almost exclusivel y fro m erosio n o f th e Bren t Group, however , a s withi n th e loca l drainag e

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area erosio n di d no t cu t dee p enoug h t o encounter othe r sand y formations . Sediments wer e disperse d int o th e Snorre- H sub-basin bot h fro m th e nort h an d eas t dow n the Snorr e hangingwal l dip-slop e an d t o a minor extent from th e Statfjor d Eas t faul t scar p to th e west . Durin g infan t stage s o f erosion , the footwal l erosiona l ape x ( = palaeodrainage divide) wa s shifte d close r toward s th e foot wall crest s tha n th e presen t bas e Cretaceou s structural apex , simpl y a s a resul t o f th e back ward retreatin g natur e o f faul t scar p ero sion. This means that a somewhat large r portio n of th e erode d materia l wa s likel y transporte d down th e hangingwall s compare d t o th e footwalls o f th e faul t blocks . I n performin g the Snorre- H sub-basi n materia l balancing ,

Fig. 19. Structura l map of the base Cretaceous in the eastern Tampen Spur area, (a) Subcrop and areas of erosion of th e Bren t Group . Note also structural map ape x an d basi n axis trajectories, (b ) Upland sediment source areas, transport routes and depositional. The structural apex trajectories have been used as basis for constructing the palaeodrainage divides .

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however, underestimatio n o f material dow n th e Snorre-H hangingwal l i s considere d largel y compensated fo r b y overestimatio n o f materia l down th e Statfjor d Eas t footwall . In a regiona l perspective, larg e volume s o f erode d materia l were shed also from th e Snorre Faul t Bloc k into to th e Visun d sub-basi n an d int o th e Marul k basin t o th e NW. Biostratigraphical reworkin g dat a suggest s that th e Bren t Grou p withi n th e establishe d drainage domai n becam e expose d t o erosio n from Lat e Oxfordian times. Accordingly, eroded Brent Grou p sediment s ma y potentiall y b e distributed withi n th e complet e Draupn e For mation thicknes s interval . A s i t becam e over stepped an d drowne d b y Lat e Volgia n times , material erode d of f th e Statfjor d Eas t footwal l implicitly must be distributed largely in the MidOxfordian-Early Volgia n successio n (Fig . 16 ; Seqs-I-II). Effectiv e hangingwal l erosio n an d sediment transpor t di d no t commenc e unti l the Late Kimmeridgia n (Fig . 8) , however , suggest ing that most o f the Brent Group eroded of f the Snorre hangingwal l dip-slop e wa s removed an d dispersed toward s th e WS W durin g Volgian Ryazanian, and may be concentrated withi n this time interval (Fig. 16 ; Seqs-III-V). Mass budget In th e Snorre- H sub-basi n sedimen t sourc e area ( = Snorre-H -f Statfjor d Eas t footwall) , the Brent Group is completely eroded across an area corresponding t o roughly 10km 2 (Fig. 19b) . The thickness o f th e Bren t Grou p i n th e Snorre- H sub-basin i s c . 200 m givin g a n erode d Bren t Group volum e of 2km 3 . I n addition , th e Bren t Group i s partl y erode d acros s approximatel y 40km 2 of the sedimen t sourc e area . Usin g an average thicknes s o f 100 m o r hal f th e norma l thickness of the Brent Group, the eroded volume corresponds t o abou t 4km 3 . Thi s give s a tota l volume o f erode d Bren t Grou p o f 6km 3 . Applying a N/ G valu e o f 0.5 , c . 3km3 o f coarse clasti c materia l (sand ) wa s erode d an d supplied fro m th e Snorre- H sedimen t sourc e area t o th e Snorre- H sub-basi n (Tabl e 4) . Th e main Statfjord footwal l has not bee n included in the abov e estimate , a s i t probabl y di d no t contribute significantl y to th e Snorre- H deposi tional basi n sedimen t supply. Volumes of mapped anomalie s ar e calculate d as follows : (i) Th e aeria l exten t o f th e Snorre- H centra l and flankin g seismi c anomalie s (shoreline / shallow marin e sandstones) , ha s bee n esti mated to approximately 14km 2 (Tables 3, 4 and Fig . 13 ; A1-A6). Using sand thickness

numbers a s give n i n Tabl e 4 a tota l san d volume of 0.5km 3 has bee n calculated, (ii) Th e seismi c anomal y t o th e sout h (basi n floor sandstones) has a mapped area l extent of abou t 35km 2 (Fig s 1 2 an d 13 ; A7) . It may, however, continue somewhat south of th e presen t surve y area . B y usin g th e mapped are a an d a seismi c thicknes s o f 44ms (62m) , a gros s san d volum e o f 0.21 km3 ha s bee n estimated . The volumetric calculations show that the two mapped seismi c anomalie s adde d u p ma y con tain about 0.7km 3 of sand, or about one quarter of the eroded volume . Doubling th e san d thick ness i n al l th e mappe d seismi c anomalies raise s the deposited san d volume only to 1.4km 3, o r a little les s tha n hal f o f th e erode d volume . This suggests tha t a substantia l volum e o f san d supplied fro m th e Snorre- H drainag e domai n to th e Snorre- H sub-basi n ha s no t ye t bee n identified. Som e o f th e exces s san d ma y hav e been deposite d a s footwal l talu s sand s of f th e Statfjord Eas t footwal l an d som e ma y b e distributed in intra Draupne sand accumulations so far undetected. It is further speculate d if some of the sand delivere d t o th e Snorre- H sub-basi n may have bypassed an d becam e deposited i n the half-graben t o th e south.

Syn-rift play models Several differen t pla y model s occu r i n associa tion wit h late Jurassic extension and rotatio n of the Snorr e Faul t Bloc k (Fig . 20) . The y ar e al l intimately linked to erosion int o th e sand pron e Brent Grou p an d subsequen t depositio n o f pinch-out sandston e reservoir s withi n Draupn e Formation shales . Th e envelopin g uppe r Jur assic Draupn e Formatio n constitute s the mai n source roc k i n th e are a an d buria l depth s ar e within th e rang e o f th e oi l window , allowin g generation an d migratio n o f hydrocarbon s int o structural/stratigraphic pinch-ou t traps . Below , various reservoi r facie s models an d distributio n of syn-rif t reservoir s i n th e Snorre- H are a ar e discussed. (1) Shallow marine I shoreline sandstones Upper Draupn e shallo w marine/shorelin e sand stones hav e bee n drille d i n th e hangingwal l of the Snorr e Faul t Block , i n th e previousl y described well s 34/7-2 1 an d -2 3 (Fig . 9) . Th e sandstones fring e th e Snorre- H sub-basi n an d their presenc e depen d o n loca l suppl y o f san d from erosio n int o th e Bren t Group , mainl y o n the Snorr e hangingwall. Good qualit y reservoir

SYN-RIFT EVOLUTIO N I N TH E SNORRE- H AREA

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Table 4. Comparison of eroded sand in the sediment source area and volumes of sand deposited i n the mapped seismic anomalies in the Snorre-H area Source area

Erosional area

Net san d (50%)

Eroded san d volume

Snorre-hw; B G completel y erode d Snorre-hw + Statfj.E-fw; B G ptly eroded Total

10km2 40km 2

100m 50m

1.0km 3 2.0km 3 3.0km 3

Depositional basi n

Depositional area

Reservoir thicknes s (assuming al l sand )

Deposited san d volume

Snorre-H centra l are a (Al ) Snorre-H centra l are a (A2 ) Snorre-H centra l are a (A3* ) Snorre-H centra l are a (A4* ) Snorre-H central are a (A5** ) Snorre-H centra l are a (A6 ) Snorre-H structura l lo w (A7***) Total

0.51km 2 5.10km 2 0.34km2 0.18km2 0.47 km2 7.21km 2 3.46km2

30m 30m 25/2 msec x 2. 1 3 m/msec ( = 2 7 m) 20/2 msec x 2. 13 m/msec ( = 2 1 m) 28/2 msec x 2.80 m/msec (=39m) 40m 44/2 msec x 2. 80 m/msec (= 62m )

0.01530km3 0.15300km3 0.00905km3 0.00383km3 0.01842km3 0.28840km3 0.21314km3 0.7km 3

Snorre-hw, Snorr e hangingwall ; Statfjor d E-fw , Statfjor d Eas t footwall ; BG , Bren t Group . (* ) Use d interva l velocity of 2.13 m/msec from 34/7-23 wells. (**) Used interva l velocity of 2.8 m/msec from 34/7-2 1 wells. (***) No well informatio n nea r thi s area; used 2. 8 m/msec interva l velocity .

sandstones occu r i n a bel t alon g th e easter n margin of the Snorre half-graben, an d in a more axial positio n t o th e nort h o f wel l 34/7-24 . However, suc h syn-rif t sandstone s hav e bee n penetrated mostl y i n updi p position s an d reservoir sandstone s ma y thicke n considerabl y down th e hangingwal l o f th e Snorr e Faul t Block. The Uppe r Draupn e Sandston e (Seqs-III-V ) was deposited durin g the late rift stage , when the rate o f faul t bloc k rotatio n wa s significantl y reduced. Thi s enable d san d t o prograd e acros s and fills the northerl y embaymen t o f the south erly plungin g Snorre- H half-graben . Earlie r syn-rift shorelin e sandstone s (Seq-I-II ) wer e probably cannibalize d completel y durin g th e forced regressiv e evolutio n o f th e Snorre- H hangingwall. I f locall y preserved , however , they likel y hav e a differen t geometry . Du e t o relatively high depositional gradient s established during rift climax , such shoreline sands probably aggraded mor e verticall y than laterally. (2) Basin floor gravity transported sandstones It i s suggested that a significant amount o f sand must hav e bypasse d th e centra l Snorre- H sub basin to be deposited furthe r to the south, likel y by gravit y flo w processes . Basi n floo r gravit y transported sandstone s hav e no t ye t bee n proven withi n th e Snorre- H half-graben , bu t the basi n physiograph y an d larg e volume s o f eroded san d tha t cannot b e accounted fo r make such a reservoi r mode l ver y likely . Suc h sand -

stones ma y b e detache d fro m th e shorefac e sandstone sequence s i n th e Snorre- H are a an d there ma y b e a stratigraphi c tra p potentia l related t o them . A significan t ba floor deposi t (Seq-II ) ha s been mappe d i n a topographi c lo w i n th e southerly reache s o f th e Snorre-Gullfak s half graben (Fig s 6 , 1 3 & 16) . Degradatio n o f th e Statfjord Eas t footwal l likel y contribute d sig nificantly t o th e suppl y o f san d t o thi s bathy metric deepening , however , an d ther e i s a ris k that th e sandstone s ar e interbedde d wit h shal e and tha t th e net/gros s i s reduced. O n th e othe r hand, san d derive d fro m footwal l erosio n probably wa s store d intermediatel y in shoreline systems fringin g th e footwal l cresta l islands , subsequent to being released an d transporte d t o the basi n floo r settin g (i.e . th e Intr a Draupn e Sandstone, Seq-II) . Thi s increase s the probabil ity of better sorted and thicker packages of basin floor sandstone reservoir facies bein g developed. Similar topographi c lo w area s ar e presen t i n the larger map outline , in the Statfjord Eas t an d northern Visun d half-grabens. Thes e ar e potential accumulatio n site s fo r basi n floo r gravit y transported san d (Fig . 18b) . Basi n floo r sand stones have been foun d in the Draupn e Forma tion i n th e 33/9-1 5 and -1 6 wells. (3) Footwall talus sandstones Footwall talu s sandstones hav e not bee n proven on the Snorre Fault Block. On the contrary, well 34/7-15S sits in a footwall position to the Snorre

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Fig. 20. Schemati c illustratio n o f various reservoi r model s acros s th e Snorr e Fault-block .

Fault Zone, an d foun d nothin g bu t shale s an d very mino r sandston e stringers . O n th e Visund Fault Block, however, thick footwall talus sandstones wer e shed of f the main fault scar p nex t to deep erosiona l incision s (Faerset h e t al. 1995) . Such incisions cannot b e seen in the footwalls of the Snorr e no r th e Statfjor d Eas t Faul t Blocks , but thi s doe s no t rul e ou t th e possibilit y tha t footwall talu s sandstones may be locally present. In fact , th e exces s erode d san d note d i n th e volumetric calculation s ma y suppor t thi s reser voir model. Because o f th e stee p footwal l gradients, san d is likel y t o hav e bee n supplie d b y roc k fal l an d gravity transpor t associate d wit h faul t scar p retreat an d woul d therefor e b e les s sorte d an d clean compared t o the easterly, dip-slope derive d sandstones tha t wer e subjec t t o shorelin e an d shallow marin e reworking. (4) Footwall degradation complexes The present stud y ha s no t looke d i n detai l int o potential fault-scar p degradatio n pla y models . The Snorre-H an d Statfjor d Eas t footwalls have been degrade d dow n throug h Bren t an d Dunli n Formations, however , an d i t is possible tha t th e footwall surface s ma y hav e bee n subjec t t o slumping and disintegratio n to for m small-scale deformation complexe s (Fig . 20 ) similar to what has been show n for the Bren t Field t o th e sout h (Coutts e t al . 1996) . Towards a methodology for evaluating syn-rift play s The present analysi s of the Snorre-H are a serve s to illustrat e a methodolog y fo r evaluatin g syn rift plays . I n evaluatin g suc h play s i t i s im portant t o extract information o n a very detailed level fro m al l availabl e data , wit h th e objectiv e to predic t th e distributio n o f reservoir an d tra p

potential i n spac e an d time . I n th e Snorre- H example, good qualit y 3D seismi c data an d wel l data fro m footwal l a s wel l a s hangingwal l locations wer e available. More commonly, vari able data type s and qualit y will b e available. The analysi s o f th e Snorre- H syn-rif t pla y involved th e followin g basic steps : (1) Mappin g i n detai l o f ke y seismi c surfaces and sequence s (2) Prope r datin g an d sedimentologica l analy sis o f th e syn-rif t successio n (3) Volumetri c mappin g an d mas s budge t estimates o f san d Time maps and fault maps of the base and top of the Snorre- H syn-rif t successio n wer e con structed fro m th e seismic data. Careful mappin g of the Statfjor d Eas t boundar y faul t syste m and the intra-faul t bloc k faul t patter n provide d ke y information o n th e Snorre- H hangingwal l basi n development an d accommodation . B y extracting various seismi c attributes , particularl y ampli tude an d reflectio n intensity , a correlatio n be tween observed san d occurrenc e i n the wells and the seismi c dat a wa s obtained . Th e ma p out line wa s extende d t o cove r th e Snorre- H an d Statfjord Eas t footwal l crestal area s i n additio n to th e Snorre- H hangingwall , i n orde r t o ma p out th e shape , ape x an d topograph y o f the synrift unconformity surface . This enabled a defini tion o f sedimen t sourc e area' s an d sedimen t transport route s int o th e Snorre- H sub-basin . Examination o f th e syn-rif t seismi c sequenc e geometries provide d th e basis for understandin g the tilt history, which is essential to establish th e rate an d magnitud e o f footwal l uplif t an d hangingwall subsidenc e throug h time . High-resolution biostratigraph y allowe d a precise datin g o f th e Snorre- H syn-rif t succes sion. Equally important, b y a systematic analysis of the reworked bioflora , informatio n on timing of erosion o f upland sandy formations as source for th e local syn-rift san d deposit s wa s obtained.

SYN-RIFT EVOLUTIO N I N TH E SNORRE- H ARE A Fades analysi s an d sequenc e stratigraph y o f cored, a s wel l a s non-cored , wel l section s allowed th e constructio n o f a detaile d deposi tional mode l fo r th e syn-rif t succession . Specia l attention wa s pai d t o th e definitio n o f ke y boundary surface s an d th e establishmen t o f a generic sequenc e architecture . The final ste p in the evaluatio n proces s included volumetri c mapping o f th e erode d los s of san d abov e th e unconformit y surfac e i n th e Snorre-H sedimen t sourc e area relative to models for san d depositio n within the Snorre- H syn-rif t succession. B y superimposing th e information on structural gradients , basi n morpholog y an d timing o f erosion , predictio n o f potentia l san d rich system s an d definitio n o f pla y model s were enabled . Conclusions The uppe r Jurassi c infil l (Heathe r an d Draupn e Formations) o f th e Snorre- H are a ha s bee n evaluated i n term s o f it s potentia l fo r syn-rif t plays. Th e petroleu m geologica l relevanc e o f the wor k include s th e developmen t o f a mode l for th e Snorre- H hangingwal l basin , a s wel l a s the erectio n o f a methodolog y fo r evaluatin g syn-rift plays . It i s conclude d fro m th e structura l analysi s that th e development o f the Snorre- H sub-basi n was closel y couple d t o th e developmen t o f th e western, basi n boundin g Statfjor d Eas t Fault . The growt h o f the Statfjor d Eas t Faul t throug h northward propagatin g segmen t linkage resulted in incremental , northwar d migratio n an d bac k stepping of the upper Jurassic syn-rift infil l along the axi s of the Snorr e hangingwal l dip-slope . The sedimentological facie s analysi s ha s linked the development of the syn-rif t successio n to th e deformatio n histor y o f th e Snorr e Faul t Block. A phase o f basin deepenin g an d increas ing sediment instability related to increased faul t block rotatio n (earl y rif t an d rif t climax ) le d t o predominantly shal e depositio n wit h increasin g evidences o f sedimen t instability . This wa s fol lowed b y reduce d rate s o f rotatio n (lat e rift) , leading to the progradation o f a shallow marine/ shoreline sandston e unit , th e Uppe r Draupn e Sandstone, cappin g th e successio n i n updi p areas. Suppresse d generatio n o f ne w relie f finally le d t o levellin g o f th e rif t terrai n an d subsequent drowning . From high-resolutio n biostratigraphy , th e Upper Draupn e Sandston e ha s been show n to b e slightl y diachronou s (Lat e Volgian Ryazanian) an d youngin g northward . Detaile d seismic sequenc e an d attribut e analyse s hav e

215

demonstrated tha t th e sandstone uni t is complex in natur e an d consist s o f distinc t lob e shape d units tha t wer e deposite d a s incrementa l pro gradational sequence s i n a n overal l retrograda tional setting . A thoroug h stud y o f th e reworke d bioflor a suggests tha t loca l erosio n o f footwal l high s started i n th e Lat e Callovian/Earl y Oxfordian . becoming mor e substantia l i n Kimmeridgia n times. A significan t increas e i n reworkin g ha s been identifie d fro m Kimmeridgia n throug h Ryazanian times , includin g pre-rif t (Bren t Group) as well as syn-rift (Heathe r an d Draupn e Formations) taxa . I t i s interprete d t o reflec t expanding uplan d topograph y an d increase d local erosiona l yiel d followin g increase d faul t block rotatio n an d footwal l uplift . By mappin g ou t th e syn-rif t depositiona l basin an d complementar y drainag e basi n a s a whole , i t ha s bee n possibl e t o demonstrat e a dynamic relationshi p betwee n uplan d erosion , sediment transpor t an d basi n deposition . Pro gressive, bu t stepwise , tiltin g o f th e Snorr e Fault-block t o som e 9 C (centra l Snorre-H area ) in th e Ryazania n led t o depositio n o f a twofol d (threefold i f Tarber t Formatio n i s included) , coarsening upwar d (sandstone)-shale-sandston e syn-rift successio n with variable reservoir potential. Th e Uppe r Draupn e Sandston e ha s bee n proved b y drilling across the Snorre-H area , bu t reservoir sandston e facie s ma y thicke n considerably dow n th e hangingwal l of th e Snorr e Faul t Block. Additiona l reservoi r potentia l i s sug gested i n the upper Draupn e Formation , related to depositio n o f basi n floo r sand s t o th e sout h on th e Snorr e Fault-block . The analysi s of th e Snorre- H are a provide s a powerful methodolog y fo r evaluatin g syn-rif t plays. By attempting t o quantif y th e parameter s controlling syn-rif t sedimentatio n int o th e Snorre-H sub-basin , predictiv e models fo r sand stone reservoi r distributio n an d hydrocarbo n play model s have been established. The wor k i s based o n result s fro m a researc h projec t carried out between the Norsk Hydr o Researc h Centr e in Berge n an d th e Universit y o f Edinburgh . Morte n Smelror, IKU , performe d mos t o f th e ne w palynolo gical analyses . Nors k Hydr o funde d th e projec t an d Ingrid Johnsru d i s gratefull y acknowledge d fo r sup porting th e work . W e als o than k Joh n Underhil l fo r stimulating discussion s an d Graha m yieldin g fo r a very constructive review of the manuscript. Finally, we like t o than k Nin a Anthu n a t th e Nors k Hydr o drafting departmen t i n Bergen , wh o prepare d th e illustrations. Note added i n revision. Sinc e thi s pape r wa s written , three ne w wells have been drilled i n the Snorre-H area , Norwegian Bloc k 34/7 . These have proven sandstone s

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in th e Uppe r Draupn e Formation, validating th e cur rent interpretatio n o f th e seismi c attribut e ma p an d model for the Snorre-H tectonostratigraphi c evolution.

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SYN-RIFT EVOLUTIO N I N TH E SNORRE- H ARE A & MARTINSEN , O . J . 1996 . Shorelin e trajectorie s and sequences : descriptio n o f variabl e deposi tional dip-scenarios . Journal o f Sedimentary Research, 66, 670-688. , ASTON , M. , L0MO , L . & STEEL , R . J . 1992 . Advance an d retrea t o f th e Bren t delta : recen t contributions t o th e depositiona l model . In : MORTON, A . C , HASZELDINE , R . S. , GILES , M. R . & BROWN , S . (eds ) Geology o f th e Brent Group. Geologica l Society , London , Specia l Pub lications, 61 , 109-127 . HOME, P . C . 1987 . Ula . In : SPENCER , A . M . (ed. ) Geology of the Norwegian Oil and Gas Fields. Graham and Trotman , London , 143-151 . HAGER, K.-J . & SMELROR , M . 1997 . Reworkin g o f dinocysts as a means of obtainin g usefu l geologi cal information : A stud y o f th e Uppe r Jurassi c Draupne Formatio n i n the Tampen area , Norwe gian Nort h Sea . American Association o f Stratigraphic Palynologists, Abstracts , 30t h Annua l Meeting, Wood s Hole .

JOHANNESSEN, E . P. , MJ0S , R. , RENSHAW , D. , DAL -

LAND, A . & JACOBSEN, T. 1995 . Northern limit of the 'Brent delta' at the Tampen Spu r - a sequence stratigraphic approac h fo r sandston e prediction . In: STEEL , R . J. , FELT , V., JOHANNESSEN , E. P . & MATHIEU, C . (eds ) Sequence Stratigraphy o n th e Northwest European Margin. Norwegia n Petro leum Society , Specia l Publication , 5 , 213-256. JOHNSON, H . D. , MACKAY , T . A . & STEWART , D . J . 1986. Th e Fulma r oi l field : geologica l aspect s o f its discovery, appraisal an d development . Marine and Petroleum Geology, 3 , 99-125. LARSEN, V. B . 1987 . A synthesis of technically related stratigraphy i n th e Nort h Atlantic-Arcti c regio n from Aalenian-Cenomania n time . Norsk Geologisk Tidsskrift, 7 , 281-294 . LEEDER, M . R . & GAWTHORPE , R . L . 1987 . Sedimen tary model s fo r extensiona l tilt block/half-grabe n basins. In : COWARD , M . P. , DEWEY , J . F . & HANCOCK, P . L . (eds ) Continental Extensional Tectonics. Geologica l Society , London , Specia l Publication, 28 , 139-152 . MAKER, C . E . 1981 . Th e Pipe r oi l field . In : ILLING , L. V . & HOBSON , G . D . (eds ) Petroleum Geology of the Continental Shelf of North-West Europe. Heyden, London , 358-370 . & HARKER , S . D . 1987 . The Claymor e oi l field . In: BROOKS , J . & GLENNIE, K . W . (eds ) Petroleum Geology o f North West Europe. Graham & Trot man, London , 835-845 . McLuRE, N. M . & BROWN, A . A. 1992 . Miller Field: a subtle Uppe r Jurassi c submarin e fa n tra p i n th e South Vikin g Graben, U K secto r Nort h Sea . In: HALBOUTY, M. T . (ed.) Giant Oil and Gas Fields o f the Decade. America n Associatio n o f Petroleu m Geologists, Memoir , 54. N0TTVEDT, A. , GABRIELSEN , R . H . & STEEL , R . J . 1995. Tectonostratigraph y an d sedimentar y architecture o f rif t basins , wit h referenc e t o th e northern Nort h Sea . Marine an d Petroleum Geology, 12 , 881-901. ODINSEN, f. , CHRISTIANSSON , P. , GABRIELSEN , R . H. , FALEIDE, J . I . & BERGE , A . M . 2000 . Th e

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geometries an d dee p structur e o f th e norther n North Se a rif t system . This volume. PARTINGTON, M . A. , MITCHENER , B . C. , MILTON , N. J . & FRASER , A . J . 1993 . Geneti c sequenc e stratigraphy fo r th e Nort h Se a Late Jurassic an d Early Cretaceous : distributio n an d predictio n o f Kimmeridgian-Late Ryazania n reservoir s i n th e North Se a an d adjacen t areas . In : PARKER, J . R . (ed.) Petroleum Geology o f Northwest Europe: Proceedings o f th e 4t h Conference. Geologica l Society, London , 347-370 . PROSSER, S . 1993 . Rift-relate d linke d depositiona l systems an d thei r seismi c expression . In : WILLIAMS, G. D . & DOBB, A. (eds) Tectonics and Seismic Sequence Stratigraphy. Geologica l Society . London, Specia l Publications, 71. 35-66. RATTEY, R . P . & HAYWARD , A . B . 1993 . Sequenc e stratigraphy o f a faile d rif t system : th e Mid dle Jurassi c t o Earl y Cretaceou s basi n evolutio n of th e Centra l an d Norther n Nort h Sea . In : PARKER, J . R . (ed. ) Petroleum Geology o f Northwest Europe: Proceedings of the 4th Conference. Geological Society , London , 215-249 . RAVNAS, R . & BONDEVIK , K. 1997 . Architecture an d controls o n Bathonian-Kimmeridgia n shallow marine syn-rift wedge s of the Oseberg-Brage area, northern Nort h Sea . Basin Research, 9, 197-226 . & STEEL , R . J . 1997 . Contrasting style s of Lat e Jurassic syn-rif t turbidit e sedimentation : a com parative study of the Magnus and Oseber g areas , northern Nort h Sea . Marine an d Petroleum Geology, 14,417-449 . & 1998 . Architectur e o f marin e rif t basi n successions. AAPG Bulletin, 82, 110-146 . , BONDEVIK , K., HELLAND-HANSEN , W. . LOMO , L., RYSETH , A . & STEEL , R . J . 1997 . Sedimenta tion histor y a s a n indicato r o f rif t initiatio n an d development: Late Bajocian-Bathonian evolutio n of th e Oseberg-Brag e area , norther n Nort h Sea . Norsk Geologisk Tidsskrift, 77 , 205-232 . , NOTTVEDT , A. , STEEL , R . J . & WINDELSTAD . J . 2000. Syn-rif t sedimentar y architecture s i n th e northern Nort h Sea . This volume. ROBERTS, A . M. , YIELDING , G . & BADLEY . M . E . 1993. Tectoni c an d bathymetri c control s o n stratigraphic sequence s withi n evolving half-gra bens. In : WILLIAMS , G . D . & DOBB , A . (eds ) Tectonics and Seismic Sequence Stratigraphy. Geological Society , London, Specia l Publications , 71, 87-121 . ROUBY, D. , FOSSEN , H . & COBBOLD , P . R . 1996 . Extension, displacement an d block rotation i n the Larger Gullfak s Area , norther n Nort h Sea . AAPG Bulletin, 80, 875-890. RONNEVIK, H . & JOHNSON , S . 1984 . Geolog y o f th e Greater Trol l Fiel d area . Oi l and Gas Journal, 82, 100-106. SCLATER, J . G . & CHRISTIE , P. A . F . 1980 . Continental stretching : a n explanatio n o f th e pos t mid Cretaceous subsidenc e o f th e Centra l Nort h Sea Basin . Journal o f Geophysical Research, 85 , 3711-3739. SCHLISCHE, R . W . & OLSEN , P . E . 1990 . Quantitative filling models fo r continenta l extensiona l basin s

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with application s t o Earl y Mesozoi c rift s o f eastern America. Journal of Geology, 98, 135-155 . SMITH, R . L . 1987 . Th e structura l developmen t o f the Clyde Field. In: BROOKS, J. & GLENNIE, K. W . (eds) Petroleum Geology o f North West Europe. Graham & Trotman, London , 523-531 . SOLLI, T . 1995 . Uppe r Jurassi c pla y concep t - a n integrated stud y i n Bloc k 34/7 , Norway . First Break, 13 , 21-30. SPEKSNIJDER, A. 1987 . The structura l configuration of Cormorant Bloc k I V i n contex t o f th e norther n Viking Graben structural framework. Geologic en Mijnbouw, 65 , 357-379. SURLYK, F . 1989 . Mid-Mesozoi c syn-rif t turbidit e systems: controls an d predictions . In : COLLINSON, J. D . (ed. ) Correlation in Hydrocarbon Exploration. Norwegia n Petroleu m Society , Graha m & Trotman, London , 231-241 . TURNER, C . C , COHEN , J . M. , CONNEL , E . R . & COOPER, D . M . 1987 . A depositiona l mode l fo r

the Sout h Bra e oilfield . In: BROOKS , J. & GLENNIE, K . W . (eds ) Petroleum Geology o f North West Europe. Graha m & Trotman , London, 853-864 . VOLSET, J . & DORE , A . G . 1984 . Revise d Jurassi c lithostratigraphy o f th e Norwegia n Nort h Sea , northern area . Norwegian Petroleum Directorate, Bulletin, 3. WHEATLEY, T . J. , BIGGINS , D. , BUCKINGHAM , J . & HOLLOWAY , N . H . 1987 . Th e geolog y an d exploration o f th e Transitiona l Shelf , an are a t o the wes t of the Vikin g Graben . In: BROOKS , J. & GLENNIE , K . W . (eds ) Petroleum Geology of North West Europe. Graha m & Trotman , London, 979-989 . YIELDING, G. 1990 . Footwall uplift associated wit h Late Jurassic norma l faultin g i n th e norther n Nort h Sea. Journal o f th e Geological Society, London, 147, 219-222 .

Cenozoic evolution of th e central an d northern North Sea wit h focu s on differential vertica l movements of th e basin floo r an d surrounding clastic source area s HENRIK JORDT, 1

12

BRI T I . THYBERG 1 & ARVI D N0TTVEDT 13

Department of Geology, University of Oslo, P.O. Box 1047 Blindern, N-0316 Oslo, Norway ^Present address: Statoil, N-4035 Stavanger, Norway 3 Present address: Norsk Hydro a.s., 1190, lll-5th Avenue SW, Calgary, Alberta, T2P 3Y6, Canada Abstract: Change s i n outbuildin g direction s an d shift s i n sedimen t provenanc e ar e ke y factors fo r understandin g th e developmen t o f th e Cenozoi c successio n i n th e centra l an d northern Nort h Sea . B y establishing a seismi c stratigraphi c framewor k fo r th e Cenozoi c succession, and throug h utilizatio n of velocity dat a and analyse s of sediment compositions , these factors are investigated. Sequence geometries and reflectio n termination s are identifie d and related t o basin-floor geometry and variations i n erosional bas e level. Mineralogical an d geochemical analyse s o f sedimen t sample s fro m selecte d well s ar e compare d wit h thei r seismic sequences . Paleogen e clasti c wedge s ric h i n smectit e wer e buil t fro m th e Eas t Shetland Platfor m an d probabl y draine d a sourc e are a wes t an d northwes t o f th e presen t North Sea . Outbuildin g fro m souther n Norwa y too k plac e durin g Lat e Paleocene-Earl y Eocene an d i n Oligocene t o mid-Miocen e time . I n mid-Miocene tim e outbuilding from th e west stoppe d an d subsequentl y sediments wit h a hig h content o f feldspar , detrita l chlorite and carbonat e wer e deposite d fro m easterl y direction s i n th e centra l an d norther n Nort h Sea. Th e seismi c velocit y o f th e Cenozoi c sediment s increase s dow n t o a mid-Miocen e unconformity, bu t th e velocitie s of Oligocen e an d Eocen e sediment s belo w ar e lowe r an d nearly unaffecte d b y buria l depth , resultin g i n a marke d velocit y inversion. Tw o tectoni c subsidence events, late Eocene-early Oligocene an d lat e Miocene-early Pliocene , have been identified subsequen t t o genera l eustati c sea-leve l lows. These subsidenc e event s cannot b e explained b y therma l subsidence , an d the y ma y hav e crucia l implication s fo r th e understanding o f hydrocarbo n generation , migratio n an d trappin g i n th e Tertiar y an d deeper stratigraphi c levels.

The North Se a Basin is an extensiona l sedimen- tha t uplift s concentrat e i n thre e periods : tary basi n o n the northwest Europea n continen - lat e Cretaceous-earl y Paleocene , lat e Eocene tal shel f (Ziegle r 1982 , 1990 ; Glenni e 1990 ) earl y Oligocene , an d lat e Pliocene-Pleistocen e (Fig. 1) . It wa s formed by lithospheric extension time . during Permo-Triassi c an d lat e Jurassic-earl y Th e Eas t Shetlan d Platfor m an d th e Mi d Cretaceous time . It i s generally agreed tha t th e Nort h Se a Hig h ar e prominen t highs , wes t an d Tertiary sediment s i n th e Nort h Se a constitut e sout h o f th e stud y area , tha t ar e separate d post-rift o r therma l subsidenc e fil l relate d t o fro m th e Centra l an d Vikin g Graben b y majo r these earlie r riftin g episode s (Glenni e 1990) . fault s (Fig . 1) . Th e Fjerritsle v Faul t Zone . Recently, however , Hal l & Whit e (1994 ) hav e whic h ma y b e considere d a s a WN W con interpreted anomalousl y hig h rate s o f Earl y tinuatio n o f th e Sorgenfrei-Tornquis t Zone , Tertiary subsidenc e i n th e norther n Nort h Sea , an d th e 0ygarde n Faul t Zone , ar e zone s o f and Jord t e t al. (1995 ) hav e suggeste d tha t a weaknes s along th e Scandinavia n Shield. Depo complex patter n o f differentia l vertica l move - centre s develope d alon g thes e Mesozoi c an d ment o f th e basi n floo r occurred . Severa l epi - olde r faul t system s during lat e Paleocene , earl y sodes o f uplif t alon g th e margin s o f th e Nort h Eocen e an d earl y Oligocen e tim e (Fig . 2) . Lat e Sea hav e bee n suggeste d b y severa l workers . Paleocen e t o earl y Oligocene , an d earl y Mio A comprehensiv e bibliography o f studie s invol - cen e outbuildin g develope d depocentre s alon g ving uplif t o f Norwa y ha s bee n give n b y Stue - boundar y fault s i n th e Centra l an d Vikin g void & Eldhol m (1996) . Thes e studie s indicat e Grabens . From: N0TTVEDT , A . e t al . (eds ) Dynamics o f th e Norwegian Margin. Geologica l Society , London , Specia l Publications, 167 , 219-243 . 1-86239-056-8/OO/ S 15.00 C Th e Geologica l Societ y o f Londo n 2000 .

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Fig. 1 . Regiona l geological setting of the Nort h Se a and Nort h Atlantic continental rift zone . Position o f Greenland restore d t o 5 3 Ma. Cenozoi c outbuilding directions are indicated alon g the northwest Europea n Atlantic margin. ESB, East Shetlan d Basin; ESP, East Shetlan d Platform; HP, Hord a Platform ; TS, Tampen Spur. (Major outbuilding directions are based on Spjeldnaes 1975 ; Stuevold et al. 1992; Cameron et al. 1993; Jordt et al . 1995) .

Sediments wer e supplie d fro m the wes t and northwest into the North Sea , and furthe r nort h into th e V0rin g margi n durin g Paleocen e an d Eocene time s (Fig. 1) . However, th e outbuildin g direction shifte d fro m wes t t o eas t durin g th e Eocene-Oligocene transitio n i n th e centra l North Se a (Michelse n e t a l 1995 , 1998) , an d the sediment supply from th e west stopped i n the M0re an d V0rin g Basin s (Fig . 1) . Sedimen t supply fro m th e east increase d considerabl y an d dominated durin g Miocen e an d Pliocen e time . Abundant sediments , whic h ma y hav e com e from th e Scandinavia n an d Balti c Shield s eas t of th e basin , wer e probabl y transporte d int o the basi n b y th e so-calle d Balti c Rive r syste m (Bijlsma 1981 ; Gibbar d 1988 ; Camero n e t a l 1993). A lat e Miocen e o r Pliocen e syste m o f

NNW-SSE trendin g ridges and trough s in the UK secto r may have formed on a pro-delta slop e (Cartwright 1995) , possibl y reflectin g th e earl y evolution o f a larg e Balti c delta . North-sout h trending incised valleys in the Middle and Uppe r Pleistocene successio n i n th e Danis h Nort h Sea (Salomonsen 1995 ) sugges t tha t thi s are a wa s part o f a deltai c syste m controlled b y the Balti c River u p t o Pleistocen e time . This paper focuses on the Cenozoic successio n in th e norther n Danis h an d th e Norwegia n North Se a basins . Th e mai n objectiv e i s t o discuss the geological evolutio n of this area, with focus o n differentia l vertica l movement s o f th e basin floor . Ou r inten t i s t o sho w tha t seismi c sequence geometrie s an d outbuildin g direction s provide informatio n abou t change s in th e basi n

CENOZOIC EVOLUTIO N O F TH E CENTRA L AN D NORTHER N NORT H SE A 22

1

Fig. 2. Structur e map o f the study area showin g the seismic data coverag e and locatio n of wells with logs and o r biostratigraphic data . Shade d oval s indicate positions o f depocentres an d number s within these refer t o CSS sequence numbe r show n i n Fig . 3 . Indicated seismi c profile s ar e use d a s examples later i n th e text . Outbuilding directions an d depocentr e location s o f CSS-1 t o CSS- 7 sout h o f 57 C N ar e base d o n Jord t (1995).

topography, an d tha t thes e are related to underlying structures, t o shifts i n provenance are a an d to sedimen t accumulatio n rates . Methods In th e study , w e interprete d abou t 17000k m o f conventional multi-channe l seismi c reflectio n profile s (Fig. 2), and integrate d mineralogica l an d geochemica l analyses of cuttings from more than 4 0 wells (Thyberg el al. 2000) . Velocit y log s fro m th e investigate d well s were analysed and use d t o convert seismic travel times to depth . Publishe d high-resolutio n biostratigraphi c data wer e use d t o dat e th e seismi c sequences . Accu mulation rate s fo r th e mapped seismi c sequence s wer e estimated an d compare d wit h sedimen t compositio n and stackin g pattern . Result s fro m th e mineralogica l

and geochemica l analyses of th e sediment s have been presented b y Thyber g e l al. (2000) . Th e establishe d seismic sequenc e framework , lithostratigraphic correlations, mappe d outbuildin g directions , an d hiatuse s indicated b y seismi c an d biostratigraphi c dat a ar e summarized i n Fig s 3 and 4 .

Constraints on the seismic interpretation Seismic mapping and interpretation s are based o n th e identification an d correlatio n o f seismi c unconformi ties tha t separat e an d boun d package s tha t hav e relatively unifor m seismi c signatures . Thes e seismi c unconformities ar e marke d b y reflectio n termination s (i.e. onlap , downlap an d toplap) , b y differences i n th e amplitude an d frequenc y spectr a o f th e adjoinin g sequences and/o r b y regionally continuous reflections .

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Fig. 3 . Summar y o f main observation s i n th e study area correlate d wit h lithostratigraphy in the U K an d Norwegian Nort h Sea . Depocentr e location s ar e indicate d i n Fig. 2 . CG, Centra l Graben; ESP , Eas t Shetlan d Platform; HP , Hord a Platform ; MB , Marulk Basin ; NDB, Norwegian-Danis h Basin ; PR, Patchban k Ridge ; TS, Tampen Spur ; RFH , Ringk0bin g Fyn High ; SB , Stord Basin ; SG, Sog n Graben ; SN , souther n Norway; UH, Utsir a High ; VG, Viking Graben .

Seismic mapping an d interpretatio n wer e initially carried ou t independentl y o f biostratigraphi c an d petro graphic data . However , in cases where several seismic interpretations wer e possible, fo r example , becaus e o f noise, the interpretation wa s guided b y available highresolution biostratigraph y an d completio n lo g strati graphic subdivision . In thi s paper, we interpret seismi c onlap to reflec t a base-discordant relationship in which initially horizontal strat a terminat e o r thi n (belo w seismi c resolution ) progressively against an initially inclined surface, o r in which initially inclined strata terminate o r thi n (below seismic resolution ) progressivel y updi p agains t a sur face o f greater initial inclination (see Mitchum 1977).

Progradational sequence s interprete d fro m th e seismic dat a outlin e th e ne t outbuildin g o f sedimentar y wedges, bu t d o no t necessaril y reflect th e directio n of sediment transport. A s an example, th e main directio n of sediment transport is parallel to the coast along nondeltaic coastlines, whereas it i s more perpendicula r o r oblique to the coast where deltas are present (Gallowa y 1989). Bot h depositiona l setting s ma y sho w ne t progradation, an d resul t i n simila r facie s association s (Galloway 1989) . Implicitly , westwar d progradatio n observed on a seismi c profil e may be the resul t of dominating northward sedimen t transport . It ma y als o b e deceptiv e t o interpre t curren t bathymetric condition s base d o n seismi c geometries,

CENOZOIC EVOLUTIO N O F TH E CENTRA L AN D NORTHER N NORT H SE A 22

3

Fig. 4 . Stratigraphi c schem e showin g correlation o f the seismi c sequence s wit h the NSR zonatio n (Gradstei n & Backstrom 1996 ) and th e depositiona l sequence s fro m th e Danis h Nort h Se a (Michelsen e t al. 1995 . 1998) . Dashed correlatio n line s indicate correlation without direc t seismi c tie to th e Danis h Nort h Sea . as th e progradatio n o f clasti c sediment s alon g shel f margins (Orto n & Readin g 1993 ) an d i n deep-se a settings (Readin g & Richard s 1994 ) ma y resul t i n similar, bu t undifferentiable , depositiona l geometries . Recognition o f onlap is of little help i n the interpreta tion o f depositiona l environment , a s it i s not possibl e to differentiat e betwee n coasta l onla p an d marin e onlap o n seismi c dat a withou t havin g cor e dat a o r palaeobathymetric informatio n (Bertra m & Milto n 1988; Robert s e t al . 1993) .

Net accumulation rates Absolute age s o f th e mappe d sequence s wer e used t o calculate ne t accumulatio n rate s (Tabl e 1 ) using th e following expression :

where 5 i s net accumulation rat e i n m Ma ' . TWT is thickness i n s two-wa y trave l time , V i s interva l velocity i n ms" 1 an d D i s th e tim e duratio n fo r for mation o f th e actua l sequence . S i s a minimu m esti mate o f th e accumulatio n rat e a s n o accoun t i s taken of erosion afte r depositio n o r mechanical compaction .

Mineralogical analyses The method s use d i n th e mineralogica l an d geochem ical analyses have been describe d i n detail by Thyber g et al . (2000 ) an d ar e onl y briefl y summarize d here . Cutting sample s fro m th e Cenozoi c successio n wer e analysed wit h X-ra y fluorescenc e (XRF ) an d X-ra y diffraction (XRD) . Th e inorgani c geochemica l dat a were use d t o estimat e whole-rock mineralogy, an d i n the assessmen t o f main an d trac e elements . The cla y mineralog y was categorize d as smectit e (including expandabl e mixed-laye r illite-smectite) .

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Table 1 . Th e calculated net accumulation rates are based on seismic velocities derived from well velocity logs (Fig. 10) and on observed seismic two-way travel times Sequence

Thickness (ms)

Duration (Ma)

Velocity (m/s)

Net accumulatio n rat e (m/Ma) Max

Min

2200

235

29

2000 2000 2000

100 71 166

11 7 16

9 2.4

2000 2000

44 250

11 41

100

20.6

2000

39

5

50

4.5

2200

195

12

Max

Min

CSS-8

800

100

3.4

CSS-7 CSS-6 CSS-5

900 500 500

100 50* 50*

9 7 3

CSS-4 CSS-3

400 600

100 100

CSS-2

800

CSS-1

800

The sequence s wer e mechanicall y compacte d an d possibl y eroded , therefor e th e accumulatio n rate s represen t minimum estimates. Minimum seismic two-way travel times are taken basinwards in the mapped sequences , and therefore probabl y reflec t a sedimentar y environment marked b y starvation an d limite d erosion . * Minimum thicknesse s fro m th e Germa n Nort h Se a (Michelsen e t al. 1998).

illite, kaolinite and chlorite. The clay mineralogy at the time o f depositio n ca n b e relate d t o th e sedimen t provenance an d sedimentar y facies. This include s not only th e natur e o f th e erosio n product s o f th e sourc e rock bu t als o th e rate s o f erosio n an d th e processe s of roc k weatherin g acting i n differen t climates . Shales contain a grea t variet y o f non-cla y mineral s (e.g. detrital feldspar), which make u p a significant fraction of thes e rocks . Th e feldspa r conten t i s a goo d indication o f a rock' s mineralogica l maturity . Th e Al2O3/SiO2 rati o i s used a s a n indicatio n o f th e clay / sand rati o becaus e A1 2O3 is commonly found in clays, feldspar an d mica .

Seismic stratigraph y The stratigraphi c framewor k establishe d i s shown i n Fig . 4 . Th e followin g discussion o n dating o f th e seismi c sequence s i s base d o n th e work b y Jordt et al. (1995). However, th e datin g has bee n improve d b y new data illustratin g th e tie betwee n th e seismi c dat a an d wel l infor mation. Biostratigraphi c dat a fro m wel l 34/7-1 , 34/8-1 an d 35/11-1 , publishe d b y Gradstei n & Backstrom (1996 ) an d Steurbau t e t a l (1991) , are include d i n Fig s 5 an d 6 t o documen t th e suggested age s o f th e mappe d seismi c sequences in th e presen t study .

Age determinatio n Palaeontological ages The CSS- 1 sequenc e seem s t o correlat e wit h th e zones NSR- 2 an d NSR- 3 (Gradstei n & Back strom 1996 ) i n wel l 34/7-1 , suggestin g a Lat e

Paleocene t o earlies t Eocen e ag e (Fig . 5) . Th e top o f sequence CSS- 1 has a n Earl y Eocen e ag e in wel l 35/11-1 (Fig. 6) , where it i s found belo w the nannoplankto n zon e NP1 1 (Steurbau t e t al. 1991). Thi s suggest s a n earlies t Eocen e ag e fo r the top of sequence CSS-1 in the northern Nort h Sea, whic h i s i n agreemen t wit h th e datin g o f this seismi c boundar y i n th e centra l Nort h Sea (Michelse n e t al . 1995 , 1998) . Th e to p o f sequence CSS- 1 is found nea r the top of a nar row interva l wit h lo w gamma-ra y value s an d high soni c velocities , whic h correspond s t o th e top o f th e Balde r tuff s (Fig . 5) . Th e to p o f sequence CSS-1.1 correlates wit h the PaleoceneEocene transitio n in wel l 35/11-1 . The seismi c boundary at th e top o f the CSS- 2 sequence correlate s wit h th e Eocene-Oligocen e transition (Fig. 4) . In wel l 34/7-1 , 34/8-1 (Fig. 5 ) and 35/11- 1 (Fig. 6) , the to p o f sequenc e CSS- 2 correlates wit h a majo r hiatu s betwee n Middl e Eocene an d Earl y Oligocen e sediments . I n th e Danish well s Inez-1 , F- l an d K-l , thi s seismi c boundary i s locate d belo w th e NP2 2 zon e (Michelsen e t al . 1998) . Steurbau t e t al . (1991 ) indicated tha t zone NP22 i s present i n Oligocen e sediments overlyin g the CSS- 2 sequenc e i n well 35/11-1. We correlate th e top o f sequence CSS- 2 with th e bas e o f th e Oligocen e sediment s an d suggest that the top of sequence CSS-2 correlate s with top Eocene . The top o f sequence CSS-2.1 is found i n a successio n tha t ha s bee n assigne d a n Early, possibl y Mid-Eocen e ag e i n wel l 35/11- 1 (Fig. 6) . W e sugges t tha t th e CSS-2. 1 sequenc e correlates wit h uni t 2 of Michelse n e t al . (1998) (Fig. 4) , indicating a n Early Eocen e age .

Fig. 5 . Well s 34/7- 1 an d 34/8-1 . Correlatio n o f seismi c data wit h hiostratigraphy an d lo g dat a fro m well s 34/7- 1 an d 34/8-1 .

Fig. 6 . Wel l 35/11-1 . Correlatio n o f seismic data wit h biostratigraph y an d lo g data.

CENOZOIC EVOLUTIO N O F TH E CENTRA L AN D NORTHER N NORT H SE A 22 7 The to p o f sequenc e CSS- 3 ha s n o direc t seismic ti e t o th e Danis h Nort h Sea . In well s 34/7-1 an d 34/8-1 , the to p o f sequence CSS- 3 is situated i n stratigraphi c uni t NSR- 7 (Gradstei n & Backstro m 1996) , suggesting a n Earl y Oligocene age (Fig. 5). A simila r ag e is supported b y the biostratigraph y fro m wel l 35/11- 1 (Fig . 6). An upwar d chang e i n seismi c signatur e fro m progradation t o marke d aggradatio n an d on lap occurs against th e top o f sequence CSS- 3 in the norther n Nort h Sea . A simila r chang e i n sequence geometry is observed acros s th e to p o f sequence 4. 2 i n th e Danis h Nort h Sea-centra l North Se a (Michelsen e t al. 1998) . On th e basi s of similaritie s i n th e seismi c signature , an d th e ages indicated b y the biostratigraphy, we believe that th e to p o f sequenc e 4. 2 o f Michelse n e t aL (1998) correlates with the top o f sequence CSS- 3 (Fig. 4) , indicating a mid-Early Oligocen e age. The seismi c boundary representin g the top of sequence CSS- 4 ha s bee n date d t o lates t Oligo cene time by Michelsen et al. (1998). The to p o f sequence CSS- 4 correlate s wit h th e transitio n between NSR-8 A an d NSR-8 B i n wel l 34/8- 1 (Fig. 5) , indicatin g a Lat e Oligocen e o r Earl y Miocene age. Although the ages indicated by the biostratigraphy ar e no t quit e i n agreement , w e suggest that the top of sequence CSS-4 is of latest Oligocene age. The Miocen e successio n i s divided int o thre e seismic sequences, CSS-5, CSS-6 and CSS-7, and the boundaries separatin g these ar e dated in the central North Sea . Th e boundaries of the CSS-5 and CSS- 6 sequence s correlat e seismicall y with the boundarie s o f Unit 5 and Uni t 6 of Michelsen e t a L (1998) , respectively . This correlatio n suggests that the top o f sequence CSS-5 is of late Early Miocene age and the top of sequence CSS 6 is of early Mid-Miocene ag e (Fig. 4). Miocen e sediments are also present in the northern Nort h Sea; however , th e biostratigraph y i s equivoca l (Fig. 7) . Eidvi n & Rii s (1992 ) indicate d tha t

Fig. 7 . Biostratigraph y fro m wel l 34/8-1 .

Early Miocen e sediment s ar e overlai n b y lat e Mid- an d Lat e Miocen e strat a i n wel l 34/8-1 . This suggest s tha t sequence s CSS- 5 an d CSS- 7 are presen t an d separate d b y a hiatu s o f lat e Early o r earl y Mid-Miocen e age. Dating o f th e same sediment s i n th e sam e well (i.e. 34/8-1) by Gradstein & Backstro m (1996 ) als o indicate s that th e Earl y Miocen e sequenc e is present an d in agreement with the biostratigraph y o f Eidvin & Riis (1992). However, Gradstein & Backstrom (1996) date d th e overlyin g sediment s t o lat e Early an d earl y Mid-Miocen e age , suggesting that sequence s CSS-5 and CSS-6 are present and that th e CSS- 7 sequence is absent in well 34/8-1 . The dating s o f Gradstei n & Backstro m (1996 ) also indicat e th e presenc e o f a majo r hiatu s of Mid- an d Lat e Miocen e ag e i n th e norther n North Sea . Steurbaut e t al . (1991 ) hav e no t identified upper Oligocene or Miocene sediments in well 35/11-1; however, as they have no datings between 675 m an d 725 m th e presenc e o f sediments of thi s age i s still possible (Fig . 6). The CSS- 8 sequenc e i s correlate d wit h Plio cene sediments . I t is , however , uncertai n whether a Lowe r Pliocen e sequenc e i s presen t in th e norther n Nort h Sea . Ag e dating s b y Gradstein & Backstro m (1996 ) i n well s 34/7- 1 and 34/8- 1 (Fig. 5 ) indicate that Lowe r Pliocene sediments ar e present . Th e dating s o f Eidvi n & Riis (1992) , however , are i n disagreemen t wit h that interpretation , indicatin g th e presenc e o f only Uppe r Pliocen e sediment s i n wel l 34/8- 1 (Fig. 7). The top o f sequence CSS-8 is correlated with the Pliocene-Pleistocen e transition in wells 34/8-1 (Fig. 5 ) and 35/11- 1 (Fig. 6) . However, in well 34/7-1 , th e uppe r boundar y o f th e CSS- 8 sequence i s o f Pleistocen e age . This ag e incon sistency i s relate d t o ambiguitie s in th e seismi c interpretation cause d b y nois e fro m sea-floo r multiples an d b y th e westwar d chang e i n toplap becaus e o f erosiona l truncation , an d t o toplap cause d b y progradation.

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low accumulatio n rat e alon g th e M0r e Basin . The uppe r sub-sequenc e o f CSS- 1 (CSS-1.2 ) i s The tim e scale devise d b y Harlan d e t at. (1990) characterized b y hig h concentration s o f carbo was use d t o conver t th e suggeste d palaeontolo - nate, smectiti c clay s an d sometime s P2O 5 Th e gical age s o f th e seismi c sequenc e boundarie s tuffaceous Balde r Formatio n cause s a marke d to absolut e age s (Fig . 4) . Th e accurac y an d increase i n th e smectit e conten t uppermos t i n precision o f tim e resolutio n var y wit h th e sequence CSS- 1 (Fig . 11) . chronostratigraphic methods . Th e achievabl e The CSS-2 sequences (Eocene) prograded fro m resolution fo r integrate d stratigraph y i n Cen - the Eas t Shetlan d Platfor m an d majo r depo ozoic marin e succession s varie s betwee n Sk a centres forme d i n th e Vikin g an d th e Centra l and 1 Ma (Mial l 1995) . The duratio n o f deposi - Grabens (Fig s 8 and 9) . CSS-2.1, the lower subtion o f th e Fu r Formatio n (Spjeldnae s 1975) , sequence o f sequenc e CSS-2 , prograde d fro m which forms the upper part of sequence CSS-1.2, southern Norwa y an d form s a mino r depo is use d her e a s a n exampl e o f th e uncertaint y centre i n the Norwegian-Danish Basi n (Fig . 8) . related t o absolut e ag e dating. Spjeldnas s (1975 ) Sequence CSS- 2 a s a whol e show s a marke d summarized tw o opposit e view s abou t th e eastward thinnin g i n th e centra l Nort h Sea , i n sedimentation rat e o f th e Fu r Formation . I f i t the same area where Michelsen et al. (1998) indiis assumed tha t th e laminatio n i n th e Fu r For - cated continuou s sedimentatio n durin g Eocen e mation represent s annua l varve s (Bond e 1973) , time. Th e Eocen e sequenc e als o thin s alon g then i t ma y represen t a perio d o f onl y 6 0 ka. the 0ygarde n Faul t Zon e furthe r t o th e nort h However, studie s o f magneti c reversal s o f (Fig. 8) . However, biostratigraphica l dat a indi the volcani c as h bed s indicat e a muc h lon - cate a hiatu s betwee n th e Eocen e an d th e ger time for depositio n o f this interval o f abou t Oligocene sequences i n that are a (Fig . 6) . 3 Ma (Sharm a 1969) . Bot h view s ar e compat The maximu m net accumulatio n rat e attain s ible wit h th e fac t tha t ther e ar e mor e tha n only 39 m Ma"1 fo r th e Eocene sequence , whic h 130 well-define d as h bed s i n th e Fu r Forma - is ver y lo w compare d wit h th e othe r Tertiar y tion, an d wit h th e observation s o n sedimentasequences (Tabl e 1) . Sequenc e CSS- 2 i s domi tion rate s o f similar recent sediment s (Spjeldnses nated b y clay, consisting mainly of smectite, and 1975). Th e perio d o f formatio n o f mos t o f with a low quartz and feldspar content (Fig. 11) . the mappe d sequence s (i.e . CSS- 1 t o CSS-8 ) The hig h conten t o f smectit e i n th e CSS- 2 (Fig. 4 ) i s beyon d th e uncertaint y leve l sequence suggest s a dominantl y volcani c influ (±lMa) relate d t o th e absolut e ag e datin g o f enced source . marine Cenozoi c successions . Fo r example , th e Although th e seismic and mineralogical infor shortest perio d o f formatio n i s indicate d fo r mation give s a regionally consistent picture of a sequence CSS- 3 o f onl y 2. 4 Ma (i.e . 35.4 - clay-dominated sequence , completion log s fro m 33 Ma), an d th e longes t i s indicate d fo r th e wells in licence block 35/ 8 indicate locally a high CSS-2 sequence , nearl y 21 Ma coverin g the time sand content. These sand-ric h Eocen e sediment s interval 56-35. 4 Ma. Th e absolut e age s o f th e reflect a local thickenin g of the CSS- 2 sequenc e mapped sequence s ar e use d i n th e discus - northwest o f th e depocentr e i n th e Vikin g sion below. Graben (Fig . 8) , correlatin g wit h a moun d a t the bas e o f sequenc e CSS- 2 o n th e norther n Horda Platfor m (Fig s 1 2 and 13) . Th e precis e palaeontological ag e o f thes e san d deposit s i s Sequence description unknown, bu t thei r locatio n a t th e bas e o f sequence CSS- 2 (Fig . 12 ) suggest s a n earl y Upper Paleocene-Eocene sequence Eocene age . Th e moun d ha s a thicknes s o f The CSS- 1 sequenc e (Uppe r Paleocene-lower - 300ms TW T (3-40 0 m) an d lengt h o f mor e most Eocene ) i s marked b y progradin g wedge s than 25km , thu s it ha s a lo w topography. I t i s that buil t ou t fro m souther n Norwa y an d fro m located o n th e cres t o f th e downthrow n bloc k the Eas t Shetlan d Platfor m (Fig s 8 and 9) . The of th e Lomr e Terrac e (Fig . 12) , an d wes t o f depocentre i s recognize d a s a majo r westwar d the 0ygarde n Faul t Zon e (Fig . 9) . North prograding wedg e off Sognefjorden . ward progradatio n o f th e CSS- 2 sequenc e i s The sediment accumulation rates for sequence reflected b y th e stackin g patter n o n th e Hord a CSS-1 vary between 1 2 and ^SmMa" 1 (Table 1, Platform (Fig . 12) . Th e bas e o f sequenc e Fig. 10) , with maximum net accumulation in th e CSS-2 i s a t 1700m s TW T dept h nort h o n th e depocentre. Sequenc e CSS- 1 ha s a relativel y Lomre Terrace , bu t onl y a t 1400m s TW T uniform thicknes s of 200 ms TWT i n th e north - depth o n th e Hord a Platfor m t o th e sout h western par t o f th e stud y area; thu s it attain s a (Fig. 12) .

Fig. 8 . Isopac h map s of Upper Paleocen e (CSS-1) , Eocen e (CSS-2 ) and Oligocen e (CSS-3 and CSS-4 ) sequences. Outbuildin g direction s an d contour s o f CSS-3 an d CSS- 4 in the Norwegian-Danis h Basi n ar e based o n Jordt (1995).

Fig. 9 . NE-S W strikin g seismic section showin g th e mappe d sequence s i n th e Norwegian-Danis h Basin . Lin e locatio n i s shown i n Fig . 2 .

CENOZOIC EVOLUTIO N O F TH E CENTRA L AN D NORTHER N NORT H SE A 23

Oligocene sequence The CSS- 3 sequenc e (lowe r Earl y Oligocene ) was buil t fro m th e Eas t Shetlan d Platfor m an d from th e 0ygarde n Faul t Zon e toward s th e centre o f th e norther n Nort h Se a (Fig . 8) , an d towards the south (Fig . 14 ) into the NorwegianDanish Basin . Sequenc e CSS- 3 i s thi n (Fig . 5 ) or absen t (Fig . 8 ) i n th e norther n par t o f th e study area . Hiatuse s a t th e Eocene-Oligocen e transition ar e indicate d i n th e Tampe n Spu r area (Fig . 5 ) an d alon g th e 0ygarde n Faul t Zone (Fig . 6) . The ne t accumulatio n rat e fo r th e CSS- 3 sequence varie s betwee n 4 1 an d 25 0 mMa"1 (Table 1) , and these rate s ar e almos t a n orde r of magnitud e highe r tha n thos e o f th e under lying Eocen e sequenc e bot h i n th e depocentr e and i n proxima l positions . Suc h a larg e differ ence ca n probabl y b e attribute d t o post-deposi tional erosio n o f th e Eocen e strata . Th e Oligocene CSS- 3 depocentre s i n th e Norwe gian-Danish Basi n an d i n th e Vikin g Grabe n attain approximatel y th e sam e maximu m thickness o f 5-60 0 ms TWT , indicatin g simila r high rate s o f ne t accumulatio n alon g thos e basin margins . The sediment s of sequence CSS- 3 contain less Co, Ni , C u an d Z n tha n th e underlyin g Eocene sediments, whereas the smectite content appear s higher tha n i n the sediment s immediately below (Fig. 11) . Thes e shift s i n th e sedimen t composi tion occu r a t th e sam e dept h a s th e suggeste d hiatus between the Eocene and Oligocene in well 34/7-1 (Fig . 5 ) an d mak e th e CSS-2-CSS- 3 boundary eas y t o detec t o n th e compositiona l data. A highe r conten t o f mic a an d illit e i n the lowe r Oligocen e tha n i n th e underlyin g Eocene sediment s occur s i n th e Norwegian Danish Basi n (Michelsen e t al. 1998) . The CSS- 4 sequenc e (uppe r Lowe r t o Uppe r Oligocene) i s characterize d b y a relativel y uni form sedimen t distributio n withou t marke d depocentres (Fig . 8) . Locally , i n th e norther n North Sea , th e thicknes s o f sequenc e CSS- 4 varies muc h becaus e o f remobilizatio n o f underlying smectite-rich clay of sequence CSS- 3 (Fig. 15) . Th e directio n o f progradatio n wa s from th e northeas t i n th e Norwegian-Danis h Basin an d mainl y from th e west in th e norther n North Sea . Progradationa l stackin g i s generally not pronounce d i n th e CSS- 4 sequenc e an d internal reflectio n termination s ar e generall y subtle. Therefore , sequenc e CSS- 4 appear s t o be a n aggradationa l sequenc e tha t onlap s against th e underlyin g CSS- 3 sequenc e an d pinches ou t towar d th e basi n margin s (Fig s 9 and 13) . Sequence CSS- 4 is possibly represente d

1

by a prograding wedg e in the northwesternmost part o f th e stud y area , an d b y northwar d pro gradation wit h topla p o n th e Patchban k Ridg e (Jordt e l a l 1995) . The ne t accumulatio n rat e fo r th e CSS- 4 sequence varie s betwee n 1 1 an d 44 m Ma"1 , significantly lowe r tha n thos e o f sequenc e CSS-3, bu t o f simila r magnitud e t o thos e o f sequence CSS- 2 (Tabl e 1) . The conten t of smectite decreases whereas the feldspar conten t appear s t o increas e upward s within sequence CSS-4 (Fig . 11) . The conten t of Co, Ni, Cu and Z n is almost constant, but lowe r than i n th e Eocen e sediments.

Miocene sequence Development o f th e Nort h Se a Basi n durin g Miocene tim e (sequence s CSS-5 , CSS- 6 an d CSS-7) i s characterize d b y a reductio n i n outbuilding fro m th e west . Th e mappe d Mio cene depocentres are all located south of 60 N in the Nort h Se a (Fig . 16) . Farthe r north , th e Miocene sequence s thi n t o clos e t o (o r below ) seismic resolution , and , consequently , i t i s no t possible t o ti e thei r boundarie s withi n th e seismic grid. The CSS- 5 sequenc e wa s buil t ou t int o th e Viking Graben are a fro m a basin margi n to th e west, and fro m th e northeast i n the NorwegianDanish Basin . A marke d basinwar d shif t i n onlap suggests that deposition wa s influenced by basin-floor topography , generated by the underlying Oligocen e deposits (Fig. 9). Sequence CSS- 5 show s a ne t accumulatio n rate betwee n 1 6 an d 166 m Ma"1 , comparabl e with that o f sequence CSS-1 (Table 1) . It attain s the highes t ne t accumulatio n rat e i n th e depo centre of the Viking Graben . The overlyin g sequences . CSS- 6 an d CSS-7 . thin northward an d may be absent i n large parts of the study area nort h o f 60 N (Fig . 16) . CSS-6 is a relativel y thin, aggradin g sequence that wa s built fro m th e Norwegian-Danish Basi n toward the south and southwest across the RingkobingFyn Hig h an d toward s th e Centra l Graben . Sequence CSS- 7 wa s buil t fro m souther n Nor way towar d th e sout h an d southwest , an d downlaps towar d th e sout h i n th e Norwegian Danish Basi n (Jordt 1995) . The CSS- 7 sequence is compose d o f westward-dipping , sub-paralle l reflections tha t onla p distall y agains t th e Mi d North Se a Hig h (Fig . 9) . Sequenc e CSS- 7 i s truncated by the base Quaternary, as shown by a marked angula r unconformity , reflectin g post Miocene uplif t an d erosio n i n th e Norwegian Danish Basi n (Fig. 9).

Fig. 10 . Soni c lo g velocities fro m th e norther n Nort h Sea . Th e velocit y data ar e base d o n soni c log s fro m 4 0 wells between 6 0 and 62°N .

Fig. 11 . Mincralogica l an d geochemica l dat a fro m wel l 34/7- 1 based o n th e established seismi c framewor k an d biostratigraph y (fro m Thyber g et al. 1999) .

Fig. 12 . N~ S striking seismi c sectio n fro m th e Hord a Platfor m i n th e norther n Nort h Se a showin g th e mappe d seismi c sequences .

Fig. 13 . N W S E strikin g seismi c sectio n showin g th e mappe d sequence s i n th e norther n Nort h Sea . Lin e locatio n i s shown i n Fig . 2 .

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Fig. 14 . DCS-3 0 N-S strikin g seismic section showin g the Oligocene and Miocen e sequences in the NorwegianDanish Basin . Lin e locatio n i s shown i n Fig. 2. The sedimen t accumulatio n rate s fo r th e sequences CSS- 6 an d CSS- 7 var y betwee n 7 and 71 m Ma"1, and 1 1 and 100 m Ma"1, respec tively (Table 1) . The maximum net accumulatio n rates are all lower than for the CSS-1 and CSS- 3 sequences, bu t highe r tha n fo r sequence s CSS- 2 and CSS-4 . The Miocen e sediment s ar e characterize d b y an upwar d increas e i n feldspa r content, reflect ing a decrease in sedimentary maturity (Fig. 11) . An upwar d increasin g amoun t o f clastic carbo nate and chlorit e reflects a n increasing supply of clastic material . Pliocene sequence The CSS-8 sequenc e is dominated by a > 1000m thick clasti c wedg e nort h o f th e Hord a Plat form (Fig s 1 6 and 17 ) which was buil t out fro m the sout h (Fig . 12 ) and eas t (Fig . 13) . Th e bas e of sequenc e CSS- 8 i s marke d b y a prominen t regional downlap surface (Fig . 17) . In th e Stor d Basin a >400 m thic k Pliocen e wedg e was built from nort h (Fig . 12 ) and east . Th e uppermos t Pliocene sedimen t succession is eroded alon g th e 0ygarden Faul t Zone , wherea s furthe r wes t the upper boundar y of sequenc e CSS-8 i s conformable, resultin g fro m continuou s progradation . Sequence CSS- 8 onlap s basinwar d agains t Mio -

cene sediment s alon g th e Eas t Shetlan d Plat form, generating a marked seismic unconformity (Jordt e t al. 1995) , and i t aggrade s an d pinche s out toward s th e Stavange r Platfor m (Fig . 16) . The CSS- 8 sequenc e build s northward s a s wel l as southwards o n the Horda Platfor m (Fig . 12) . The overlyin g Quaternar y sequenc e show s a thickening in th e sam e area . Pliocene sedimen t accumulatio n rate s var y between 2 9 an d 235 m Ma"1, whic h i s muc h higher than those o f the Miocen e (Table 1) . The Pliocene maximu m net accumulatio n rat e ha s a similar magnitud e t o tha t fo r th e CSS- 1 an d CSS-3 sequences . The sedimen t composition o f sequenc e CSS-8 is characterize d b y a n immatur e mineralog y reflected b y upwar d increasin g amount s o f feldspar, moderat e amounts of clastic carbonate and detrita l chlorit e (Fig . 11) . Discussion Outbuilding of the CSS- 1 sequenc e alon g the basin margin s indicate s tha t th e landmasse s o f southern Norwa y an d th e U K secto r existe d during Lat e Paleocen e time . The sediment com position reflect s erosio n o f area s wit h mixe d lithologies. Th e hig h ne t accumulatio n rate s (Table 1 ) and thickenin g of sequence CSS-1 into

Fig. 15 . N W S E strikin g seismi c sectio n showin g remobili/e d Oligocen e clay.

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Fig. 16. Isopac h maps of the Miocen e (CSS-5 to CSS-7 ) and Pliocen e (CSS-8) sequences. Outbuilding directions in th e Norwegian-Danis h Basi n are base d o n Jord t (1995).

CENOZOIC EVOLUTIO N O F TH E CENTRA L AN D NORTHER N NORT H SE A 23 depocentres (Fig . 8 ) may sugges t tha t sedimen t input was focused along older structure elements (i.e. Sognefjorden ) int o a basi n possibl y deepe r than 1000m . A s th e inheren t post-Cretaceou s bathymetry alon g th e Nort h Se a basi n margi n areas wa s no t significant , th e accommodatio n and developmen t o f th e thic k Uppe r Paleocen e wedge probably resulte d fro m subsidenc e o f th e basin floo r combine d wit h increase d sedimen t supply relate d t o tectoni c uplif t an d subsequen t erosion o f souther n Norway . Suc h uplif t ma y have occurre d i n response t o th e openin g o f the North Atlanti c rif t zone , possibl y relate d t o spreading an d emplacemen t o f a mantl e plum e beneath th e lithosphere (Skogseid e t al. 1999), as the zon e o f uplif t i s concentrate d alon g th e flank o f th e rift . Risin g an d spreadin g o f mel t theoretically would hav e caused initia l inversion of th e centra l rif t zone , followe d b y progressive uplift toward s th e surroundin g flan k areas . To wha t exten t intra-plat e stres s induce d b y the Alpin e collisio n playe d a n importan t rol e is unknown. The higher smectiti c conten t i n the upper par t of th e CSS- 1 sequenc e (Fig . 11 ) indicate s weathering o f a basalti c sourc e rock . Th e mos t likely provenanc e are a fo r thes e sediment s i s the Lat e Paleocene-Earl y Eocen e basal t pro vince alon g th e Atlanti c rif t zone . Aggradatio n of sequenc e CSS-1.2 (Jordt e t al . 1995 ) indicates that suppl y fro m Norwa y wa s significantl y reduced i n earlies t Eocen e time , compare d wit h Late Paleocen e time . Reduce d runof f cause d by a climatic change, changes in land vegetation, or tectoni c quiescenc e o r erosiona l levellin g o f southern Norway , are possible mechanism s tha t could explai n a sudde n reductio n i n sedimen t supply. However, progradation o f the lower part of sequenc e CSS- 2 o n th e Hord a Platfor m an d deposition o f san d alon g th e easter n margi n o f the basi n (Fig . 12 ) indicate continue d sedimen t supply fro m th e eas t followin g depositio n o f sequence CSS-1.2 . Thi s suggest s tha t a tempor ary climati c deterioration , o r a chang e i n vegetation, ma y hav e le d to th e reduce d suppl y from th e eas t i n th e ver y earlies t Eocen e time . Subsequently, highe r runof f an d renewa l o f supply alon g th e basi n margin s ma y hav e caused renewe d outbuilding . Eastward thinnin g o f sequenc e CSS- 2 an d termination o f seismi c outbuildin g fro m th e east an d northeas t reflec t a marke d decreas e in sedimen t suppl y fro m Norwa y int o th e Norwegian-Danish Basi n durin g Earl y Eocen e time (Fig . 8) . Th e unifor m litholog y an d com position o f th e Eocen e sediment s o n th e Hord a Platform an d northward s (Rundber g 1989 ; Thyberg e t al . 2000) , an d th e velocit y distribu -

9

tion o f sequenc e CSS- 2 (Fig . 10) , sugges t a rather uniform source and sedimen t supply fro m an are a dominate d b y basalti c rocks , i.e . th e Atlantic basal t province t o the northwest. Over all, it seems that sedimen t suppl y from basemen t erosion eas t o f th e stud y are a wa s insignificant . It i s suggested , therefore , tha t th e reduce d sediment suppl y int o th e Norwegian-Danis h basin an d Nort h Se a fro m mainlan d Norwa y during Earl y Eocen e tim e was probably a resul t of tectoni c quiescenc e an d erosiona l levellin g of the provenance areas . Middle Eocen e sediment s ar e overlai n b y Lower Oligocen e sediment s i n th e norther n North Se a (Fig s 5 an d 6) , reflectin g a regiona l hiatus (Fig s 3 and 4) . A lowe r eustati c se a level during Lat e Eocen e tim e (Abre u & Anderso n 1998) ma y hav e enhance d erosio n i n th e north ern Nort h Sea , wherea s th e Eocen e successio n in th e Norwegian-Danis h Basi n appear s unaf fected b y erosion . Lowe r Oligocen e sediment s were supplie d fro m bot h th e eas t an d wes t an d thick (>500m ) wedge s wer e deposite d i n th e northern Vikin g Graben and i n the Norwegian Danish Basin . A highe r conten t o f smectit e in th e CSS- 3 sequence , compare d wit h th e sediments immediatel y belo w (Fig . 11) . sug gests inpu t fro m a ne w sourc e are a i n Earl y Oligocene time . The observe d change s i n directio n o f out building (Fig . 8 ) an d increas e i n sedimen t accumulation rat e (Tabl e 1 ) fro m Eocen e t o Early Oligocen e time , a s wel l a s th e Uppe r Eocene hiatus in the northern North Sea . cannot be easil y explaine d by eustati c sea-leve l changes or by changes in the sediment supply rate caused by climatic variations alone. I t is our interpreta tion, therefore, that there was a significan t uplif t of souther n Norwa y i n Lat e Eocene-Earl y Oligocene times , whic h wa s accompanie d b y a lower eustati c se a level . A lat e o r post-Eocen e uplift o f southern Norwa y is in accordance with the observation s b y Rii s (1996 ) of a n uplifted , deeply weathere d Eocen e erosiona l peneplai n surface acros s Fennoscandia . I t is suggested tha t lithospheric stres s induce d b y ridg e pus h force s from th e Nort h Atlanti c rif t zon e an d b y th e Alpine collisio n probabl y playe d a n impor tant role . T o wha t exten t continue d deple tion an d movemen t o f magm a belo w th e litho sphere alon g th e Atlanti c rif t playe d a rol e i s not known . Subsequent rapi d accumulatio n o f Lowe r Oligocene sediment s (Tabl e 1 ) occurred durin g a relativel y hig h eustatic sea level. However , the magnitude o f th e sea-leve l rise i n Earl y Oligo cene time was only c. 100m (Abre u & Anderson 1998), compare d wit h th e >500 m o f sediment s

Fig. 17 . NW-S E trending seismic sectio n showin g th e mappe d sequence s i n th e northernmos t Nort h Sea . Lin e location i s shown i n Fig . 2 .

CENOZOIC EVOLUTIO N O F TH E CENTRA L AN D NORTHER N NORT H SE A that wa s deposite d i n th e basin . Th e inheren t Late Eocen e bathymetr y i n th e basina l area s o f the norther n Nort h Se a wa s probabl y no t sig nificant, an d a s sedimen t loadin g ca n accoun t for onl y part of the excess accommodation spac e needed, i t i s suggeste d tha t rapi d tectoni c sub sidence occurre d i n th e centra l an d norther n North Se a during Earl y Oligocene time . No majo r change s i n compositio n o f th e sediments occurre d i n th e transitio n t o Lat e Oligocene time , durin g depositio n o f th e CSS- 4 sequence (Fig . 11) , an d n o majo r entr y point s have bee n identified . Sequenc e CSS- 4 slowl y filled in an d smoothe d th e basin-floo r topogra phy generate d b y cla y mobilizatio n an d defor mation o f sequence CSS- 3 (Fig . 9) . Incision o f Miocen e channel s int o Oligocen e strata along th e eastern margi n (Rundber g e t al. 1995; Jord t e t al . 1996 ; Gregerse n 1998 ) indi cates a shallowin g o f th e norther n Nort h Se a during Miocen e time . Th e shallowin g may b e related t o uplif t i n th e norther n Nort h Sea , a s a continuous Oligocen e t o Miocen e successio n i s preserved furthe r t o th e sout h i n the stud y area (Fig. 9) . The presenc e o f a >500m s TW T thic k CSS-5 sequenc e progradin g fro m th e west , a northward pinchou t o f sequenc e CSS- 6 an d subsequent depositio n o f sequenc e CSS- 7 fro m the east (Fig. 16) , indicate that an Early Miocen e reduction i n accommodation space was followed by a n increas e i n accommodatio n spac e dur ing lat e Miocen e time . Th e eustati c se a leve l shifted fro m a relativel y hig h leve l durin g Early an d Mid-Miocen e time s to a significantl y lower leve l i n Lat e Miocen e tim e (Abre u & Anderson 1998) . Dating o f th e Miocen e successio n i n th e northern Nort h Se a i s uncertai n (Fig . 7) , but , according to Gradstei n & Backstrom (1996), the upper Miocen e sequenc e appears to be absent in this area . Eustati c se a leve l wa s relativel y lo w during depositio n o f the CSS-7 sequenc e (Abre u & Anderso n 1998) , supportin g tha t th e uppe r Miocene sequenc e ma y b e absen t an d tha t eustatic variation s ma y hav e controlled , o r a t least influenced , deposition an d erosio n o f th e Miocene sediment s (i.e . lo w eustatic se a leve l v. deposition o f sequenc e CSS-7) . However , eus tatic variation s canno t adequatel y explai n th e change i n outbuildin g tha t occurre d durin g Miocene tim e (Fig . 16) , no r ca n the y explain the chang e fro m a matur e t o a n immatur e sediment (Fig . 11) , whic h suggest s a significan t shift i n provenanc e area . If , o n th e othe r hand , Upper Miocen e sediment s ar e present , a s indicated b y th e mixe d sedimen t composi tion (Fig . 11) , the n th e eustati c sea-leve l variations wer e ou t o f phas e wit h deposition , an d

241

tectonic movement s o f th e basi n floo r ma y have bee n a prim e contro l o n th e Miocen e development. The Pliocen e sequenc e i s characterize d b y outbuilding fro m Norwa y an d depositio n o f a major clasti c wedg e i n th e norther n Nort h Se a (Figs 1 6 an d 17) . I t is , however , uncertai n i f Lower Pliocen e sediment s ar e presen t (Fig . 7) . The thicknes s of th e Pliocen e wedg e indicates a major increas e i n accommodatio n spac e befor e outbuilding. A s depositio n o f sequenc e CSS- 8 occurred durin g a genera l eustati c sea-leve l fal l (Abreu & Anderso n 1998) , th e increas e i n ac commodation spac e canno t b e explained purely by eustac y no r b y sedimen t loading, suggesting that a significan t tectoni c subsidenc e o f th e basin floo r occurre d durin g (early ) Pliocen e times. The tectonic subsidenc e was accompanie d by a dramati c increas e i n sedimen t supply , a s a result o f glacia l erosion (Rii s 1996) . The immatur e mineralogy of sequenc e CSS- 8 (Fig. 11 ) reflect s weatherin g o f basemen t rock s that suggest s contemporaneous uplif t eas t o f the study area , o f th e Norwegia n mainland . I t i s debated i n the literature to what extent the uplif t was a resul t purely of glacial rebound, o r i f also had a tectonic component (Rohrma n e t al. 1995; Riis 1996) . Before depositio n o f th e CSS-1 0 sequence , the norther n Hord a Platfor m wa s uplifte d (Fig. 12) , causin g erosio n int o Pliocen e an d Miocene sediments . Th e timin g o f thi s relative uplift i s uncertain ; southward-dippin g reflec tions i n th e Pliocen e an d Miocen e sequence s may indicat e tha t i t occurre d befor e depositio n of sequenc e CSS-8 . Summary an d conclusions In th e presen t study , we have integrate d seismi c stratigraphic interpretatio n an d velocit y ana lyses with analyses of sediment composition. We propose tha t th e base s o f th e CSS-1 , CSS-3 , CSS-7 an d CSS- 8 sequence s correspon d t o periods o f tectonic uplif t i n th e norther n Nort h Sea. I t appear s tha t th e northern Nort h Sea was more affecte d b y uplift s tha n th e centra l Nort h Sea, suggestin g tha t th e mechanism s causin g uplift ma y hav e bee n relate d t o th e openin g o f the Nort h Atlanti c rift zone . The unconformitie s at the bas e of the CSS- 3 and CSS- 8 sequence s bot h wer e generate d during period s wit h relativel y lowe r eustati c sea levels . However , th e subsequen t sea level rise s wer e no t o f a sufficien t magnitud e (<100m, Abreu & Anderson 1998 ) to allo w the rapid depositio n o f thic k (>500m ) overlyin g

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sequences durin g earl y Oligocen e an d Pliocen e time, respectively. This suggests that, subsequent to th e thre e o r fou r uplif t events , th e basi n subsided an d th e compositio n o f th e sediment s changed a s a resul t o f uplif t i n th e sedimen t source areas . Th e mos t significan t subsidenc e event too k plac e durin g Pliocen e time , coin cident wit h a genera l eustati c sea-leve l fall . However, depositio n o f th e >500 m thic k CSS 3 sequenc e (lowe r Oligocene ) abov e a n erode d Eocene successio n i n th e norther n Nort h Se a may possibly reflect a subsidence event of similar magnitude. This interplay , between tectonic uplif t durin g eustatic low s an d subsequen t tectoni c subsi dence, canno t b e explaine d i n term s o f passiv e thermal subsidenc e followin g th e Jurassi c rif t event. Thus , a Mackenzie-typ e riftin g mode l cannot adequatel y describ e th e Tertiar y evolu tion o f the central an d norther n Nort h Sea . Th e mechanism causing the observed complex, tectonically controlle d basin-floo r movement s is probably composite . However , lithospheri c stres s induced b y ridg e pus h force s fro m th e Atlan tic rif t zon e an d b y th e Alpin e collisio n probably playe d a n importan t role , an d deple tion an d movemen t o f magm a belo w th e litho sphere alon g th e Atlanti c rif t ar e possibl e other factors . The researc h wa s funde d b y th e Commissio n o f th e European Unio n and the Norwegian Researc h Council (NFR) i n th e framewor k o f th e DGXII-Joul e Programme: Energ y fro m fossi l sources ; Integrate d Basi n Studies (IBS ) project. Nors k Hydr o a.s. , Sag a Petro leum a.s . an d Statoi l a.s . provided dat a an d stimulu s for thi s study . W e than k Nope c fo r givin g permissio n to publish their seismic sections. We are also especially indebted t o fello w colleague s withi n th e above-men tioned oi l companies wh o share d thei r knowledg e an d thoughts with us and gav e valuabl e comments.

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CAMERON, T . D . J. , BULAT , J . & MESDAG , C . S . 1993. Hig h resolutio n seismi c profil e throug h a Late Cenozoi c delt a comple x i n th e souther n North Sea . Marine an d Petroleum Geologv, 9 . 591-599. CARTWRIGHT, J . 1995 . Seismic-stratigraphica l analysi s of large-scale ridge-trough sedimentary structures in th e Lat e Miocen e t o Earl y Pliocen e o f th e Central Nort h Sea . In : FLINT , A . G . (ed. ) Sedimentary Fades Analysis. Specia l Publications of th e Internationa l Association o f Sedimentolo gists. 22 , 285-303. DEEGAN, C . E . & SCULL . B . J . 1977 . A Standard Lithostratigraphic Nomenclature for the Central and Northern North Sea. Norwegia n Petroleu m Directorate Bulletin , 1 . EIDVIN. T . & Rns , F . 1992 . En biostratigrafis k o g seismostratigrafisk analys e a v tertiasr e sedimente r i nordlig e dele r a v Norskerenna . me d hovedvek t pa ovr e pliocen e vifteavsetninger . Norwegian Petroleum Directorate Contribution. 32 . 40. GALLOWAY. W . E . 1989 . Geneti c stratigraphi c sequences i n basi n analysi s I : Architectur e and genesi s o f flooding-surfac e bounde d deposi tional units ; II : Applicatio n t o Northwes t Gul f of Mexic o Cenozoi c Basin . AAPG Bulletin. 73 . 125-154. GIBBARD. P . L . 1988 . The histor y o f th e grea t northwest Europea n river s during th e pas t thre e million years . Philosophical Transactions o f th e Royal Society o f London, Series B , 317. 559-602. GLENNIE. K . W . J990 . Outlin e o f th e Nort h Se a history an d structura l framework . In : GLENNIE . K. W . (ed. ) Introduction t o the Petroleum Geology of th e North Sea. Blackwell, Oxford . 34-77 . GRADSTEIN, F . & BACKSTROM . S . 1996 . Cainozoi c biostratigraphy an d palaeobathymetry , norther n North Se a an d Haltenbanken . Norsk Geologisk Tidsskrift. 76 . 3-32. GREGERSEN, U . 1998 . Upper Cenozoi c channel s an d fans o n 3 D seismi c data i n th e norther n Norwegian Nort h Sea . Petroleum Geoscience. 4, 67-80. HALL. B . D. & WHITE . N . 1994 . Origin o f anomalou s Tertiary subsidenc e adjacen t t o Nort h Atlanti c continental margins . Marine an d Petroleum Geology. 10 , 702-713. HARLAND. W . B. . ARMSTRONG . R . L. . Cox . A. V. . CRAIG, L . E.. SMITH . A. G . & SMITH , D. G . 1990. A Geologic Time Scale 1989. Cambridge Uni versity Press , Cambridge . ISAKSEN, D . & TONSTAD , K . 1989 . A Revised Cretaceous and Tertiary Lithostratigraphic Nomenclature for th e Norwegian North Sea. Norwegian Petro leum Directorat e Bulletin , 5. JORDT, H . 1995 . Regional Cenozoi c uplif t an d sub sidence patterns in the southeastern North Sea. In: MICHELSEN, O . (ed. ) Proceedings o f th e 2n d Symposium on Marine Geology. Geology of the North Sea and Skagerrak, Aarhus University 1993. Danmarks Geologisk e Undersogels e Seri e C . 12 . 53-68. The Cenozoic Evolution of the Central and Northern North Sea Based on Seismic Sequence Stratigraphy. Doctora l thesis . Universit y of Oslo.

CENOZOIC EVOLUTIO N O F TH E CENTRA L AN D NORTHER N NORT H SE A ,, & IBRAHIM , M . T . 1995 . Cenozoi c sequence stratigraph y o f th e Centra l an d North ern Nort h Se a basin: tectoni c development , sedi ment distributio n an d provenanc e areas . Marine and Petroleum Geology, 12 , 845-879. KNOX, R . W . O . & HOLLOWAY , S . 1992 . Paleogene o f the Central an d Northern North Sea. Britis h Geological Survey , Keyworth. MIALL, A . D . 1995 . Sequenc e stratigraph y an d chronostratigraphy: problem s o f definitio n an d precision in correlation, an d thei r implications for global eustacy . Geoscience Canada, 21 , 1-26 . MICHELSEN, O. , DANIELSEN, M., HEILMANN-CLAUSEN, C., JORDT , H. , LAURSEN , G . V . & THOMSEN , E . 1995. Occurrenc e o f major sequenc e stratigraphi c boundaries i n relatio n t o basi n developmen t i n Cenozoic deposit s o f the southeaster n Nort h Sea . In: STEEL , R. J., FELT , V. L., JOHANNESEN , E. P. & MATHIEU, C . (eds ) Sequence Stratigraphy o n th e Northwest European Margin. Norwegia n Petro leum Societ y Specia l Publication , 5, 415-427. ,, , , & 1998 . Cenozoi c sequence stratigraph y i n th e easter n Nort h Sea . In: D E GRACIANSKY , P.-C , HARDENBOL , J. , JACQUIN, T . & VAIL , P . R . (eds ) Mesozoic an d Cenozoic Sequence Stratigraphy of Western European Basin. SEP M Specia l Publication , 60 , 91-118. MITCHUM, R . M. 1977 . Seismic stratigraphy and global changes of sea level, Part 10 : glossary of terms use d in seismi c stratigraphy . In : PAYTON , C . E . (ed. ) Seismic Stratigraphy - Applications to Hydrocarbon Exploration. America n Associatio n o f Petroleum Geologists, Memoir, 26 , 205-212. ORTON, G . J . & READING , H . G . 1993 . Variability of deltaic processes in terms of sediment supply, with particular emphasi s o n grai n size . Sedimentology, 40,475-512. READING, H . G . & RICHARDS , M . 1994 . Turbidit e systems in deep-water basi n margin s classifie d by grain siz e an d feede r system . AAPG Bulletin, 78, 792-822. Rus, F . 1996 . Quantificatio n o f Cenozoi c vertica l movements o f Scandinavia b y correlation of mor phological surface s with offshore data . Global and Planetary Change, 12 , 331-357. ROBERTS, A . M., YIELDING, G . & BADLEY, M. E . 1993 . Tectonic an d bathymetri c control s o n strati graphic sequences withi n evolvin g half-graben . In: WILLIAMS , G . D . & DOBB , A . (eds ) Tectonics and Seismic Stratigraphy. Geologica l Society , London, Specia l Publications , 71 , 87-121.

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domal uplif t inferre d fro m apatit e fissio n trac k thermochronology. Tectonics, 14 , 704-718. RUNDBERG, Y. 1989 . Tertiary and sedimentary history and basin evolution of the Norwegian Sea between 60° an d 62°N. A n integrated approach. Dr.ing . thesis, Universit y o f Trondheim . , OLAUSSEN, S. & GRADSTEIN, F . 1995 . Incision of Oligocene strata ; evidence for northern Nort h Sea Miocene uplif t an d ke y t o th e formatio n o f th e Utsira sands . Geonytt, 22 , abstract. SALOMONSEN, I . 1995 . Origi n o f a dee p burie d valle y system i n Pleistocen e deposit s o f th e easter n North Sea . In : MICHELSEN , O . (ed. ) Proceedings of the 2nd Symposium on Marine Geology. Geology of the North Sea and Skaggerak, Aarhus University 1993. Danmarks Geologiske Unders0gelse , Serie C , 12 , 7-19 . SHARMA, P . V . 1969 . Earl y Tertiar y fiel d reversal s recorded i n volcani c as h layer s in norther n Den mark. Meddelelser fr a Dansk Geologisk Forening, 19, 218-223. , PLANKE , S. , FALEIDE , J . I. , PEDERSEN , T. , ELDHOLM, O . & NEVERDAL, F . 1999 . NE Atlanti c continental riftin g an d volcani c margi n forma tion. This volume. SPJELDN^S, N . 1975 . Palaeogeograph y an d facie s distribution i n th e Tertiar y o f Denmar k an d sur rounding areas . In : WHITEMAN , A. , ROBERTS , D . & SELLEVOLL , M . A . (eds ) Petroleum Geology and Geology of the North Sea and Northwest Atlantic Continental Margin. Norge s Geologisk e Under sokelse Bulletin , 29, 289-311. STEURBAUT, E. , SPIEGLER , D. , WEINELT , M . & THIEDE, J . 1991 . Cenozoic Erosion and Sedimentation on the Northwest European Continental Margin. Geomar , Kiel . STUEVOLD, L. & ELDHOLM, O. 1996 . Cenozoic uplif t of Fennoscandia inferre d fro m a stud y o f th e mid Norwegian margin . Global and Planetary Change, 12, 359-386. , SKOGSEID , J . & ELDHOLM , O . 1992 . Post Cretaceous uplif t event s o n th e V0rin g continen tal margin . Geology, 20, 919-922. THYBERG, B . I., JORDT, H., BJORLYKKE, K. & FALEIDE , J. I . 1999 . Relationshi p betwee n sequenc e strati graphy, mineralog y an d geochemistr y i n th e Cenozoic sediment s o f th e norther n Nort h Sea . This volume. ZIEGLER, P . A . 1982 . Geological Atlas o f Western an d Central Europe. Elsevier , Amsterdam . 1990. Tectoni c an d palaeogeographi c develop ment o f th e Nort h Se a rif t system . In: BLUNDELL, D. J. & GIBBS , A. D. (eds ) Tectonic Evolution of the North Se a Rifts. Clarendo n Press , Oxford , 1-36.

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Relationships between sequence stratigraphy , mineralog y and geochemistry in Cenozoic sediment s of th e northern North Se a B. I . THYBERG, 1 2 H . JORDT, 1

12

K. BJ0RLYKKE1 & J . I . FALEIDE 1

Department of Geology, University of Oslo, P.O. Box 1047 Blindern, N-0316 Oslo, Norway 2 Present address: Statoil, Box 300, 4001 Stravanger, Norway Abstract: Cutting s o f Cenozoi c sediment s fro m selecte d well s in th e norther n Nort h Se a have bee n analyse d bot h mineralogicall y and geochemically , an d compare d wit h seismi c sequence stratigraphi c units . Variation s i n th e compositio n wit h respec t t o cla y minerals , feldspars, majo r element s and selecte d trac e elements hav e bee n use d to establish change s in the provenance of the Cenozoi c sequences. Th e Eocen e sequence consists o f a thicknes s o f several hundred s o f metre s o f almos t exclusivel y smectiti c clays with high N i an d Z n con tents and littl e or no quartz. This represents a strong input from subaeria l volcanism related to the rifting an d initial spreading of the Norwegian-Greenland Sea. The Eocene-Oligocene smectitic mudstone s hav e high porosity an d low seismic velocity (<2 km s"1) compared with the underlyin g Paleocene sediment s an d overlyin g Miocene-Pliocene sequences . Th e Pli o cene an d Pleistocen e mudstone s ar e coarse-graine d clay s wit h a n immatur e mineralogy , which compac t muc h mor e rapidl y tha n th e underlyin g smectitic sediments . Th e seismi c velocity ma y the n reac h 2.5-3. 0 km s"1 an d thu s caus e a stron g velocit y inversion . Thi s sequence i s partly glacial an d mineralogically immature , bu t may contain some kaolinite and smectite, whic h indicat e reworking o f lowe r Tertiar y an d Mesozoi c sediments . Ou r obse r vations indicate that sequenc e boundarie s ar e closely related t o tectoni c movements and t o changes i n th e sedimen t suppl y fro m adjacen t areas . Bot h mineralogica l an d geochemica l data ca n b e used a s additiona l tool s for correlatio n i n Cenozoic sediment s i n the norther n North Sea . Th e mineralogy o f these sediment s als o determine s th e rate of compaction an d other factor s relevant for basi n modelling.

The Cenozoi c successio n in the North Se a Basin was one of the main examples used when seismic stratigraphy was introduced (Vai l et al. \971a, b) and severa l sequenc e stratigraphi c studie s fro m this are a hav e subsequentl y been publishe d (e.g. Nielsen e t al. 1986 ; Stewart 1987 ; Rundber g 1989; Gallowa y e t al . 1993 ; Jone s & Milto n 1994; Michelsen 1994 ; Jordt et al. 1995 ; Michelsen e t al . 1995) . Th e initia l interpretatio n o f these dat a relate d th e seismi c stratigraphic units to change s i n eustati c se a leve l (Vai l e t al . 19770, b). However , recen t studie s of th e Nort h Sea Basin suggest that th e seismic sequences and sequence boundarie s ar e mainl y controlle d b y tectonic processe s (e.g . Rundber g 1989 ; Gallo way e t al . 1993 ; Jordt e t al . 1995 , 1999) . Thes e tectonic processe s caus e subsidenc e (accommo dation space) , a s well a s uplif t an d erosion , an d thus t o a larg e extent control th e sedimentation into a basin suc h as the North Sea . B y contrast, sediment suppl y i s no t directl y relate d t o sea level changes . I t i s dominantl y influence d b y environmental factors , includin g relief , climate , the drainage syste m and th e biogenic productivity o f th e ocean . Th e clasti c sedimen t suppl y is

controlled b y th e rat e o f erosio n an d weather ing i n th e sourc e area , itsel f a functio n o f th e drainage an d relie f generated b y tectoni c uplift . The type of clay mineral found in a shale is thus a functio n o f provenanc e (roc k type) , climat e (weathering), transpor t (facies ) an d diageneti c history o f the sequenc e (Weave r 1989) . The main objectiv e o f this paper i s to analys e and establis h the relationshi p between sediment composition an d seismi c sequenc e stratigraphi c units. Specia l emphasi s ha s bee n place d o n th e relationship betwee n th e sedimen t compositio n and th e Cenozoi c seismi c sequence stratigraphi c framework establishe d b y Jord t e t al . (1995) . The presen t pape r als o provide s a contributio n to th e discussio n o f th e origi n o f th e strati graphic sequence s an d thei r development . The boundarie s o f seismi c sequence s ar e identified b y physical variables, such a s changes in th e acousti c impedance , define d a s th e prod uct o f seismi c velocity and bul k density , and b y differences i n the amplitude an d frequenc y of the signal. A s th e seismi c velocit y i s significantl y influenced b y primary mineral composition, texture an d diageneti c processes , whic h ar e als o

From: NOTTVEDT , A . e t al . (eds ) Dynamics o f th e Norwegian Margin. Geologica l Society , London , Specia l Publications, 167 , 245-272 . 1-86239-056-8/OO/ S 15.00 © Th e Geologica l Societ y o f Londo n 2000 .

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reflected i n fine-grained sediments , th e natur e of seismic reflection s i s a direc t functio n o f th e lithology (e.g . Rundber g 1989) . I n spit e o f this , sequence stratigraphi c studie s hav e almos t exclusively bee n base d o n seismi c profiles , bio stratigraphic dat a an d petrophysica l wel l logs . Clay mineralogy has been use d a s a basis fo r th e stratigraphy o f barren unit s of Triassic, Permia n and Devonia n re d bed s i n U K Quadran t 9 an d as a stratigraphi c too l t o distinguis h Mesozoi c and Palaeozoi c roc k unit s (Glasman n & Wilkinson 1993) . Fishe r & Jean s (1982 ) divide d th e Permo-Triassic re d bed s int o eigh t cla y mineral zones, an d date d an d correlate d a Permo Triassic sequenc e expose d o n th e sout h Devo n coast an d a sequenc e o f re d bed s penetrate d in BNO C 72/10- 1 A, som e 500k m awa y i n th e Western Approaches , b y usin g facies-relate d clay minera l assemblages . I n addition , Henso n (1974) describe d variou s compositiona l change s in th e cla y mineral assemblage s relate d t o litho facies. Rundber g (1989 ) integrate d geophysica l data an d mineralogica l analyse s t o reconstruc t

climatic an d tectoni c change s i n th e Ceno zoic sectio n i n th e norther n Nort h Sea . Triassi c sediments fro m seve n well s i n th e Centra l Graben als o sho w clearl y identifiabl e geochem ical differences , whic h have bee n relate d t o shif t in source s an d transpor t regime s (Race y e t al 1995). Th e presen t stud y compare s seismi c sequence stratigraphi c unit s wit h th e mineralo gical compositio n o f th e fine-graine d Cenozoi c sediments.

Database an d methods Mineralogical an d geochemical dat a obtained b y X-ray diffractio n (XRD ) an d X-ra y fluorescence (XRF) o f washed dril l bi t cutting s of wells 24/61, 26/4-1 , 30/3-3 , 34/7-1 , 34/7- 2 an d 34/7- 6 ar e presented in this paper (Fig. 1) . They ar e selected from a very large mineralogical and geochemica l database establishe d durin g thi s study . Th e sampling interva l wa s betwee n 3 0 an d 100m . XRD analyse s are based o n th e bul k sample s a s

Fig. 1 . Structura l ma p o f th e stud y are a showin g location o f wells analyse d mineralogicall y an d geochemically . Seismic data coverage and depocentr e locations also shown. Arrows indicate outbuilding directions (Jordt er al. 1995). Th e number s refe r t o seismostratigraphi c unit s denned i n Fig . 2.

SEQUENCE STRATIGRAPH Y I N CENOZOI C SEDIMENT S well a s th e <2/m i fraction , whic h hav e bee n taken ou t t o indicat e th e cla y minera l composi tion o f some selected wells. Unoriented powder s of bul k sample s wer e analyse d o n a Phillip s PM1700/1710 difTractomete r wit h CuK a radia tion (1.54060) . Diffractio n diagram s wer e col lected fro m 2 t o 50°2# . Additiona l slo w sca n runs (26.0-28.5°2# ) wer e carrie d ou t routinely , to obtai n a bette r identificatio n o f feldspars . Peak intensitie s used fo r estimatin g weight percentages ar e fro m th e diffracte d intensitie s (DI) listings given by the Phillips APD software . Peak intensities fo r 'al l clays' , quartz , feldspars , car bonates an d pyrite s are weighted using intensity factors fro m Ram m (1991) . The clay fraction (<2 /mi) was separated fro m the materia l b y suspensio n settlin g followin g dispersal b y ultrasoni c disaggregation. Oriente d samples of the < 2 /mi fraction were prepared b y filtering th e sample s throug h a Millipor e filte r and the n transferrin g th e materia l t o a silic a slide b y inverting . XR D diagram s fro m 2 t o 50°2# wer e obtained , afte r th e followin g treat ment; sample s were Mg saturate d an d air-dried , then ethylene glycolated an d heate d t o 550° C in 2h. Slow-sca n diagram s betwee n 2 4 an d 26°20 on air-drie d specimen s wer e obtaine d routinel y to separat e the d(002) kaolinite and d(004) chlorite reflections . The cla y mineral s wer e categor ized int o 'expandabl e clays ' (<14 A mineral s on ethylen e glycolate d runs ; referre d t o belo w as smectite) , illite , kaolinit e an d chlorite . Th e weight factor s use d wer e base d o n Pearso n & Small (1988 ) and Pearso n (1990) . The inorgani c geochemica l analyse s obtaine d by XRF o f the cuttings were also included in this study. The samples were analysed with respect to major elements such as SiO2, A12O3, MnO, MgO , CaO, Na 2O, K 2O, TiO 2 an d P 2O5, an d los s of ignition (LOI ) ha s been calculated , expresse d in weight percen t o f th e oxide . Th e su m o f majo r elements is in most case s 100 % but , mainl y as a result o f contaminatio n o f drillin g mud, value s of <100 % hav e bee n obtained . Sample s wer e analysed fo r trac e element s (V, Cr, Co , Ni , Cu , Zn, Rb , Sr , Y Zr, Nb, Pb , Th and U) , presented in parts per million (ppm) . In this study we have presented th e Co , Ni , C u an d Z n content s i n selected wells. Major element data wer e obtained from crushe d materia l drie d a t 100° C fo r a t least 12h , mixe d with lithium tetraborate i n th e ratio o f 9:1 , an d the n fuse d t o 1000° C ( 1 h). The trac e elemen t conten t wa s determine d b y using pressed pellet s made b y mixin g 1 0 g finel y ground materia l wit h 2m l Elvacit e solution . Analysis wer e performe d i n a Phillip s PW240 0 spectrometer. Th e USG S (U S Geologica l Sur vey) standard s wer e used fo r calibration .

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Cenozoic seismi c stratigraphy The results of the geochemical analyses have been correlated wit h th e Cenozoi c seismi c sequenc e stratigraphic framewor k establishe d b y Jord t et al. (1995) , wh o divide d th e Cenozoi c suc cession o f th e Nort h Se a Basi n int o te n main sequences , CSS- 1 t o CSS-1 0 (Fig . 2) . Th e CSS-1 sequenc e i s o f Lat e Paleocene-earlies t Eocene ag e (60.5-5 6 Ma). Th e to p CSS- 2 i s dated t o th e Eocene-Oligocen e transitio n (35.4 Ma). Oligocen e sediment s correspon d t o the CSS- 3 an d CSS- 4 sequences . Th e Miocen e succession i s divide d int o thre e seismi c se quences: CSS-5, CSS-6 and CSS-7 . Top CSS- 5 is of lat e Earl y Miocen e ag e an d to p CSS- 6 is dated t o earl y Mid-Miocene . A Mid - t o Lat e Miocene age is assigne d the CSS- 7 seismi c sequence. CSS- 8 i s o f Pliocen e age , an d th e Quaternary sediment s correspon d t o CSS- 9 an d CSS-10. Approximate location s of the Cenozoi c depocentres i n the norther n Nort h Se a and cor responding outbuildin g direction s o f th e differ ent seismi c sequences , base d o n Jord t e t al . (1995), ar e show n i n Fig . 1 .

Cenozoic sedimen t composition The feldspa r conten t o f bot h shale s an d sand stones i s a goo d indicatio n o f a rock' s miner alogical maturity , which may be relate d t o roc k type, weatherin g an d transpor t a s detrita l feld spar. K-feldspa r i n particula r may , however , be dissolved b y dee p burial , whic h reduce s it s effectiveness a s a provenanc e indicator . The Al 2O3/SiO2 rati o i s used a s an indicatio n of th e clay/san d rati o (e.g . Bjorlykk e 1974) , a s A12O3 i s mostl y foun d i n clays , feldspa r an d mica. Th e iro n conten t o f th e sample s ca n b e related t o iron-ric h chlorite , pyrit e o r siderite . Iron-rich smectit e ha s als o bee n reporte d i n Paleocene an d Eocen e sediment s (Tyridal 1994) . The amoun t o f calciu m presen t i n silicat e min erals i s generall y smal l i n sediment s fro m th e North Sea , an d sandstone s withou t carbonat e minerals contai n les s than 2-3% CaO . The cla y mineralogy at th e tim e of depositio n can b e relate d t o th e sedimen t provenanc e an d facies (e.g . Weaver 1989) . This includes no t onl y the nature of the erosion o f the sourc e rock , bu t also its rate and the processes of rock weathering acting i n differen t climates . Kaolinit e form s i n soils develope d unde r abundan t rainfall , goo d drainage and i n shallow or coastal marin e sand stones durin g meteori c wate r flushing . I n drie r climates, illit e an d smectit e becom e important , although volcani c as h an d othe r volcani c rock s

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Fig. 2. Cenozoi c seismic sequence stratigraphic units (Jordt et al. 1995 , 2000) correlated wit h Chronostratigraphy and lithostratigraph y fro m Deega n & Scul l (1977) . are additiona l importan t source s o f smectite . Illites in the basin, however, are probably derive d from weatherin g o f pre-existin g muscovite s o r illites, bu t ca n als o b e forme d diageneticall y b y the transformatio n o f smectit e t o illit e o r kaoli nite t o illite . Chlorit e form s fro m pre-existin g chlorites bu t ma y b e forme d diageneticall y with smectite o r kaolinit e a s a precurso r mineral . The cla y minera l siz e distributio n relativ e t o the proximit y o f th e sourc e ha s been use d b y several workers (e.g. Porrenga 1967 ; Millot 1970; Gibbs 1977) . Some of the principles are summar ized int o a schemati c an d simplifie d representa tion o f the theoretica l mineralogica l respons e t o the sequenc e stratigraphi c framewor k i n mud rocks, reflectin g periods o f highstan d an d low stand se a leve l i n a sequenc e developmen t (Fig. 3) . Thi s illustrate s th e distributio n o f th e clay mineral s during progradation , wit h period s of highstan d ideall y characterize d b y a highe r concentration o f fine-grained clay minerals, suc h

as illit e an d smectite , wherea s kaolinit e i s enriched i n a mor e proxima l facies , a s th e kaolinite crystal s ten d t o b e large r tha n thos e of illit e an d smectite . A starve d interva l i n th e basinward directio n i s evidence d i n th e mos t fine-grained par t o f th e sequence . Condense d sections (e.g . Louti t e t al . 1988 ; Galloway 1989 ) are laterall y extensive thi n layers . A condense d sequence ma y b e represente d b y a dark , fine grained shal e or marl , often with a hig h content of biogenic carbonate an d silica . Relative enrichments o f phosphoru s an d manganes e (Fig . 3 ) probably indicat e ver y lo w sedimentation rates .

Upper Paleocene-lowermost Eocene sequence (CSS-1) The cla y mineralog y distributio n o f CSS- 1 shows a hig h amount o f smectite , bu t als o chlo rite, wit h smaller amounts o f illit e an d kaolinit e

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Fig. 3. Conceptua l model fo r latera l variation s in clay mineralogy in mudstones. Illit e an d smectite , bein g finer grained, are enriche d i n th e dista l facie s compare d wit h kaolinite . detected b y XR D (Fig s 4 an d 5) . Rundber g (1989) also documente d thi s i n Paleocen e sedi ments fro m th e sam e region . Th e CSS- 1 sediments ar e variabl e wit h respec t t o thei r feldspar conten t an d th e K-feldspar/plagioclas e ratio, althoug h plagioclas e i s dominant ove r K feldspar i n most o f the analysed wells (Fig. 6) , as is also reflecte d i n th e K 2O/Na2O ratio (Fig . 7). Close t o th e top o f CSS-1, a n increas e in plagioclase conten t i s ofte n observed . Th e erosio n o f a plagioclase-rich basement sourc e could explai n this. However , som e plagioclas e ha s volcani c origin, a s Mal m e t al. (1984 ) an d Morto n & Knox (1990 ) hav e reporte d o n plagioclase-ric h xeno-crystals i n th e Balde r tuff . CSS-1 sediment s hav e relativel y hig h Fe 2C>3 and TiO 2 content s (Fig . 8) ; mea n value s oi f Fe2O3 an d TiO 2 ar e 8. 7 an d 1.17wt % respec tively. Th e averag e Mg O conten t i s 2.6wt% . Carbonate-cemented interval s ar e also rathe r characteristic (Fig s 9 an d 10) , an d som e o f th e high iro n conten t ca n b e relate d t o anker ite layer s confirme d b y XRD . Relativel y hig h metal concentration s ar e see n i n thi s sequenc e (Fig. 11) , an d th e averag e value s for Ni , C u an d Co ar e onl y lowe r tha n thos e o f CSS- 2 an d CSS-3, an d ar e muc h highe r tha n i n th e upper most sequences . A n increas e i n th e concen tration o f thes e trac e element s clos e t o th e

CSS-l-CSS-2 boundary , especiall y wit h respec t to nicke l and zinc , can be see n (Fig . 11) . Thi s is not foun d i n wel l 26/4-1 , however , a s CSS- 1 contains mainl y san d an d sample s analyse d ar e not fro m tuff-beds . Th e silic a conten t i n th e shales mainl y varie s betwee n 4 5 an d 60wt% , which i s clos e t o typica l value s fo r mafi c rock s (e.g. Eart h 1951) . However , i n th e sand-ric h units o f CSS- 1 i n well s 26/4- 1 an d 24/6-1 , a higher silic a content i s observed . Paleocene sediment s underlyin g th e Balde r Formation ma y hav e a relativel y hig h smectit e content, a s well a s high Fe , Mg , Cu , N i an d Z n content, indicatin g mafi c sourc e rocks . Thes e sediments ma y b e derive d fro m basi c basemen t rocks o r erosio n o f Mesozoi c sediment s fro m surrounding areas . A suppl y o f volcaniclasti c material fro m th e N W ma y als o explai n thi s distribution. The content o f chlorite an d a weak increase i n illit e i n CSS- 1 (Fig s 4 an d 5 ) ca n be attributed t o diageneti c transformation s wit h smectite a s th e precurso r mineral , o r th e disso lution o f roc k fragments . Top CSS- 1 probabl y indicate s a condense d section, wit h hig h concentration s o f MnO , carbonate an d a hig h concentratio n o f fine grained 'expandable ' clay s an d sometime s P 2O5 near th e CSS-l-CSS- 2 boundary . Thes e dat a indicate period s o f lo w clasti c sedimen t suppl y

Fig. 4. Cla y minera l distribution of some wells in the norther n Nort h Sea . Th e seismi c sequence stratigraphi c units are characterized b y different cla y mineral composition and particularl y wit h respec t t o th e smectit e content . A lo w smectit e an d hig h kaolinil e conten t i n th e Lowe r Oligocen e sequenc e (CSS-3 ) i n easter n wel l 30/3- 3 ma y indicate a mor e proxima l fades . Locatio n o f well s is show n i n Fig . 1 .

Fig. 5 . Cla y mineral distributio n o f well s 24/6- 1 an d 26/4-1 , correlated wit h seismic section CNST82-1 0 an d th e Cenozoi c seismi c sequence stratigraphi c framework Depths ar e relativ e t o mea n se a leve l accordin g to Jord t e t al. (1995) . Locatio n o f well s an d seismi c line i s shown i n Fig . 1 .

Fig. 6. Feldspa r conten t analysed by XRD an d relate d to the seismic sequence stratigraphic framework. Locatio n of well s is shown i n Fig . 1 . Clasti c feldspa r ma y b e use d a s a n indicato r o f provenanc e an d weathering . Th e K-feldspar conten t i s low in all wells except i n the Pliocen e sediments, where it reflects a colder climate. The albit e content i s also lo w i n Miocen e an d olde r sediment s excep t i n wel l 26/4-1 . I t i s possible tha t albit e to som e extent ma y b e authigeni c i n th e Eocen e sediment s wit h a volcani c source , mrkb . metre s belo w Kell y bushing.

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and a starved depositional environmen t (Fig. 3). The enrichmen t o f manganese, phosphoru s an d zinc i s therefore probably du e t o a hig h relative organic productivity , particularl y o f siliceou s and calcareou s organisms , thoug h S0rense n & Nielsen (19810-c ) have interpreted th e presenc e of rhodochrosit e i n Paleocen e sediment s a s a result o f the suppl y o f Mn durin g volcanism . The Balde r Formation , correspondin g t o th e upper part of sequence CSS-1 , contains siliceou s shales and numerou s tuffaceous zone s and layers (Jacque & Thouveni n 1975 ; Malm e t al. 1984) . These interbedde d shale s and tuf f layer s contain a hig h amoun t o f microcrystallin e quart z an d have a ver y lo w conten t o f detrita l mineral s (Malm e t al . 1984) . Explosive volcanis m in th e British an d th e Faeroe-Greenlan d Tertiar y volcanic province s hav e bee n regarde d a s th e source fo r th e volcani c ashe s foun d widesprea d in th e Nort h Se a (Mal m e t al . 1984 ; Kno x & Morton 1988 ) and onshore Denmark (Spjeldnae s 1975; Nielse n & Heilmann-Clausen 1988) .

Eocene sequence (CSS-2)

Fig. 6 (continued}

Eocene (CSS-2 ) sediment s consis t mostl y o f smectite i n th e stud y are a (Fig s 4 an d 5) . Petrographical observation s sho w ver y lo w quartz content . Silic a content ha s a mea n value of 50 wt% (Fig . 7 ) and th e XRD dat a show also little quart z (mainl y les s tha n 20wt% ) an d feldspar (0-1 0 wt%) . Magnesium has an average value o f 2.5wt % MgO . Th e Eocen e mudrock s are ric h i n aluminiu m (15wt % A\ 2O^), iro n (9.5 wt% Fe 2O3) an d titaniu m (1.14wt % TiO 2). This i s typica l o f basi c volcani c sediments , an d the dry bul k compositio n i s not fa r fro m that o f Icelandic basalt s (e.g . Barth 1951 ) and indicate s that volcaniclasti c materia l represent s a domi nant par t o f th e sedimentatio n i n CSS-2 . Th e high conten t o f th e trac e metal s suc h a s nickel , cobalt, zin c an d coppe r (Fig . 11 ) i s als o con sistent wit h a volcani c origin. Eocene sediments are sodiu m ric h an d thi s ma y possibl y b e be cause of the presence of authigenic albite formed in volcani c sediments . Enrichmen t o f phos phorus and/o r manganes e i s als o see n clos e t o the CSS-2-CSS-3 boundary o r within the CSS-2 sequence in som e wells , th e bes t examples being wells 24/6-1 and 34/7-6 . The fine-graine d CSS- 2 sediments an d indication s o f lo w rate s o f non volcanic clasti c sedimen t suppl y indicat e a rela tively dee p marin e facie s during CSS- 2 time s in the Nort h Se a Basin. The Paleocen e an d Eocen e sediment s i n th e North Sea are highly smectitic (Figs 4 and 5) , and

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Fig. 7 . Distributio n o f some majo r element s (SiO2 + A1 2O3 and Na 2O + K 2 O) based on XRF analyse s of cuttings and relate d t o th e seismi c sequence stratigraphi c framework . Locatio n o f well s is shown i n Fig . 1 . (Not e th e lo w silica conten t an d hig h aluminiu m content i n th e smectite-ric h Eocen e sediment s (CSS-2). )

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Fig. 7 (continued)

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Fig. 7 (continued)

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SEQUENCE STRATIGRAPH Y I N CENOZOI C SEDIMENT S this has also by previous workers been related t o a volcani c sourc e (Nielse n 1974 ; Karlsson e t al . 1979; Sorense n & Nielse n 1981a-c ; Rund berg 1989 ; Pearso n 1990 ; Hugge t 1992) . Th e quartz content woul d b e expected t o b e higher if erosion o f the adjacent landmasse s wa s the main source o f th e basi n sediments . Th e cla y mineral analysis an d th e trac e an d mai n elemen t dat a also sho w that thi s is not th e case. A higher ratio of kaolinite-illit e + smectite i n sandstone s i n well 24/6-1 indicates this well to be located closer to th e sedimen t sourc e tha n th e othe r well s i n the stud y are a (Ziegle r 1982 ) (Figs 3 and 5) .

Lower Oligocene sequence (CSS-3) The cla y minera l distributio n (Fig s 4 an d 5 ) shows tha t smectit e remain s th e mai n cla y mineral componen t i n th e Lowe r Oligocen e sedi ments (CSS-3). However, the content of kaolinite is higher, particularly in wells 30/3-3 and 24/6-1, which represen t a mor e proxima l facie s clos e to a n easter n an d wester n source , respectivel y (Fig. 1) . Well 26/4-1 , locate d mor e i n th e dista l part o f th e basin , ha s a correspondingl y highe r content o f smectit e compare d wit h kaolinite . Thus, th e regiona l cla y minera l distributio n i n CSS-3 tim e is also controlle d b y the basi n development an d sedimen t transpor t directio n a s indicated i n Fig . 3 . In wel l 30/3- 3 a pronounce d shift i n cla y minera l assemblage s nea r th e CSS-2-CSS-3 boundary occurs, fro m a smectiterich sedimen t t o a cla y mineral composition en riched i n kaolinite above th e sequence boundar y (Fig. 4) . However , Rundber g (1989 ) indicated a higher smectit e conten t i n wel l 30/3- 3 i n th e depth interval corresponding t o CSS- 3 i n Fig. 4. The K-feldspa r an d plagioclas e conten t i s broadly similar and low throughout the sequence (Fig. 6) . Th e silic a conten t ha s value s clos e t o 59wt% an d A1 2O3 conten t clos e t o 13wt% . A sligh t increas e i n silica , an d correspondin g decrease in alumina content, is usually seen close to th e transitio n fro m Eocen e t o Oligocen e sediments (e.g . wel l 30/3- 3 (Fig . 7b ) an d 34/7- 2 (Fig. 7d)) , thoug h thi s i s not alway s the case , a s illustrated b y wel l 34/7- 1 (Fig . 7c) . Th e sedi ments i n th e lowe r par t o f CSS- 3 i n wel l 34/7-1 have silic a conten t o f abou t 48wt% , whic h indicates a basi c composition . I n wel l 24/6- 1 (Fig. 7a) , th e highe r silic a conten t i s du e t o a more sand y litholog y compared wit h th e under lying section . Th e iro n conten t i n th e investi gated well s varie s widely , betwee n 4. 4 an d 8.27wt% Fe 2O3, an d i s enriche d i n smectiti c mudstones (Fig . 8) . Thi s variatio n i s als o exhibited b y titanium . Aluminiu m an d titaniu m

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contents are , however , lowe r than i n the Eocen e sediments (Fig . 8) . The trac e meta l value s (Ni , Co , C u an d Zn ) are almos t alway s lowe r tha n i n th e under lying CSS- 2 (Eocene ) sediments , bu t ar e stil l relatively hig h (Fig . 11) . However , th e presenc e of smectite , an d th e iron , titaniu m an d trac e element (Ni , Co, C u and Zn ) contents indicate a high volcani c componen t i n th e CSS- 3 (Lowe r Oligocene) sediments . Continue d inpu t o f vol canic ash fall s o r reworkin g of volcanic material is probabl y th e explanatio n fo r thes e observe d trends. Th e sedimentatio n rat e increase d durin g Early Oligocen e time , probabl y i n respons e t o Oligocene uplifts an d erosio n of soft Eocen e an d Paleocene sediments . Thi s also explain s the hig h sedimentation rate s foun d b y Jordt e t al . (1995, 1999) fo r thi s period . N i an d Z n sulphide s ar e easily oxidize d when th e sediment s ar e exposed , and wil l quickl y b e depleted . Th e reduce d amount o f the trac e metal s therefor e ma y b e a n indication o f reworked volcanic material. In well 34/7-1, trac e elements , iro n an d titaniu m values are simila r t o thos e obtaine d fro m CSS- 2 sequence (Fig. 1 Ic). The highe r content o f kaolinite (Fig . 4 ) may indicat e more sedimen t suppl y from land . Extensivel y reworke d Earl y Paleo cene and Earl y Eocene as h fall s an d lavas , yielding glass-ric h sediments , thu s occu r throughou t Late Eocen e t o Miocen e tim e (Hugge t 1992) , and Pearso n (1990) interpreted sediment s from a volcanic sourc e a s fa r u p a s Lowe r Oligocene .

Lower-Upper Oligocene sequence (CSS-4) Smectite i s a majo r componen t i n th e cla y fraction i n well s 34/7-1 , 34/7-2 , 34/7-6 , 34/7- 2 (Fig. 4 ) and 26/4- 1 (Fig. 5) . Kaolinite i s also a n important constituen t i n thes e sediments , parti cularly i n th e proxima l part s o f th e basi n (wel l 24/6-1, Fig . 5 ) identifie d b y progradatio n fro m the wes t in CSS- 4 tim e (Figs 1 and 5) . However, in wel l 34/7- 1 thes e sediment s hav e lowe r smec tite an d highe r illit e content s (Fig . 4) . Th e feld spar conten t i s approximatel y th e sam e a s i n the underlyin g Oligocene an d Eocen e sediment s (Fig. 6) , wit h respec t t o bot h plagioclas e an d K-feldspar. Th e sand y litholog y i n wel l 24/6- 1 has a relativel y hig h silic a conten t (Fig . 7a) ; otherwise th e silic a conten t i s abou t 50wt% . Differences i n lithology and facie s ar e als o indi cated i n variable s suc h a s aluminiu m an d iro n contents, wit h the highes t iro n conten t foun d i n well 34/7-6 , which als o ha s th e highes t accumu lation o f smectit e (Fig . 4) . Th e lowes t iro n content i s i n wel l 24/6-1 . Otherwis e averag e iron conten t fo r CSS- 4 i s abou t 4.4-4.9wt %

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Fig. 8 . Distributio n o f som e majo r element s (Fe 2O3 + Mg O an d TiO 2) base d o n XR F analyse s o f cuttings and related t o th e seismi c sequenc e stratigraphi c framework . Location o f wells is shown i n Fig. 1 . (Note th e hig h iron content o f th e smectite-ric h mud.)

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Fig. 8 (continued]

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Fig. 8 (continued)

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Fig. 9. Relationship s between seismic stratigraphic units, velocity and cla y mineralogy in well 30/3-3. Location o f well and seismi c line is shown in Fig. 1 . (Note changes i n velocity nea r th e boundar y betwee n seismi c units CSS- 3 and CSS-4 , probabl y cause d b y th e difference s i n smectit e content. )

Fig. 10 . Relationship s betwee n seismi c slruligruphi c units , velocit y an d cla y mineralog y i n wel l 34/7-1 . Locatio n o f wel l and seismi c lin e is shown i n Fig . I . I t shoul d h e noted tha i th e Pliocen e sediments, which arc mineralogicall y i m m a t u r e with lo w smectite content, sho w a stron g increas e i n velocity as a resul t of progressiv e compaction . This i s no t th e cas e wit h th e Eocen e an d Oligoccn e smectite-ric h sediments .

SEQUENCE STRATIGRAPH Y I N CENOZOI C SEDIMENT S Fe2O3. Th e hig h Ca O conten t i n wel l 30/3- 3 is due t o carbonate-cemente d intervals , a s indi cated b y XRD (Fig . 9) . Both clos e t o th e CSS-3-CSS- 4 sequenc e boundary an d withi n th e CSS- 4 sequenc e ther e is a n upwar d decreas e i n Fe , T i an d th e trac e metals Ni, Co , C u an d Zn , an d increas e in silica content (Fig s 7 , 8 an d 11) . Th e trac e metal s Co, Ni , C u an d Z n hav e generall y slightl y lower value s i n th e CSS- 4 tha n th e underly ing CSS-3 (Fig . 11) ; this finding also reflect s tha t CSS-4 ha d a simila r sedimen t compositio n t o that of CSS- 3 throughou t the Oligocen e epoch . However, th e mos t pronounce d reductio n i n trace metal s suc h a s Co , Ni , C u an d Z n ca n b e seen in well 34/7-1 (Fig. 1 Ic), which also shows a decreasing smectit e content clos e t o th e CSS-3CSS-4 boundar y (Fig s 4 and 10) . The highes t concentratio n o f biogenic silic a is found i n the basa l par t of sequence CSS- 4 (wells 30/3-3, 34/7- 1 an d 34/7-6) . Th e biogeni c silic eous facie s i s mainly dominated b y diatoms, bu t sponge spicule s ar e als o commo n an d radio laria ar e present a s minor component s (Thyber g et al 1999) . Thyberg e t al . (1999) , on th e basi s of detaile d diato m flor a i n CSS-4 , considere d that th e siliceou s sediment s ma y no t represen t the same stratigraphi c horizons , indicate d b y the presence o f diatoms bot h in Eocene i n the northernmost studie d wel l (36/1-2) , an d a low diatom content, mostl y fragments , i n CSS- 5 (Lowe r Miocene) i n wel l 26/4-1 . A n exampl e o f th e diatom assemblage s an d commonl y observe d 'silica aggregates ' i s shown i n Fig . 12 . The sequenc e CSS- 4 ha s a ric h diato m flor a with the dominanc e of Paralia species , Paralia thybergii (Stabel l 1996) , possibl y indicatin g a nearshore settin g compare d wit h th e dominan t Paralia sulcata foun d i n coasta l setting s i n general an d i n the Nort h Se a during the presen t (Stabell 1985 ; Stabell & Lange 1990) . Diatomac eous facie s interbedde d wit h glauconiti c facie s (Fig. 13 ) therefore sugges t that shallower marin e conditions prevaile d i n CSS- 4 tim e tha n i n earlier Cenozoi c time . Poorl y sorted , imma ture clasti c an d angula r grain s observe d i n wel l 34/7-1 also indicate a short-distance transport of clastic material , whic h interfinger s wit h shallow marine biogeni c and glauconiti c sediments . Thi s input o f clasti c fragment s indicate s structura l local high s within , o r clos e to , th e basin. Rund berg (1989 ) suggeste d a n easter n sourc e fo r th e sediments, o n th e basi s o f a sand y prograda tional syste m in the uppe r par t o f the Oligocene sequence i n th e Aga t area . An eve n sedimen t thicknes s i n th e norther n North Se a Basin has been interprete d t o indicat e low clasti c suppl y fro m th e lan d an d a reduce d

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topographic relie f onshor e (Jord t e t al . 1995 , 1999). Thi s i s supporte d b y th e stud y o f th e sediment compositio n i n th e CSS- 4 sequence , which indicate s sediment-starve d depositio n i n the basin . Thus , th e mineralog y o f CSS- 4 an d the aggradin g seismi c reflectio n patter n bot h indicate reduce d clasti c sedimen t supply . Th e basin configuratio n during Lat e Oligocen e tim e was an important facto r controlling the regional distribution o f siliceous sediments, couple d wit h organic productivity , lo w dissolutio n rate , p H and a reduce d clasti c sedimen t supply .

Lower Miocene sequence (CSS-5) A pronounced upwar d shif t i n the mineralogica l and geochemical compositio n occur s clos e to the base o f CSS-5 (Fig s 4 and 5) . Chlorite increase s up-section, concomitan t wit h a reductio n i n th e concentration o f smectit e i n th e cla y fractio n (Fig. 4) . Smectite-ric h sediment s are , however , characteristic i n wel l 34/7-6 , where CSS- 5 sedi ments hav e simila r cla y mineralog y t o th e underlying CSS- 4 (Fig . 4) . The smectit e conten t in this well, however, seems to decrease upwards , as wel l a s i n th e CSS- 5 sequenc e i n wel l 30/3-3 . In th e southernmos t wells , 24/6- 1 an d 26/4- 1 (Fig. 5) , upwar d enrichmen t o f kaolinit e an d reduction o f smectit e close t o th e bas e o f CSS- 5 is als o observed . I n thre e wells , 34/7-1 , 34/7-2 and 34/7-6 , th e plagioclas e conten t i n Lowe r Miocene sediment s i s relativel y hig h (Fig . 6 ) compared wit h that i n the underlying sequences . This is , however , no t foun d i n wel l 30/3- 3 (Fig. 6b ) or wel l 26/4- 1 (Fig. 6a) . The variation s in K 2 O/Na 2 O rati o als o reflec t thes e regiona l variations wit h respect t o th e feldspars. The silic a conten t i s high, wit h a mea n valu e of 75wt% , an d th e A1 2O3 conten t i s clos e t o 6.85 wt% i n CSS- 5 (Fig . 7) . The Ti , M g an d F e contents ar e generall y lo w (Fig . 8) , wit h mea n values o f 0. 4 wt% TiO 2, 1. 0 wt% Mg O an d 4.3wt% Fe 2O3, respectively . Trace metal s (Co, Ni , Z n an d Cu ) ofte n sho w a wea k reductio n i n concentratio n clos e t o th e base o f th e Lowe r Miocen e sequenc e (Fig . 11) . However, thi s i s no t observe d i n wel l 26/4- 1 (Fig. lib) . Lo w value s sugges t tha t volcani c influence i s low i n thes e sediments . Eastward decreasin g kaolinit e content i n wells 24/6-1 an d 26/4- 1 (Fig . 5 ) indicates tha t clasti c sediments wer e source d fro m th e Eas t Shetlan d Platform i n CSS- 5 time , i n accordanc e wit h th e seismic data markin g outbuilding fro m th e west in the area aroun d 60°N , into the Viking Grabe n (Fig. 1) .

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Fig. 11 . Distributio n of trace elements (Co, Ni , Cu an d Zn ) base d o n XRF analyse s of cuttings. Location o f wells is shown i n Fig. 1 . (Note the relativel y high content s o f most of these element s i n the partly volcani c Eocen e an d Oligocene sediment.)

SEQUENCE STRATIGRAPH Y I N CENOZOI C SEDIMENT S

Fig. 11 (continued)

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Fig. 12 . Scannin g electron micrograph s o f th e mos t commo n diatom s i n Uppe r Oligocen e sediment s in well 30/3-3 .

Middle-Upper Miocene sequence (CSS-6 and CSS-7) CSS-6 is mapped onl y in the southern part o f the Norwegian Nort h Sea , an d thin s belo w seismi c resolution i n a northwar d direction . Som e bio stratigraphic investigation s hav e indicate d tha t sediments o f CSS- 6 ag e ar e absen t i n th e northern Nort h Se a (e.g . Eidvi n & Rii s 1992) ; however, othe r studie s (e.g . Gradstei n & Ba'ck strom 1996 ) have indicated tha t CSS- 6 sediment s actually ar e presen t there . Th e northwar d thin ning o f CSS- 6 indicate s tha t CSS- 6 i s absen t in th e norther n Nort h Sea ; therefore , w e hav e grouped possibl e CSS- 6 sediment s togethe r with CSS-5 . The bas e o f CSS- 7 i s characterize d b y a pronounced downla p surfac e generated b y sediments buildin g ou t fro m th e eas t i n th e centra l North Sea . Th e thicknes s o f CSS- 7 i s clos e t o seismic resolutio n farthe r t o th e north , bu t biostratigraphic evidenc e fro m quadran t 3 4 (Steurbaut e t al. 1991 ; Eidvi n & Rii s 1992 ) indicates tha t i t i s presen t locally . W e have , however, no t bee n abl e t o distinguis h betwee n these sediments of CSS-7 age and th e underlying CSS-5 o n the seismic data in the northern Nort h

Sea. However , gamma-ra y lo g respons e i n th e investigated well s ofte n show s a marke d shif t in th e uppe r par t o f th e Miocen e succession , suggesting a subdivisio n int o tw o sequence s (Fig. 10) . W e us e thi s shif t i n lo g respons e t o indicate th e bas e o f CSS-7 . W e hav e als o use d gamma-ray an d soni c lo g data , i n additio n t o biostratigraphic data , t o identif y th e to p o f th e Miocene sequence ; thi s i s often recognize d b y a marked upwar d increas e i n gamma-ra y lo g response (Fig s 9 an d 10 ) and i n som e well s b y an increas e i n th e seismi c velocity causin g pro nounced velocit y inversio n a t th e bas e Pliocen e (Fig. 10) . W e hav e n o mineralogica l an d geochemical dat a fro m CSS- 7 i n thi s study.

Pliocene (CSS-8) and Pleistocene sequences (CSS-9 and CSS-10) The CSS- 8 sequenc e ha s bee n investigate d i n wells 30/3-3 , 34/7-1 , 34/7- 2 an d 34/7-6 . W e therefore hav e only mineralogical an d geochem ical dat a fro m th e northernmos t well s (Fig . 1) . The littl e Pleistocene sedimen t analyse d appear s to be similar to the mineralogical composition of the Pliocen e sediments .

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Fig. 13 . Glauconit e san d grain s fro m wel l 34/7- 1 indicating slow sedimentation , possibl y o n a submarin e high .

The Pliocen e an d Pleistocen e sequence s ar e 'coarse-grained mudstones ' wit h littl e smectit e and enrichmen t o f chlorite, illit e an d t o a certain degree kaolinit e (Fig . 4) . Muc h o f th e kaolinit e found ma y b e take n a s a n indicatio n o f th e degree o f reworkin g o f olde r sediments , a s very little kaolinit e woul d hav e bee n produce d whe n the climate was cold an d erosion rate s were high. Much o f wha t i s recorde d a s illit e b y XR D i n these layer s i s actuall y mic a derive d fro m rela tively unweathere d basemen t rocks . A relativel y high chlorite content i n the sediments is evidence of limite d weatherin g i n a col d climat e i n a region undergoin g glacia l erosion . In th e lowe r Tertiar y sequence , muc h o f th e chlorite ma y b e presen t a s authigeni c mineral s formed b y th e alteratio n o f volcani c glas s an d smectite o r basi c roc k fragments , and ca n there fore b e a possible sourc e fo r CSS- 8 sequences . However, th e iro n conten t i n CSS- 8 i s normally relatively lo w compare d wit h tha t i n th e lowe r part o f the Cenozoi c sectio n (Fig . 8c-f) , a s was also observed b y Rundberg (1989) . Chlorite with a lo w iron and hig h Mg content ma y represent a metamorphic sourc e (Curti s e t al. 1985) . Th e Pliocene an d Pleistocen e sediment s hav e prob ably bee n derive d fro m metamorphi c basemen t rocks fro m wester n Norway . Karlsso n e t al .

(1979) hav e als o reporte d detrita l chlorit e i n a study o f wel l 2/11- 1 i n th e centra l Nort h Sea . The highes t silic a conten t i s foun d i n th e Pliocene and Pleistocene sediment s and , couple d with lo w alumin a (Fig . 7d-f) , give s als o stron g indications tha t Pliocen e an d Pleistocen e sedi ments hav e a differen t compositio n fro m thos e of th e lowe r Tertiar y sequence . Generally , th e highest conten t (>20% ) o f plagioclase i s almost always limite d t o CSS- 8 (Fig . 6c-e) . Lowe r amounts o f K-feldspa r (generall y <10% ) ar e also characteristic for CSS-8 . Trac e elemen t (Ni, Co, C u an d Zn ) value s i n th e sequenc e ar e generally low (Fig. 1 Ic and d) , an d als o show a n increasing trend toward s th e to p o f CSS-8. CSS 8 contain s relativel y fres h feldspa r (e.g . albite) , poorly sorted , crystallin e rock fragments , volcanic roc k fragments , mica , amphiboles , pyrox enes an d hornblend e gneisse s withi n a mudd y matrix an d Dania n chal k fragment s (F . Grad stein, pers . comm.) . Kaolinit e an d smectit e indicate reworkin g o f lower Tertiar y an d Meso zoic rock s o r erosio n o f th e basement . Th e Pliocene sediment s ar e plagioclas e rich , prob ably fro m albit e gneisse s an d basi c rock s fro m Western Norway . Calcite i s th e dominan t carbonat e mineral , although siderit e and ankerite-dolomite are also

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Fig. 14 . Clas t o f chal k i n Pliocen e sedimen t i n wel l 36/1-2 . Th e clas t contain s spherica l calcareou s fossil s (calcisphera) typica l o f Dania n chal k (fro m Sing h 1996) .

present i n significan t quantities . Thi s i s partl y as contemporaneou s biogeni c carbonat e bu t includes reworke d Dania n chal k fragments , microfossils (e.g . Foraminifera) , shel l debri s and fragments . Occurrenc e o f Dania n chal k fragments (Fig . 14 ) as fa r nort h a s wel l 36/1- 2 shows tha t chal k facie s extende d int o th e north ern Nort h Se a an d covere d area s wher e lat e Neogene erosio n reache d int o the basement. Th e high conten t o f carbonat e recorde d b y XR D (Figs 9 and 10 ) is therefore no t mainl y carbonate cement, bu t occur s mostl y a s reworke d carbo nate clasts . The immatur e grains , an d th e angula r an d poorly sorte d mineralog y o f th e CSS- 8 t o CSS-10 sequence s indicat e hig h erosio n rate s i n a relativel y cold climat e prevailin g i n the north ern Nort h Sea . CSS- 8 ha s a hig h conten t o f glacial-induced erode d materia l an d represent s glacimarine deposits , wit h a proxima l sourc e area. Eidvi n & Rii s (1992 ) observe d glacia l material i n Pliocen e sediment s i n th e norther n North Sea , an d suggeste d tha t thes e sediment s are i n clos e proximit y t o th e provenanc e area . The hig h sedimen t inpu t rat e ha s bee n relate d to glacia l erosio n contemporaneou s wit h tec tonic uplif t o f Norwa y i n Pliocen e tim e (Eidvi n

& Rii s 1992 ; Stuevol d e t ai 1992 ; Jord t e l al 1995, 2000).

Mineralogical change s associated with sequence boundarie s The mineralogica l an d geochemica l dat a hav e been compare d wit h seismi c dat a fo r selecte d wells (Figs 5 , 9 and 10) ; the cla y mineralogy an d carbonate distributio n has bee n correlate d wit h the gamma-ra y and velocit y logs o f well s 30/3- 3 and 34/7-1 . The Paleocen e t o Oligocen e (CSS- 1 t o CSS-4 ) sediments i n wel l 26/4- 1 hav e a generall y high content o f smectit e (Fig . 5) . A marke d shif t i n clay minera l compositio n occur s clos e t o th e CSS-4-CSS-5 sequenc e boundary , showin g a n upward reduction in smectite content. The CSS-1 to CSS- 4 sequence s ar e al l smectit e rich , although ther e are som e mino r variation s in th e clay minera l compositio n correlatin g wit h th e seismic sequenc e boundarie s (Fig s 4 an d 5) . However, some minor shifts i n clay mineral compositions als o occu r withi n th e sequences , indi cating a composite litholog y of the sourc e area .

SEQUENCE STRATIGRAPH Y I N CENOZOIC SEDIMENT S Sediments correspondin g t o CSS-1 , CSS- 2 and partl y CSS- 3 i n wel l 24/6-1 (Fig. 5 ) have a relatively hig h conten t o f smectite . Kaolinit e is, however, als o a n importan t constituen t i n th e lower par t o f th e Tertiar y sectio n (CSS- 1 t o CSS-4). A reductio n i n smectit e conten t occur s close t o th e CSS-3-CSS- 4 sequenc e boundary . The CSS-2-CSS- 3 sequenc e boundar y ca n t o a certain degree als o b e correlate d wit h a shif t i n clay mineral composition indicated by gradually increased kaolinit e content . A latera l decreas e i n th e kaolinite/smectit e ratio i n CSS-1 to CSS- 4 occurs fro m wel l 24/6- 1 to well 26/4-1 (Fig. 5) , indicating that th e Uppe r Paleocene t o Lowe r Miocen e sediment s o f th e most easterl y wel l (26/4-1 ) wer e deposite d i n a more dista l position. Hig h kaolinite-smectit e t o illite ratio s i n th e proxima l facie s ar e consistent with th e propose d genera l relationshi p between clay mineralogy an d facie s and sea-leve l changes (Fig. 3) , an d wit h th e outbuildin g pattern indi cated b y th e seismi c dat a (Fig . 5) . Th e cla y mineral distributio n o f CSS- 1 progradin g fro m western Norwa y show s a similar , increasin g kaolinite-smectite t o illit e rati o toward s th e depocentre (Rundber g 1989) , indicatin g tha t these sediment s wer e derived fro m th e east . An increase d smectit e an d carbonat e conten t is observed nea r th e base of CSS-1 in well 30/3-3 (Fig. 9) . Clos e t o th e sequenc e boundar y be tween CSS- 2 an d CSS-3 , ther e i s an apparentl y marked decreas e i n smectit e an d a n increas e i n the kaolinit e conten t upwards . W e find , how ever, les s smectit e i n ou r sample s fro m wel l 30/3-3 tha n reporte d b y Rundber g (1989) . A marke d increas e i n smectit e conten t i s observed acros s th e CSS-3-CSS- 4 sequenc e boundary. Marked change s i n th e cla y minera l assem blages* across seismi c sequenc e boundarie s ar e also observe d i n wel l 34/7- 1 (Fig. 10) . Th e to p of CSS- 1 correlate s wit h a smectit e pea k an d the to p o f CSS- 2 i s reflecte d b y a decreas e i n smectite conten t i n wel l 34/7-1 . CSS- 3 ha s a generally high content of smectite; a marked fal l is, however , indicate d nea r th e bas e o f th e overlying CSS-4 . Th e bas e o f CSS- 4 i s reflected by a marke d upwar d increas e i n th e conten t o f kaolinite. Chlorite is found at severa l levels in the lower Tertiary sequenc e o f th e investigate d well s (Figs 4 , 5 , 9 an d 10) , an d occur s apparentl y independently o f seismi c sequenc e boundaries . However, the chlorite content o f the Miocene t o Pleistocene sequence s canno t b e explaine d b y diagenetic changes , an d th e variation s i n th e occurrence o f chlorit e correlat e wit h th e map ped sequenc e boundarie s (Fig s 4 , 5 , 9 & 10) .

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An increas e i n clasti c chlorit e o f th e Miocen e and Pliocene-Pleistocen e sediment s i s probabl y related t o provenanc e changes , hig h erosio n rates and a reduced rat e of weathering, reflecting a colde r climat e i n th e sourc e area . Fo r o f th e same reasons, CSS-8 , an d sometimes , the CSS-5 sequence boundarie s ca n als o b e rathe r easil y identified o n th e basi s o f th e feldspa r content , as increase d plagioclas e conten t clos e t o bas e CSS-8-CSS-5 is seen in the norther n North Sea . A correlatio n betwee n th e majo r elemen t distribution an d th e sequenc e boundarie s i s demonstrated b y the Si , Al, F e an d T i contents, and t o a certai n exten t b y th e N a an d K variations (Fig s 7 an d 8) . I n wel l 24/6- 1 th e increased concentratio n i n Si is seen clos e t o th e CSS-2-CSS-3 an d CSS-3-CSS- 4 sequenc e boundaries (Fig . 7a) . Th e upwar d increas e i n Si conten t correspond s t o th e CSS-3-CSS- 4 and CSS-5-CSS- 8 sequenc e boundarie s i n wel l 34/7-1 (Fig . 7d) , an d i n wel l 30/3- 3 (Fig . 7c ) is close t o th e CSS-4-CSS- 5 sequenc e boundary . The A l conten t show s a n upwar d decreas e i n concentration, a s exemplifie d b y wel l 24/6- 1 (Fig. 8a ) an d wel l 30/3- 3 (Fig . 7c) . A reduce d concentration i n A l conten t i s see n clos e t o th e CSS-3-CSS-4 boundary in well 34/7- 1 (Fig. 7d) . The highest concentratio n o f Fe and T i is found in CSS-2 , CSS- 3 an d sometime s als o i n CSS-4 , and concentratio n variation s i n F e an d T i ar e almost always following the same trend through out th e Cenozoi c sectio n an d acros s sequenc e boundaries (Fig . 8) . A reductio n i n T i an d als o sometimes Fe content close to the CSS-3-CSS-4 boundary i s demonstrate d i n wel l 34/7- 1 (Fig. 8d ) and clos e t o th e bas e CSS- 5 boundar y in well 30/3-3 (Fig. 8c) . The difference s i n Na 2O and K 2 O concentratio n als o occu r acros s th e sequence boundaries , a s demonstrate d i n wel l 34/7-1 (Fig . 7d ) and wel l 30/3-3 (Fig. 7c) . Significant variation s in trace elements , exemplified b y Ni, Co, Cu and Zn , ca n be observed in the Tertiary sediments , a s shown in Fig . 11 , an d they largel y follo w th e sam e tren d throughou t the Cenozoi c section , probabl y becaus e thes e elements ar e enriche d i n volcani c material . Variations i n th e concentration s o f Ni , Co , C u and Z n occu r acros s th e sequenc e boundaries , and hig h Ni and Z n value s are observed clos e to the CSS-l-CSS- 2 boundary . Suc h ver y high Zn values ma y als o b e relate d t o enrichmen t o f silica-producing organisms . Enrichmen t o f bio genie silica i n the Balde r Formation (to p CSS-1) has bee n reporte d i n th e literatur e (e.g . Mal m et al. 1984) . Clos e t o th e to p o f th e Balde r Formation, accumulatio n o f carbonat e ma y be du e t o hig h organi c productivit y (Fig s 9 and 10) , particularly of siliceou s and calcareou s

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organisms. Th e upwar d decreas e i n trac e ele ment concentratio n (Co , Ni , Z n an d Cu ) i s particularly demonstrate d clos e t o th e CSS-2 CSS-3 boundar y i n wel l 30/3- 3 (Fig . l i b ) an d well 26/4-1 , an d acros s th e CSS-3-CSS- 4 boundary i n wel l 34/7-1 (Fig . lie) . The lowermos t siliceou s facie s i n th e Oligo cene sequence s clos e t o th e CSS-3-CSS- 4 boundary represent s a starvatio n wit h respec t to clasti c suppl y (Thyber g e t al. 1999) . The Pliocen e mudstone s ar e coars e graine d and poorl y sorted wit h an immature mineralogy, and therefor e probabl y rathe r permeabl e an d have undergon e rapi d compactio n compare d with th e Eocen e an d Oligocen e smectiti c mud stones. Thi s i s probabl y th e mai n reaso n fo r the hig h soni c velocit y (wel l 34/7-1 , Fig . 10) . Smectitic mudstone s ar e extremel y fin e graine d and hav e hig h tortuosity , hig h specifi c surface and lo w permeability , and compac t ver y slowl y (Weaver 1989) . Th e lo w soni c velocitie s (Fig s 9 and 10 ) may b e partly du e t o th e relativel y high porosity an d partl y t o th e natur e o f th e grai n contacts in such fine-grained rocks. Overpressure primarily occur s where ther e i s a thic k sequenc e of smectiti c mudstone s withou t laterall y exten sive sandstone bodies to provide lateral drainage. This i s als o apparen t fro m publishe d pressur e data fro m th e Nort h Se a (Buhri g 1989) . Th e gamma-ray log s sho w a pronounce d reduction , which correlate s wit h dramati c increas e i n velocity clos e t o th e CSS- 1 boundary . Thi s shif t has als o bee n reporte d regionall y (Nielse n e t al . 1986; Isakse n & Tonstad 1989 ; Morton & Knox 1990). Carbonat e cementatio n an d th e transformation of opal-A to opal-CT may cause the high seismic velocities.

Summary an d conclusion s The mai n factor s influencin g th e sedimen t composition i n basin s suc h a s th e Nort h Se a are th e clasti c suppl y t o th e basi n an d th e bio genie production. Th e volcaniclastic input is very dominant i n th e Eocen e sediment s an d thi s i s related t o volcanis m associate d wit h th e break up an d initia l openin g o f th e Norwegian Greenland Sea , and a transgression ove r western Norway reducin g th e clasti c sedimen t supply . Paleocene (CSS-1 ) sediments ar e probably , t o a large extent , derived fro m th e erosio n o f Meso zoic sediments . Th e volcani c componen t i n th e Eocene sequenc e (CSS-2 ) i s rathe r dis tinctive an d ther e i s als o som e evidenc e fo r sediment starvation . An increase d rat e o f inpu t o f coarser-graine d material i n Earl y Oligocen e (CSS-3 ) tim e occurred i n th e Nort h Se a Basin , indicate d b y

an increas e i n kaolinit e and highe r SiO 2 values. However, th e volcani c inpu t i s still distinctive in CSS-3 an d CSS- 4 time, perhaps partl y representing reworke d volcani c Eocen e material . Durin g Oligocene time (CSS-4 and CSS-3), the basin configuration changed , an d i n late Oligocene time , a glauconitic facie s i n CSS- 4 sediment s indicate s a transgressio n an d sedimen t starvation . Significant shift s i n th e sedimen t compositio n occurred a t th e Miocene-Pliocen e transition , which coincide s wit h a chang e toward s a colde r climate. Lowe r Miocen e (CSS-5) , Pliocen e an d Pleistocene sediment s i n the northern par t of the study are a posses s a relativel y immature miner alogy, including clastic chlorite and hig h content of plagioclase , reflectin g a colde r climat e and a reduced rat e o f weathering in the source regions . A significan t kaolinit e an d smectit e conten t i n these sediment s i s probably no t du e t o contem poraneous weatherin g bu t ma y indicat e th e reworking o f lowe r Tertiar y o r Mesozoi c rock s during Lat e Cenozoi c time . Predictio n o f regio nal variations in geophysical properties and rate s of compactio n ca n b e carrie d ou t onl y i f th e primary mineralogica l compositio n i s known. In thi s pape r w e hav e establishe d a corre lation betwee n th e seismi c sequenc e strati graphic framewor k o f th e Cenozoi c successio n in th e norther n Nort h Se a and variation s i n th e mineralogical an d geochemica l composition . Changes i n th e provenanc e area s an d sedimen t supply, an d selectiv e erosion o f differen t sourc e rocks an d sediment s ma y explai n th e variation s in sedimen t composition . Thi s ca n b e relate d t o tectonic processe s affectin g th e Nort h Se a Basin and surroundin g areas (Jord t e t al. 1995 , 2000). The regiona l cla y minera l distributio n i s con trolled b y sourc e an d facie s change s throughout the developmen t o f th e Nort h Se a Basin , which is probabl y controlle d b y tectoni c movement s and change s i n sedimen t suppl y an d distanc e from th e source . The mappe d seismi c sequence s d o no t onl y reflect depositiona l processes . Th e primar y mineralogical compositio n o f th e sediment s an d burial diageneti c processe s strongl y influenc e the seismi c velocity and seismi c attributes. Th e rate o f compaction als o influence s th e geometry of th e seismi c sequences . An integratio n o f mineralogy , geochemistry and sequenc e stratigraph y has therefor e the po tential for correlation in the Nort h Se a Basin on local an d regiona l scales . This wor k wa s funde d b y th e Commissio n o f th e European Unio n and th e Norwegia n Researc h Coun cil (NFR ) i n th e framewor k o f th e DGXII-Joul e Programme, sub-programme : Energ y Fro m Fossi l

SEQUENCE STRATIGRAPH Y I N CENOZOI C SEDIMENT S Sources: Hydrocarbons , Integrate d Basi n Studie s The Dynamic s o f th e Norwegia n Margin . Nors k Hydro AS , Saga Petroleum AS , Statoil and Norwegia n Petroleum Directorat e ar e gratefully acknowledge d fo r providing samples . R . G . Simmon s provide d inter pretation o f cla y mineralog y o f well s 26/4-1 , 30/3- 3 and 34/7-2 . W e woul d lik e t o than k B . Stabel l regarding discussio n o f diatoms . V . Cadma n kindl y corrected th e manuscript .

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Cenozoic tectoni c subsidenc e fro m 2 D depositiona l simulation s o f a regional transec t in the norther n North Se a basi n RUNE KYRKJEB0, 1 MARTI N HAMBORG, 2 JA N ING E FALEIDE, 3 HENRIK JORDT 3 & PETE R CHRISTIANSSON 3 1

Geological Institute, University of Bergen, AI legal en 41, N-5007 Bergen, Norway 2 SINTEF, Petroleum Research, N-7034 Trondheim, Norway * Department of Geology, University of Oslo, P.O. Box 1047 Blindern, N-0316 Oslo, Norway Abstract: Th e Cenozoi c depositiona l histor y alon g a regiona l E- W profil e acros s th e northern North Sea has been simulated using a forward process-based simulatio n program of dynamic-slope type . I t involve s a depth-dependent , dual-litholog y diffusio n equatio n tha t handles transport, erosion and deposition of sediments. The data used in the simulation were derived fro m a seismi c lin e calibrate d agains t wells , an d fro m th e regiona l literatur e concerning the norther n North Sea . Th e mos t importan t o f the factor s used are : th e initia l basin form (Paleocene bathymetry) , tectoni c subsidence , isostati c variables , sediment suppl y (sand-shale), sedimen t compactio n (porosity-dept h relationship s fo r sand-shale ) an d eustatic sea-leve l changes . Th e interactio n betwee n th e dat a value s extracte d fro m th e literature coul d no t reproduc e a cross-sectio n simila r t o th e observe d cross-sectio n fro m seismic data. Therefore, th e subsidence pattern and th e initial basin form were reconsidered. The resultin g model gav e a n anomalou s Cenozoi c subsidenc e pattern, differen t fro m th e expected post-rif t therma l subsidence, with deviations corresponding t o Paleocen e and Lat e Miocene-Pliocene times . Th e model-derive d Paleocen e subsidenc e migh t hav e bee n overestimated b y usin g an over-shallo w palaeobathymetric value, although a deepenin g of the basi n i s also indicate d b y biostratigraphi c data . Th e pronounce d Neogen e subsidenc e created accommodatio n spac e fo r a thic k Pliocen e sequence , derive d fro m th e uplifte d eastern sourc e area .

Computer simulation s ar e increasingl y impor tant fo r th e improve d understandin g o f basi n development an d infilling , an d severa l forwar d models tha t simulat e sedimentar y processe s i n basins ar e available . Kendal l e t al. (1991 ) summarized variou s approache s i n forwar d model ling an d conclude d tha t th e significanc e an d utility o f an y particula r mode l i s a matte r o f need, compute r hardwar e an d programmin g resources. However , th e qualit y o f th e simu lation depend s o n th e exactnes s o f th e inpu t values (Aigne r e t al . 1990) . DEMOSTRAT (Rivenae s 1993) , whic h i s a for ward computer-simulatio n mode l fo r siliciclastic basin-fill i n tw o dimensions , ha s bee n use d t o simulate th e Cenozoi c seismi c stratigraphi c framework alon g a regiona l E- W transec t i n the norther n Nort h Se a region . Th e progra m DEMOSTRAT i s a process-base d dynamic-slop e type mode l fo r clasti c settings , wit h a dual lithology, depth-dependen t diffusio n algorithm , which ha s successfull y demonstrate d realisti c erosion an d depositio n (Syvitsk i e t al . 1988 ; Flemings & Jordan 1989 ; Sinclair etal. 1991; Riv-

enaes 1992) . The process-elements in the progra m are performed as logically linked events in a timestep loop (Helland-Hanse n e t al. 1988; Lawrence et al . 1990 ; Rivenae s 1992) . The mos t importan t elements i n th e time-ste p loo p tha t wil l b e dis cussed her e are : tectoni c subsidenc e histor y (user-specified subsidenc e rates) ; eustati c sea level variation s (user-define d sea-leve l curve) ; erosion an d depositio n (user-define d transpor t coefficients tha t describ e ho w easil y erosion an d transport occur , an d a user-define d rat e o f sedi ment san d an d shal e supply) ; compactio n o f sediment (user-define d empirical porosity-dept h relations); flexural isostatic subsidence caused b y loading (o r unloading ) o f sedimen t an d wate r (user-defined siz e of the flexura l rigidit y describing th e lithospheri c strength) . Fo r detaile d des criptions o f th e element s i n th e program , th e reader i s referred t o Rivenae s (1993). The presen t stud y i s base d o n th e seismi c lines SG8043-10 1 an d NSDP84- 2 (transec t 2 ) (Jordt e t al . 1995 , 2000 ; Christiansso n e t al . 2000; Odinse n e t al . 2000), adjacen t well s and a documented regiona l Cenozoic seismi c sequenc e

From: NOTTVEDT , A . e t al . (eds ) Dynamics o f th e Norwegian Margin. Geologica l Society , London , Specia l Publications, 167 , 273-294 . 1-86239-056-8/OO/ S 15.00 © Th e Geologica l Societ y o f Londo n 2000 .

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analysis (Jord t e t al. 1995 , 2000) . I n addition , information ha s bee n extracte d fro m th e litera ture concernin g th e Nort h Sea . Th e mai n objec tives o f th e presen t simulatio n hav e bee n t o reproduce th e Cenozoi c successio n a s see n i n transect 2 b y applyin g a se t o f realisti c inpu t values an d t o discus s th e tectoni c subsidenc e predicted b y th e simulation .

Geological setting The Nort h Se a Basi n (Fig. 1 ) is an intracratoni c basin situate d a t th e northwester n margi n o f

the Europea n Platform . Sinc e th e crusta l accre tion complete d durin g earl y Devonia n time , the Nort h Se a are a ha s experience d severa l episodes o f lithospheri c extensio n (Devonian Carboniferous, Permo-Triassic , lat e Jurassic early Cretaceous) , eac h followe d b y a stag e o f thermal subsidenc e (Ziegle r 1982# ; Glenni e 1984; Giltne r 1987 ; Badley e t al . 1988 ; Gabriel sen e t al . 1990) . Th e Devonian-Carboniferou s part o f th e structura l histor y i s no t ver y wel l understood i n th e stud y area , mainl y becaus e of lac k o f dat a (Gabrielse n e t al . 1990) . Th e timing o f th e Permo-Triassi c rif t phas e i s stil l a matter o f debate , an d bot h Permia n (Eyno n

Fig. 1 . Regiona l map o f th e norther n Nort h Se a wit h mai n structura l elements and th e locatio n o f transec t 2. Transect 2 a combination o f the conventional SG8043-10 1 lin e and th e deep seismi c line NSDP84-02. crosses the Horda Platfor m (HP), Viking Graben (VG), Eas t Shetlan d Basi n (ESB) and th e Shetland Platform (SP) . and ha s a tota l lengt h o f 240k m i n thi s stud y (fro m Christiansso n e t al . 1999) .

CENOZOIC TECTONI C SUBSIDENC E 1981; Badle y et al 1988 ; Gabrielsen e t al. 1990 ; Faerseth e t al , 1995 ) an d Triassi c (Beac h e t al . 1987; Giltne r 1987 ; Robert s e t a l 1995 ) age s have bee n suggested . Mos t o f th e Permo Triassic stretchin g occurre d betwee n th e 0ygarden Faul t Zon e t o th e eas t an d th e Shetlan d Platform an d th e Hutto n alignmen t t o th e west (Odinsen e t a l 20006) . Modellin g result s (te r Voorde e t a l 2000 ) sho w tha t n o hea t acces s from th e broa d extende d Permo-Triassi c rif t

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event wa s lef t a t th e onse t o f th e mor e focuse d late Jurassic-earl y Cretaceou s riftin g (Odinse n et a l 20006) , whic h le d i n tur n t o th e develop ment o f th e mai n structura l element s o f th e Viking Grabe n area . In earl y Cretaceou s time , therma l subsidenc e commenced; thi s was marked b y the cessation of fault bloc k tilting . Th e followin g post-rif t sedi mentation (Cretaceous-Cenozoic ) fille d i n an d buried th e rif t topography . I n th e norther n

Fig. 2. Correlation chart showing the seismic seuence stratigraphic framework. connected to chronostratigraphy and lithostratigraphy (from Jordt et al. 1995. 2000).

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North Sea , sout h o f 61 : N, earl y Cretaceou s deposits ar e mainl y restricte d t o th e Vikin g Graben an d mostl y la p ont o th e basi n margins , but lowe r Cretaceou s sediment s ar e als o foun d in th e Stor d Basi n an d a t th e Hord a Platform . Late Cretaceou s sedimentatio n burie d th e rif t topography an d Cenozoi c sedimentatio n over stepped th e grabe n boundaries . Period s o f in creased subsidenc e associate d wit h faultin g ar e detected i n the post-rift stage , particularly in late Cretaceous-early Tertiar y tim e (Thorn e & Watts 1989 ; White & Lati n 1993 ; Hal l & Whit e 1994; Leper q & Gaulie r 1996 ; Nadin & Kusznir 1996). bu t thi s cannot b e considered a rif t phas e (Badley e t al. 1988 ) The Cenozoi c basin-fil l wa s derive d fro m tw o main sources : th e Eas t Shetlan d Platfor m t o th e west an d th e Fennoscandia n Shiel d t o th e east . Deposition i n Tertiar y tim e wa s controlle d b y several phase s o f uplif t affectin g th e sourc e areas, combine d wit h relativ e sea-leve l change s and therma l subsidenc e o f the basi n (Jord t e t al. 1995. 2000) . Th e wester n sourc e area s domi nated th e sedimen t influ x fro m lat e Paleocen e until earl y Miocen e time s (Stewar t 1987 ; Gallo way e t a l 1993 ; Jord t e t al . 1995 , 2000) . Th e basin shallowe d fro m relativ e deep-marin e conditions durin g th e lat e Paleocen e time , t o shallow-marine an d locall y fluvia l condition s i n late Earl y Miocen e tim e (Rundber g e t al . 1995 ; Jordt e t al . 1995 , 2000) . Th e sedimen t influ x from th e wes t cease d i n Earl y Miocen e tim e and mos t o f Mid - an d Lat e Miocen e time , wa s characterized b y non-depositiona l conditions . In Pliocen e time , a majo r uplif t o f th e Fennos candian landmas s cause d a majo r sedimen t influx fro m th e east . I n lat e Pliocen e an d Pleistocene time , th e glaciation s o f Fennoscan dia caused th e removal o f proximal Tertiar y an d older deposit s fro m th e basin' s easter n margin .

CSS-1 t o CSS-1 0 (Cenozoi c Seismi c Sequences) . The correlatio n char t i n Fig . 2 connect s th e sequence stratigraphi c interpretatio n wit h th e standard chrono - an d lithostratigraphi c subdivisions o f Deega n & Scul l (1977), revise d b y Isak sen & Tonstad (1989 ) fo r th e Norwegia n secto r and b y Kno x & Hollowa y (1992 ) fo r th e U K sector. A depocentr e ma p wit h indicate d out building direction s i s show n i n Fig . 3 . Figur e 4 shows th e ful l cross-section , wherea s th e dis tribution an d thicknes s (depth-converted) of th e Cenozoic sequence s ar e give n i n Fig . 5 . CSS-1 (o f Lat e Paleocene-earlies t Eocen e age), whic h correlate s wit h th e Rogalan d Group (Fig . 2) . ha s it s mai n outbuildin g direction fro m th e west , wher e i t wa s source d b y th e uplifted Shetlan d Platform . As can b e see n fro m Fig. 5 . westerly derived sediment s reache d fa r t o the east . Th e sediment s wer e deposite d a s submarine fa n complexe s i n a relativel y dee p basin, wit h wate r depth s suggeste d t o b e middle bathyal (500-100 0 m) i n wel l 3/25- 1 (locate d close t o transec t 2 i n th e U K sector ) (Gradstei n et al . 1994) . I n th e east , sedimentar y wedge s prograded westward s toward s a depocentr e probably situate d nort h o f th e transect , wes t o f Sognefjorden (Fig . 3) . I t i s probable tha t CSS- 1 covered a large r area , an d tha t part s o f western Norway wer e transgressed i n late Paleocene tim e (Jordt e t al . 1995) . However, thes e deposit s wer e removed b y pos t Paleocen e erosio n durin g th e Late Pliocen e an d Quarternar y glaciations .

Cenozoic seismi c stratigraphy The seismic-stratigraphi c framework use d i n this study wa s establishe d b y Jord t e t al . (1995) , on th e basi s o f a regiona l analysi s o f th e Ceno zoic i n the centra l an d norther n Nort h Sea . Th e sequence description s ar e no t repeate d here : instead, th e importan t constraint s fo r th e com puter modellin g ar e examined . From eas t t o west , transec t 2 (Fig . 1 ) crosse s the mai n structura l feature s o f th e norther n North Se a Basin , namely , th e 0ygarde n Faul t Zone, th e Hord a Platform , th e Vikin g Grabe n and th e easter n par t o f th e Eas t Shetlan d Plat form. Th e Cenozoi c succession in the North Sea has bee n subdivide d int o te n sequences , labelle d

Fig. 3. Depocentr e location s an d outbuildin g directions indicated by arrows for each of the Cenozoic seismic sequence s give n b y number s (fro m Jord t et al . 1999) .

Fig. 4. Transec t 2 in full basi n scale and locatio n o f key well 30/11-2 (from Christiansson e t al. 2000). The transec t show s the Paleogen e an d Neogen e succession s relativ e to deepe r structure s and th e mai n structural elements. Lin e locatio n i s indicated i n Fig . 1 .

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Fig. 5. Cenozoi c basin fill (CSS-1 to CSS-10) along transect 2. Arrows show main outbuilding directions, and th e interpreted borde r betwee n easterl y an d westerl y derive d sediment s i s marked fo r sequence s CSS-1 and CSS-2 . CSS-2 (o f Eocen e age ) correlate s wit h th e lower par t o f th e Hordalan d Grou p (Fig . 2) . This sequenc e ha s bee n subdivide d int o tw o units: a lowe r Eocen e sequence , CSS-2.1 , foun d mainly t o th e east , an d a middle-uppe r Eocen e sequence, CSS-2.2 , wit h it s depocentr e i n th e Viking Grabe n (Fig . 3) . CSS-2. 1 i s predomi nantly eas t derived , an d it s depositio n wa s ter minated wit h the correspondin g submergenc e o f Fennoscandia (Rundber g 1989 ; Sterbau t e t al. 1991; Jord t e t al . 1996) . Th e depositiona l styl e of CSS- 2 seem s t o b e simila r t o tha t o f CSS-1 , both fo r westerly and easterl y derived sediments. CSS-3 (of Early Oligocene age ) correlates with the middle part o f the Hordaland Grou p (Fig. 2). At th e Eocene-Oligocen e transition , uplif t o f Fennoscandia cause d a relativ e fal l o f se a level and a n outbuildin g o f a sand y wedg e fro m th e east, sout h o f transec t 2 . The depocentr e o f thi s sandy unit is located north of the transect (Fig. 3). Another depocentre i s located i n the Viking Graben, an d wa s source d fro m th e wester n hinter land (Jord t e t al . 1995) . CSS-4 (o f lat e Early-Lat e Oligocen e age ) correlates with the middle part o f the Hordalan d Group (Fig . 2) . I t prograde d eastwards , an d has a depocentr e alon g transec t 2 (Fig . 3) . Another depocentr e i s locate d sout h o f th e transect, wher e CSS- 4 prograde d t o th e north west. A n irregula r surfac e o n th e to p o f CSS- 4 (Fig. 5 ) was probabl y cause d b y th e remobilization o f clay s (Jordt e t al . 199 5 and 1996) . CSS-5 (o f Lates t Oligocene-Earl y Miocen e age) correlate s wit h th e uppe r par t o f th e Hordaland Grou p (Fig. 2) , and is the uppermost preserved strat a o f th e grou p i n th e stud y are a (see CSS-6). CSS-5 prograde d fro m th e west, but formed a north-trendin g shee t i n th e middl e o f the basin (Figs 3 and 5) . No sedimen t influ x fro m the east has been recognize d i n this sequence. Th e irregular top o f CSS-5 is interpreted a s an incised valley surface , probabl y forme d b y subaeria l erosion. This latter feature is more distinct in the

seismic sections taken further nort h i n the North Sea. Rundber g e t al . (1995 ) an d Jord t e t al . (1996) suggeste d tha t th e subaeria l exposur e i n this are a wa s primaril y controlle d b y tectoni c uplift o f the northernmos t Nort h Sea . CSS-6 (o f lates t Early-earl y Mid-Miocen e age) i s th e uppermos t uni t correlativ e wit h th e Hordaland Grou p (Fig. 2). It is only found south of th e Vikin g Grabe n an d o n th e Hord a Plat form. I n th e presen t stud y area i t i s represented by a hiatu s (o r belo w seismi c resolution) . CSS-7 (o f Mid-Lat e Miocen e age ) correlates with th e lowe r par t o f th e Nordlan d Grou p (Fig. 2). This sequence prograded fro m the east as a resul t o f th e uplif t o f Fennoscandi a relativ e to th e subsidin g basi n i n Mid-Miocen e time . The depositio n o f thi s sequenc e marke d a shif t of th e mai n sourc e are a fro m a westerl y t o an easterl y domain . I n th e transec t area , th e sequence is represented b y a relatively thin depo sitional wedg e (Fig. 5) , with th e majo r depocen tres locate d nort h an d sout h o f th e transec t (Fig. 3) . Thi s sequenc e als o correlate s wit h th e Utsira Formatio n (Fig . 2) . However , correla tions o f seismi c an d wel l dat a show ? tha t th e sandy unit s of bot h CSS- 5 an d CSS- 8 hav e als o been interprete d as th e Utsir a Formatio n (Jord t et al . 1995) . CSS-8 (o f Pliocen e age ) correlate s wit h th e Nordland Grou p (Fig . 2 ) an d wa s deposite d i n response t o erosion of a strongly uplifted Fenno scandian hinterland , togethe r wit h Nort h Se a basinal subsidenc e providin g accommodatio n space (Jord t e t al. 1995 , 2000). The CSS- 8 clastic wedge prograded to the west, forming two depocentres (Fig . 3) : a smalle r wedge wa s locate d o n the transec t i n th e Stor d Basin , and a large r on e in th e Tampe n Spu r area . CSS-9 (o f earl y Quaternar y age ) correlate s with th e uppe r Nordlan d Grou p (Fig . 2) . but i s represented b y a hiatu s i n transec t 2 . CSS-10 (o f lat e Quaternar y age ) correlate s with the uppermost par t o f the Nordland Grou p

CENOZOIC TECTONI C SUBSIDENC E (Fig. 2) . This sequence show s an angular unconformity a t th e lowe r boundar y (Fig . 5) , espe cially in the eastern part o f the study area, where the olde r sequence s ar e truncate d b y CSS-10. Input factor s The modellin g was performe d b y systematically changing inpu t value s t o reproduc e th e geom etry observe d i n th e seismi c reflectio n data . We starte d wit h a fixe d se t o f value s fo r initia l basin form , amoun t o f tectoni c subsidenc e (uplift), rate s o f sedimen t supply , sea-leve l curve an d value s fo r isostati c calculations , whereas th e value s o f th e transpor t coefficient s were varied. The fixed input values are summarized i n Tabl e 1 an d th e choic e o f value s i s discussed below.

Initial basin form Water depth is a critical factor in subsidence and burial analysis, because it represents the amoun t of th e basin' s 'underfill' . Larg e palaeobathy Table 1 . Initial input factors summarized for th e postrift modelling Variable

Source

Initial basi n form

Water depth s alon g transec t 2 based o n wor k b y Rocho w (1981), Barto n & Wood (1984) , Ziegler (1990 ) an d Gradstei n et al. (1994) Haq e t al . (1987) long-term curve; age based o n Harlan d e t al . (1990) tim e scal e Derived fro m backstripping ; supply fro m bot h side s According t o th e typ e curves shown i n Fig . 7

Sea-level function Sediment suppl y Transport coefficient function Surface laye r thickness Compaction

Isostasy Time span , number o f time step s Basin length , number o f columns Tectonic movements

30cm Sclater & Christie (1980) curves ; the compactio n o f th e bas e is measured an d give n togethe r with tectoni c subsidenc e Flexural rigidit y D = 1. 0 x 10 22 Nm; n o axia l stress 60.5 Ma, 20 0 time steps , whic h gives each tim e ste p a duratio n of 30 2 500 a 240km, 20 0 column s Thome & Watts (1989 )

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metric value s a t th e en d o f Cretaceou s tim e provided muc h o f th e accommodatio n spac e for Tertiar y sedimentation . O n th e basi s o f a n analysis o f fauna l trend s alon g palaeoslop e transects i n th e norther n Nort h Sea , Gradstei n et al . (1994) and Gradstei n & Backstrom (1996 ) estimated conservativel y that th e earl y Palaeo cene wate r dept h wa s betwee n 25 0 an d 500m , and tha t th e Lat e Paleocen e wate r dept h wa s between 50 0 and 1000 m i n th e UK-wel l 3/25-1 , which i s clos e t o th e modelle d transect . Othe r wells (Fig . 6 ) analyse d i n th e are a sho w tha t water depth s i n earl y Paleocen e tim e wer e be tween 50 0 an d 750 m i n UK-wel l 9/23- 1 an d between 20 0 and 500 m i n UK-well 9/13-1. Lat e Paleocene wate r depth s wer e betwee n 50 0 an d 1000m i n UK-wel l 9/13-1 , an d betwee n 75 0 and 1500 m i n UK-wel l 9/23-1 (Gradstein e t al . 1994, Gradstei n & Backstrom , 1996) . Palaeo bathymetric maps (Barto n & Wood 1984 ) based on informatio n obtaine d fro m wel l samples , combined wit h th e palaeogeographica l map s of Ziegler (1981 , 19820) , show a 500 m contour fo r the Paleocen e wate r depths , indicatin g tha t water depth s i n th e deepes t par t o f th e basi n must hav e exeede d 500m . Accordingly , a max imum initia l basi n dept h i n thi s stud y wa s estimated t o b e c. 550m. The initial basin configuration along the transect at the beginning of Late Paleocene tim e was constructed b y combining the work of Barton & Wood (1984 ) (500 m contours) , palaeogeogra phical map s o f Rocho w (1981 ) an d Ziegle r (1990) (Fig . 6) , an d th e presen t bas e Tertiar y relief. Th e initia l for m o f th e basi n (Fig . 7 ) i s believed t o hav e bee n nearl y symmetrical , with the wester n flan k reachin g se a leve l o n th e Eas t Shetland Platfor m (Rochow 1981 ; Ziegler 1990), whereas parts of the eastern flank on th e Hord a Platform wer e submerged (Ziegle r 1990) .

Sediment input The depth-converte d sejsmi c profil e wa s back stopped t o determin e th e decompacte d cross sectional are a o f eac h sequenc e afte r removin g the overburden . Th e backstrippin g proces s wa s carried out using the Balancing Section Progra m (XBSP; Midlan d Valle y Exploratio n Ltd) . Th e exponential porosity-dept h relationship s o f Sclater & Christi e (1980 ) fo r shale s an d sand s were use d i n th e backstrippin g procedure . Th e lithological compositio n o f eac h sequenc e wa s established b y investigatin g the completio n lo g from wel l 30/11- 2 (Fig . 4) ; th e shale-to-sand ratio wa s estimate d fro m th e gamma-ra y log . The volumetri c sand fractio n (Tabl e 2 ) appear s

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Fig. 6 . Integratio n o f Lat e Paleocen e palaeogeographica l map s b y Rocho w (1981) . Barto n & Wood (1984 ) an d Ziegler (1990 ) wit h interpreted fauna l trends fro m well s location s b y Gradstei n e t al . (1994 ) an d Gradstei n & Backstrom (1996) .

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Fig. 7 . Th e initia l form o f the basi n i n the beginnin g Lat e Paleocen e tim e based o n Fig . 6 . The maximu m wate r depth i s suggested t o b e 550m . Th e basi n for m i s symmetrical wit h shallo w wate r depth s a t th e flanks .

to b e relativel y high fo r mos t o f th e sequences . The estimates ar e admittedly uncertain , becaus e of limitation s o f th e metho d an d th e uncer tainty o f the log-derive d values . Because th e sectio n i s 2D , th e decompacte d sediment-flux rate s ( C2D volumes') fo r san d an d shale are given in square metres per year (m2 a"1 ) (see Fig . 8) . Changes i n sediment-flu x rate s ar e closely relate d t o tectoni c movement s an d t o changes i n the sedimen t suppl y system ; they ar e synchronous wit h th e developmen t o f deposi tional sequenc e boundarie s (Jord t e t al. 1995) . To tes t th e sensitivit y of compactio n t o th e lithological variations, a backstripping was done for 100 % shal e an d 100 % san d content . Th e difference betwee n the shale and san d volumes in the basina l cross-sectio n wa s foun d t o rang e from 0 to 20%, with the greatest difference i n the oldest (mos t compacted ) units , an d decreasin g upwards. Hence , th e sensitivit y analysi s indi cates tha t th e error s i n bul k volum e cause d b y departures i n lithological composition ar e small. The lack o f information o n the distribution of lithologies along th e profile mad e i t necessary to simplify thi s par t o f th e problem . A s th e litho facies architectur e o f th e basina l cross-sectio n Table 2 . Interpreted sand fraction fo r each sequence applied in the simulation Sequence

Sand fractio n

CSS-1 CSS-2 CSS-3 CSS-4 CSS-5 CSS-7 CSS-8 CSS-10

0.76 0.62 0.47 0.35 0.80 0.99 0.49 0.72

was not th e main objective of the simulation, the compaction-related error s caused b y lateral lithological variation s hav e bee n neglected . Instead , the modellin g focuse d o n th e bul k geometrica l relationships o f eac h seismi c sequence. Some sequences , especiall y those derived fro m the easter n source , hav e bee n erode d fro m th e flank. This erosional truncation is shown in Fig. 5. As th e backstrippin g take s onl y th e volum e o f sediment present in the profile (2D ) into accoun t and n o out-of-plan e transpor t o f sediment s i s possible in the forwar d simulation , a redistribution patter n fo r th e erode d sedimen t ha d t o b e suggested. The volum e of the redistribute d sediment wa s calibrated t o matc h th e overal l thickness framework , an d th e tota l sedimen t bul k volume. Th e patter n o f redistributio n woul d obviously have to reflec t th e palaeogeographica l situation. Sequence s CSS-1 , CSS- 2 an d CSS- 4 therefore ha d t o includ e additiona l sedimen t i n the easter n par t t o fi t th e pre-Miocen e palaeo geography. Thi s wa s accomplishe d b y extrapo lating th e observe d sequence s eastwards ; th e amount o f volum e adde d wa s calibrate d t o th e related volume s o f sequence s CSS- 7 an d CSS-8 . While CSS- 7 an d CSS- 8 wer e bein g deposited , the additiona l sedimen t volum e i n th e eas t i s assumed t o hav e bee n erode d an d redeposited . These depositiona l event s correspon d t o th e uplift o f th e easter n sourc e area . The sediment s are assume d t o ente r th e basi n at poin t source s o n eac h sid e o f th e cross section. O n th e seismi c reflectio n dat a i t i s pos sible to distinguish between easterly and westerly derived sediments (see location o f depocentres i n Fig. 3 ) for som e of the sequence s (CSS-1, CSS-2 , CSS-4 and CSS-10) (Jordt e t al 1995) , and therefore calculat e th e decompacte d cross-sectiona l area t o ente r fro m eac h sid e o f th e section . Because o f restricte d knowledg e o f th e sourc e

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Fig. 8 . Sedimen t suppl y (irra ' ) o f san d (continuou s line ) an d shal e (dashe d line ) fro m wes t (a ) an d eas t (b). Volumes ar e calculate d fro m backstripping . Thes e volume s hav e 0 % porosity , whic h i s the inpu t for m i n DEMOSTRAT. Th e progra m wil l conver t inpu t volume s to maximu m porosities. 0.49 fo r san d an d 0.6 3 fo r shale , according t o th e empirica l porosity-dept h relation s o f Sclate r & Christi e (1980 ) fo r Nort h Se a settings .

area, erosio n o f sourc e are a topograph y i s no t specified i n detai l i n th e simulation .

Compaction The function s of Sclate r & Christi e (1980 ) hav e commonly bee n use d t o describ e th e porosity depth relationshi p i n th e Nort h Se a region . DEMOSTRAT use s on e curv e fo r eac h lithology , whether san d o r shale , wherea s XBS P (Midlan d Valley Exploratio n Ltd ) uses a singl e lithology dependent curve . An importan t aspec t i s th e volum e contro l on th e sediment s belongin g t o eac h sequence . By using the same empirical porosity-depth relationship an d san d fractio n value s bot h i n th e

inverse an d forwar d modelling , th e volum e i s expected t o b e th e sam e i n th e simulate d an d in th e observe d section . Th e porosit y value s in the simulate d cross-sectio n ar e no t necessar ily th e sam e a s i n th e observe d section . Thes e deviations ar e du e t o uncertaint y relate d t o the usag e o f empirical porosity-depth functions . However, a s th e erro r woul d b e th e sam e fo r the invers e an d forwar d simulations , th e un certainty o f th e porosity-dept h relationshi p i s not relevant . The underlyin g Cretaceous an d olde r deposit s have a non-unifor m thickness distribution along the cross-sectio n (Fig . 4) , bein g thickes t i n th e Viking Grabe n an d thinnin g greatl y toward s the basi n flanks . Thi s probabl y cause d differ ential compactio n an d base-Tertiar y subsidence .

CENOZOIC TECTONI C SUBSIDENC E which i t i s no t possibl e t o represen t i n th e DEMOSTRAT simulation . I n th e decompactio n process based o n XBS P (Midland Valle y Exploration Ltd) , th e Cretaceou s strat a were include d and th e resultin g value s wer e adde d t o th e tectonic subsidenc e i n th e DEMOSTRA T simula tion. I n al l presentations o f tectoni c subsidenc e curves, thes e compactio n value s ar e subtracted . These are , how - ever, minimu m value s fo r th e compaction o f th e underlyin g sediments , a s n o sediments olde r tha n Cretaceou s hav e bee n taken int o account in the decompaction proces s and i n the depositiona l modelling .

Transport coefficients In DEMOSTRA T (Rivenaes 1992, 1993) , the depth dependent diffusio n dual-litholog y algorith m handles erosio n an d deposition . Diffusio n ca n be characterized a s a time-dependent smoothin g process. I n geologica l term s thi s ca n b e erosio n of topographic high s an d depositio n o f the ero sional product in sedimentary basins . Th e model has severa l limitations . I n th e presen t case , th e most importan t ar e lac k o f transvers e sedimen t transport (ou t o f th e plan e o f th e section ) an d inability t o mode l turbidites . Input data are values for transport coefficients , termed K fo r san d an d shale . Th e transpor t coefficients giv e a measure of the transport effici ency o f sediments , tellin g u s ho w easil y erosio n and depositio n occu r (Rivenaes 1992). Sediments are supplie d fro m point source s o n eac h sid e of the cross-section . I n general , dept h variation s of K value s i n th e simulatio n ar e adapte d a s

283

follows (Fig . 9) : a highe r transpor t efficienc y above se a level , a n exponentia l decreas e nea r and belo w se a leve l whe n th e loa d enter s th e basins, an d belo w a certai n leve l th e transpor t efficiency i s constant an d 'low' . Rivenaes (1993) discussed the quantification of the transpor t coefficient s an d summarize d som e measured an d som e synthetic K values propose d by variou s workers , showing tha t K value s ar e a functio n of the scal e of investigation an d tha t K value s ar e dependen t o n environment . Rive naes (1992) suggested syntheti c lvalues rangin g from 300m 2 a"1 i n mountai n environment s t o 1.2 x 10 4 m2 a"1 fo r nonmarin e mud. Kenyo n & Turcotte (1985 ) have measured ^ values for the Mississippi delt a fron t o f 5. 3 x 10 5 m2 a"1. Unfortunately, n o values for the transport coefficients fo r the Cenozoic basin-fil l in the North Sea ar e available , an d i n th e presen t stud y i t has bee n necessar y t o assum e th e transpor t coefficients representin g th e contributin g pro cesses invoke d i n sedimentatio n an d erosion . The applie d K values , therefore , ar e take n a s those tha t reproduc e th e latera l distributio n o f sediments. Th e K curv e patter n use d i n thi s simulation i s complex , becaus e o f change s i n depositional patter n bot h i n tim e (difference s in sediment supply for eac h sequence) and spac e (differences i n sedimen t suppl y fo r th e wester n and easter n part , an d differenc e i n depositiona l regime with depth) . The sequence s CSS- 1 t o CSS- 5 ar e mainl y derived from a western source, with limited sediment supply from the east. This suggests a higher transport efficienc y fo r th e westerl y source d sediments. CSS-1 , whic h wa s source d fro m th e

Fig. 9. Th e transport coefficient function s used in the simulation. The values decrease to a depth of 40m an d ar e constant a t deepe r levels . Difference s i n transport efficiency ar e neede d t o reproduc e the observe d geometrie s in Fig . 5. A hig h valu e above sea level, decreasin g t o a hig h valu e a t 20 m belo w se a level, an d constan t high further dow n (lin e a ) ar e adapte d for th e wester n derive d CSS-1 . A lo w value both above and belo w se a level, describing a period of non-deposition, and tryin g to avoi d erosion, is used for CSS- 6 to CSS- 9 from th e west. Intermediate value s to thos e of line a an d lin e c are applie d fo r th e res t o f the sedimentar y record.

284

R. KYRKJEB 0 E T AL .

west, ha s th e larges t sedimen t suppl y an d a higher A T value i s assumed . Furthermore , CSS- 1 consists partl y o f submarin e fa n deposit s an d therefore a hig h K valu e a t grea t depth s ha s been applie d t o matc h th e observe d geometrie s (Fig. 9 , lin e a) . I n period s o f erosio n o r non deposition (CSS- 6 t o CSS- 9 fro m th e wester n source), K values are assume d t o hav e bee n very low (c . 1.0) (Fig . 9 , lin e c) . Thi s migh t no t b e correct considerin g th e erosio n observe d i n un derlying sequence s an d th e northward-directe d transport o f erosiona l product s (Jord t e t al. 1995). bu t becaus e o f th e 2 D limitation s w e ar e not abl e t o transport sediment s ou t o f the plane. The CSS- 8 sequence , source d fro m th e east , ha s high K value s t o simulat e erosio n o f previousl y deposited CSS- 1 to CSS- 4 on the easter n flank (Fig. 9 , lin e b) . We mak e n o attemp t t o predic t a detaile d lithological facie s distributio n alon g th e cross section. Therefore , n o particula r sedimen t sort ing (sand-shale ) i s applie d t o differentiat e between th e K values . Thi s i s du e t o restricte d knowledge o f sedimen t compositio n alon g th e profile, time-ste p resolutio n (302.50 0 a), an d t o the necessaril y limite d scop e o f thi s study. CSS-1 an d CSS- 2 ar e partl y submarin e fa n deposits (turbidites ) an d eac h turbidit e even t cannot b e regarde d t o hav e a diffusiv e nature , but th e su m o f turbidit e event s withi n eac h sequence ca n b e assume d t o b e diffusive .

Tectonic movements The genera l hypothesi s o f Cenozoic basin devel opment i n th e norther n Nort h Se a i s tha t th e basin wa s dominated b y thermal subsidenc e tha t followed th e Lat e Jurassic-Earl y Cretaceou s rifting (e.g . Sclate r & Christi e 1980 ; Badle y et al . 1988 ; Gabrielse n e t al . 1990) . Severa l workers hav e quantified and explaine d th e caus e of thi s subsidence . Thorn e & Watt s (1989) , White & Lati n (1993) . Hal l & Whit e (1994) , Lepercq & Gaultier (1996 ) and Nadin & Kusznir (1996); hav e inferre d a highe r subsidenc e rat e in Earl y Tertiar y tim e an d a lowe r rat e i n th e rest o f th e Tertiar y time . Th e subsidenc e dat a obtained b y Thorn e & Watt s (1989 ) (Fig . 10 ) were applied i n the initial simulation run . This is considered t o be a minimum subsidenc e pattern , because i t wa s a resul t o f backstrippin g usin g Airy isostasy, which favours isostatic compensa tion. However , Thorn e & Watt s (1989 ) appar ently di d no t conside r th e effect s o f uplif t o r subsidence o f basi n flanks . Therefore , tectoni c movements o n th e flank s wer e connecte d

(linked) t o th e qualitativ e interpretatio n o f Rundberg (1989 ) an d Jord t e t al . (1995). Only on e faul t i n th e basina l cross-sectio n i s thought t o hav e substantiall y influence d th e basin geometry . Th e faul t i s 4 6 k m fro m the wester n end o f the section an d cut s sequenc e CSS-1. wherea s sequenc e CSS- 2 i s unaffected . As thi s faul t ha s affecte d th e depositiona l pattern, i t i s include d i n th e simulation . Th e relative tectoni c dro p o f th e easter n par t o f the sectio n i s taken t o b e about 350m . an d th e faulting i s assume d t o hav e occurre d i n Lat e Paleocene time .

Eustasy The mode l require s th e inpu t o f sea-leve l variations through the geological time . The long-term curve of Haq e t al. (1987), therefore, was chosen to reflec t se a level . Th e curv e wa s sample d an d converted t o the Harland e t al. (1990) time scale . However, eustati c sea-leve l change s ar e no t within th e mai n scop e o f th e presen t study , a s the rat e o f sea-level chang e i n terms o f the long term curv e (Ha q e t al . 1987 ) i s to o lo w fo r it s effects t o be represented i n the depositional style . An exceptio n i s th e larg e fal l i n se a leve l fro m 15 Ma t o th e present, whic h affects th e sequence s CSS-7 t o CSS-10 . T o b e comparabl e wit h th e possible effect s o f th e short-ter m eustati c curv e of Haq e t al. (1987), the simulation ha s been ru n with 20 0 tim e steps . Th e result s confir m tha t there i s n o significan t difference wit h respec t t o the mai n simulatio n base d o n th e long-ter m eustatic curve , probabl y becaus e o f th e combi nation o f the depositional condition s in the dee p basin an d th e larg e scal e o f th e stratigraphi c approach. A n increas e i n th e numbe r o f tim e steps an d horizonta l resolution (numbe r of columns) migh t giv e greate r detai l o f change s i n a shallow-marine depositional system with respec t to th e short-ter m curv e of Ha q e t al . (1987) .

Isostasy The lithospher e i n th e Nort h Se a regio n wa s more stabl e i n Cenozoi c tim e tha n durin g Mid Jurassic t o Cretaceou s times . Th e regiona l tec tonic stres s i n Cenozoi c time s wa s apparentl y low, with the strain n o greater tha n th e develop ment of minor features. Therefore, n o axial stress has been applied i n the simulation, and a flexural rigidity o f th e lithospher e o f 1 . 0 x l 0 2 2 N m was used (Fjeldskaa r 1994) . The matri x densities for th e isostati c compensatio n hav e been take n as follows : 2.65gcm" 3 fo r sand , 2.72gem" 3 for shale , an d l.OSgcm" 3 fo r wate r (th e sam e

CENOZOIC TECTONI C SUBSIDENCE

285

Fig. 10 . Cenozoi c tectonic subsidence pattern give n by Thorne & Watts (1989). The subsidenc e analysis is based on backstrippin g wit h Airy isostasy (flexura l rigidit y zero) o f wells in the norther n Vikin g Graben area , an d suggests a n accelerate d subsidenc e i n Paleocene tim e followe d b y gentle subsidenc e rate s i n the res t o f th e Cenozoic period , reflectin g post-rif t thermal subsidence (McKenzi e 1978). In th e simulation , the subsidence data obtaine d b y Thorne & Watts (1989 ) were applied fo r modellin g case 1 . A flexural rigidity of 1. 0 x 10 22 N m was used .

286

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as used b y Sclate r & Christi e (1980)) . Fo r the crus t an d th e uppe r mantle , densitie s of 2.8gcm~ 3 an d 3.4gcm~ 3 , respectively , hav e been assumed .

Modelling results, case 1 The us e o f th e value s specifie d i n Tabl e 1 resulted i n a n overfil l o f th e basi n b y Earl y Oligocene time , an d subaeria l depositio n o f sequences younge r tha n Earl y Oliogocen e (se e Fig. 1 1 a). Therefore , i t i s no t possibl e t o com pare th e modelle d sequenc e geometrie s wit h th e observed geometries . Furthermore , th e model ling result s place d th e bas e Tertiar y to o shallo w at presen t time . Th e effec t o f introducin g Airy isostasy (flexura l rigidit y zero ) wa s tested , bu t was no t sufficien t t o avoi d earl y overfillin g (see Fig lib) , whic h occurre d i n Mid-Oligocen e times i n thi s model , o r t o improv e th e positio n of th e presen t dept h leve l o f th e bas e Tertiary . It i s eviden t fro m th e simulate d cross-sectio n

that th e basi n mode l underestimate s eithe r th e total subsidenc e o r the initial water depths alon g the cross-section . Accordingly , a ne w initia l form o f th e basi n (Fig . 12) , with a n asymmetri cal form , enhance d initia l wate r depth s i n it s eastern part , an d a sligh t increas e i n th e maxi mum initia l wate r dept h wa s suggested . The overall objective in this study was to carry out a simulatio n tha t coul d matc h th e observe d cross-section wit h respec t t o th e geo-histor y related t o th e basin-fillin g processe s an d th e observed geometrica l relationships . The eustatic sea-level, isostas y an d sedimen t inpu t wer e considered t o b e documente d value s an d i t was assume d tha t th e softwar e handle s erosio n and depositio n properly . Th e factor s tha t ar e associated wit h th e larges t uncertaintie s wer e therefore tectoni c subsidenc e an d wate r depths . We could, by an increas e in the maximu m initial water dept h o f th e basi n t o a n orde r o f 1000m , still use the subsidence rate s proposed by Thorne & Watt s (1989) . Thi s i s als o i n agreemen t wit h the wate r depth s fo r earl y Paleocene tim e in th e

Fig. 11 . (a ) The interactio n betwee n th e chose n value s faile d t o reproduc e the cross-section. The basi n wa s overfilled i n Early Oligocene time, (b) Use of D = 0, Airy isostasy, could no t improv e the results, as it also led to an overfille d basi n i n Lat e Oligocen e time .

CENOZOIC TECTONI C SUBSIDENCE

287

Fig. 12 . O n th e basi s o f the assumptio n tha t th e overfille d basi n i n Earl y Oligocene time in the mode l was mainly caused b y a n underestimate d wate r depth o f the initia l basin, a revise d initia l for m o f the basi n i n Lat e Paleocene time with increased wate r dept h along the cross-section was required. Also, the tectonic subsidence pattern was reconsidered t o obtai n agreement o f the modelled an d observe d geometrica l relationships .

northern Nort h Se a propose d b y Bertra m & Milton (1989 ) and Nadi n & Kusznir (1995) , bu t differs fro m wate r depths proposed b y Barton & Wood (1984 ) an d Gradstei n e t al (1994 ) fo r early Paleocen e time . However , Gradstei n e t al. (1994) an d Gradstei n & Backstro m (1996 ) suggested a rapi d deepenin g o f th e basi n fro m Early (200-50 0 m) to Lat e Paleocen e tim e (500 1000m). I t wa s therefor e decide d t o kee p th e maximum initia l water dept h a t a maximu m of c. 700m at the beginning of Late Paleocene time . Paleocene uplif t i n northwes t Britai n include s the shel f area s northwestwar d t o Shetland , which wer e associate d wit h th e earl y Tertiar y opening o f th e Nort h Atlanti c (Bot t 1975 ; Rochow 1981) . Delta-to p lignit e sequence s i n the Mora y Firt h Basi n suggest marginal marine to subaeria l sedimentatio n i n this area (Rocho w 1981; Ziegler 1990 ; Nadin & Kusznir 1995) . N o corresponding Paleocen e uplif t even t ha s bee n detected fo r th e Fennoscandia n hinterland , which probabl y wa s partl y submerge d durin g this perio d (Ziegle r 1990 ; Jordt e t al 1995) . I n principle, th e sedimentatio n patter n shoul d reflect th e basi n topograph y (Jord t e t al . 1995) . The depositiona l patter n o f th e firs t fou r sequences, CSS- 1 t o CSS-4 , i s characterize d by sediment suppl y predominantl y fro m th e wes tern source , wit h th e sediment s reachin g fa r t o the east . Thi s give s reaso n t o assum e tha t th e basin floo r wa s dippin g slightl y t o th e eas t during th e deposition o f the first four sequences . Therefore, a n asymmetrica l basi n form , wher e the westernmos t par t o f th e cross-sectio n i s characterized b y shallow water depths as a result of the uplifted sourc e area to the west, and water depths ar e abou t 7.00 m i n th e deepes t par t o f

the norther n Vikin g Graben, i s assumed fo r th e beginning o f Lat e Paleocen e time . Th e easter n part o f th e basi n i s anticipate d t o hav e ha d relatively dee p wate r depth s a t thi s time , indi cating submergin g o f th e westernmos t par t o f the Norwegian mainland . Several simulatio n run s usin g th e revise d initial for m o f th e basi n showe d tha t a n earl y overfill o f th e basi n persisted . Th e rapi d deepening o f th e basi n fro m Earl y (200-500 m) to Late Paleocen e tim e (500-1000m) recognized on biostratigraphi c dat a b y Gradstei n e t al . (1994) an d Gradstei n & Backstro m (1996) , i s assumed t o hav e ha d a tectoni c origin . There fore, i t was decided t o keep the maximum initial water dept h a t 700m , an d instea d increas e th e tectonic subsidenc e rate s compare d wit h thos e given b y Thorne & Watts (1989) . The tectoni c subsidenc e suggeste d b y Thorn e & Watts (1989) and Hal l & White (1994) is based on backstrippin g wit h th e respons e t o sedimen tary loading assumed t o be of local Airy-isostasy type (flexural rigidit y D = 0), thus giving a minimum value . T o improv e th e model , a litho spheric flexura l rigidit y o f D = 1. 0 x 10 22 Nm was applied . Therefore , subsidence rates should be increase d compare d wit h thos e give n b y Thorne & Watts (1989) . Geological studie s (Jord t e t al . 1995 , 1996 , 2000) suggest that th e basin was filled at th e end of sequenc e CSS- 5 time and locall y subjecte d t o episodic subaeria l erosion i n Mid-Late Miocen e time. Thi s scenari o ha s le d to a modification of the tectoni c subsidenc e (Fig . 13 ) patter n sug gested b y Thorn e & Watt s (1989) , wit h th e purpose o f avoiding complete basin filling before the depositio n o f the CSS- 5 sequence.

Fig. 13 . Tectoni c subsidenc e pattern s wit h a flexura l rigidit y of 1. 0 x 10 2 2 Nm a t selecte d position s alon g th e transect a s predicted b y the simulation . An accelerate d tectoni c subsidenc e pattern is detected fo r Paleocen e tim e along th e entir e cross-section. I n Lat e Miocene-Pliocen e time , tectonic subsidenc e o f th e centra l an d wester n part o f the basi n is detected, corresponding t o th e uplif t o f the eastern sourc e area (Fennoscandia) . Bot h o f these subsidence events are anomalou s relativ e to post-rif t therma l subsidenc e (McKenzie 1978) .

CENOZOIC TECTONI C SUBSIDENC E

Fig. 14 . Simulate d basi n geometr y an d stratigraph y a t selecte d tim e step s i n th e Cenozoi c basi n evolution .

289

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Modelling results, cas e 2 Figure 1 3 show s simulated tempora l subsidence for selecte d localitie s alon g th e cross-sectio n resulting fro m th e simulation . Accelerate d sub sidence is seen for the Paleocene CSS-1 sequenc e at al l positions alon g th e cross-section , an d th e subsidence rates are highest in the central part of the basi n (a t 90km , Fig . 13) . I n th e wester n (Fig. 13 , Okm an d 45km ) an d centra l part s o f the basi n (Fi g 13 , 90km) subsidence rate s i n the Eocene-Oligocene CSS- 2 t o CSS- 4 tim e sho w reduced, bu t stil l hig h rates . I t i s note d tha t sedimentation rate s outpace d subsidenc e i n th e western par t o f the basi n and th e Eocen e CSS- 2 sequence buil t u p t o se a leve l (se e Fig . 14) . For th e centra l par t o f th e basin , Oligocen e sedimentation rate s als o outpace d th e subsi dence rate s an d durin g Lates t Oligocene-Mid Miocene time, CSS-5 built up t o se a level at th e same tim e tha t subsidenc e ceased (se e Fig . 14) . During CSS-6 time, parts o f the basin were subject t o subaeria l erosion . I n th e easter n centra l part o f th e basi n (Fig . 13 , 130km) , subsidenc e rates wer e low i n Eocen e time . I n Lat e Eocen e time, subsidence ceased. In the eastern part of the basin (Fig . 13 , 180k m an d 240km) , subsidence ceased i n Eocen e CSS- 2 time . I n th e wester n (Fig. 13 , Okm an d 45km ) an d centra l part s o f the basin , a ne w phase o f subsidenc e started i n Mid-Miocene tim e (CSS-7 ) and laste d through out Pliocen e tim e (CSS-8) . Subsidenc e rate s fo r

Fig. 15. Modelle d (a) and observe d (b) cross-section.

the CSS- 7 tim e ar e relativel y gentle . Durin g deposition o f CSS-8, a considerable acceleratio n in subsidenc e occurred an d provide d accommo dation fo r th e Pliocen e CSS- 8 (Fig . 14 ) wedge, especially i n th e centra l par t o f th e basi n (Fig. 13 , 90 km and 13 0 km). The eastern part of the cross-sectio n (Fig . 13 , 180k m an d 240km ) shows uplif t throug h Oligocene-Miocen e tim e (CSS-3 t o CSS-7) , an d a prominen t uplif t i n Pliocene tim e (CSS-8). The uplift cause d erosio n of previousl y deposite d sequence s CSS- 1 t o CSS-4 from th e flank and deposition in a rapidly subsiding basi n i n th e west . Th e fina l cross section wit h time-lines is shown i n Fig . 14 . A compariso n o f th e simulate d an d th e observed basina l cross-section s (Fig . 15 ) shows some geometrica l discrepancies , mainl y cause d by smoothin g processe s resultin g fro m th e diffusive algorith m use d t o handl e erosio n an d deposition. Otherwis e ther e i s a strikin g similarity betwee n the two . The similarit y is sufficientl y high to render the simulation realistic and justif y the mai n conclusion s derive d fro m thi s study . This pertain s particularl y t o th e patter n o f tectonic subsidenc e o r uplif t (Fig . 13 ) implied by th e simulatio n an d th e patter n o f source area activity. Discussion The se t of factors used in the simulatio n is not a unique solution . Othe r combination s o f factors

CENOZOIC TECTONI C SUBSIDENC E could undoubtedl y giv e th e sam e results . Fo r instance, woul d inclusio n o f th e sourc e are a reveal a totall y differen t patter n o f transpor t coefficients? A differen t sea-leve l fluctuatio n curve would affec t th e depositiona l patter n an d the tectoni c subsidenc e pattern , bu t amon g th e published sea-leve l curves there are not curve s of such a radica l patter n t o contradic t ou r mai n results. An increase d time-ste p resolution would make i t possibl e t o simulat e a highe r orde r o f sea-level variation s an d t o predic t lithofacie s distributions. W e assum e tha t ou r documen tation an d choic e o f factor s i s reasonable , an d that th e trend s o f th e mai n results , therefore , are adequate . The subsidenc e curve s fo r differen t localitie s along th e transec t (Fig . 13 ) sho w a n anoma lous subsidenc e pattern , differen t fro m th e expected curv e o f post-rif t therma l subsidenc e (McKenzie 1978) , both fo r Lat e Paleocen e an d Late Miocene-Pliocen e times . Thorn e & Watts (1989), Whit e & Lati n (1993 ) Hal l & Whit e (1994) and Leperc q & Gaultier (1996 ) have als o pointed t o th e anomalou s subsidenc e pattern i n Paleocene time , wit h apparent subsidenc e accel eration. Th e Lat e Miocene-Pliocen e subsidenc e increase, u p to 1 km in magnitude, i s anomalous , for i t i s no t predicte d b y th e existin g mode l of lithospheric stretching , unles s a lat e phas e o f Tertiary stretchin g is assumed t o hav e occurred . However, ther e is little regional evidence of an y significant norma l faultin g in late Tertiary time. In thi s model , i t wa s assume d tha t som e o f the Paleocen e subsidenc e i s incorporated i n th e initial basi n form , becaus e th e simulatio n com mences wit h th e Lat e Paleocen e development . In Fig. 16 , the Cenozoic subsidence pattern for a pseudo-well, a t 110. 7 k m fro m th e section' s lef t

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margin (west ) (30/11-2) , ha s bee n plotte d i n combination wit h th e subsidenc e analysi s b y Hall & Whit e (1994 ) fro m wel l 30/9-1 , locate d relatively close t o the pseudo-well site . The plot, showing th e subsidenc e patter n fro m Jurassi c time t o th e present , indicate s anomalou s sub sidence rates in Paleocen e (1 ) (CSS-1 sequence) , and Lat e Miocene-Pliocen e times (2 ) (CSS-7 t o CSS-8 sequences) . The Paleocene subsidence rate may not be real, but rathe r a n artefac t o f the backstrippin g pro cedure, which may have underestimated palaeo bathymetries. Accordin g t o Nadi n & Kuszni r (1995, 1996) , n o mechanis m i s require d fo r the formation of Tertiary accommodation spac e other tha n post-Jurassi c therma l subsidenc e 'buffered' b y a n even t o f transien t Paleocen e uplift. Bertra m & Milto n (1989 ) came t o th e same conclusion : i f a wate r dept h o f 1 km i s assumed fo r th e Cretaceou s an d Tertiar y development o f th e norther n Nort h Se a basin, the n no post-Jurassi c riftin g need s t o b e invoke d i n the basi n t o explai n it s subsidenc e pattern . Th e latter workers also suggested an episode of uplif t in th e Paleocen e whic h affecte d muc h o f th e basin's northwes t margin . Thi s uplif t wa s sub sequently eliminate d b y th e margi n collapse , except fo r th e Inne r Mora y Firt h area . If th e Paleocen e subsidenc e rat e fro m ou r simulation is considered, and w e let it follow th e curve of post-rift thermal subsidence (McKenzi e 1978) (se e uppe r dashe d lin e i n Fig . 16) , th e reduced subsidenc e create s th e nee d fo r deepe r water depth s i n th e basin . I n thi s case , a maximum basin depth o f c. 1000m would account fo r the anomalous subsidence pattern o f Late Paleocene time. Water depths of that orde r hav e been suggested b y Jone s (1988 ) an d Gradstei n e t al.

Fig. 16 . Subsidenc e pattern for Earl y Jurassic time to th e presen t fro m a pseudo-well on transec t 2 . Anomalous subsidence events , which d o no t follo w th e expecte d post-rif t therma l subsidenc e pattern, ar e detecte d i n Paleocene (1 ) and Lat e Miocene-Pliocene (2) times.

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(1994), albeit thos e workers have also postulate d a rapid deepenin g of the basin in Paleocene time . The presen t simulatio n show s tha t th e basin , after th e Miocen e episod e o f subaeria l sedimentation, require s a n anomalou s amoun t o f sub sidence t o provid e sufficien t accommodatio n space fo r Lat e Miocene-Pliocene sedimentation. To ou r knowledge , such a n accelerate d Pliocen e subsidence ha s no t bee n recognize d i n th e northern Nort h Sea . A possibl e coeva l even t has bee n suggeste d b y Kooi e t al (1989 ) fo r th e North Sea' s Centra l Grabe n t o th e south, where the increase d subsidenc e i s attribute d t o a late stage puls e o f compression . The presen t stud y doe s no t revea l th e causa l mechanism t o explai n th e Paleocen e an d Lat e Miocene-Pliocene anomalou s subsidence , bu t indicates that suc h a subsidenc e pattern (Fig s 1 3 and 16 ) is required accordin g t o the input values, which are considered t o be realistic. The need fo r the anomalou s Paleocen e subsidenc e ca n b e avoided onl y i f th e initia l dept h o f th e basi n i s taken t o hav e bee n greater .

Conclusions The mai n conclusion s o f thi s stud y ar e a s follows. The numerica l simulatio n ha s produce d a lithostratigraphic cross-sectio n tha t matche s fairly wel l th e observe d basina l cross-sectio n (see Fig s 1 4 and 15) . The geometri c departure s are mino r an d d o no t significantl y affec t th e similarity o f th e simulate d cross-sectio n t o the observe d one . Th e visua l tes t render s th e simulation resul t satisfactory. The simulatio n indicate s anomalou s tectoni c subsidence, differen t fro m th e patter n o f post rift therma l subsidenc e predicte d b y genera l models, fo r Lat e Paleocen e an d Lat e Miocene Pliocene time . Thi s forme r episod e o f acceler ated subsidenc e ma y partl y b e a n artefac t o f the basin' s underestimate d initia l wate r depth , but i t i s mor e difficul t t o attribut e th e latte r episode t o a mistake n inpu t value . A n applica tion o f a modifie d sea-leve l curve wil l no t caus e any majo r chang e o f th e simulatio n result . A decrease d flexura l rigidit y will caus e a lowe r degree o f tectoni c subsidenc e durin g th e whol e time interva l considered . In short , th e resul t o f the simulatio n seems t o be vali d an d require s furthe r lithostratigraphi c analysis o f th e basin . We than k O . Hanse n o f IK U fo r hi s valuabl e contribution t o the computer simulation. We gratefully acknowledge th e constructiv e comment s mad e b y A. Ryseth. J. C. Rivenaes, W. Nemec and R. Gabrielsen .

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NE Atlanti c continenta l riftin g an d volcanic margi n formatio n JAKOB SKOGSEID, 1'3 SVERR E PLANKE, 1 JA N ING E FALEIDE, 1 TOM PEDERSEN, 2 OLA V ELDHOLM 1 & FLEMMIN G NEYERDAL 1 1

Department of Geology, University of Oslo, P.O. Box 1047 Blindern, N-0316 Oslo, Norway 2 Institute for Energy Technology, P.O. Box 40, N-2007 Kjeller, Norway ^Present address: Saga Petroleum ASA, P.O. Box 490, N-1302 Sandvika, Norway (e-mail: [email protected]) Abstract: Dee p seismic data fro m th e Hatton-Rockall region, the mid-Norway margin an d the SW Barents Sea provide images o f the crustal structure that mak e it possible to estimat e the relative amount s o f crustal thinnin g for the Late Jurassic-Cretaceous and MaastrichtianPaleocene N E Atlanti c rif t episodes . I n addition, plat e reconstructions illustrate the relativ e movements between Eurasia an d Greenland bac k to Mid-Jurassic time. The NE Atlantic rif t system developed a s a result of a series of rift episode s fro m the Caledonian orogeny t o early Tertiary time . The Late Palaeozoi c riftin g i s poorly constrained, particularly with respect t o timing. However, rifte d basi n geometries, inferre d t o b e of this age, are observed a t depth in seismic dat a o n th e flank s o f th e younge r rif t structures . Intra-continenta l riftin g i n Lat e Jurassic-Cretaceous time s cause d c . 50-70 km o f crusta l extensio n an d subsequen t Cretaceous basi n subsidenc e fro m th e Rockal l Trough-Nort h Se a area s i n th e south , t o the S W Barents Se a in the north . I n lat e Earl y t o earl y Lat e Cretaceous times, ne w riftin g occurred i n th e Rockal l Troug h an d Labrado r Se a associate d wit h th e northwar d propagation o f Nort h Atlanti c sea-floo r spreading . Whe n sea-floo r spreadin g wa s approached i n the Labrado r Se a the Rockal l rif t apparentl y becam e extinct . Th e fina l N E Atlantic rift episod e was initiated near the Campanian-Maastrichtian boundary , lasted until continental separatio n nea r th e Paleocene-Eocen e transition , an d cause d c . 140km extension. Th e lat e syn-rif t an d th e earlies t sea-floo r spreadin g period s wer e affecte d b y widespread igneous activity across a c. 300 km wide zone along the rifted plat e boundary. Th e deep seismi c dat a provid e lower-crusta l structura l geometrie s tha t represen t boundar y conditions fo r a better mappin g an d understandin g o f the extensional thinnin g o f the crust. The crustal geometrie s question extensio n estimates previously made fro m basi n subsidenc e analysis, an d ai d i n th e definitio n o f bodie s o f magmati c underplatin g beneat h th e oute r volcanic margins .

The openin g o f th e N E Atlanti c Ocea n a t th e base d o n interpretatio n o f al l availabl e magne Paleocene-Eocene transitio n marke d th e culmi - ti c dat a fo r th e N E Atlanti c (Fig . 1) . Finally, nation o f a c . 340 Ma histor y o f extensiona l de- a presentatio n o f differen t aspect s o f plume formation an d sedimen t basi n formatio n sinc e lithospher e interactio n an d associate d vertica l the en d o f th e Caledonia n orogen y (Fig . 1) . motio n durin g th e margi n formatio n i s linke d In this paper we utilize previously published cru- wit h a discussio n o f exploratio n challenge s a t stal transects betwee n th e Hatton-Rockal l mar- volcani c margins i n general, gin an d th e wester n Barent s Se a to evaluat e th e Dee p seismi c reflectio n profile s provid e degree o f crustal thinnin g in the various regions image s o f bot h Moh o an d intra-crusta l geome with respect t o th e rifting history of the area . tries , whic h i n particula r allo w u s t o bette r The evaluatio n include s firs t a discussio n o f constrai n th e dimension s o f th e Maastrichtian the observed structure s an d associated extensio n Paleocen e break-up-relate d rift , an d th e Lat e estimates mad e fo r separat e rif t episode s wit h Jurassic-Cretaceou s intra-continenta l rift . Inte respect t o a relativel y simpl e tectoni c model , gratio n o f deep seismi c reflection dat a wit h deep We the n attemp t t o illustrat e th e spatia l an d seismi c refractio n transect s provide s als o th e temporal geologica l evolutio n b y usin g exten - velocit y distribution o f th e lowe r crus t (Mutte r sion estimate s i n plat e reconstruction s an d e t al. 1984 , 1988 ; Gudlaugsson e t al. 1987 ; Hinz transect restorations . Th e reconstructio n t o e t al. 1987; White et al. 1987; Jackson e t al. 1990; magnetic anomal y 24 b tim e (5 3 Ma), th e oldes t Faleid e e t al . 1991 , 19936 ; Makri s e t al . 1991 ; magnetic spreadin g anomal y i n th e region , i s Olafsso n et al. 1991; Planke et al. 1991; Larsen & From: N0TTVEDT , A . e t al . (eds ) Dynamics o f th e Norwegian Margin. Geologica l Society , London , Specia l Publications, 167 , 295-326 . 1-86239-056-8/00/S15.0 0 © Th e Geologica l Societ y o f London 2000 .

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Fig. 1 . Magneti c sea-floo r spreadin g anomalie s (numbere d 5-7 , 13 , 20, 24) and structura l setting in th e N E Atlantic region. Transects 1- 7 discussed i n text are located b y bold lines . The outline of the flood basalt province in the regions off Britain and i n vicinity of the Faeroe Islands is based o n Hitche n & Ritchie (1993) and Boldree l & Andersen (1994) . BS, SW Barents Sea ; GB, Grea t Britain ; MM, Mor e margin ; NS. Nort h Sea ; RT. Rockal l Trough; VM , Vorin g margin ; HM , Hatton-Rockal l margin . Marcussen 1992 ; Shanno n e t al. 1993 ; Eldhol m & Grue 1994 ; Skogsei d & Eldholm 1995 ; Mjeld e et al . 1997 ; Christiansson e t al . 1999) . Such dat a reveal th e basi n geometries , an d allo w quantita tive evaluation an d modellin g o f the variou s rift episodes. Non e th e less , th e dee p seismi c data base i s limited . Moreover , multipl e rifting , extensive floo d basalt s an d sil l intrusion s alon g the oute r margi n mak e th e interpretatio n o f th e crustal structur e ambiguou s i n larg e part s o f the NE Atlanti c stud y area .

Regional geologica l framewor k Late Palaeozoic-Early Mesozoic phase The locatio n an d structura l expressio n o f th e late Palaeozoi c N E Atlanti c rif t syste m withi n the Caledonia n orogeni c domai n wa s influence d by Caledonian , an d possibl y pre-Caledonian , structures. Betwee n Norwa y an d Greenlan d

the rif t syste m followe d th e NE-oriente d Caledonides int o th e S W Barent s Sea , wher e north-trending structure s sugges t a structura l connection t o th e Arcti c rif t system . Betwee n Britain an d Greenlan d basi n trend s indicat e a possible reactivatio n o f th e Appalachia n defor mation system . A phas e o f lat e orogeni c extensiona l collaps e during lates t Silurian-Earl y Devonia n tim e apparently initiate d the formatio n o f a serie s of large half-grabe n basins , whic h subsequentl y were fille d wit h thic k succession s o f mainl y intra-continental deposit s (e.g . Ziegle r 1988 ; Coward 1993 ; Hart z & Andrese n 1995) . A n orogenic collaps e o f th e Barent s Se a Caledo nides southeas t o f Bj0rn0y a ha s bee n suggested , on th e basi s o f structure s beneat h th e flank s o f the S W Barent s Se a Permo-Carboniferou s rif t system (Gudlaugsso n e t al . 1994) . I f we conside r orogenic collaps e mainl y a syn-orogeni c phe nomenon, w e assum e tha t riftin g i n respons e

NE ATLANTI C CONTINENTA L RIFTIN G to lithospheri c extensio n betwee n Eurasi a an d Greenland wa s initiated a t th e end o f Devonia n time. Th e mai n Lat e Palaeozoi c rif t episode s took plac e i n Mid-Carboniferous , Carbonifer ous-Permian an d Permian-Earl y Triassi c time s (e.g. Ziegle r 1988).

Late Jurassic-Cretaceous phase During Mid-Jurassi c tim e th e developmen t o f oceanic crus t i n the central Atlanti c marke d th e onset o f a ne w kinematic regime in the Atlanti c domain (e.g . Larse n 1987 ; Ziegle r 1988) . Fol lowing the separation between Africa an d Nort h America, sea-floor spreading was confined t o the region sout h o f th e Azore s Fractur e Zon e fo r c. 50 Ma. Riftin g persisted , however , int o th e North Atlanti c domain , a s documente d b y rift associated sediment s o f Lat e Jurassi c ag e commonly foun d i n Nort h Atlanti c basins . Late Jurassic-Cretaceou s riftin g ha s i n thi s context bee n considere d a precurso r t o th e progressive continenta l separation i n th e south ern par t o f th e Nort h Atlantic , wher e th e final break-up betwee n Iberi a an d th e Gran d Bank s has bee n give n a Hauterivia n ag e (c . 134 Ma) (Srivastava & Tapscott 1986 ; Keen & de Voog d 1988; Whitmarsh et al. 1993). Thus, at the end of the Jurassi c perio d dee p rif t basin s probabl y existed fro m th e Rockal l Troug h t o th e S W Barents Sea, with a separate branch in the North Sea, wit h a possibl e linkag e t o spreadin g i n th e Tethys Ocea n (Lundi n & Dor e 1997 ) (Fig . 1) . Sea-floor spreadin g progresse d northward , an d the opening betwee n Newfoundland an d Goban Spur-Porcupine Ban k bega n i n lat e Aptia n to earl y Albia n (c . 112 Ma) time s accordin g to D e Graciansk y e t al . (1985 ) and Srivastav a et al . (1988). Further north , th e Labrado r Se a an d th e Rockall Troug h develope d contemporaneousl y with riftin g an d driftin g t o th e south . I n th e Labrador Se a gabbroi c dyke s o f Berriasian Valanginian (133-13 8 Ma) ag e i n southwes t Greenland ar e relate d t o incipien t riftin g (Wat t 1969; Larse n e t al . 1999) . O n th e basi s o f dat a from commercia l well s of f Canada , th e oldes t known rif t deposit s here are coarse-grained nonmarine sandstone s o f Barremia n age . Thes e sediments lie , however , partl y o n Neocomia n basalts (Berrasian-Barremian, Alexi s formation) (Balkwill et al. 1990) . The basalts are thus of the same ag e a s th e Greenlan d dyke s an d ma y suggest tha t eve n Uppe r Jurassi c sedimentar y rocks ma y occup y th e deep , undrille d part s o f some rif t basins . Th e younges t rif t deposit s sampled in wells are Coniacian in age. Extension

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in th e Labrado r Se a continue d t o continenta l separation betwee n Nort h Americ a an d Green land i n Lat e Cretaceou s (9 2 Ma, Roes t & Sri vastava 1989 ) to earlies t Tertiar y time s (62 Ma, Chalmers e t al . 1993) . The developmen t o f th e Rockal l Troug h i s more dispute d (Robert s 1974 , 1915a,b; Russe l 1976; KristofTersen 1978 ; Russel & Smythe 1978; Srivastava 1978 ; Robert s e t al . 1981 ; Smyth e 1989). Irrespectiv e o f th e natur e o f th e crus t flooring the troug h (oceani c o r continental ) i t is clear tha t significan t amounts o f extension hav e been take n u p alon g th e structur e (Joppe n & White 1990 ; Makri s e t al . 1991 ; Dor e 1992 ; Keser Neis h 1993 ; Knot t e t al . 1993 ; O'Reilly et al. 1996). There seems to be consensus about a Cretaceous ag e fo r th e las t majo r extensiona l phase i n th e area . Fro m plat e reconstructions , Srivastava & Verhoef (1992 ) assumed tha t 65 % of th e trough' s presen t widt h can b e accounte d for b y Lat e Jurassi c an d Cretaceou s extension . Further nort h the Late Jurassic boundary fault s along th e Wes t Shetlan d Platfor m wer e mildly reactivated i n Aptia n time , an d i n th e Faero e Basin, whic h ma y b e considere d th e norther n extremity o f th e Rockal l Trough , th e Earl y Cretaceous perio d wa s characterize d b y in creased subsidenc e rates (Duinda m & Van Hor n 1987; Hitchen & Ritchie 1987 ; Mudge & Rashid 1987; Nelson & Lamy 1987). Off Norwa y the M0re and V0rin g basins were formed primaril y a s a resul t o f th e Lat e Jurassic-Early Cretaceou s rifting , wherea s i t may b e questione d whethe r o r no t large-scal e extension too k plac e i n this region contempora neous wit h riftin g t o break-u p i n th e Labrado r Sea. Accordin g t o Surly k e t al . (1981) , coars e clastic submarin e fan s associate d wit h deep water shale s sugges t activit y alon g th e majo r boundary fault s i n centra l Eas t Greenlan d per sisting int o Aptian-Albia n time , an d coeva l activity ha s bee n reporte d alon g th e flank s o f main faul t zone s in the More and V0rin g basin s off Norwa y (Blysta d e t al . 1995 ; Dor e e t al . 19970). Furthermore , accordin g t o Dallan d (1981) Albian-Aptia n tectonis m i s reporte d from Andoya , norther n Norway , an d thre e Early Cretaceou s tectoni c phases , Berriasian Valanginian, Hauterivian-earl y Barremia n an d Aptian-Albian, hav e bee n reporte d fro m th e SW Barents Se a (Faleide e t al . \993a,b).

Late Cretaceous-Tertiary phase The fina l Maastrichtian-Paleocen e rif t episod e lasted fo r c . 20 Ma, leadin g int o continenta l separation an d onse t o f sea-floo r spreadin g a t

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the Paleocene-Eocen e transition . Thi s rif t epi sode forme d a mor e tha n 300k m wid e zon e associated wit h lithospheri c thinnin g an d post break-up subsidenc e (Skogsei d 1994) . In Paleo cene tim e (63-6 2 Ma) th e rif t wa s affecte d b y the impingemen t o f th e Icelan d mantl e plum e beneath th e thinne d lithosphere , whic h induce d considerable regiona l uplift. Subsequen t igneou s activity characterize d th e rif t history . The mag matism wa s governe d b y th e lithospher e relief , which provide d pressur e condition s (thin-spots ) for voluminou s mel t generatio n (Thompso n & Gibson 1991 ; Skogsei d e t al . 19920) . Th e con tinental separatio n an d initia l sea-floo r spread ing wer e associate d wit h a 2- 3 Ma perio d o f massive volcani c activit y alon g th e mor e tha n 2600km lon g ne w plat e boundar y (Eldhol m & Grue 1994 ) (Fig. 1) . Stretching estimate s In thi s stud y th e amoun t o f crusta l thinnin g is estimated fro m th e presen t thicknes s o f th e continental crus t alon g th e transect s i n Fig . 1 .

The thinnin g i s compare d wit h result s fro m independent estimate s base d o n standar d basi n subsidence analyses , i.e . unifor m extensio n models (McKenzi e 1978 ; Steckle r & Watt s 1978), applie d t o th e N E Atlanti c margin s (Skogseid e t al . 19926 ; Skogsei d 1994 ; Robert s et al . 1997 ; Walke r e t al . 1997) . Al l method s have limitations , and ideall y requir e inpu t dat a that ar e difficul t t o obtai n alon g mos t transects. Nevertheless, w e clai m tha t th e integratio n o f subsidence modellin g an d crusta l structur e evaluation provides the best platform for analysing crusta l extensio n an d fo r furtherin g ou r understanding of the extensional developmen t of the region . With referenc e t o th e riftin g histor y w e subdivide th e sedimentar y successio n i n th e crustal transect s int o thre e mega-sequence s (Figs 2 , 3 an d 4) . Th e pre-Cretaceou s strat a represent a n unspecifie d tectoni c an d basi n subsidence histor y fo r whic h littl e informatio n exists i n th e offshor e regions . Th e Cretaceou s sequence outline s th e basin s forme d afte r th e Late Jurassic-Cretaceou s rifting ; wherea s the Tertiar y sequenc e an d th e presen t wate r

Fig. 2 . Crusta l transect s 2 and 3 across th e norther n an d souther n V0rin g margi n (modifie d fro m Skogsei d & Eldholm 1995 ) with a n alternativ e interpretatio n o f the bas e Cretaceou s horizo n fro m Blysta d e t al. (1995) . Crustal velocitie s beneat h th e Voring Marginal Hig h (VMH ) ar e annotate d i n kms" 1 . COB , Continent-ocea n boundary; FG-GR , Fenris Graben-Gjallar Ridge ; HG , He l Graben; HHA, Helland-Hanse n Arch ; HT . Halte n Terrace; NH, Ny k High ; NR , Nordlan d Ridge ; NS , Nagrind Syncline ; TB , Troena Basin; VS , Vigrid Syncline . Location show n i n Fig . 1 .

NE ATLANTI C CONTINENTA L RIFTIN G depth approximat e th e amoun t o f subsidenc e since th e Maastrichtian-Paleocen e rif t episode . It shoul d b e noted , however , tha t a larg e Plio Pleistocene sedimentar y wedg e i s par t o f th e Tertiary mega-sequenc e i n th e easter n par t o f the transects . Thes e deposit s ar e relate d t o continental uplif t an d glacia l erosio n o f Fen noscandia (Rii s & Fjeldskaa r 1992 ; Stuevol d et al 1992) . The basi n analysi s depend s o n th e definition of th e bas e Cretaceou s an d th e bas e Tertiar y levels, separatin g th e thre e mega-sequences . These horizons , whic h correspon d t o majo r hiatuses o r condense d sequence s o n th e basi n flanks, ma y b e boundin g o r foun d withi n thic k sedimentary unit s reflectin g rapi d differentia l subsidence i n th e deepes t basins . I n fact , th e base Cretaceou s unconformit y doe s no t mar k the transition fro m syn-rif t t o post-rift sediments in th e Nort h Atlanti c rift , wher e riftin g con tinued well into Earl y Cretaceous times . Ideally, tectono-structural consideration s shoul d hav e

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been mad e wit h reference t o the top o f the Mid dle Jurassi c shallo w marin e sandstones , whic h marks th e onse t o f Lat e Jurassi c riftin g an d rapid subsidence . I t is , however , difficul t t o ascertain thi s leve l becaus e of the lac k of stratigraphic contro l i n th e dee p rif t basins . Similarly, th e Campanian-Maastrichtia n hori zon shoul d b e used t o boun d th e las t rif t episode. Non e th e less , her e th e near-bas e Tertiary rif t unconformit y is at presen t th e bes t approximation i n th e seismi c data . Erosio n across th e to p o f rotate d faul t block s beneat h this horizo n i s furthe r use d a s a n indicato r o f near sea-leve l conditions i n Paleocene time . The uncertainties in the seismic interpretation, which ar e subsequentl y implemente d i n th e modelling, ca n b e illustrate d bot h i n th e S W Barents Se a an d o n th e V0rin g margin . I n th e SW Barent s Se a th e bas e Cretaceou s horizo n represents a mai n rif t unconformit y o n th e flanks o f th e dee p Troms 0 Basin , wherea s i t i s located 1- 2 s tw t abov e th e to p Middl e Juras -

Fig. 3 . Crusta l transect s 1 and 4 across th e southwester n Barents Sea and M0r e margins, respectively (modifie d from Olafsso n et al . 1991 ; Faleide e t al . 19936 ; Breivi k e t al . 1998) . Crustal velocities are annotate d i n kms" 1 . COB, Continent-ocea n boundary ; HB , Hammerfes t Basin ; SB, S0rvestnaget Basin ; SFZ, Senj a Fractur e Zone ; SR, Senj a Ridge ; TB , Troms 0 Basin . Location show n in Fig . 1 .

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sic leve l i n th e basi n itsel f (Faleid e e t al. \993b). O n the V0ring margi n th e structural an d stratigraphical continuit y associate d wit h th e base Cretaceou s horizo n fro m th e Ra s an d Traena basin s westwar d int o th e Vigri d an d Nagrind synclines , respectively , i s poorl y con strained. Th e mai n reaso n i s tha t th e seismi c data beneat h th e Tertiary Helland-Hanse n Arc h are generall y poor , an d tha t igneou s intrusion s make th e structura l settin g difficul t t o interpre t with confidence . T o illustrat e th e variatio n i n interpretation w e hav e implemente d th e sug gested bas e Cretaceou s horizo n fro m Blysta d et al . (1995 ) (Fig . 2) . However , ou r interpreta tion i s aide d b y th e observe d lower-crusta l geometries. Th e dee p seismi c reflectio n dat a show a c . 45-70 km wid e zon e o f pronounce d crustal thinnin g beneath th e Ra s Basin , wherea s the Moh o interface lies significantly deepe r bot h eastward an d westwar d o f thi s zone . Thi s con figuration resemble s th e structura l settin g i n the Rockal l Trough , th e Mor e an d Troms o basins, an d th e norther n Nort h Sea , wher e a

similar spatia l correlatio n exist s betwee n th e upper-crustal rif t an d area s o f pronounce d crustal thinnin g (Fig s 2-4) . I n ou r opinio n w e find tha t a relativel y narro w centra l rif t basi n therefore i s more plausibl e tha n th e ver y broa d deep basi n suggeste d b y Blysta d e t al . (1995) . Furthermore, ou r interpretatio n allow s for thick units o f pre-Cretaceou s strat a beneat h th e Voring Basin , whic h i s consisten t wit h ob served stratigraphica l succession s bot h beneat h the adjacen t Trondela g Platfor m (Fig . 2 ) an d in Greenland . Voring margin transects We us e th e Vorin g margi n transect s (Fig s 2 . 5 and 6 ) t o illustrat e th e evaluatio n procedur e for th e crusta l thinnin g derive d extension . We interpre t th e c . 45 km lon g reflecto r segment at 8.5- 9 stwt (c . 20km) a s Moh o beneat h th e Ras Basin . O n eithe r sid e Moh o deepen s t o 10-1 Is twt . i.e . c.25k m (Skogsei d & Eldhol m

Fig. 4 . Crusta l transect s 5 and 6 and associate d degre e o f crustal thinnin g (3) acros s th e Rockal l Troug h (modified fro m Makri s e t al . 1991 ; Keser Neis h 1993 ; Shannon e t al . 1993) . Crustal velocitie s are annotate d i n kms -1 . (Note the Felativel y rapid deca y i n width o f the Rockal l Troug h betwee n th e two transects , whic h ar e located c . 200km apart. ) Locatio n show n i n Fig . 1 .

Fig. 5 . (a ) Transect 2 across th e norther n V0rin g margin , wher e th e thre e mai n tectono-stratigraphi c units , the base Cretaceou s interpretatio n fro m Blysta d e t al. (1995) and th e lowe r crusta l geometrie s ar e outlined . Abbreviations a s in Fig. 2 . (b) Transect backstrippe d t o th e base Cretaceou s leve l and isostaticall y balanced. Th e stretching factor , /3tot , illustrate s the tota l pos t mid-Jurassi c (bas e Cretaceous ) thinnin g calculated a s th e rati o between depth s t o isostati c an d backstrippe d Moho . (c ) Same a s i n (b) , but wit h the bas e Cretaceou s interpretation fro m Blysta d et al (1995). (Note th e extreme thinning beneath th e western portions o f the transect associated wit h this interpretation. ) (d ) A summar y o f extension estimate s made alon g th e transect . /9tot , tota l post mid-Jurassi c thinning ; (32 and /?3 , previously made subsidenc e derive d stretchin g estimate s fo r th e Lat e Jurassic-Cretaceous an d th e Maastrichtian-Paleocen e rif t episode s (Skogsei d e t al . 19926) ; /32_resulting , ne w estimates of Late Jurassic-Cretaceous thinnin g derived a s the ratio betwee n /?to t and /33. Where (3 3 is larger tha n /?tot, /32_resultin g i s se t t o 1.0 .

Fig. 6 . (a ) Transect 3 across th e souther n V0ring margin, where the thre e mai n tectono-stratigraphic units, the base Cretaceou s interpretatio n fro m Blysta d e t al. (1995) and th e lower-crusta l geometrie s ar e outlined . Abbreviations a s i n Fig . 2 and labe l as i n Fig . 5 . (b) Transect backstrippe d t o th e bas e Cretaceou s leve l an d isostatically balanced . Th e stretchin g factor , j3, illustrates the tota l pos t mid-Jurassi c (base Cretaceous ) thinning calculated a s the rati o betwee n depths t o isostati c and backstrippe d Moho . (c ) Same as in (b), bu t wit h the bas e Cretaceous interpretatio n fro m Blysta d et al . (1995). (Note th e broa d (c . 250km) area affecte d b y high thinning factors i n contrast t o the two more distinc t zone s o f thinning in (b).) (d) A summary of extension estimate s mad e along th e transect . /?tot , tota l pos t mid-Jurassi c thinning; (32 and /33 , previousl y made subsidenc e derived stretching estimates fo r th e Lat e Jurassic-Cretaceou s an d th e Maastrichtian-Paleocen e rif t episode s (Skogsei d et al . 1992/7) ; /32_resulting, ne w estimates o f Lat e Jurassic-Cretaceou s thinning derived a s describe d i n Fig . 5 . (Note th e larg e differences tha t exis t between /32_resulting an d (3 2 beneath th e dee p Ra s Basin. )

NE ATLANTI C CONTINENTA L RIFTIN G 1995). Th e overlyin g dee p Cretaceou s rif t trough i s betwee n 5 0 an d 70k m wide , where as th e Cretaceou s Vorin g Basi n i s c . 300 km wide. To the north, transect 2 reveals two Cretaceous depocentres , th e Traen a Basi n an d th e Hel Graben . Betwee n thes e depocentres , th e Nagrind Synclin e contain s a 3- 7 km thicknes s of Cretaceous strata above a very thick pre-Late Jurassic section . Skogsei d & Eldhol m (1995 ) suggested a shallower base Cretaceous horizon in this region, resulting in an even thicker pre-Late Jurassic succession . Recen t commercia l drilling on the Nyk High , well 6707/10-1, demonstrated , however, tha t Lat e Cretaceou s sediment s stil l exist at 503 9 m depth below sea level. The bas e o f th e crus t i s determine d b y integrating dee p seismi c refractio n an d wide angle reflectio n dat a (Plank e et al. 1991 ; Mjelde el a l 1996 , 1997) . Eas t o f th e Nagrin d an d Vigrid synclines refraction Moho (Skms" 1 ) corresponds i n dept h t o a dee p crusta l reflecto r (Fig. 2 ) (Skogsei d & Eldhol m 1995) . Towards the west , however , refractio n Moh o i s locate d deeper than this observed reflector. Her e the two levels boun d a 7 + kms"1 velocit y bod y inter preted t o represen t magmati c underplatin g emplaced durin g riftin g an d break-up . Thus , the reflecto r probabl y reflect s th e origina l base of th e crust , a ke y crusta l thinnin g factor . Th e underplating explains why the thinned crust ha s not subside d t o for m a dee p earl y Tertiar y rift basin , lik e th e dee p Cretaceou s Ra s Basin , which forme d a s a resul t o f th e Lat e Jurassic Cretaceous extensio n episode . Th e crus t again thicken s beneat h th e V0rin g Margina l High, presumabl y reflecting emplacemen t o f a n expanded igneou s sequenc e approachin g th e continent-ocean boundar y (COB) . The tota l amoun t o f crusta l thinnin g ca n b e derived fro m th e observe d crusta l thicknes s i f we ar e abl e t o determin e a reasonabl e esti mate o f th e initial , or pre-rift , crusta l thickness. In standar d basi n analysi s i t i s commonl y assumed tha t a 3 2-3 5 km thic k crystallin e crust of densit y 2.8gem" 3 , an d a 125-130k m thic k lithosphere i s balance d a t se a leve l (McKenzi e 1978). Hence, sediments will be deposited belo w sea level associated with lithospheric thinning or flexural bending. I n th e N E Atlantic , however , eclogites in Greenland an d Norway indicate synorogenic an d mayb e earl y post-orogenic crusta l thicknesses o f 80-9 0 km (Anderse n & Jamtveit 1990; Brueckne r e t al . 1999) . Furthermore , between th e Caledonia n front s i n Norwa y an d Greenland th e Lat e Palaeozoic-earlies t Meso zoic sediments are generall y of intra-continental and/or lacustrin e facies . Marin e deposit s wer e regionally firs t introduce d i n Permian-earl y

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Triassic times . Th e easter n Barent s Se a and th e southern Nort h Se a have a somewha t differen t palaeo-environmental history , bot h area s probably receiving large amounts of detritus from th e high-standing Caledonia n mountai n bel t (e.g . Coward 1993) . Our preferre d margi n stratigraph y include s very thic k sequence s o f uppe r Palaeozoi c an d lowest Mesozoi c sediments . I n term s o f th e erogenic an d earl y post-orogeni c histor y th e question o f reference crustal thicknes s i s important, an d standar d assumption s applied in basin modelling ar e considere d no t applicabl e t o th e pre-Late Jurassi c basi n history . Mor e impor tantly, th e so-calle d pre-rif t sediment s d o no t reflect amount s o f rift-relate d tectoni c subsi dence, a ke y facto r i n calculatio n o f extensio n factors fro m subsidenc e analysi s also fo r th e later rif t episode s (L e Pichon & Sibuet 1981) . None th e less , crustal thinnin g estimates can be mad e fo r th e Lat e Jurassic-Cretaceou s an d Maastrichtian-Paleocene rif t episodes , assum ing: (1 ) a regiona l marin e depositiona l environ ment i n Mid-Jurassi c time ; (2 ) a correctl y identified bas e Cretaceou s horizon ; (3 ) wate r depths wer e small an d ma y b e ignored ; (4 ) th e lithosphere ha d coole d ove r sufficientl y lon g time sinc e precedin g rif t episodes ; (5 ) isostati c conditions. B y backstrippin g th e crus t t o th e base Cretaceou s (Middl e Jurassic ) horizon , a n image o f th e pre-Lat e Jurassi c basi n an d to p basement configuratio n i s achieve d (Fig . 5b) . Isostatic balancin g o f thi s sectio n (assumin g an averag e -porosity o f 20 % i n sediment s with matrix densit y o f 2.7gm~ 3 , an d crystallin e crustal an d mantl e densitie s o f 2.7 5 an d 3.3gcm~3, respectively ) result s i n a n 'isostati c Moho', whic h i s use d a s ou r referenc e crustal thickness. Thus , th e tota l post-Mid-Jurassi c crustal thinnin g i s image d b y th e differenc e between th e isostati c Moh o an d th e backstrip ped presen t Moho . Thi s approach , lik e th e traditional subsidenc e analysis , account s fo r only compactio n o f th e sediments , wherea s i t does no t assum e extensiona l thinnin g o f th e sedimentary column . Th e latte r facto r ma y no t be importan t fo r smal l stretchin g factor s (/?
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et al. (1995). Th e averag e stretchin g factor s an d the integrate d amoun t o f stretchin g (ove r th e 325km portio n o f th e transec t wit h stretchin g factors >1.0 ) fo r th e tw o alternativ e interpreta tions ar e 1. 7 and 127k m fo r ou r interpretation , v. 2. 8 an d 200k m fo r th e interpretatio n o f Blystad e t al . (1995) . Similarly , th e souther n V0ring margi n transec t (390k m portio n wit h stretching factors >1.0 ) yields values of 1.7 5 and 170km and 2. 1 and 194km , respectively (Fig. 6). These result s d o no t alon e exclud e an y o f th e interpretations. W e shoul d kee p i n mind , how ever, tha t als o th e conjugat e eas t Greenland margin exhibits rifted basi n geometries reflecting extension tha t wil l ad d t o thes e estimates if they are related t o the same rif t episodes . In any case, the ver y hig h value s o f integrate d stretchin g indicate that th e width of the present margi n was increased almos t b y a facto r o f tw o durin g th e 170-55 Ma period . I n additio n t o ou r previou s comments o n th e upper - an d lower-crus t struc tural correlation , w e believ e th e extrem e max -

imum thinnin g ( 0 betwee n seve n an d eight ) suggested b y the Blysta d et al. (1995) interpreta tion argue s agains t th e possibilit y tha t Cretac eous sediment s ar e restin g directly on basement . To distinguis h betwee n th e Maastrichtian Paleocene an d Late Jurassic-Cretaceous compo nents of thinning we prefer to us e the previously subsidence define d stretchin g estimate s fo r th e Maastrichtian-Paleocene rif t episode . Ther e i s general consensus that th e Maastrichtian-Paleocene stretchin g affected a 150-20 0 km wid e zone landward of the continent-ocean boundary, with factors approachin g 2.0-2. 5 acros s th e Fenri s and He l graben s o n th e oute r margi n (Skogsei d et al . 19920 ; Skogsei d 1994 ; Roberts e t al . 1997 ; Walker e t al. 1997). The estimates fro m Skogsei d et al . (19920 ) an d Skogsei d (1994 ) wer e derive d from subsidenc e analyse s correcte d fo r reduce d subsidence becaus e o f magmati c underplating . The averag e stretchin g factor s an d amoun t o f extension fro m the stretching distribution show n in Figs 5d and 6 d (j33) ar e 1. 6 and 10 7 km an d 1. 6

Fig. 7 . Crusta l transects across th e V0rin g (a) an d Hatton-Rockal l (b ) margins illustrating crustal geometries, velocity structur e (annotation in kms" 1 ) an d modelle d thicknes s of'magmatic underplating' . COB , Continent ocean boundary . Location a s transect s 3 and 6 in Fig . 1 . Subsidenc e (continuous curve) an d crusta l thinning (stippled curve ) derived stretchin g distribution for th e Maastrichtian-Paleocen e riftin g episod e ar e shown . The modelled thickness of underplating is derived from th e difference betwee n the thinning derived and th e subsidence derived stretchin g factors as described b y Skogsei d (1994) . The Hatton-Rockal l transec t is modified fro m White et al . (1987 ) and Makri s e t al . (1991) .

NE ATLANTI C CONTINENTA L RIFTIN G and 104k m fo r th e souther n an d norther n transects, respectively . I t shoul d b e note d tha t the thicknes s o f underplatin g i s th e measure d 7 + km s-1 velocity body on the northern V0ring margin transect , whereas the thickness is derived from modellin g o n th e souther n Vorin g margi n transect (Skogsei d 1994 ) (Fig . 7) . Althoug h i t may b e questione d whethe r th e 7 + kms-1 velocity bod y ma y contai n som e intrude d lowe r continental crust , we consider thes e estimates as well constrained base d o n the present knowledge of the crustal structure , stratigraphic record an d the subsidence history. In fact, they are compar able wit h stretchin g estimate s o n typica l non volcanic margins (Ginzburg et al. 1985), but ten d to be high compared wit h estimates derived fro m amounts o f extensio n fro m faul t analysi s (Roberts et ai 1997 ; Ren et al. 1999). On bot h V0rin g margi n transect s th e Maas trichtian-Paleocene stretchin g estimate s corre late wit h th e tota l pos t Mid-Jurassi c crusta l thinning factor s acros s th e wester n province s (Figs 5 an d 6) . Alon g th e souther n V0rin g margin transec t thi s compariso n ha s n o signifi cance, a s the thickness o f modelled underplatin g (Fig. 7 ) wa s define d b y th e differenc e betwee n

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observed an d theoretica l tectoni c subsidenc e based o n th e crusta l thicknes s (Skogsei d 1994) . On th e northern transect , an d th e More margin transect (Fig . 8) , however , th e correlation s indicate tha t mos t o f th e observe d crusta l thin ning beneat h th e Ny k High-He l Grabe n an d western Mor e Basi n relate s t o th e las t rif t episode only . The Lat e Jurassic-Cretaceou s componen t o f the thinning , /?2_resulting , ca n b e compare d with th e subsidence-derive d estimate s o f Skog seid e t al . (19920,6) - W e not e tha t significantl y greater thinnin g seem s t o hav e take n plac e beneath th e Ra s an d Mor e basin s tha n pre viously suggeste d (Fig s 6 d and 8) . In these dee p basins, n o notabl e differenc e i s introduce d b y various interpretation s o f th e bas e Cretaceou s level, an d i t appear s tha t les s subsidenc e ha s taken plac e tha n woul d b e expecte d fro m th e crustal structure . Our interpretatio n o f onl y tw o mai n rif t epi sodes, th e Lat e Jurassic-Cretaceou s an d Maas trichtian-Paleocene phases, contrasts wit h those of Blysta d e t al . (1995 ) an d Lundi n & Dor e (1997). Thes e worker s argue d tha t additiona l extensional deformation affecte d th e regio n als o

Fig. 8 . Transect s 1 an d 4 acros s th e Wester n Barent s Se a an d th e M0r e margin s (modifie d fro m Faleid e et al . (I993b) an d Olafsso n e t al . (1991), respectively ) backstrippe d t o th e mid-Jurassic-bas e Cretaceou s level . Extension estimate s are mad e wit h sam e procedure a s described fo r Fig s 5 and 6 . /3tot, tota l post mid-Jurassi c thinning; (3 2 and /?3 , previousl y mad e subsidenc e derive d stretchin g estimate s fo r th e Lat e Jurassic-Cretaceous and th e Maastrichtian-Paleocen e rif t episode s (Skogsei d e t al . 1992/7) ; /?2_resulting , ne w estimate s o f Lat e Jurassic-Cretaceous thinning derive d a s described i n Fig . 5 . (Note large difference s between /?2_resultin g and (32 beneath th e Mor e Basin , wherea s /fto t an d (3 3 show ver y consisten t value s toward s th e wes t (continent-ocea n boundary).)

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in Lat e Cretaceous time , either as a separat e rif t episode o r a s continuous stretching . N o distinc t rift structure s or basin s of this age are identified, however. O n th e othe r hand , wit h referenc e t o the Troms 0 Basi n w e shoul d us e a broade r definition o f th e Cretaceou s rif t basins , indicat ing tha t substantia l amount s o f Lat e Jurassi c sediments ma y constitut e th e deepes t part s o f the rif t structures , an d tha t riftin g ma y hav e continued als o toward s lat e Cretaceou s time . Late lithospheri c thinnin g can , i n fact , partl y explain bot h th e relativel y larg e discrepanc y between subsidence and crusta l thinning derived stretching factor s beneat h th e deepes t Cretac eous basins (i.e. the apparen t lac k o f subsidence with referenc e to crusta l thickness ; /52-resultin g -32), a s wel l a s som e o f th e discrepanc y ob -

served betwee n th e extensio n derived fro m nor mal fault s an d fro m subsidenc e fo r th e las t rift episode .

Hatton-Rockall, More and SW Barents Sea transects The tota l crusta l thinnin g has als o bee n calcu lated fo r th e othe r transect s i n Fig . 1 , and i t i s shown tha t significantl y mor e pos t Mid-Jurassic extension ha s affecte d th e Hatton-Rockal l margin, includin g th e Rockal l Trough , tha n i s possible t o envisag e of f Norwa y an d i n th e western Barent s Se a (Fig . 9). The norther n Nort h Se a rif t ha s bee n eval uated b y Christiansso n e t aL (2000) , usin g a

Fig. 9. Crusta l thinning estimates for all transects stacked along the axis of the deep Late Jurassic-Cretaceous rif t zones. Th e fairl y consistent rif t dimension s fo r th e Maastrichtian-Paleocen e rif t episod e (/?3 ) on al l transect s should b e noted, an d likewis e for th e Lat e Jurassic-Cretaceou s rif t episode , /32_resulting , fo r al l transects fro m the M0r e margi n an d northwards . The Rockal l Troug h transects , o n th e othe r hand , sho w totally different rif t dimensions. ,3tot i s the tota l pos t mid-Jurassi c thinnin g derived a s describe d i n Fig . 5 .

NE ATLANTI C CONTINENTA L RIFTIN G comparable approach . The y showe d tha t th e basin wa s mainl y affecte d b y Lat e Jurassi c extension, an d tha t i t i s characterize d b y a narro w centra l rif t zone , 30-5 0 km, withi n a 250 km wide region affecte d b y Cretaceous post rift subsidence . Th e rif t follow s th e Lat e Palaeozoic-Early Mesozoi c basi n trend , an d i s surrounded b y rift flan k terraces . Th e estimate d stretching ove r th e tw o profile s discusse d b y Christiansson e t al. (2000 ) show s average s o f 1.2 an d 1.25 , respectively , resultin g i n c . 4050 km extension . West o f Britain , th e wid e Rockal l Troug h appears t o follo w th e SW-oriented Appalachia n trend. With reference to the above procedure, n o backstripping i s needed o f the Rockall transects , as th e existenc e o f pre-Cretaceou s sediment s i s speculative an d accordingl y no t include d i n this analysis. I n Fig . 4 a rapi d deca y i n th e mag nitude o f extensio n i s observe d fro m sout h t o north i n th e trough . Ove r th e distanc e o f c. 200km betwee n th e tw o transect s th e widt h of th e Rockal l Troug h i s reduced b y c . 100 km. Extension estimate s base d o n crusta l thicknes s show tha t th e averag e stretchin g facto r an d magnitude o f extension ar e 2. 9 and 315k m an d 2.9 an d 203k m fo r th e souther n an d norther n transects, respectively , assuming a 30 km pre-rif t crustal thickness . Farthe r wes t th e c . 300 km wide Hatton-Rockall Basin shows average thinning factor s o f c . 1.7, interprete d a s 126k m o f Maastrichtian-Paleocene extension according t o Skogseid (1994 ) (Fig . 7) . The dee p M0r e an d V0rin g basin s represen t the northwar d continuatio n o f th e Rockal l Trough an d th e Faeroe-Shetlan d Basi n trend . On th e M0r e margi n th e averag e stretchin g factors an d tota l amoun t o f extensio n ar e 1. 9 and 150km , o f whic h 87k m relate s t o th e Maastrichtian-Paleocene extensio n (Fig . 8) . Roberts e t al . (1990 ) mad e simila r extensio n estimates fo r th e Vikin g Graben , M0r e an d Faeroe-Shetland basins . Althoug h the y onl y estimated Lat e Jurassi c stretchin g (assumin g i t to b e c . 25% o f th e tota l Lat e Jurassic-Earl y Cretaceous episode) , the y conclude d tha t mos t of th e extensio n in th e M0re Basi n had t o hav e occurred wes t of Shetland. In th e S W Barent s Se a th e Lat e Jurassic Cretaceous crusta l thinnin g wa s severe , wit h maximum crusta l thinnin g value s approachin g four i n th e Troms 0 Basi n (Fig . 8) . Her e th e depth t o the Middle Jurassic horizon i s c. 13 km, and i n th e westernmos t basin, th e S0rvestnage t Basin (Fig . 3) , th e dept h i s comparabl e o r greater. Th e extensiona l development o f thes e two basin s appear s t o b e similar , givin g rise t o thick Uppe r Jurassic-Lowe r Cretaceou s syn-

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rift sequence s (Breivi k et al. 1998) . The fina l rif t phase i n the Late Jurassic-Cretaceous episode is well constraine d t o b e o f Aptia n ag e i n thes e western Barent s Se a basins , wherea s th e sub sequent perio d wa s characterized b y rapid post rift subsidenc e an d sedimentation . Mos t o f th e structural relie f wa s covere d b y Cenomania n time (Faleid e e t al . 1993<2 , b). Th e tota l crusta l thinning derive d stretchin g distributio n show s an averag e valu e o f 2.2 , yieldin g 100k m ex tension ove r th e <200k m lon g transect, o f which the Maastrichtian-Paleocene episode may account fo r c . 30 km. I n th e Troms 0 Basi n removal o f materia l belo w th e Middl e Jurassi c horizon b y sal t diapiris m ha s amplifie d th e subsidence. Our thinning estimates are therefore considered maximu m values . Restored basin geometries and palaeogeography In thi s stud y w e hav e attempte d t o quantif y the amoun t o f relativ e movement s betwee n th e continents associate d wit h th e respectiv e exten sional episodes , an d spatiall y illustrat e thi s o n maps an d alon g transects .

Break-up setting A plate reconstruction t o magnetic Chro n 24n.3 (c. 53 Ma) represent s th e firs t tim e slic e i n thi s restoration, an d show s a palaeogeographica l view o f th e N E Atlanti c configuratio n shortl y after break-u p (Fig . 10) . Th e 18-2 0 M a Maas trichtian-Paleocene rif t episod e pre-datin g th e final continental separation wa s characterized b y normal faultin g ove r broa d regions , wherea s during th e lat e syn-rif t perio d mor e focuse d extensional deformatio n wa s associate d wit h uplift an d erosio n o f th e centra l rif t zon e (Skogseid e t al . \992a). Eventually , break-u p was accompanie d b y massiv e volcani c activity, which covere d ver y larg e province s generall y along th e break-u p axis . Th e distributio n o f flood basalts , a s show n i n Fig . 10 , presumably outline the regions that both had a high potential for mel t generation (e.g . thinnes t lithosphere ) and ha d dominantl y a subaeria l depositiona l environment during break-up. The extensio n estimate s fo r th e Maastrich tian-Paleocene rif t episod e ar e ver y consisten t from on e margi n segmen t t o th e other , an d viewed in a conjugate margin setting the amount and distributio n o f stretchin g see m t o resembl e estimates made across typica l non-volcanic margins (Skogsei d 1994) . Alon g th e N E Europea n

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Fig. 10 . Plat e reconstructio n at 5 3 Ma (magneti c Chron 24n.3 ) showing the NE Atlanti c rift syste m shortly afte r continental break-up . Emphasi s is placed on the Late Palaeozoic-Early Mesozoic. Lat e Jurassic-Cretaceous and the Maastrichtian-Paleocene rift zones . The las t i s generally illustrated b y the Earl y Cenozoic volcani c overprint along th e youn g plat e boundary . Th e smal l an d larg e circles, respectively , illustrat e the suggeste d siz e of th e Iceland mantl e plume befor e and afte r spreadin g beneat h th e N E Atlanti c lithosphere. assuming a symmetrical distribution. The circles are centred a t the projected locatio n o f the Iceland mantl e plume according to Lawve r & Miiller (1994) . (For other abbreviations , se e Fig. 1.) margins th e area s tha t hav e bee n affecte d b y increased Cenozoi c subsidenc e ar e c . 200 250km wide , an d coincide wel l wit h th e regio n affected b y extensiona l deformatio n b y faultin g and igneou s activit y (Skogseid e t al. \992a). Th e discrepancies tha t exis t betwee n extensio n esti mates mad e fro m structura l modellin g o n on e hand, an d fro m crusta l structure an d subsidenc e analysis o n th e other , ma y b e ascribe d t o limitations i n seismi c resolution , whic h caus e small fault s t o b e lef t ou t (Robert s e t al . 1997 ; Walker e t al . 1997 ; Re n e t al . 1999) . Mor e important, however , ar e th e seismi c imagin g problems associate d wit h th e igneou s rocks . If we assume break-u p too k place near the centre of the rif t zon e w e can conclud e that between 50 and 15 0 km o f th e centra l rif t structur e (o n th e European side , only ) i s a t presen t covere d b y volcanic rocks . I n thi s context , th e Fenri s an d

Hel grabens , o n th e Vorin g margin , represen t the exception s withi n th e N E Atlanti c realm , being locate d i n a regio n wher e the floo d basal t province i s particularl y narro w (c . 50km) (Fig . 10). Here , ver y well imaged extensiona l terrane s are locall y dominate d b y low-angl e norma l faults an d associate d lower-crusta l extensiona l unroofing (Lundi n & Dor e 1997 ; Re n e t al . 1999). Wit h respec t t o th e observe d abrup t crustal thinnin g beneat h thes e faul t zone s w e interpret the m a s th e mai n boundar y fault s o f the centra l break-up-relate d rif t zone .

Cretaceous setting; 75 Ma The plat e restoratio n t o 7 5 Ma (Fig . 11 ) and th e respective shortenin g implemente d o n th e trans ects (Fig s 12-14 ) compensate s fo r th e amoun t

NE ATLANTI C CONTINENTA L RIFTIN G

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Fig. 11 . Plat e reconstruction a t 7 5 Ma (c . magnetic Chron 33n ) based o n pre-break-up stretchin g estimates according t o Fig. 9. Emphasis is placed o n th e Lat e Jurassic-Cretaceous central rif t zone s and th e Late Palaeozoic-Early Mesozoi c basins. (Note the close spatial correlation between the Late Jurassic-Cretaceous central rif t zon e on th e present N E Greenlan d margin and thos e at th e outer V0ring margin and i n the SW Barents Sea. Also note the pronounced fan-shap e of the Rockal l Trough.) Locations of crustal transects (2, 3, 4 and 7 ) are shown . (For other abbreviations , see Fig. 1.)

of extensio n derive d fo r th e Maastrichtian Paleocene rif t episod e (Fig . 9) . Th e widt h o f the Cretaceous basins has, thus, been reduced by c. 100km. Th e reconstructe d transect s ar e gen erated assumin g pur e shea r thinnin g o f th e crystalline crust, whereas the 'pre-rift ' sediment s are decompacte d only , i n genera l agreemen t with the way the stretching estimates were made. We realiz e that thi s restoratio n approac h mus t be considered wit h caution, bu t believ e that th e restored section s sho w mor e realisti c basi n configurations an d regiona l rif t geometrie s than thos e observe d o n unrestore d sections . I n combination wit h th e map s the y ma y carr y important information on provenance and basin fill relation s i n th e area . Th e structura l an d stratigraphic detail s are unreliable , however. The transect s sho w tha t th e Lat e Jurassic Cretaceous centra l rift zones had widths between 30 an d 60km , wherea s th e post-rif t regiona l

subsidence forme d Cretaceou s basin s 150 250km wide . Th e M0r e an d souther n V0rin g basins ar e fairly symmetrica l about thei r central rift zones , wherea s th e norther n V0rin g Basi n reveals three main depocentres. Here , the central depocentre, the Nagrind Syncline , is interprete d as a deep synform betwee n two 'real ' rif t basins, the Traena Basin and the Hel Graben (Figs 2 and 14). I n th e Nagrin d Synclin e post-Campania n sediments seem to represent a significant portio n of the Cretaceou s succession . The reconstructio n (Fig . 11 ) demonstrate s our understandin g o f ho w th e dee p Lat e Jurassic-Cretaceous centra l rif t structure s wer e located, locally revealing overlapping rif t zones . Off N E Greenlan d a 'N E Greenlan d mar gin basin ' i s interprete d fro m aeromagneti c data (Fig . 15) . Here , a magneti c quie t zon e obliquely truncates coastal parallel N-S trendin g anomalies. Th e latte r ar e interprete d t o reflec t

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Fig. 12 . Simpl e restoration of transect 4 across the Mor e margin, where shortening and thickenin g of the crystalline crust are determined accordin g to extension estimates for the Maastrichtian-Paleocene rift episod e (/33, 7 5 Ma), and th e Lat e Jurassic-Cretaceous rif t episod e (/32_resulting , 17 0 Ma), respectively. The sediment successions are decompacted, only . The 17 0 Ma restoration is combined with transect 7 across the Jameson Land Basin (JB) in Greenland (modifie d from Larsen & Marcussen 1992 ; Fig. 1 for location). Here, very thick Devonian to Jurassic deposits fill in a large half-graben basin bounded eastward by the Liverpool Land (LL) basement block. a Palaeozoic-Earl y Mesozoi c structura l grain . The change i n trend betwee n the quiet zone an d the coastal parallel structures is similar to what is observed betwee n th e Palaeozoic-Earl y Meso zoic grain and the Late Mesozoic basins of! mid-

Norway. Th e dee p basin s of f Norwa y were , i n fact, als o firs t suggeste d base d o n thei r smooth magnetic signature (Am 1971) . With respec t t o th e dimensio n o f th e dee p Late Jurassic-Cretaceou s rif t structures , i t i s

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Fig. 13 . Simpl e restoration o f transect 3 across th e souther n V0ring margin, where shortening an d thickenin g of the crystalline crust are determined accordin g t o extension estimates for the Maastrichtian-Paleocene rif t episod e (/?3, 7 5 Ma), and th e Lat e Jurassic-Cretaceou s rif t episod e (/?2_resulting , 17 0 Ma), respectively. The sedimen t successions ar e decompacted , only . difficult t o envisag e tha t a northwar d continua tion fro m th e V0rin g t o th e Barent s Se a basins coul d tak e plac e withou t involvin g th e present N E Greenlan d margin , an d i t therefor e gives confidenc e tha t th e relocatio n o f th e 'N E Greenland margi n basin ' merge s wit h th e He l Graben i n th e norther n V0rin g Basin , th e Harstad Basi n of f norther n Norway , an d th e Troms0 Basi n i n th e S W Barent s Sea . Th e obliquely cuttin g lin e o f continenta l break-up ,

thus, subsequentl y cause d a potentiall y larg e Cretaceous basi n to be found on the present N E Greenland margin . With respec t t o th e Labrado r Se a an d th e Rockall Trough rif t history , we propose a model where th e southwar d fan-shap e o f th e Rock all Troug h reflect s a faile d rif t formation . Th e trough wa s forme d a s a relativel y shor t rif t branch withi n th e large-scal e Nort h Atlanti c rift system , wher e th e Labrado r rif t too k ove r

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Fig. 14 . Simpl e restoration o f transect 2 across th e norther n Vorin g margin, wher e shortening and thickenin g of the crystalline crust ar e determined accordin g t o extension estimates for the Maastrichtian-Paleocen e rif t episod e (.33, 7 5 Ma), and th e Lat e Jurassic-Cretaceou s rif t episod e (,32_resulting , 170 Ma), respectively . The sedimen t successions ar e decompacted , only . "all" tectoni c activit y betwee n Coniacia n an d Campanian-Maastrichtian times . I n thi s mode l the mid-Lat e Cretaceou s perio d appear s no t t o have represente d a n epoc h o f renewe d large scale basin-formin g extension betwee n Norwa y and Greenlan d nort h o f the Rockall Trough. We recognize, a s outline d above , tha t som e tectonic activity occurre d alon g som e o f th e majo r faul t zones (e.g . Faleid e e l al. 19936 ; Blysta d e t al. 1995; Lundi n & Dore 1997) .

Pre-Late Jurassic selling; 170 Ma In th e reconstructio n t o 17 0 Ma w e hav e com pensated fo r c . 60 km o f Lat e Jurassic-Cretac eous stretchin g (Fig . 16) . W e not e tha t th e remaining Lat e Palaeozoic-Earl y Mesozoi c basin provinc e i s les s tha n 400k m wid e i n east-west direction , locate d betwee n th e Cale donian front s i n Greenlan d an d Norwa y within a rang e o f 650-110 0 km. Thes e dimensiona l

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Fig. 15 . Magneti c anomaly wiggle plot from th e NE Greenlan d continental margin. Positive magnetic anomalies (grey), interpreted coastal paralle l lineations (thick continuous lines) and a zone o f 'smooth magnetic signature' (blue area) ar e identifie d beneat h the shelf , whereas sea-floor spreadin g anomalies 13 , 20, 22 and 2 4 are show n in the ocea n basi n (red lines). GFZ, Greenland Fractur e Zone .

relations ar e importan t wit h respec t t o ou r understanding o f th e Caledonia n orogen y itself, an d t o provenanc e relation s bot h durin g the earl y stag e non-marin e environmen t an d subsequently durin g th e developmen t o f marin e transgressions fro m bot h nort h an d south . Basin formatio n associate d wit h the orogeni c collapse of the Caledonide s is generally reflected by dee p half-grabe n structure s o f mainl y Devo nian ag e an d regiona l depositio n o f continental red sandstone s (e.g . Hossach 1986 ; Duindam & van Hoor n 1987 ; Marcusse n e t al 1987 ; Ziegle r 1988; Hart z & Andrese n 1995) . Subsequen t episodes o f lithospheri c extensio n hav e bee n reported, o f Late Carboniferous-Early Permian , Permo-Triassic, Mid-Late Triassic and TriassicEarly Jurassi c age s (e.g . Surly k e t al . 1984; Ziegler 1988 ; Surlyk 1991) . However , th e earl y non-marine depositional setting , the general lack of good stratigraphi c correlation, and th e totally unknown structur e o f th e orogeni c crus t an d lithosphere mak e basi n modellin g and extension estimates for thes e basin s a speculativ e task. In th e S W Barent s Se a th e Lat e Palaeozoi c tectonic development gav e rise to a 300 km wide rift zone , th e Nordkap p Basin , suggeste d b y Gudlaugsson e t al . (1994 ) t o exten d a t leas t 600km northeastward int o th e Barents Sea. The

rift forme d mainl y durin g Mid-Carboniferou s times, probabl y linkin g up wit h Carboniferous Early Permia n extensio n a s fa r southwestwar d as th e Jameso n Lan d Basi n i n Greenlan d (Surlyk 1991) , and causin g th e establishmen t o f a marin e depositiona l environmen t alon g thi s axis durin g Lat e Permia n time . There i s onl y littl e direc t informatio n o n th e pre-Mid-Jurassic basi n histor y beneat h th e deep S W Barents Se a basins, althoug h th e dee p seismic reflectio n an d refractio n dat a indicat e a considerable sectio n o f ol d sediment s (Fig . 3). In th e Troms 0 Basi n th e dee p seismi c dat a indicate a crystallin e basemen t a t 18-2 0 km (Gudlaugsson e t al . 1987 ; Jackson e t al . 1990; Faleide e t al . 19936) . Regiona l consideration s combined wit h th e presenc e o f sal t i n th e Tromso an d Sorvestnage t basin s sugges t tha t the pre-Middl e Jurassi c sequenc e probabl y includes thic k Triassi c an d Jurassi c clasti c sediments an d Permo-Carboniferou s mixe d carbonates, evaporite s an d clasti c deposit s (Faleid e et a l 19930 , b\ Gudlaugsso n e t al . 1994) . Basin characteristics, sediment thicknesse s an d th e salt diapirism can be compared wit h observations on the NE Greenlan d shel f (Hinz et al. 1991 ; Escher & Pulvertaf t 1995) , wherea s simila r sal t diapir s are no t reporte d of f mid-Norway.

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Fig. 16 . Reconstructio n a t 17 0 Ma base d o n tota l pos t mid-Jurassi c stretchin g estimates accordin g t o Fig . 9 revealing a c.400k m wid e basin provinc e betwee n Greenland an d Eurasia . Locatio n o f crustal transect s (2 , 3, 4 and 7 ) discussed i n thi s respec t ar e shown . The remainin g fan-shaped Rockal l Troug h i s left shade d partl y t o illustrate th e siz e of th e regio n w e believe may hav e a separat e histor y in term s o f riftin g and/o r driftin g a s described i n text . LI , Lofote n Islands . (Fo r other abbreviations , se e Fig. 1.)

In contras t t o th e Lat e Carboniferous-Earl y Permian marin e transgressio n i n th e Barent s Sea, an d it s presumabl e continuatio n off - an d onshore East Greenland, a generally non-marin e depositional environmen t persiste d of f mid Norway an d i n th e norther n Nort h Se a unti l Late Triassi c tim e (e.g . Eid e 1989 ; Stee l 1993) . With respec t t o Fig . 16 , we may thu s infe r tha t a structural and/o r depositiona l boundar y existe d west o r southwes t o f th e Lofote n Islands , sepa rating a Carboniferous-Permia n marin e trans gression an d associate d evaporit e basi n nort h and wes t fro m th e area s t o th e sout h an d east characterize d b y continuou s non-marin e conditions. In th e transec t restoration s (Fig s 12-14 ) removal o f th e Lat e Jurassic-Cretaceou s exten sion provide s image s o f th e pre-Lat e Jurassi c basin configuration , whic h demonstrate s ver y thick loca l depocentre s a s part s o f a regiona l

basin province . Th e M0r e margi n transec t (Fig. 12 ) is relocate d t o b e almos t conjugat e t o the Jameso n Lan d Basi n i n Greenlan d (wher e the Ja n Maye n Ridg e appear s t o hav e repre sented a n intermediat e basemen t high) . Alon g the transec t onl y moderat e thicknesse s o f pre Late Jurassic sediment s can b e imaged, wherea s in th e Jameso n Lan d Basi n as much a s 16k m of Devonian t o Jurassi c sediment s hav e bee n sug gested, o n th e basi s o f seismi c reflectio n dat a (Larsen & Marcusse n 1992 ) ( a significantly smaller thicknes s o f 8-1 0 km is , however , inferred fro m seismi c refractio n dat a accord ing t o Fechne r & Joka t (1996)) . Accordin g t o Marcussen e t al. (1987), the successio n generally consists o f Devonia n braide d rive r an d aeolia n sediments foun d deposited in a rapidly subsidin g intra-mountain basi n setting , underlying Carboniferous t o Lowe r Permian coarse-graine d sand stones, re d siltstone , organic-ric h blac k shal e

NE ATLANTI C CONTINENTA L RIFTIN G and coa l bed s o f fluviatile and lacustrin e origin . The Uppe r Permia n sequenc e show s conglom erate an d blac k shale , an d thi n interbedde d layers o f evaporit e ar e noted . Th e Lowe r Tri assic formation s consis t o f finel y laminate d marine sil t an d sandstone , wherea s Middl e an d Upper Triassi c sequence s ar e continenta l t o shallow marin e i n origin . Along th e souther n V0rin g margi n transec t (Fig. 13 ) ver y thic k succession s o f pre-Lat e Jurassic sediment s exis t beneat h th e Tr0ndela g Platform are a clos e to the Norwegian mainland , and beneat h th e Vigri d Syncline , in th e centra l parts o f th e transec t (Fig s 2 an d 13) . Th e stratigraphic interpretation s assum e a substan tial accumulatio n o f Permo-Triassi c sediments , but Devonia n sediment s may also be present in a number o f deep half-grabe n basins . The transec t connects t o th e Greenlan d margi n betwee n th e Triall 0 an d Hol d wit h Hope , a n are a tha t is characterize d b y thic k sequence s o f Carboni ferous t o Jurassi c unit s overlai n b y thin ner squence s o f Cretaceou s shale s generall y found o n th e footwal l sid e o f majo r Lat e Jurassic faul t block s (e.g . Surly k 1991 ; Pric e & Whitham 1997) . Further north , th e norther n V0rin g margi n transect stil l reveal s thic k unit s o f pre-Lat e Jurassic sediment s i n wha t ma y appea r t o b e deep half-grabe n structure s (Fig s 2 an d 14) . In vie w o f th e plat e reconstruction s (Fig s 10 , 11 and 16 ) it i s temptin g t o us e th e dee p Jameso n Land Basi n (Fig . 12 ) a s a n analogu e fo r thi s basin configuration . Th e easter n flan k o f th e Nagrind Syncline , th e Utgar d High , exhibit s a crystalline basement hig h adjacent to the deepest pre-Late Jurassi c basi n (Plank e e t al. 1991 ; Skogseid e t al . 19926) . W e interpre t thi s t o indicate a simila r dee p half-grabe n basi n a s o n East Greenland , wher e th e Utgar d Hig h i n this context resemble s th e Liverpoo l Lan d basemen t block. Westwar d th e norther n V0rin g margi n transect connect s t o Greenlan d wher e th e shel f starts t o becom e broader , bu t wher e Tertiar y lavas inhibi t goo d seismi c imagin g o f th e sub stratum (Larsen 1990 ; Escher & Pulvertaft 1995) . Onshore, i n th e Wollaston e Forelan d province , however, th e stratigraph y i s characterize d b y Upper Mesozoi c sediment s deposite d i n Lat e Jurassic-Cretaceous extensiona l half-grabens directly ont o Caledonia n basemen t (Surly k 1991) . The magnetic lineations on the NE Greenlan d margin (Fig . 15 ) may b e correlate d wit h base ment structures , which , furthermore , ma y resemble metamorphi c rock s suc h a s thos e o f the Lofote n peninsula r islands , th e Utr0s t an d Utgard high s of f mid-Norway , an d th e pro nounced basement-relate d structure s alon g

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the easter n Mor e margi n an d beneat h th e More Basi n (i.e . th e las t generall y reflecte d b y gravity an d magneti c lineations , e.g . Dore et al . 19976). The interpretatio n o f th e Nagrin d Synclin e (Fig. 2 ) a s primaril y a pre-Jurassi c basi n con trasts severel y with th e interpretatio n presente d by Blysta d e t al . (1995) . I n thei r interpretatio n no pre-Cretaceou s strat a hav e existe d i n thi s region, indicatin g a s muc h a s 1 5 k m o f Cretac eous sediment s in the syncline. We argue agains t this interpretatio n primaril y o n th e basi s o f the crusta l configuratio n combinin g reflectio n and refractio n dat a a s discusse d previously . We further thin k the observatio n o f pre-Jurassic strata bot h beneath th e Trondelag Platform an d offshore Greenlan d indicate s tha t th e Vorin g Basin ha s als o bee n par t o f th e pre-Jurassi c basin provinc e (Fig . 16) . Th e possibilit y o f a somewhat deepe r leve l fo r th e bas e Cretaceou s reflector tha n presente d i n ou r interpretatio n can, however , no t b e rule d out . Wit h referenc e to th e Troms0 Basi n in the western Barents Sea , it may be that thi s inter-basinal syncline trapped significant thicknesse s o f lat e Jurassi c an d possibly earl y Cretaceous sediments . In th e norther n Nort h Se a Lat e Palaeo zoic rifting i s best documented i n marginal area s with respec t t o th e Vikin g Graben . Th e seis mic dat a displa y a tilte d fault-bloc k geometr y where sediment s probably o f Permian ag e fill in the tectoni c relie f an d ar e subsequentl y burie d by Triassic deposits. The initiatio n of this extensional tectonis m i s no t dated , bu t fro m obser vations mad e fro m th e seismi c dat a combine d with mor e regiona l information , i t i s generally assumed t o hav e bee n durin g Permia n tim e (Faerseth e t al . 1995 ; Faleide e t al . 2000) . Fur thermore, seismic and gravit y data indicat e thick Upper Palaeozoi c strat a o f probabl e Devonia n age to exis t beneat h bot h th e Eas t Shetlan d an d the Hord a platform s (Platt 1995 ; Christiansson et al . 2000 ; Faleid e e t al . 2000). The reconstructio n (Fig . 16 ) shows a remaining fan-shape d regio n i n th e Rockal l Trough , which w e hav e lef t shade d partl y t o illustrat e the siz e o f th e are a w e believe may hav e a sepa rate histor y i n term s o f riftin g and/o r driftin g (i.e. Smyth e 1989 ; Srivastava & Verhoe f 1992 ; Knott e t al . 1993) . Th e rapi d deca y i n widt h of the troug h fro m sout h t o nort h i s interpreted t o suggest tha t 'extra ' crustal thinning affected thi s rather shor t rif t segment , i.e . c . 800-1000 km. It should be noticed, however, that furthe r nort h neither the Lat e Jurassic-Cretaceous 'major' rif t episode no r th e Maastrichtian-Paleocen e riftin g to continenta l break-u p wer e abl e t o creat e a similarly wid e area o f thi n crust .

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As a n alternativ e t o previousl y presente d interpretations fo r th e formatio n o f th e troug h (i.e. Permia n t o Cretaceous) , o r rathe r a s a pos sible addition t o these , ther e are indications tha t the Maastrichtian-Paleocen e extensio n episod e and subsequen t sea-floo r spreadin g ma y hav e influenced th e area . Thi s assumptio n come s from th e plat e tectoni c reconstructio n betwee n Greenland an d Eurasi a (Fig . 10) . T o generat e good closur e betwee n conjugat e pair s o f magnetic Chron s 2 1 t o 24n. 3 sout h o f Iceland , the Hatton-Rockal l 'plate ' ha s t o b e rotate d anti-clockwise relativ e to Eurasia , yieldin g postbreak-up (i.e . pos t 53-4 6 Ma) movement s pos sibly centre d alon g th e trough . Wit h respec t t o the significan t Earl y Tertiar y igneou s activit y in bot h th e Rockal l Troug h an d th e Porcu pine Basi n (Tat e e t al. 1993 ) thi s coul d b e a n attractive explanation .

Volcanic margi n formation and exploration challenge s Plume-lithosphere interaction and Paleocene vertical movements The Maastrichtian-Paleocen e rif t episod e i n th e NE Atlanti c regio n wa s strongl y influence d b y the arriva l o f th e Icelan d mantl e plum e a t lithospheric levels . Regionally , Earl y Tertiar y igneous activit y initiate d in West Greenlan d an d in th e Britis h Tertiary Igneou s Provinc e (BTIP ) about 62-6 3 Ma (Musset t e t al . 1988 ; Larse n et al . 1992 ; Hitche n & Ritchi e 1993) . Th e age s from th e BTI P have , however , recentl y bee n questioned b y Jolle y (1997) , wh o propose d a n age o f c . 58 Ma. o n th e basi s o f ne w radiometri c dating. Th e massiv e volcanis m alon g th e Maas trichtian-Paleocene N E Atlanti c rift zon e starte d some millio n year s later , a t c . 55-54 Ma, usin g the Cand e & Ken t (1992 ) tim e scal e (Robert s et al . 1984 ; Larsen & Wat t 1985 ; Eldholm e t al . 1989; Skogsei d e t al . 19920) . Th e lat e introduc tion o f th e igneou s activit y wit h respec t t o th e initiation o f rifting ha s bee n relate d t o th e timing of the emplacemen t o f the Icelan d mantl e plum e beneath th e prot o N E Atlanti c lithosphere . On th e basi s o f our presen t understanding , earl y rifting (75-6 3 Ma) occurre d withou t th e plume related therma l anomal y a t lithospheri c levels , and thu s withou t muc h mel t formation . Th e key to thi s interpretatio n i s foun d i n th e Vorin g Basin, wher e a thic k Paleocen e sedimen t pack age awa y fro m th e centra l rif t zon e allow s u s t o resolve th e tim e relation s betwee n initiatio n o f rifting, uplif t an d magmatis m a s describe d b y Skogseid e t al . (19920) .

Thermal plumes , whic h transpor t energ y towards th e Earth' s surfac e becaus e o f buoy ancy, contai n a larg e hea d o f ho t materia l tha t ascends throug h th e mantle. Upo n impingin g on the lithospher e th e plum e hea d spread s laterally. The arriva l o f plume s therefor e result s i n a n episode o f surfac e uplif t ove r larg e areas . According t o Griffith s & Campbel l 0990 ) th e average temperatur e anomal y carrie d b y th e plume hea d wit h respec t t o th e adjacen t mantl e is c . 100 CC. Thi s magnitud e o f th e therma l anomaly i s i n genera l agreemen t wit h th e asthenospheric temperature s derive d fro m depth an d gravit y anomalie s (Cochra n & Tai wan! 1978 ) and analysi s o f crusta l thicknes s i n the realm of the Iceland plume (i.e. 10.3 ± 1. 7 km. from Whit e e t al . 1992) . It i s presumed tha t th e hot plum e material preferentially fills traps at the base o f the lithosphere created eithe r by previous plate deformations o r b y continuing lithospheric thinning (e.g. Artyshkov et al. 1980: ; Sleep 1997) . It is , however , no t obviou s fro m th e variou s plume model s a t wha t velocit y th e plum e material flow s laterall y beneat h th e lithosphere . In terms of dimensions. Hill (1991) has estimated that a n origina l plume head , wit h a diamete r o f 1000km, wil l for m a n ellipsoida l dis c 2000k m wide, stil l wit h a maximu m thicknes s o f 160k m centrally. The initiatio n o f magmatis m a t 62-6 3 Ma i n the N E Atlanti c region , suggeste d t o b e corre lated wit h th e impingemen t o f th e plum e head o n th e lithosphere . wa s centre d beneat h Greenland a t c . 63 Ma accordin g t o Lawve r & Miiller (1994) . A plum e 3000k m i n diamete r (after spreading ) i s require d t o encircl e th e North Atlanti c Volcani c Province , includin g the Britis h Tertiar y Volcani c Provinc e an d th e Labrador Se a volcani c province s (Fig . 10) . The shap e an d thicknes s o f th e plum e bod y underneath th e lithospher e i s stil l speculative , but shoul d i n principl e b e possibl e t o estimat e from observation s of relative vertical movements in th e area . Here, w e wan t t o outlin e th e first-orde r patterns o f uplif t expecte d i n relatio n t o th e plume emplacement . T o achiev e this , thicknes s models o f the lithospher e and th e plume bod y a t the tim e o f plum e emplacemen t (c . 63 Ma) ar e constructed. Th e lithospher e thicknes s mode l i s based o n ou r break-u p reconstructio n an d knowledge abou t th e regiona l geolog y i n th e NE Atlanti c region , bu t canno t b e considere d well constraine d (Fig . 17) . The mode l generall y reflects th e mai n rif t configuration s withi n th e area, wher e i n particula r th e Lat e Jurassic Cretaceous rif t zone s an d th e Maastrichtian Paleocene rif t alon g th e subsequen t break-u p

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Fig. 17 . N E Atlanti c lithospher e thicknes s model suggeste d for c . 63 Ma, bu t implemente d o n th e ma p showin g the plat e reconstructio n t o tim e o f break-up (Fig . 10). The circl e wit h a 150 0 km radiu s illustrate s th e siz e of a flattened plume hea d larg e enoug h t o encircl e the entir e Earl y Tertiary N E Atlanti c volcanic province, an d centred approximatel y a t th e projected Icelan d plum e location accordin g to Lawve r & Muller (1994). The yellow straight line show the location o f the plum e profile i n inset frame i n Fig. 18. (For other abbreviations, see Fig. 1.) axis hav e thinnes t lithospher e (i.e . 55-10 0 km). In contrast , th e mor e stabl e cratoni c area s ma y have thicknesse s a s hig h a s 160k m (Suhadol c et al 1990) . The shap e o f th e plum e bod y beneat h th e lithosphere i s based o n th e assumptio n tha t th e volume o f a spherica l plum e head , 1200k m i n diameter, spread s ou t radiall y an d tha t th e thickness o f plum e material s i s a functio n o f both th e distance fro m th e plume centre and th e structure o f th e overlyin g lithosphere (Fig . 18) . It is, thus, eviden t tha t the thin lithospher e area s govern a significantl y thicke r plum e bod y tha n the cratoni c regions . The emplacement of a thermal plume beneath the lithospher e results in an isostaticall y defined uplift, wher e th e magnitud e o f movement s i s a functio n o f th e thicknes s o f plum e bod y and densit y differenc e betwee n th e plum e mate rial an d th e origina l asthenosphere. W e have assumed a n averag e therma l anomal y o f A7 " = 100°C in the plume material , and tha t a therma l

expansion coefficien t o f a = 0.00003 K ! i s applicable fo r plume-relate d materials . A T - c t thus characteriz e th e densit y differenc e betwee n the plume body an d th e adjacent asthenosphere . The plume-lithosphere interaction demonstrate s the probabilit y o f th e establishmen t o f a n elongated uplif t are a alon g th e activ e rif t zones, a s wel l a s inversio n o f region s wit h thi n lithosphere. Th e sam e region s are thos e tha t b y nature are prone t o decompressio n mel t generation, an d fro m Fig . 1 8 we not e tha t ther e i s a close spatia l relationshi p betwee n th e uplif t region an d th e distributio n o f subaeriall y emplaced floo d basalts . Formation , migratio n an d eventually pondin g o f igneou s melts at th e bas e of th e crus t i s a secondar y proces s o f central rif t inversion (Pederse n e t al . 1996) , an d th e ascending melts , causin g significan t additiona l weakening o f th e lithospher e preferentiall y along th e rif t zone , ma y hav e drive n th e fina l continental separation (c. 55 Ma) (Skogsei d et al. 19920).

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Fig. 18 . Theoretica l isostati c lithospher e uplif t relate d t o a n averag e temperatur e anomal y o f 10 0 C i n a plum e head where the thickness distribution is a function o f both the distance from th e plume centre and th e relief at th e base o f th e overlyin g lithospher e (see plume profil e i n inse t fram e locate d b y yello w straigh t lin e o n map). (For othe r abbreviations , se e Fig. 1.) The central rift inversio n described abov e ma y cause th e development o f a relativel y short-live d land are a alon g th e rif t axis , which becaus e i t is prone t o erosio n ma y represen t a provenanc e region fo r sediment s bein g she d int o adjacen t shallow-water basins . Erosio n i s therefor e a third mechanis m o f amplifie d uplift . Thic k Paleocene sedimentar y sequence s ar e mappe d along th e oute r margin s of f mid-Norwa y an d west o f th e Shetlan d Island s (e.g . Hitche n & Ritchie 1987 ; Skogseid & Eldhol m 1989) . According t o Griffith s e t al. (1989) , th e emplacement o f mantl e plume s beneat h conti nents appear s t o sho w les s uplif t tha n i n area s where plumes ar e locate d beneat h oceani c litho sphere. Th e experiment s o f Griffith s e t al . demonstrate tha t thicke r surfac e layer s resul t in reduce d swel l heights. Th e effec t i s describe d as a buoyanc y response , an d applie s t o a dyna mic syste m only . Th e spreadin g o f th e plum e beneath th e lithospher e shoul d als o b e consid ered dynamically , whic h durin g th e perio d o f plume hea d collaps e yield s a situatio n wher e

the thickness , an d thu s weight , o f th e litho sphere play s a n importan t rol e wit h respec t to th e magnitud e o f vertica l movements . Thus , in region s wher e th e lithospher e i s thic k wit h respect t o th e thicknes s o f th e plume , i.e . Greenland an d region s aroun d th e margi n o f the plume , suc h a s th e Canadia n margi n o f th e Labrador Se a (Balkwil l e t al . 1990) . th e Nort h Sea (e.g. White & Lati n 1993 ; Jordt e t al . 1999). mainland Norwa y an d Fennoscandi a (e.g. Tynni 1982; Rii s 1996) , th e Wester n Barent s 'Se a (Faleide e t al . 1993<:/,/?) , Svalbar d (e.g . Nottvedt et al. 1992) , etc., significan t and rapi d subsidenc e (<500m) ma y b e expecte d durin g th e plum e emplacement, i.e . durin g lat e Paleocene-earl y Eocene times . Thi s subsidenc e shoul d theoreti cally b e followed by a rapid reboun d resultin g in net uplif t a s th e plum e movemen t ceases . The c . 20 Ma rif t developmen t associate d w 7ith break-up i n the NE Atlantic , apparently starting before a plum e establishmen t a t lithospheri c levels, contrast s t o som e degre e wit h th e activ e rifting mode l of Campbell & Griffiths (1990 ) and

NE ATLANTI C CONTINENTA L RIFTIN G Hill (1991) . Bas e lithospher e erosio n durin g plume emplacemen t may , o n th e othe r hand , be on e mechanis m tha t ca n explai n observe d discrepancies i n magnitud e o f upper-crusta l v . lithospheric thinning . Melt generatio n an d sub sequent migratio n throug h th e lithospher e wil l further caus e dramati c mantle-lithospher e weakening an d probabl y facilitat e break-up . The apparentl y lo w temperatur e anomal y (50-100°C) require d t o explai n volume s o f 'observed' igneou s rock s acros s th e V0rin g margin (Pederse n & Skogsei d 1989) , th e geo chemistry o f th e extrusiv e complexe s a t som e distance fro m th e centra l plum e trai l (e.g . Zehnder e t al. 1990) , an d th e thicknes s o f th e adjacent oceani c crus t (Whit e et al. 1992 ) within the N E Atlanti c larg e igneou s provinc e (LIP) , should als o ope n a broade r discussio n o n tectono-magmatic model s v . plum e an d non plume hypotheses . Holbroo k e t al . (19940 , b) suggested a non-plum e mechanis m fo r th e igneous activit y o n th e U S Eas t Coas t margi n primarily becaus e n o obviou s plum e trai l i s present. Wilso n (1997) , o n th e othe r hand , indicated tha t a c . 200 Ma large-scal e plum e head cause d a widesprea d tholeiiti c magmati c event o n th e Wes t Africa n margin , whic h bot h temporally an d spatiall y correlate s wit h th e Central Atlanti c riftin g an d opening . Through out th e pas t 20 0 Ma th e centre s o f th e Cap e Verde, Canar y an d Madeir a plume s hav e bee n located o n th e Africa n plat e tha t hav e move d relatively littl e i n term s o f absolut e movement s (Watts & Marr 1995 ; Demets e t al. 1990) . From the assumptio n tha t mantl e plume s ar e fixed points wit h respec t t o th e Eart h an d tha t th e lithospheric plate s ar e moving , a plum e trai l should no t exis t o n th e Nort h America n plate . In thi s discussio n i t ma y als o b e importan t t o note that , accordin g t o Brook s e t al . (1996) , volcanic rock s a t dista l region s o f th e Icelan d plume sho w a smal l degre e o f partia l meltin g and provid e sample s fro m th e mantl e unconta minated b y reactio n wit h continenta l crust , partly i n contras t t o som e o f th e tholeiiti c volcanism. Thes e rock s sho w a mantl e ric h i n trace element s and volatiles ; this finding may b e of significanc e als o t o th e muc h mor e volumi nous volcanis m o f th e region , wher e suc h characteristics ar e maske d b y larg e degree s o f partial melting. Exploration challenges Volcanic margin s ar e considere d t o b e mor e common the n previousl y though t (Coffi n & Eldholm 1994) . T o broade n ou r knowledg e o f

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passive margins , i n general , an d volcani c pas sive margins , i n particular , comparativ e studie s between th e N E Atlantic , th e Namibian , th e US Atlantic and the NW Australian margins are considered importan t a s references or analogue s (Eldholm et al. 2000). The main results from thi s comparative wor k ar e tha t al l margins see m t o have ha d a protracte d rif t developmen t befor e continental separation , tha t significan t crusta l thinning i s observed adjacen t t o th e continent ocean boundary , an d tha t th e mai n puls e o f igneous activit y i s centre d aroun d th e tim e of break-up . One importan t observatio n i s the recognitio n of that magmatic underplating (characterized b y 7 + kms~ 1 velocity ; Furlon g & Fountai n 1986 ) is found up t o 100-15 0 landward fro m th e CO B (Fig. 7) . These rock s represen t actua l growt h in crustal thickness , whic h lead s t o rif t inversio n during magm a emplacemen t an d significantl y reduced post-rif t subsidenc e compare d wit h a-magmatic rifts . Thus , th e widt h o f volcani c margins, i n term s o f th e are a affecte d bot h b y lithospheric thinnin g an d b y much magmatism , may b e increase d significantl y compare d wit h the 'narro w rift ' hypothesi s initiall y applie d o n volcanic margin s b y Mutte r e t al . (1988) . Thi s hypothesis i s ofte n repeate d whe n subsidenc e analysis acros s th e volcani c construction s i s made withou t applicatio n o f a mor e regiona l perspective (e.g . Clif t e t al . 1995) , o r wher e the lac k o f a plum e trai l i n th e vicinit y o f the volcani c margi n make s i t difficul t t o poin t to a plum e sourc e fo r th e exces s magmat ism (Holbroo k e t al . \994a, b; Kee n & Bouti lier 1995) . Thus, i n additio n t o th e larg e variabilit y an d complexity associated wit h continental riftin g i n general, a numbe r o f tectono-magmati c event s characterize volcani c margin s an d mak e the m particularly interestin g bot h fo r scientifi c research an d a s exploration targets . The expression o f thes e tectono-magmati c event s is mostly associated wit h vertica l motion s durin g exten sion; first , b y alteratio n o f th e temperatur e structure (plume-lithospher e interaction ) dur ing plum e emplacement ; second , b y centra l rif t inversion, erosio n an d associate d sedimen t transport an d redepositio n primaril y linke d with th e magm a generation , migratio n an d emplacement (Whit e & McKenzi e 1989) ; third, by relativel y little post-rift subsidenc e compare d with non-volcani c margin s (e.g . Bot t 1992) . Rift-related footwal l uplif t an d possibl e verti cal motio n relate d t o differentia l lithospheri c stretching ar e additiona l factor s tha t ma y pla y important role s locally . W e stres s tha t fro m a palaeo-environmental, a n ocea n circulatio n and

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a basi n histor y poin t o f vie w i t i s importan t t o recognize tha t durin g riftin g an d break-u p th e present oute r margin s forme d part s o f a wid e land-bridge stretchin g alon g th e rif t axis . I n th e NE Atlanti c regio n thi s lan d are a probabl y reached fro m th e Charli e Gibb s Fractur e Zone , at c . 55°N, to th e S W Barents margin , a t c . 75°N (Fig. 18) . With respec t t o exploratio n o n volcani c margins w e hav e t o understan d th e complexit y in th e geologica l developmen t an d t o b e abl e t o evaluate th e heat-flo w histor y an d temperatur e effects associate d wit h th e igneou s rock s a t crustal levels . Here , th e igneou s characteristic s of volcanic margins represen t a specia l challeng e by means o f seismic definition. On seismi c reflec tion dat a th e floo d basal t provinc e i s normall y characterized b y a serie s o f seaward-dippin g reflectors formin g hug e wedges o f volcanic rock s and interbedde d sediment s (<10k m thick) (Hinz 1981; Mutte r e t al 1984 ; Skogsei d & Eldhol m 1987; Whit e e t a l 1987 ; Hoppe r e t a l 1992 ; Holbrook e t a l 19940 , b\ Plank e 1994 ; Planke & Eldholm 1994) . Fro m a tectono-magmati c poin t of vie w w e assum e them , i n general , t o b e associated wit h th e actua l continenta l break-u p and firs t phas e o f sea-floo r spreadin g whe n ful l asthenosphere decompressio n i s reached . A s a result o f the syn-rif t uplif t referre d t o above , th e rift zon e ma y b e a t o r abov e sea leve l at tim e o f break-up, resultin g i n a significan t increas e i n basalt flowability . Th e basalt s may , thus , over flow larg e portion s o f th e centra l rif t zon e an d associated extensiona l structures . Th e interna l seismic characteristics o f the floo d basal t wedge s are generall y resolvable, wherea s th e interpreta tion of the sub-basaltic parts of the seismic section is more challengin g (Planke e t al. in press). Landward o f th e mai n wedge s o f volcani c rocks a wid e (50-10 0 km) zon e o f th e margi n may b e affecte d b y sill s an d dyke s i n th e pre break-up sedimentar y unit s (Skogsei d e t al . \992a). Th e sill s ar e ofte n eas y t o interpre t b y their seismi c signatur e and abrup t latera l termi nation. The y hamper , however , th e interpreta tion o f th e sedimentar y sequence s the y ar e intruded in , an d mak e structura l interpretation more difficul t a s the y partl y ar e crossin g fault s without reflectin g the faul t offsets . The thir d typ e o f igneou s feature s associate d with volcani c margin s ar e th e bodie s o f mag matic underplating , whic h fro m a n exploratio n point o f vie w ma y b e considere d th e mos t important facto r i n heat-flo w histor y an d hydrocarbon maturatio n (Pederse n e t al . 1996) . Factors influencin g the volume s o f mel t t o b e formed ar e primaril y th e magnitud e o f th e temperature anomal y (McKenzi e & Bickl e

1988) an d th e magnitud e an d duratio n o f lithospheric stretchin g (Pederse n & R o 1992) . The molte n rock s ar e emplace d beneat h th e crust, withi n th e crus t a s dyke s o r sills , o r a t the surfac e a s floo d basalts . A typica l thicknes s of 3- 5 km underplatin g wil l resul t i n uplif t o r reduced subsidenc e i n the range betwee n 0. 8 and 1.5km. Thi s uplift , whic h i s ascribe d t o a permanent re-creatio n o f crusta l thickness , i s subsequently identica l wit h th e relativ e lac k of therma l subsidenc e associate d wit h volcani c margins wit h respec t t o non-volcani c margins . The uplif t associate d wit h rifting , plum e emplacement, mel t formatio n an d mel t emplace ment wil l take place over th e same perio d of time during margi n formation , wil l b e focuse d alon g the centra l rif t zones , an d ma y su m t o magni tudes o f severa l kilometres . Exploration o n volcani c margin s need s t o consider th e proble m relate d t o th e therma l effects fro m th e igneou s unit s o n maturatio n o f organic matter . Th e tectono-magmati c history , including th e variou s surfac e uplift-erosio n mechanisms an d th e tempora l developmen t o f the igneou s complexes , need s t o b e treate d a s input t o th e calculatio n o f hea t flo w (Pederse n et al. 1996) . Traditionally the heat-flo w history is related t o extensiona l thinnin g o f th e litho sphere. A t a passive margin , wit h high extension factors, thi s backgroun d hea t flo w wil l b e substantially increase d ove r relativel y lon g periods. However , th e hea t flo w fro m a mag matic bod y a t th e bas e o f th e crus t ca n yiel d a three-fold increas e i n hea t flo w ove r a relativel y short (<15Ma ) tim e interval , thu s spikin g th e temperatures i n th e sediment s b y a n amoun t mainly dependin g o n th e thicknes s o f th e underplate. th e duratio n o f mel t emplacement , the distanc e fro m th e underplat e t o th e sedi ments, an d th e sedimen t conductivit y (Pedersen et al . 1996) . O n th e basi s o f models , wit h th e mid-Norwegian margi n a s th e stud y area , th e present temperature s ar e generall y lower withi n the pre-rif t sectio n tha n the y wer e shortl y afte r the igneou s emplacement (Pederse n e t al . 1996) . Additionally, sill s intrude d i n th e sedimentary units produc e positiv e temperatur e anomalies . Thin sill s ( < 1 0 m ) wil l quickl y die ou t thermally and increas e th e leve l o f maturatio n o f organi c matter withi n distances n o mor e tha n thei r own thickness (Jaege r 1964) . Thick sill s (>100m) . o n the othe r hand , ma y yiel d substantia l increase in maturation a t distance s u p t o 3- 4 time s thei r thickness. Floo d basalts , o n th e othe r hand , are probabl y deposite d a s severa l batche s o f magma (<50 m thick) . Thei r mos t importan t effect i s that the y cause buria l o f th e underlying sediments.

NE ATLANTI C CONTINENTA L RIFTIN G

Conclusions The integrate d approach , usin g dee p seismi c data an d plat e restoration , ha s allowe d a defini tion o f th e rif t zones , th e degre e o f crusta l thinning associate d wit h rifting , an d a n evalua tion o f the spatial relations between basins at th e time o f formation . W e believ e tha t th e lower crustal geometrie s togethe r wit h th e upper crustal structura l settin g demonstrates th e valu e of ou r approach . W e reiterat e i n particula r th e overall spatia l correlatio n betwee n Moh o relie f and upper-crusta l structur e beneat h th e mai n (deep) rif t basins , a s wel l a s th e extrem e crusta l thinning unde r th e deep central rif t troughs . In the conjugate margi n perspective, c. 140km of stretchin g occurre d betwee n Greenlan d an d Eurasia durin g th e 75-5 5 Ma Maastrichtian Paleocene riftin g an d continenta l break-up , whereas th e c . 170-130 Ma Lat e Jurassic-Cre taceous rif t episod e cause d c . 60 km o f stretch ing. I n addition , th e Rockal l Troug h seem s t o have bee n affecte d b y 100-20 0 km additiona l stretching wit h a southwar d fan-shape d signa ture. Afte r plat e restoratio n accountin g fo r th e above stretching , th e pre-Lat e Jurassi c basin s province i s les s tha n 400k m wide . Th e genera l lack o f detaile d stratigraphi c informatio n an d knowledge abou t th e post-Caledonia n crusta l and lithospheri c structur e does no t allo w furthe r quantification o f th e degre e o f stretchin g asso ciated wit h thes e basins . The crusta l structur e beneat h th e Lat e Jurassic-Cretaceous rift zones , i.e . the Mor e and southern V0rin g basins , show s degrees o f crustal thinning whic h b y fa r overste p th e stretchin g estimates derive d fro m subsidenc e analysis . This discrepancy ma y occu r becaus e w e hav e under estimated th e duratio n o f th e Lat e Jurassic Cretaceous extensio n episode , an d thu s no t included th e ful l subsidenc e histor y i n th e region, and/o r becaus e th e relativel y narro w rift basin s resis t subsidenc e owin g t o th e elasti c strength o f th e lithosphere . A simila r discre pancy doe s not exis t beneath th e Maastrichtian Paleocene rif t zone s t o th e west . Here , sub sidence an d crusta l thinnin g estimates ar e fairl y consistent, wherea s the y ar e significantl y large r than th e result s fro m structura l analysis . Thes e difference ma y b e relate d t o resolutio n limit s in the seismi c data , t o th e fac t tha t mos t o f th e central rif t zon e i s covered b y volcanic rocks an d therefore wa s no t include d i n th e structura l analysis, an d finally , becaus e o f differentia l lithosphere thinnin g by base lithospher e erosio n during plum e emplacement . Theoretical estimate s o f uplift associate d wit h the Icelan d mantl e plum e sho w a clea r correla -

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tion wit h th e region s tha t subsequentl y wer e affected b y subaeria l volcani c activity . Th e for mation o f thi s short-live d centra l rif t lan d are a may have caused conditions for drainage system s and subsequen t depositio n o f san d sequence s o f Late Paleocen e ag e (63-5 5 Ma), an d is , thus , considered particularl y important fo r hydrocar bon exploratio n alon g th e outer margins . The geometri c definitio n an d a n understand ing o f th e emplacemen t histor y o f th e unit s o f igneous materials at crustal level s have relevance for calculatio n o f heat-flo w history . I n particu lar, th e bodie s o f underplate d material s ca n spike hea t flo w b y a facto r o f thre e ove r tim e spans o f 5-1 5 Ma, an d accordin g t o model s from th e mid-Norwegia n margin , ma y hav e resulted i n palaeo-temperature s i n th e pre-rif t sequences tha t wer e highe r tha n presen t tem peratures. Th e magnitud e an d duratio n o f th e temperature increas e i s able t o caus e maturatio n of organi c matter , whic h canno t b e predicte d from presen t hea t flo w measurement s o r fro m traditional basi n modelling . We acknowledg e dat a mad e availabl e b y th e Norwe gian Petroleu m Directorate , Bundesanstal t fu r Geo wissenschaften un d Rohstoffe , an d Statoi l AS . We than k th e PLATE S Project , Institut e for Geophy sics, Universit y o f Texas a t Austin , for acces s t o thei r plate rotatio n software . W e als o than k R . England , P. Haremo and L . N. Jense n for comments an d critical review o f the manuscrip t and L . M. Larse n fo r addin g details and fo r fruitfu l discussions . The wor k ha s bee n supported b y th e Norwegia n Researc h Council , th e IBS (Integrated Basin Studies) project, par t o f Joule I I research programm e funde d b y th e Commissio n o f European Communitie s (Contract JOU2-CT-92-0110 ) and b y 'University of Oslo - Passiv e Margin Researc h Group" funde d b y Statoil.

References ANDERSEN, T . B . & JAMTVEIT , B . 1990 . Uplift o f dee p crust durin g orogeni c extensiona l collapse ; a model base d o n fiel d studie s i n th e Sogn Sunnfjord regio n o f wester n Norway . Tectonics, 9, 1097-1111 . ARTYUSHKOV, E . V. , SCHLESINGER , A . E . & YANSIN , A. L . 1980 . Th e origi n o f vertica l movement s within lithospheri c plates . In: BALLY , A . W. , BENDER, P . L. , MCGETCHIN , T . K . & WALCOTT , R. I . (eds ) Dynamics o f Plate Interiors. America n Geophysical Union , Geodynamic s Series , 1 , 37-51. BALKWILL, H . R. , MCMILLAN , H . J. , MACLEAN , B. , WILLIAMS, G . L . & SRIVASTAVA , S . P . 1990 . Geology o f th e Labrado r Shelf , Baffi n Ba y an d Davis Strait . In: KEEN , M . J . & WILLIAMS , G . L . (eds) Geology o f th e Continental Margin o f Eastern Canada. Geologica l Societ y o f America , Th e Geology o f Nort h America , 1-1 , 293-348 .

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The tectonic evolution of th e Norwegian Sea Continenta l Margin with emphasi s on the Vering an d Mere Basins HARALD BREKK E Norwegian Petroleum Directorate, P.O. Box 600, 4001 Stavanger, Norway (e-mail: [email protected]) Abstract: Th e Norwegian Se a continental margi n i s dominated b y two major basin s wit h a very thick Cretaceous basi n fill: the V0ring and M0r e Basins . The basin s are flanked b y the uplifted mainlan d an d th e Cretaceous Tr0ndelag Platfor m to th e east and b y the M0re an d V0ring Marginal High s capped by Eocene lava s to the west. The tectonic development of the area i s controlle d b y tw o structura l trends : NE-S W an d NW-SE . Th e are a ha s bee n tectonically activ e from Carboniferous t o Lat e Pliocen e tim e with the main tectoni c phase s in Late Palaeozoic, lat e Mid-Jurassic-Early Cretaceou s and Lat e Cretaceous-Early Tertiary time. Th e genera l tectoni c development comprised a lon g period o f extensio n and riftin g that ended in Early Eocene tim e by continental separation , major volcanism and subsequen t sea-floor spreadin g in the Norwegian-Greenland Sea . In Carboniferous to Early Cretaceous time the extensional tectonics were related to within-plat e continental rifting. Th e tectonics of the Lat e Cretaceou s and th e Tertiary period s wer e controlled b y the relative movement s along plat e boundaries . Th e overal l NE-S W structural grain i s constituted b y fault s an d basin axe s tha t probabl y originate d i n Lat e Palaeozoi c tim e an d wer e activ e durin g al l subsequent tectoni c phases . Th e transverse NW-S E trend i s expressed a s major lineament s that probabl y reflect th e old, Precambrian grai n of the basement . These lineaments , two of which ar e th e continuatio n int o th e continenta l crus t o f majo r oceani c fractur e zones , controlled th e tectonic activity throughout Cretaceous and Tertiar y time and constitut e the boundaries betwee n th e majo r structura l province s o f th e area . Th e differentiatio n int o the Cretaceou s basin s an d th e boundin g platform s an d margina l high s starte d b y the lat e Mid-Jurassic-Early Cretaceou s extensiona l phase. Th e subsequen t Cretaceous subsidenc e history, wher e th e basi n flank s forme d b y flexurin g rathe r tha n faulting , resulte d i n a n exceptionally thic k basi n fill . I n th e Vorin g Basi n th e Cretaceou s developmen t comprise d an earl y therma l subsidenc e phas e an d a post-Cenomania n phas e o f tectonicall y driven subsidence involvin g intermitten t phases o f norma l faulting an d compressio n an d folding . The V0ring Basi n was tectonically activ e also durin g Tertiary tim e with the main phases o f strike-slip-compression coincidin g wit h th e Alpin e orogenie s i n Lat e Eocen e an d MidMiocene time. Within the V0ring Basin there is evidence of the formation of a fossil opa l A— opal-CT transitio n an d extensiv e regiona l marin e erosio n i n Mid-Miocen e an d Lat e Pliocene times. In contrast, th e M0r e Basi n was generally tectonically quiet throughout the Cretaceous an d Tertiar y periods , experiencin g mainly continuous subsidence .

The continenta l margi n o f th e Norwegia n Se a 1972 , 1977 ; Bot t 1973 ; Huseb y e t al. 1975 ; has bee n th e subjec t o f acceleratin g geologica l Talwan i e t al. 1976 ; Eldholm & Thiede 1980) . and geophysical studie s ever since the Norwegian Th e first exploration licence s in th e are a wer e authorities acquire d th e firs t seismi c dat a fo r allocate d i n 198 0 and th e nex t 1 5 year s sa w a n exploration purpose s i n 1969 . Durin g th e nex t increase d activit y o f exploratio n an d scientifi c decade th e Norwegia n Petroleu m Directorat e studie s fro m bot h th e oil industry an d academi c (NPD) steadil y increased it s database o f seismic, institutions . Th e industr y focuse d o n th e Lat e gravimetric an d magneti c data . I n th e sam e Palaeozoi c an d Mesozoi c geologica l develop period th e Dee p Se a Drillin g Projec t (DSDP ) men t an d base d man y o f it s regiona l tectoni c and academi c institution s started t o unrave l th e model s o n th e shallo w shel f area s of f Norwa y details o f th e plat e tectonic s an d th e sea-floo r an d th e mainlan d geolog y o f Eas t Greenlan d spreading histor y o f th e Norwegian-Greenlan d (e.g . R0nnevik & Navrestad 1977 ; Bukovics et al. Sea (e.g . Johnso n & Heeze n 1967 ; Aver y e t al . 1984 ; B0e n e t al . 1984 ; Gabrielse n e t al . 1984 ; 1968; Meye r e t al . 1972 ; Talwan i & Eldhol m Pric e & Ratte y 1984 ; Surlyk et al . 1984 ; Brekke From: N0TTVEDT , A . e t al . (eds ) Dynamics o f th e Norwegian Margin. Geologica l Society , London , Specia l Publications, 167 , 327-378. 1-86239-056-8/OO/ S 15.00 © Th e Geologica l Societ y o f Londo n 2000 .

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Fig. 1 . Simplifie d structura l ma p o f th e Norwegia n Se a continental margin . GIH , Gisk e High ; GNH . Gnause n High; SH . Selj e High .

TECTONIC EVOLUTIO N O F CONTINENTA L MARGI N & Rii s 1987 ; Surly k 1990 ; Dore 1991 ; Torsk e & Prestvik 1991) . Th e tectoni c framework , crusta l structure an d th e tectono-magmati c develop ment o f th e oute r part s o f th e continenta l margin wer e mainly treated b y academic institu tions (Hin z e t al 1982 , 1984 ; Mutte r e t al. 1984 ; Skogseid & Eldhol m 1987 , 1989 , Eldhol m e t al . 1989; Plank e e t al . 1991 ; Skogsei d e t al . 1992 ; Skogseid 1994) . From 199 4 the industr y also changed it s focus to th e deep-wate r oute r par t o f th e continenta l margin a s a moder n seismi c gri d wa s mad e available and th e first concession roun d i n thos e areas was imminent. The intention of the present author i s t o giv e a descriptio n base d o n a n updated interpretatio n o f th e presen t regiona l seismic grid wit h a n emphasi s o n the oute r part s of the continental margin , but with due reference to th e presen t stat e o f knowledge of the geolog y from th e shallow part o f the shelf. Less emphasis is place d o n tryin g t o mak e a n overal l regiona l tectonic mode l t o explai n al l observations .

Tectonic settin g The overal l tectoni c framewor k o f the continen tal margi n betwee n 62 ° and 69° N consist s o f a central are a o f NE-S W trendin g dee p Cretac eous basins , th e V0rin g an d Mor e Basins , flanked b y palaeo-high s an d -platform s an d th e elevated mainland (Fig s 1 & 2). The platforms t o the wes t ar e terme d th e Mor e an d Vorin g Marginal High s an d ar e characterize d b y thick , Early Eocene basal t flows overlying an unknow n substrate. Th e boundar y betwee n th e margina l highs an d th e basi n are a i s forme d b y th e Faeroe-Shetland Escarpmen t t o th e sout h an d the V0rin g Escarpmen t t o th e north . The centra l par t o f th e easter n flank s i s dominated b y th e Lat e Jurassic-Earl y Cretac eous Trondela g Platform . Nort h an d sout h o f the Tr0ndelag Platfor m th e basi n are a is flanked to th e eas t b y the erode d mainland . To th e north, th e main basi n are a i s bounde d by th e NW-S E trendin g Bivros t Lineament , which separate s th e wide and dee p V0rin g Basin from th e narro w an d tectonicall y uplifte d continental margi n aroun d th e Lofote n Ridge . To th e south , th e southeaster n boundar y o f the Mor e Basi n i s locate d wher e th e NE-S W trending M0re-Tr0ndela g Faul t Comple x trun cates th e N- S an d NNE-SS W trend s o f th e northern Nort h Sea. To the southwest, the More Basin border s upo n th e Faeroe-Shetlan d Basi n along th e NW-S E trendin g Erlen d Platfor m (Pharaoh e t al . 1996) .

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Apart fro m th e mai n NE-S W structura l grain, th e continenta l margi n i s subdivide d b y two NW-S E trendin g lineament s int o thre e structural provinces : a northern , a middl e an d a souther n province . Thes e lineaments , th e Bivrost an d Ja n Maye n Lineaments , continu e into th e oceani c crus t a s th e Bivros t an d Ja n Mayen Fractur e Zones , respectively. The tectonic activity that has given the present structural configuratio n o f th e Norwegia n Se a continental margi n ma y b e trace d bac k int o Permo-Carboniferous tim e (Fig . 3 ) (Bukovi s et al. 1984 ; Gabrielsen e t al. 1984; Price & Rattey 1984; Surly k e t al . 1984 ; Brekk e & Rii s 1987 ; Blystad et al. 1995 ; Dore & Lundin 1996) . Three main riftin g episode s ar e identified : i n Car boniferous t o Permian , i n lat e Mid-Jurassi c to Earl y Cretaceou s an d i n Lat e Cretaceou s t o Early Eocen e times . I n Carboniferou s t o Earl y Cretaceous tim e th e extensiona l tectonic s wer e related t o continenta l rifting . Th e tectonic s o f the Late Cretaceous phas e and , i n particular, the Tertiary extensiona l regim e wer e mor e directl y influenced b y the relative movements along plat e boundaries just befor e an d durin g th e continental break-u p an d onse t o f sea-floo r spreading i n the Nort h Atlantic . Locally , thi s is expressed a s prominent compressiona l structures .

Database, stratigraph y an d seismic tie s The presen t mappin g an d descriptio n o f th e Voring an d Mor e Basin s an d adjacen t area s ar e mainly based o n the interpretation o f the seismic data acquire d b y th e NP D fro m 198 5 to 1992 , supplemented b y well dat a an d literatur e on th e geology o f East Greenlan d (Larse n 1984 ; Surlyk et al . 1984 ; Surly k 1990 , 1991 ; Larse n & Mar cussen 1992 ; Larsen e t al . 1999) . Most o f th e well s i n thi s par t o f th e Norwe gian shel f hav e bee n drille d o n th e Trondela g Platform an d th e Halte n Terrace , an d th e post Lower Triassi c stratigraphy o f the platform an d terrace sequence s i s wel l know n (Dallan d e t al . 1988). B y analog y t o Eas t Greenlan d i t i s assumed tha t th e platfor m an d terrace s als o comprise sedimentar y rock s o f Carboniferou s and Permia n ag e on top o f crystalline basement. So far , onl y seve n conventiona l exploratio n wells hav e teste d part s o f th e Cretaceou s basi n fill sequences. Thes e wells , 6505/10-1 , 6607/5-1 , 6607/5-2, 6607/12-1 , 6704/12-1 , 6706/11- 1 an d 6707/10-1, wer e drille d i n th e northeaster n par t of the Voring Basin and terminate d i n the Uppe r Cretaceous sequence , provin g a ver y thic k Upper Cretaceou s sequenc e of marine shales and intebedded sandstone . Part s o f th e sandston e

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Fig. 2 . Simplifie d structura l ma p o f th e Norwegia n Se a continenta l margi n wit h lin e location s fo r seismi c and geoseismic section s i n Fig s 5 , 6 , 8-12 , 14 , 17 , 18 , 22, 23 , 25 , 2 6 an d 30 .

TECTONIC EVOLUTIO N OF CONTINENTA L MARGIN

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Fig. 3 . Tectoni c histor y o f the Norwegia n Se a continental margin . sequences ar e thic k an d o f goo d reser voir qualit y (Kittilse n e t ai 1999 ; Brekke e t al . 1999). Th e pre-Cenomania n sequence s withi n the V0rin g an d M0r e Basin s are stil l unknown . The Paleocen e an d Eocen e unit s o f th e west ern part o f the Voring Basin show an expansio n in thicknes s compare d wit h th e easter n par t (Fig. 5) . DSD P an d Ocea n Drillin g Progra m (ODP) well s in the area hav e only penetrated th e Miocene sequenc e (Talwan i et al. 1976 ; Eldholm et al . 1987) , an d th e lithologie s o f th e Lowe r Tertiary sequenc e ar e lef t t o speculation . From wells and seismi c surveys a total of nine regional unconformitie s ar e identifie d i n th e stratigraphy of the area (Fig . 3). The three oldest unconformities ma y b e identifie d onl y i n th e platform areas : thes e ar e o f late Earl y Permian , Mid-Triassic an d lat e Mid-Jurassi c age . Th e dating o f thes e unconformitie s i s base d o n

exploration wel l dat a an d analogie s wit h Eas t Greenland (Surly k et al . 1984) . The six younger regional unconformities in the stratigraphy ar e identifie d in both th e basi n an d platform-terrace areas : thes e are of base Cretaceous, to p Cenomanian , bas e Tertiary , Uppe r Eocene-Lower Oligocene , Middl e Miocen e an d intra Uppe r Pliocen e position s i n the sequence . The bas e Cretaceou s unconformit y i s wel l defined a s a n onla p surfac e i n th e easter n basi n flanks, platfor m an d terrac e areas . Withi n th e basin area s itsel f th e bas e Cretaceou s positio n lies very deep an d i s difficult t o interpre t an d ti e up becaus e o f latera l variatio n i n th e qualit y of the seismic data. However , in the Ras Basin and partly i n th e Traen a Basi n seismi c dat a clearl y show th e bas e Cretaceou s unconformit y a s a strong reflector between 6 and 8 s twt (see Figs 6 and 7 , and Fig s 24, 25, 28 and 30 , below). Othe r

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Fig. 4. Colou r legen d o f geoseismic section s in Fig s 5, 6, 8-12 . 14 . 17 . 22 an d 23 .

workers i n th e are a (Skogsei d & Eldhol m 1989 ) have place d th e bas e Cretaceou s positio n a t shallower depth s wes t o f th e Fle s Faul t Com plex, correspondin g t o th e to p Cenomania n unconformity i n th e presen t interpretation . In ou r interpretatio n th e bas e Cretaceou s posi tion remain s dee p als o t o th e wes t o f th e Fle s Fault Complex . Thi s i s base d bot h o n a clea r onlap surfac e (take n t o b e th e bas e Cretaceou s level) dippin g eastward s alon g th e flank s o f th e Gjallar Ridg e i n the southwester n V0rin g Basin , and o n a very good, continuous doubl e reflecto r of probabl e Aptia n ag e whic h put s limit s t o the positio n o f th e bas e Cretaceou s level . Thi s Aptian reflecto r i s found throughou t th e V0rin g and Mor e Basins . A dee p positio n o f th e to p Cenomanian unconformit y (an d henc e th e dee p position o f th e bas e Cretaceou s unconformity ) west o f th e Fle s Faul t Comple x i s no w con firmed b y th e recen t wel l 6707/10- 1 o n th e Ny k High (Kittilse n e t cil. 1999) . The to p Cenomania n unconformit y i s identi fied as a tilte d an d faulte d surfac e onlappe d b y

younger strat a alon g bot h flank s o f th e Vorin g Basin (e.g . seismi c line s NPD-VB-2-8 7 an d NPD-VB-4-87). an d i n th e souther n part s o f the Trondela g Platfor m a s note d b y Brekk e & Riis (1987). Seismic dat a sho w tha t th e bas e o f th e Tertiary sequenc e i s a n erosiona l unconformit y in th e Vorin g Basi n (e.g . seismi c lin e NPD-VB 15-90. Fig . 8 ) and alon g th e flank s o f th e Mor e Basin. This i s confirmed b y well data in the basi n (e.g. wel l 6607/5-1 ) an d basi n flan k positio n (Gradstein & Backstro m 1996) . The Uppe r Eocene-Lowe r Oligocen e uncon formity i s identifie d a s a n onla p surfac e alon g the Vorin g Escarpmen t (e.g . seismi c lin e NPD VB-15-90 as indicated i n Fig . 8) , on th e flank s o f the mai n compressional structure s in th e Voring Basin, an d a s a hiatu s o r erosio n surfac e al l over th e easter n par t o f th e Vorin g Basin . Halten Terrace , Trondela g Platfor m an d th e North Se a (Eidvi n & Riis ~ 1991 ; Gradstei n & Backstrom 1996) . The Middl e Miocen e unconformit y i s seen a s an importan t erosiona l brea k al l ove r th e Nor wegian shelf . Muc h effor t ha s bee n pu t int o dating thi s unconformity bot h i n the Norwegia n Sea an d i n th e Nort h Se a (Eidvi n & Rii s 1991 . 1992; Eidvi n e t cd. 1993 ; Rundber g e t cil. 1995 ; Gradstein & Backstro m 1996) . Gradstei n & Backstrom (1996 ) argued tha t thi s unconformit y is o f Lat e Miocen e age . bu t pointe d ou t tha t this i s base d o n a differenc e i n opinio n wit h Eidvin & Riis (1991, 1992 ) on th e time-transgres sive nature o f one o f th e critica l biozones . The intr a Uppe r Pliocen e unconformit y i s the bas e o f th e thick , regiona l sedimen t wedg e that buil t ou t westwar d fro m th e elevate d main land acros s th e entir e shel f durin g lat e Pliocen e time (Eidvi n & Rii s 1991 . 1992 ; Rii s & Fjeld skaar 1992) . The mappin g o f th e Vorin g an d Mor e Basin s is base d o n th e interpretatio n o f te n seismi c reflectors. Al l th e interprete d reflector s ar e tie d to well s i n th e area . Th e mos t importan t well s for th e Vorin g an d M0r e Basin s ar e 6607/5-1. 6607/5-2, 6707/10- 1 (Fig . 1 ) an d 650611- 1 (Fig. 9) . an d well s o n norther n Tampe n Spu r and alon g th e M0re-Tr0ndela g coast. I n several places reference s ar e mad e t o specifi c seismi c lines tha t ar e no t include d a s figures as the y ar e not ye t released .

The boundin g lineament s The Ja n Maye n Lineamen t separate s th e Mor e Basin t o th e sout h fro m th e V0ring Basi n t o th e north. Thi s lineamen t i s define d b y a distinc t

Fig. 5 . Geoseismi c sectio n EE'. (Se e Fig . 2 for lin e location an d Fig . 4 for colou r legend. ) Modifie d fro m Blysta d e t al. (1995).

Fig. 6 . Geoseismi c sectio n FF' . (Se e Fig . 2 for lin e locatio n an d Fig . 4 fo r colou r legend. ) Modifie d fro m Blysta d e t al . (1995).

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Fig. 7. Two-way travel time map of the base Cretaceous unconformity of the Voring and More Basins.

sinistral shif t o f basi n axe s an d basi n flanks . It also coincide s wit h th e souther n boundar y o f the Trondelag Platform. A simila r sinistra l shif t is see n i n th e magneti c spreadin g anomalie s i n the ocea n crus t alon g th e Ja n Maye n Fractur e Zone (Fig . 1) . The Bivros t Lineamen t separate s the wide and deep V0ring Basin from the narrow and tectonicall y uplifte d continenta l margi n around Lofote n t o th e north . Th e lineamen t i s further define d a s a n apparen t dextra l shif t i n basin axe s an d flank s an d coincide s wit h th e northern terminatio n o f the Trondelag Platform. On th e othe r hand , th e magneti c spreadin g anomalies i n th e ocea n crust , a s interprete d by Skogseid & Eldhol m (1987) , alon g th e Bivros t

Fracture Zon e displa y a sinistra l shif t (Fig . 1) . In recen t compilations of magnetic data (Olesen et al, 1997 ) thi s shif t i s no t obviou s an d th e Bivrost Fractur e Zon e rathe r appear s t o b e the tectoni c boundar y betwee n th e thic k crus t of th e Vorin g Margina l Hig h an d th e oceani c crust t o th e northeast . The tw o lineament s probabl y reflec t a n ol d structural grain in the crystalline basement. Thi s is substantiate d b y th e NW-S E strik e o f th e fjords an d majo r fractures wher e the lineaments meet th e mainland . I n th e regiona l gravit y anomaly patter n th e Ja n Maye n Lineamen t continues as the most prominent NW-SE trending gravit y lo w observe d i n th e Scandinavia n

Fig. 8 . Geoseismi c sectio n GG'. (Se e Fig. 2 for lin e location an d Fig . 4 for colou r legend. )

Fig. 9 . Gcoseismi c sectio n JL'L . (Se c Fig . 2 fo r lin e locatio n an d Fig . 4 fo r colou r legend. )

TECTONIC EVOLUTIO N O F CONTINENTA L MARGI N Peninsula (Fairhea d e t aL 1994) . Th e observe d apparent latera l displacement s o f mai n struc tures acros s th e lineament s canno t b e matche d with similar displacements in the structural grain of the mainland Caledonia n geology . The lateral shifts ar e henc e no t du e t o an y large-scal e post Caledonian transcurren t tectoni c movements . However, th e wa y i n whic h thes e inherite d NW-SE lineament s divide the continenta l margin int o structura l province s strongl y indicate s that the y have had a majo r contro l o n th e post Caledonian tectoni c development . Torsk e & Prestvik (1991 ) describe d a mode l wher e thes e lineaments wer e th e locatio n o f th e controlling transfer zone s i n a simpl e shea r extensio n a t crustal scal e betwee n Norwa y an d Greenland . The position , an d possibl y th e formation , o f the Trondela g Platfor m wer e probabl y con trolled b y th e Bivros t an d Ja n Maye n Linea ments. Whethe r simila r platfor m area s existe d across th e no w deepl y erode d mainlan d t o th e north an d sout h o f th e tw o lineament s i s unknown. Th e ol d NW-S E structura l grai n probably als o controlle d th e Eocen e spreadin g geometry o f th e Norwegian-Greenlan d Sea , where th e fractur e zone s perpendicula r t o th e early spreadin g ridge s initiate d alon g these trends. Th e Ja n Maye n Lineament-Fractur e Zone pai r i s the mos t importan t o f thes e struc tures, as it has been the location of the prominent ridge offse t tha t ha s bisecte d th e Norwegian Greenland Se a throughout it s spreadin g histor y (see Fig s 3 3 and 34 , below). Through th e whol e of Lat e Cretaceou s an d Tertiar y tim e th e Ja n Mayen Lineament was a tectonic barrier between the technically very active V0ring Basin and th e quiet an d non-activ e M0r e Basin , confirmin g the significanc e o f thi s lineament.

The middle province between the Jan Mayen and Bivrost Lineaments The majo r structura l element s i n th e middl e province ar e the Trondelag Platform, th e V0ring Basin an d th e V0rin g Margina l High .

The Trondelag Platform and adjacent terraces In th e interna l part s o f the Trondela g Platfor m the mai n tectoni c episod e too k plac e i n th e Carboniferous t o Lat e Permia n tim e (Fig . 5) . The horsts an d half-graben s a t dept h ar e of that age. Th e majo r fault s wer e als o activ e throug h much o f Triassi c time , givin g ris e t o severa l

337

en echelon , NE-S W trendin g basin s fille d wit h Triassic an d Uppe r Palaeozoi c sediments . Th e southernmost an d bes t define d o f these basin s is termed th e Froa n Basin . Th e Vinglei a Faul t Complex form s th e northwester n boundar y o f this basi n an d wa s reactivate d i n bot h Jurassi c and Cretaceou s times . Th e Froy a Hig h mus t have bee n a basemen t hig h a t leas t fro m Lat e Permian times (Fig. 10) . In the high, a condense d succession o f probabl e Triassi c an d Lat e Per mian ag e unconformably overlies deeply erode d basement an d remnant s o f a n olde r basi n sequence. This sequence i n turn lie s unconform ably on the basement and show s internal angular unconformities, testifyin g t o th e comple x Lat e Palaeozoic tectoni c history . The boundar y faul t between th e Froy a Hig h an d th e Froa n Basi n was activ e throug h muc h o f Triassi c time , a s can b e see n b y th e marke d expansio n o f th e Triassic sequenc e i n th e hangin g wal l (Fig . 10) . Towards th e sout h th e Froa n Basi n become s progressively shallower , a s a resul t o f a combi nation of an original thinning of basin sequences and a late r uplif t an d erosio n i n lat e Mid - t o Late Jurassi c time . During th e tectoni c episod e i n lat e Mid Jurassic t o Earl y Cretaceou s tim e th e interio r parts o f the Trondela g Platfor m wer e subject t o minor faultin g partl y b y reactivation s o f olde r faults i n th e Vinglei a Faul t Complex , an d b y faulting alon g th e flank s o f th e Helgelan d Basi n in Earl y Cretaceou s time . Th e uplif t o f th e Nordland Ridg e o n th e platfor m edg e wa s initiated i n thi s period , a s ma y b e see n b y th e onlap o f th e Earl y Cretaceou s basi n fil l o f the Helgelan d Basi n (Fig . 11) . Alon g th e plat form edge s an d i n the futur e terrac e area s o f th e Halten an d D0nn a Terrace s ther e wa s intens e fault activit y i n thi s tectoni c episod e (Fig . 5) . The subdivisio n into th e dee p basin s to th e west and th e platfor m are a t o th e eas t wa s probabl y initiated i n th e earl y stag e o f th e tectoni c episode. Th e terrace s a s individua l element s were initiate d b y faultin g i n Lat e Jurassi c time , which resulte d i n a n expande d Uppe r Jurassi c succession alon g th e Vingleia , Bremstei n an d Revfallet Faul t Complexe s alon g th e easter n margin o f th e terraces . However , i t is important to not e tha t throughou t th e late Mid-Jurassic t o Early Cretaceou s tectoni c episode , th e Halte n Terrace stayed clos e to the same elevatio n a s the Trondelag Platfor m relativ e to th e basin s t o th e west, an d th e mai n subsidenc e o f th e terrace s relative t o th e Trondela g Platfor m di d no t tak e place unti l late r i n Cretaceou s time . Hence , i n Late Jurassi c an d earlies t Cretaceou s time , th e high alon g th e wester n edg e o f th e Halte n Terrace, th e Sklinn a Ridge , wa s a t th e sam e

Fig. 10 . Gcoseismi c sectio n KK' . (Se c Fig . 2 fo r lin e locatio n an d Fig . 4 fo r colou r legend. ) Modifie d fro m Blysta d et al. (1995) .

Fig. 11 . Geoseismi c sectio n DD' . (Se e Fig . 2 for lin e location an d Fig . 4 fo r colou r legend. ) Modifie d from Blysta d e l al. (1995).

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elevation a s th e high s alon g th e edg e o f th e Trondelag Platform , i.e . th e Nordlan d Ridg e and th e Fr0y a High . Al l thes e elevate d area s along th e platfor m edg e wer e deepl y erode d i n Late Jurassi c time , leavin g the crystallin e basement a t shallo w depth s an d i n place s eve n subcropping th e uppermos t Jurassi c and/o r th e Cretaceous unit s (Fig s 6 . 1 0 an d 11) . Th e fla t tops o f thes e erode d high s wer e the n al l part o f the commo n Lat e Jurassi c peneplai n o f th e Trondelag Platfor m (se e also Rii s 1996) . The expansio n o f th e Cretaceou s sequenc e across th e Vingleia , Bremstei n an d Revfalle t Fault Complexe s alon g th e easter n boundarie s of th e Halte n an d Donn a Terrace s show s tha t these fault s controlle d th e final separation o f th e terraces fro m th e platfor m durin g tw o phase s o f subsidence; on e i n Earl y Cretaceou s tim e an d the last , an d majo r one , i n post-Cenomania n time (Figs 5 and 6) . The latter is part o f the main Late Cretaceou s tectoni c episod e see n i n th e Voring Basi n (see below). Durin g thi s phase, th e southern par t o f th e Nordlan d Ridg e wa s subjected t o uplif t an d faulting , whic h destroyed the Lat e Jurassi c peneplai n i n tha t area . Thi s part o f th e Nordlan d Ridg e wa s furthe r uplifted during severa l phase s i n Tertiar y tim e s o tha t Triassic t o Jurassi c rock s no w subcro p Paleo cene t o Pliocen e strata.

The Voring Basin area The typica l featur e o f bot h th e Vorin g Basi n area an d th e Mor e Basi n i s th e enormou s thickness o f th e Cretaceou s sequenc e (Fig s 5 and 7) . I n places , th e bas e o f th e Cretaceou s units is as deep a s >9s twt . and th e axial parts of the basin s are generally deeper tha n 7 s twt. This means tha t th e bas e o f th e Cretaceou s sequenc e in th e centra l part s o f th e basin s lie s a t depth s bet ween 900 0 an d 13000m . The Vorin g Basin area i s bounded t o th e west by th e Vorin g Escarpmen t alon g th e Vorin g Marginal High , an d t o th e eas t b y th e faul t complexes alon g th e edg e o f th e Trondela g Platform. Th e basi n are a i s bisecte d b y th e Fles Faul t Complex , whic h runs alon g th e basi n axis from the Jan Maye n Lineamen t i n the sout h to the Bivrost Lineament i n the north. Th e othe r important tectoni c boundar y withi n the basi n is the Sur t Lineament , includin g th e Ry m Faul t Zone, i n the northern hal f of the basin area . Thi s feature run s paralle l t o th e Ja n Maye n an d Bivrost Lineament s an d i s obviousl y controlle d by th e sam e ancien t NW-S E structura l grai n o f the basement. Th e western part o f the basin are a is dominated b y the basin paralle l Gjalla r Ridge .

The developmen t o f th e Vorin g Basi n are a should b e describe d i n term s o f thre e mai n phases; from lat e Mid-Jurassi c t o Lat e Cenoma nian time , fro m Lat e Cenomania n t o Earl y Paleocene time , an d fro m Earl y Paleocen e tim e to th e present . Late Mid-Jurassic to Late Cenomanian phase. Th e Fle s Faul t Comple x probabl y origi nated a s a Mid - t o Lat e Jurassi c zon e o f norma l faults wit h a n eastwar d polarit y alon g it s northern NE-S W trendin g segment , an d a westward polarit y alon g it s souther n N- S trending segment . B y several phases o f reactivations, th e Fle s Faul t Comple x ha s playe d a central rol e i n th e tectoni c developmen t o f th e area sinc e Earl y Cretaceou s time . Th e Ra s an d Traen Basin s wer e establishe d a s dee p depocen tres between th e Fle s Faul t Comple x t o th e wes t and th e terrace s t o th e eas t (Fig s 5 . 6 an d 11) . The are a t o th e west o f the Fle s Faul t Complex , comprising the presen t Vigri d and Nagrin d Syn clines, th e Hel Graben an d th e Ny k High , probably constituted a singl e broad Earl y Cretaceou s basin somewha t shallowe r tha n th e Ra s an d Traena Basin s (e.g . Fig . 6) . Researcher s wh o prefer a shallower correlation for th e bas e o f th e Cretaceous sequenc e t o th e w ?est o f th e Fle s Fault Comple x hav e a regiona l high , informall y termed th e 'Full a Ridge' , i n place o f th e present author's wester n Earl y Cretaceou s basi n (Skog seid & Eldhol m 1989;'Skogsei d el al. 1992) . The locatio n o f th e wester n flan k o f th e western Earl y Cretaceou s basi n i s uncertain , but th e presen t author' s interpretatio n o f th e base o f th e Cretaceou s sequenc e implie s tha t the Lower Cretaceous unit s thin and onla p older rocks i n th e souther n parts o f th e Gjalla r Ridg e (Figs 6 an d 12) . T o th e nort h o f th e Gjalla r Ridge, alon g th e wester n margi n o f th e He l Graben. th e Lowe r Cretaceous sequence is interpreted t o thi n i n a simila r w ?ay (Fig . 11) . Thi s i s taken t o indicat e tha t th e wester n flan k o f th e Early Cretaceous basi n was situated more or les s parallel t o th e presen t Gjalla r Ridge , bu t probably offse t t o th e wes t o f th e ridg e alon g its central segments. This Early Cretaceous basi n formation i s believed t o b e th e resul t o f therma l subsidence afte r th e majo r riftin g episod e i n late Mid-Jurassic t o Earl y Cretaceou s time . Late Cenomanian to Early Paleocene phase. Through Cenomania n tim e th e differen t depo centres o f Early Cretaceous time an d th e Halte n and D0nn a Terrace s merge d int o on e larg e basin, th e Vorin g Basi n sensu stricto. This basi n was mainly formed b y the renewed tectonic s tha t

Fig. 12 . Geoseismi c sectio n II' . (Se e Fig . 2 for lin e location an d Fig . 4 fo r colou r legend. ) Modifie d fro m Blysta d e t al. (1995) .

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started i n lates t Cenomania n t o earlies t Tur onian tim e an d th e associate d subsidenc e tha t reached it s maximu m i n Campania n time . Hence, th e V0rin g Basi n sensu stricto i s defined at the post-Cenomanian Cretaceous levels (Blystad e l al 1995) . At th e transitio n int o Turonia n time , th e Gjallar Ridg e starte d t o ris e an d forme d th e western flan k o f th e V0rin g Basin . Thi s i s demonstrated b y th e thinnin g an d onla p o f the post-Cenomania n strat a toward s th e tilte d Top Cenomania n unit s o n th e easter n flan k o f the Gjalla r Ridg e (Fig s 5 an d 6) . Thi s Lat e Cenomanian-Early Turonian eastwar d tiltin g of the Gjalla r Ridg e i s mirrored b y a simultaneou s westward til t o f th e souther n par t o f th e Trondelag Platform ; acros s th e Fr0y a Hig h and Froa n Basin the post-Cenomanian sequenc e thins an d onlap s th e tilte d To p Cenomania n units (Fig . 10) . Thi s symmetrica l basinwar d tilting o f th e basi n flank s indicate s tha t th e development o f th e V0rin g Basi n starte d wit h a phas e o f renewe d o r accelerate d subsidence . At th e sam e time , o r jus t afterwards , faultin g activity starte d i n th e Gjalla r Ridg e an d con tinued throug h Campania n time , creatin g th e complex pattern o f strongly rotated faul t block s that characterize s the ridg e (Figs 5, 6, 7 and 12) . In th e sam e period , renewe d faultin g activit y took plac e i n th e Revfalle t Faul t Comple x along th e Nordlan d Ridg e (Fig . 5 ) and o n th e eastern boundar y fault s o f th e Halte n Terrac e (the Bremstei n an d Vinglei a Faul t Complexes ) (Fig. 6) , acros s whic h th e post-Cenomania n sequence demonstrate s a marke d growth . Thes e fault complexes then acted as the eastern flank of the Voring Basin , which implies tha t th e Halte n and Donn a Terrace s mus t b e regarde d a s part s of th e Vorin g Basi n sensu stricto. In th e norther n part o f th e Vorin g Basin , the Surt Lineamen t acted a s a tectonic hing e zone in Late Cretaceou s time . Th e post-Cenomania n sequence expand s greatl y t o th e nort h o f th e lineament. Th e lineament , includin g th e Ry m Fault Zone , constitute d th e southwester n flan k of th e He l Grabe n i n post-Cenomania n time s (Fig. 12 ) an d als o mark s th e transitio n zon e between th e very thick post-Cenomania n depos its in the Nagrind Synclin e and th e thinner post Cenomanian successio n o f th e Vigri d Syncline . The post-Cenomania n t o Campania n thicknes s is mor e o r les s constan t acros s th e Nagrin d Syncline. Ny k Hig h an d th e He l Graben . Thi s indicates tha t durin g mos t o f th e perio d fro m Cenomanian t o lates t Campanian-Maastrich tian tim e th e are a nort h o f th e Sur t Lineamen t was a majo r depocentr e befor e th e formatio n o f the syncline s and high s (Fig . 11) .

The fac t tha t th e Lat e Cretaceou s phas e o f subsidence i n th e Vorin g Basi n starte d slightl y before th e observe d faultin g activity , an d tha t the maximu m subsidenc e rat e (i n Campania n time) seem s t o hav e coincide d wit h th e max imum faultin g activity , indicates that th e forma tion o f th e Lat e Cretaceou s Vorin g Basi n wa s not th e resul t of a simple sequence of rifting an d consequent therma l subsidence . Th e Lat e Cre taceous subsidenc e wa s probabl y drive n b y active tectoni c stresse s rathe r tha n passiv e thermal relaxatio n (see discussion below) . South o f th e Sur t Lineament , th e Vorin g Basin seems to be regionally folded into a central anticline alon g th e Fle s Faul t Complex , flanke d by tw o syncline s (Fig . 5) . The wester n synclin e is terme d th e Vigri d Syncline, whereas th e east ern synclin e coincides wit h th e Ra s Basin . Thi s regional foldin g suggest s a probabl y regiona l compressional tectoni c phase involvin g reactivation alon g th e Fle s Faul t Complex . Th e rol e of the Fle s Faul t Comple x i s well demonstrate d i n section i n Fig . 8 b y th e mor e intens e interna l folding o f the Ra s Basi n to th e eas t o f th e faul t complex compare d wit h th e broad , ope n fol d o f the Vigri d Synclin e t o th e west . Th e smaller scale interna l fold s o f th e Ra s Basi n hav e a NE-SW strik e an d ma y b e see n o n th e ma p o f the To p Cenomania n unit s (Fig . 13) . The fold s are bes t develope d i n th e regiona l ben d wher e the strik e o f th e faul t comple x change s fro m N—S to NE-SW . indicating space problem s dur ing th e reactivatio n o f th e Fle s Faul t Complex . North o f the Surt Lineament the Voring Basin is folde d int o thre e NE-S W trendin g synclines (the Traen a Basin , th e Nagrin d Synclin e an d the Hel Graben) separate d b y two anticlines (the Utgard an d Ny k Highs ) (Fig . 11) . Th e fold s affect th e whole of the Cretaceou s sequence , bu t in th e uppermos t par t o f th e Uppe r Cretaceou s (Maastrichtian?) sequenc e th e Nagrin d Syncline becomes progressivel y les s tight , indicatin g a latest Campanian-Maastrichtia n ag e fo r th e folding. Th e Ny k an d Utgar d High s sho w n o internal onlap s and . therefore , mus t als o hav e been forme d i n lates t Cretaceou s time . A pos sible minor growt h of the uppermost Cretaceou s sequence i n th e He l Grabe n indicate s tha t th e faulting activit y i n th e Ny k Hig h an d alon g the Ry m Faul t Zon e tha t establishe d th e He l Graben initiate d jus t befor e th e folding . Thi s indicates tha t th e overal l anticlina l shap e o f th e Nyk Hig h an d it s positio n betwee n th e tw o synclines i s th e resul t o f changin g phase s o f both extensio n an d compression . Thinnin g of the Lower Cretaceou s sequenc e indicate s that the Utgar d Hig h wa s probabl y initiate d a s a marginal hig h alon g th e Trasn a Basi n alread y in

TECTONIC EVOLUTIO N O F CONTINENTA L MARGI N

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Fig. 13 . Two-wa y trave l tim e map o f the to p Cenomania n unconformit y i n the V0rin g and M0r e Basins . Contours labelled i n milliseconds.

Early Cretaceou s tim e (Fig . 11) . However , it s present shap e as a faulte d anticlin e between two synclines i s the resul t of a combinatio n of latest Cretaceous extensio n an d subsequen t compres sion similar to the development of the Nyk High. In additio n t o bein g a tectoni c transfe r an d hinge zone , th e Sur t Lineamen t als o probabl y marks a chang e i n th e substrat e t o th e Cretac eous an d lates t Jurassi c basi n fill . Sout h o f th e lineament th e basi n fill probably overlie s Upper Palaeozoic t o Jurassi c strata , wherea s t o th e north th e basi n fill rests directly on th e crystal line basemen t o r deepl y erode d Uppe r Palaeo zoic rock s (Fig s 12 and 14) . In the north , thi s statement is based on a westwar d extrapolatio n of th e geologica l relationship s reveale d b y wells and seismi c mapping o f the northern part o f the Trondelag Platfor m (e.g . well s 6609/7- 1 an d

6609/10-1 i n Fig . 11) . A westwar d thinning , caused b y pre-Cretaceou s erosion , o f th e base ment cove r acros s th e platfor m edg e an d basi n flanks indicate s a ver y shallo w basemen t als o beneath th e norther n V0rin g Basi n (Fig . 11) . To th e south , on e ma y als o argu e for a shallow basement unde r th e basi n are a fro m th e ero sional thinnin g o f th e basemen t cove r i n th e Fr0ya Hig h an d Sklinn a Ridg e alon g th e plat form an d terrac e edges (Figs 6 and 10) . Influx o f sand fro m th e wes t int o th e wester n Halte n Terrace durin g Mid-Jurassi c tim e als o indicates that th e V0ring Basin area was subject t o stron g uplift an d erosio n the n (Gjelber g e t al. 1987) . However, section s suc h a s Fig s 5 and 1 4 show that Jurassi c an d Triassi c rock s ar e presen t i n many parts of the basin margin along the terrace areas, indicatin g tha t Triassic , an d perhap s

Fig. 14 . Geoseismi c sectio n HIT . (Se e Fig . 2 fo r lin e locatio n an d Fig . 4 fo r colou r legend. ) Modifie d fro m Blysta d ct al. (1995) .

TECTONIC EVOLUTIO N O F CONTINENTA L MARGI N Jurassic, deposit s ma y als o b e present withi n th e adjacent part s o f the basin (se e also Brekk e at al. 1999). Sub-basin fill internal parallel reflectors in the Slettringe n an d Gjalla r Ridge s probabl y represent sedimen t layers . Across th e whol e o f th e Vorin g Basin , th e base o f th e Tertiar y sequenc e i s a regiona l angular unconformit y constitutin g a generall y even surfac e no t involve d i n th e foldin g o f th e underlying Cretaceou s strat a (Fig s 5 an d 8) , substantiating th e lates t Cretaceou s ag e fo r th e folding. Th e majo r highs , suc h a s th e Gjalla r and Nordlan d Ridges , th e Ny k an d Utgar d Highs an d th e centra l anticlin e alon g th e Fle s Fault Complex , ar e see n t o hav e bee n erode d (Figs 5 and 11) . This mean s tha t th e whole basin was uplifte d a t th e en d o f th e Cretaceou s an d beginning o f th e Tertiar y period .

345

Early Paleocene time t o the present. I n th e tw o large synclines , th e Vigri d an d Nagrin d Syn clines, wea k foldin g probabl y persiste d int o Paleocene tim e an d th e syncline s wer e infille d by a Paleocen e wedg e tha t onlap s th e synclin e limbs (Fig s 5 an d 11) . Thi s indicate s tha t th e whole are a staye d uplifted , givin g erosio n o f the mos t elevate d high s throug h mos t o f Paleocene time . Th e uplif t increase d westward s s o that th e Gjalla r Ridg e wa s deepl y erode d an d was subsequentl y onlappe d b y th e Paleocen e sequence and progressivel y oversteppe d b y early Eocene deposits (Figs 4 and 15) . Even to the east some o f the highes t area s wer e als o onlappe d b y the Paleocen e sequenc e an d oversteppe d b y younger sediments ; these area s includ e the Ny k High, th e Nordlan d Ridg e an d a n are a i n th e northern Ra s Basi n close t o th e Sur t Lineament

Fig. 15 . Two-wa y trave l tim e ma p o f the nea r to p Paleocen e unconformit y i n th e V0rin g an d M0r e Basins . Contours labelle d i n milliseconds.

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(Fig. 15) . This area must hav e been a dome dur ing Paleocen e times , an d subsequentl y becam e inactive an d subsided . During Tertiar y time , th e Cretaceou s easter n basin margin , th e entir e Tr0ndela g Platfor m and probably large parts of the present mainland were transgressed . Th e mai n Paleocen e t o Pliocene depocentr e seem s t o hav e bee n t o th e west, bounde d b y th e Sur t an d Ja n Maye n Lineaments, th e Vorin g Escarpmen t an d th e reversed Fle s Faul t Comple x (Fig s 6 an d 12) . In Neogen e time , th e mainlan d an d adjacen t regions were strongly uplifted, and huge amounts of Plio-Pleistocen e sediment s wer e transporte d westwards, buildin g the shel f edge t o it s present position (Fig s 6 and 11) . The tectoni c activity i n latest Cretaceous tim e continued int o Paleocene time and was probably part o f th e riftin g tha t wa s centre d wes t o f th e

Voring Escarpmen t an d le d t o th e break-u p o f the continen t an d separatio n o f Norwa y an d Greenland (Skogsei d & Eldhol m 1989 ; Skog seid 1994) . Late r tectoni c phase s wer e probabl y linked t o majo r relativ e plate boundar y move ments. A s i n Lat e Cretaceou s time , th e Tertiary tectonics wer e controlle d b y reactivation s o f the majo r lineament s and faul t complexes . Th e Voring Escarpment , however , wa s a ne w ele ment, whic h wa s initiate d i n Eocen e tim e an d established itsel f a s a n importan t additiona l feature o f tectoni c contro l throughou t Tertiar y time. Th e Vorin g Escarpmen t i s a combination of flexurin g an d faultin g (Fig . 8) . Th e mai n faulting too k plac e i n Earl y Eocen e tim e an d established th e separatio n betwee n th e Vorin g Marginal Hig h an d th e Vorin g Basi n area . The margina l hig h wa s additionall y uplifte d (relative t o th e basin ) durin g tw o phase s o f

Fig. 16 . Two-wa y travel time map o f the bas e Tertiary unconformity i n the Vorin g and Mor e Basins . Contours labelled i n milliseconds.

Fig. 17 . Geoseismi c section JJ'. (Se e Fig. 2 for lin e location an d Fig . 4 for colou r legend. ) Modified from Blysta d et al (1995) .

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flexuring alon g th e V0rin g Escarpmen t i n th e latest Eocene-Earl y Oligocen e an d Lat e Mio cene times , respectively. This i s indicated b y th e onlap o f post-Eocen e strat a ont o th e flexured top o f th e Eocen e sequence , an d a simila r rela tionship between the middle Miocen e unconfor mity an d Lat e Miocen e strat a (Fig . 8) . The sam e phase s o f movemen t ar e als o recorded a s th e mos t importan t Tertiar y reacti vation phase s o n th e majo r tectoni c lineaments: the Jan Maye n and Bivros t Lineaments, the Fles Fault Comple x an d th e Sur t Lineamen t including the Ry m Fault Zone . Th e activity resulted in large domes an d arche s tha t ar e well defined o n the maps o f the bas e Tertiary unconformity and of th e Middl e Miocen e unconformit y (Fig s 1 6 and 19) . Reverse reactivation s o f th e Fle s Faul t Com plex governe d th e growt h o f the very prominen t Helland-Hansen Arc h (Fig s 6 and 17) . A similar compressional structure , th e Modgun n Arch , was controlle d b y reactivation s o n th e Ja n Mayen Lineamen t (Fig . 17) . South o f 65° N th e strata across the arch are continuous; this makes it possibl e t o correlat e th e Cretaceou s reflector s from wel l contro l i n th e eas t acros s th e Fle s Fault Comple x int o th e wester n part s o f th e basin are a (Fig . 9) . A stud y o f th e interna l stratigraphic relationship s o f th e Helland Hansen Arc h i n thi s are a support s th e datin g of th e tectoni c events . Th e Middl e Miocen e unconformity define s th e present arc h surfac e of the Helland-Hanse n Arc h (Fig . 18) . The Uppe r Miocene and Pliocen e strata onla p an d overste p this surface . Th e unconformit y i s probabl y erosional an d truncate s th e underlyin g strata . The oldes t strat a overlyin g thi s unconform ity hav e bee n date d t o lates t Mid-Miocen e time i n nearb y wells , an d th e younges t strat a beneath th e unconformit y hav e bee n date d t o Early Miocen e tim e (Eidvi n & Rii s 1991) . Thi s implies tha t th e lates t phas e o f foldin g o f the arc h too k plac e i n Mid-Miocen e time , according t o th e datin g o f Eidvi n & Rii s (1991, 1992). (Th e datin g o f Gradstei n & Backstro m (1996) woul d plac e th e even t i n Lat e Mio cene time.) Flattening th e unconformit y stil l leave s a n arch structur e a t th e Uppe r Eocene-Lowe r Oligocene surface . Ther e ar e som e indication s of thinnin g an d onla p ont o thi s surfac e o n th e flanks of the arch, but thes e relationships ar e not conclusive (Fig . 18) . However , combine d wit h the clea r onlap s ont o th e flexur e o f th e sam e surface alon g th e Vorin g Escarpmen t (se e above), thes e relationship s indicat e onse t o f folding o f th e arc h alread y i n lates t Eocen e t o earliest Oligocen e time .

In th e norther n par t o f the V0rin g Basin , th e Eocene an d Miocen e tectonic s ar e controlled b y the revers e reactivation s alon g th e Fle s Faul t Complex an d th e Sur t Lineament , whic h initiated foldin g an d domin g o f th e Middl e Miocene Unconformity . The prominent Naglfar Dome i s a n inversio n o f th e He l Grabe n con trolled b y th e Ry m Faul t Zon e an d th e fault s along th e Nyk Hig h (Figs 1 1 and 14) . The Vema Dome i s situated on th e intersection between th e Surt Lineamen t an d th e fault s alon g th e Ny k High (Fig . 1) . Th e Traen a Basi n wa s inverte d along th e Fle s Faul t Complex , givin g ris e t o minor dome s an d arche s (Fig s 1 1 an d 16) . In addition , th e tectoni c activit y gav e ris e t o further elevatio n o f th e Nordlan d Ridge , th e Nyk an d Utgar d High s (Fig s 5 and 11) . The Lowe r t o Middl e Miocen e hiatu s i s recorded al l ove r th e Norwegia n Continenta l Shelf (Eidvi n & Rii s 1989 , 1991 . 1992 ; Eidvi n et al 1993 ; Jord t e t ai 1995 ; Gradstei n & Backstrom 1996) . Thi s implie s tha t jus t befor e the foldin g i n Miocen e time , ther e mus t hav e been a perio d o f erosio n o r non-depositio n t o give th e observe d hiatus . A s onl y deep-wate r faunas ar e recorde d immediatel y belo w an d above th e unconformity , thi s mus t hav e bee n a submarine erosion , a s a n uplif t o f th e whol e basin are a t o abov e wav e base seem s excluded. However, Gradstei n & Backstro m (1996 ) stated that th e basi n flank s see m t o hav e reached wave base an d bee n considerabl y eroded . Th e mech anism behin d thi s regiona l even t o f submarin e erosion is not known . However, there is evidence of a regiona l uplift , lastin g fo r abou t 7 Ma. o f the whole continental margin and th e surrounding mainland s i n Earl y t o Mid-Miocen e tim e (Anderton e t al . 1979 ; Jordt e t al . 1995) . The Mid-Miocene folding created a significan t submarine topograph y tha t laste d throug h Lat e Miocene time . The ma p o f the Middl e Miocen e unconformity show s that th e summit of the large domes an d arche s lie s 1000-1800m s tw t abov e the adjacen t low s (Fig . 19) , corresponding t o a sea-bottom topograph y o f th e orde r o f 1000 2000 m at tha t time . Substantial deposit s o f flatlying Upper Miocene sediments filled in the lows, onlapping an d buryin g th e technicall y create d topography. A regiona l hiatu s i s also demon strated a t th e bas e o f th e Uppe r Pliocen e sequence alon g th e continenta l margi n (Eidvi n & Riis 1989; Rii s & Fjeldskaar 1992 ; Eidvin et al. 1993). On the mainland and i n the shallow part s of th e continenta l margi n thi s hiatu s i s a n unconformity relate d to uplif t an d dee p erosion. In th e deeper-wate r part s o f th e margi n i t i s mainly a surface o f non-deposition, downlapped by th e deposit s o f a hug e sedimentar y wedg e

Fig. 18 . Lin e interpretation o f seismic line MB-6453-91 . (a ) The presen t structura l geometry , (b ) The Middl e Miocen e unconformit y surface restored t o horizonta l (flattening), no t compensate d fo r compaction . (Se e Fig . 2 for lin e location.) Seismi c sequences: A, Paleocene ; B , Eocene; C , Oligocene , an d Lowe r Miocene ; D , Uppe r Miocene an d Lowe r Pliocene; E, Uppe r Pliocen e and Pleistocene .

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Fig. 19 . Two-wa y travel time map o f the Middl e Miocene unconformity i n the V0ring and M0r e Basins. Line of intersection betwee n Upper Pliocene and Middl e Miocene unconformities delineate s areas where sediments of Middle and Upper Miocene and Lower Eocene are missing. Within these areas contours are on the base Upper Pliocene unconformity. Contour s labelled in milliseconds.

prograding from th e eroded mainland an d shelf. However, i n place s thi s hiatu s i s a n erosiona l unconformity als o withi n th e basi n area . I n th e V0ring Basi n th e Pliocen e unconformit y i s relatively flat , an d Uppe r Pliocen e sediment s overstep th e Middl e Miocen e unconformit y o n the summit of the large domes and arches (Figs 6 8 and 9) . In these elevated areas th e Middle an d Upper Miocen e sequenc e i s missin g becaus e o f tectonic uplift , erosio n an d non-deposition : these area s includ e th e Helland-Hanse n Arch , the Naglfa r an d Vem a Domes , th e Ny k an d Utgard Highs , th e easter n edg e o f th e V0rin g Marginal High , an d area s involve d i n th e Ja n Mayen Lineamen t (Fig . 19) . Th e mechanis m behind th e regional , submarin e erosio n i n Plio cene time is also unknown, but i s clearly related

to the onset o f glaciations and rapid uplift o f the surrounding mainlan d (Rii s & Fjeldskaar 1992 ; Riis & Jensen 1992 ; Riis 1996) .

The Voting Marginal High The margina l hig h is situated t o th e wes t of th e V0ring Escarpment betwee n the Jan Maye n and Bivrost Lineaments . Th e hig h comprise s Ter tiary sediment s o n to p o f thic k Lowe r Eocen e flood basalts , whic h ar e probabl y underlai n by continental crus t tha t progressivel y thin s an d becomes transitiona l t o oceani c crus t toward s the west . Earl y worker s place d th e boundar y between continental an d thickened oceanic crust beneath th e observe d se t o f seaward-dippin g

TECTONIC EVOLUTIO N O F CONTINENTA L MARGI N

351

Fig. 20. Mode l fo r th e development of the V0ring Escarpment an d th e formation of the 'inne r flows'. The same model i s inferred fo r th e Faeroe-Shetland Escarpment. (Se e text for furthe r discussion. ) reflectors tha t was identified a s volcanic flows by well ODP-64 2 (Talwan i e t al 1983 ; Hinz e t al 1984; Eldholm e t al 1989 ) (Fig. 6) . However, th e position o f thi s ocean-continen t boundar y ha s not bee n substantiate d b y relevan t detaile d geophysical data . Thi s transitio n i s probabl y a wide comple x zon e involvin g intrusion s an d tectonic thinnin g an d fragmentatio n o f th e continental crust . Th e magneti c anomalie s 24 A

and 24B in the V0ring Marginal High, as define d by Skogseid & Eldholm (1987), do not appea r a s typical sea-floo r spreadin g anomalie s i n curren t compilations (Olese n e t al . 1997) . The magneti c anomaly pattern i n the high rather resembles the signature of the shallow basement ridges that are common i n th e continenta l margin s aroun d th e Norwegian-Greenland Se a (e.g . th e Lofote n and Utr0s t Ridges) .

Fig. 21 . Geoseismi c sectio n CC' . (Se e Fig . 2 fo r lin e locatio n an d Fig . 4 fo r colou r legend. ) Modifie d fro m Blysta d et ai (1995) .

TECTONIC EVOLUTIO N O F CONTINENTA L MARGI N As described above , th e presen t easter n flan k of the V0ring Marginal Hig h was probably clos e to th e wester n margi n o f th e V0rin g Basi n throughout Cretaceou s time . Therefore , th e marginal hig h probabl y wa s the western bound ing platfor m an d sedimen t sourc e are a t o th e Cretaceous Vorin g Basi n are a (se e als o Brekk e et al. 1999) . A s th e presen t boundar y element , the V0rin g Escarpment , i s a post-Paleocen e feature, th e regiona l bas e Tertiar y erosio n sur face deepl y truncatin g th e Gjalla r Ridg e an d Fenris Grabe n in the western V0ring Basin must have continued a s an unbroke n surfac e into th e area of the present margina l hig h (Fig. 20) . Thi s implies that, in latest Cretaceou s time , the western platform include d th e Gjallar Ridg e and th e Fenris Graben . I n lates t Paleocen e t o Earl y Eocene time , larg e volume s o f tholeiiti c floo d basalts flowe d acros s th e erode d platfor m an d in place s covere d bot h th e Fenri s Grabe n and the Gjallar Ridge . When the V0ring Escarpment wa s initiated , th e earl y flow s eas t o f th e escarpment wer e separate d fro m th e emergin g marginal hig h an d wer e probabl y submerge d and ar e now termed 'th e inner flows' (Hinz et al. 1982; Talwan i e t a l 1983 ) (Fig . 20) . Becaus e of th e water-fron t the n establishe d alon g th e V0ring Escarpment , th e younge r flows of flood basalt fro m th e fissur e eruption s t o th e wes t stopped o n reachin g th e escarpmen t an d a prominent coastlin e lava front wa s built, enhancing th e topograph y o f th e escarpment . Th e substrate t o th e floo d basalt s i n th e margina l high i s unknown , bu t probabl y consist s o f deeply erode d Cretaceou s o r olde r rocks . The norther n provinc e betwee n th e Bivros t Lineament an d the Jennegga Hig h The continental margin in the northern province is ver y narrow . Th e mai n element s ar e th e Vestfjorden an d Ribba n Basins and th e elevated area o f th e Lofote n an d Utr0s t Ridge s (Fig. 1) . The Eocene flood basalt lavas onlap the flank of the Utr0s t Ridg e an d cove r th e adjacen t deep water area , makin g i t ver y difficul t t o ma p th e underlying geology (Mjeld e et al . 1993) . The are a aroun d th e Jennegg a Hig h o n th e northernmost par t o f Utr0st Ridg e constitutes a deeply erode d transvers e ridg e separatin g th e Norwegian Se a continental margin t o th e sout h from th e Barent s Se a margin t o th e north . Th e Utr0st Ridge , Ribban Basin , Lofoten Ridge and Vestfjorden Basi n constitut e a characteristi c regional horst-and-grabe n syste m (Fig . 21) . The Ribba n Basi n an d th e Vestfjorde n Basi n are NE-SW striking half-grabens wit h 4-5 s twt

353

of Cretaceou s basi n fil l an d wit h mutuall y opposite polarities . Th e basi n boundar y fault s run alon g bot h flank s o f th e narro w basemen t horst o f th e souther n Lofote n Ridge , whic h separates th e tw o basin s (Fig . 21) . The Utros t Ridge an d th e mainlan d ar e th e rotated , non faulted flank s o f th e Ribba n an d Vestfjorde n Basins, respectively . I n th e ver y nort h o f th e Ribban Basi n th e boundin g fault s chang e t o a n easterly polarity . Th e Ribba n Basi n i s divide d into tw o sub-basins , th e Skomvae r Sub-basi n i n the sout h an d th e Havbae n Sub-basi n i n th e north. Th e genera l hors t an d grabe n geometr y was initiated because o f late Mid-Jurassic-Earl y Cretaceous crusta l stretchin g an d late r therma l subsidence an d mino r reactivations . Seismi c mapping show s tha t ther e wa s a phas e o f rotation an d subsidenc e i n th e Ribba n Basi n in the Mid-Cretaceou s time . Truncatio n o f th e Cretaceous successio n (Fig . 21 ) indicate s tha t the whol e o f thi s are a wa s uplifte d an d deepl y eroded i n lates t Cretaceou s time , a phas e o f uplift als o recorde d i n th e V0rin g Basi n t o th e south (see above). However, the Lofoten-Utr0s t area experience d a stronge r uplif t tha n th e V0ring Basin , and th e Bivros t Lineament acte d as th e hing e line between the tw o areas . Phase s of renewe d uplif t an d erosio n probabl y too k place i n Tertiar y time , bu t th e detail s ar e no t known becaus e o f th e erosio n o f th e Tertiar y sequences. However, by reference t o the tectonic development i n th e V0rin g Basi n t o th e sout h (see above ) an d i n th e adjacen t Barent s Se a (Brekke & Rii s 1987 ; Gabrielsen e t al . 1990) , i t seems likel y tha t th e mai n phase s too k plac e i n Eocene, Miocen e and Pliocen e time. As a result of the erosion, crystallin e basement rock s ar e a t present exposed i n the Lofoten Ridg e and a t th e sea be d o n loca l high s o n th e Utr0s t Ridg e (the R0s t an d Jennegg a Highs) . The southern province betwee n 62°N and the Jan Maye n Lineamen t The mai n elements of the souther n provinc e ar e the M0r e Basin , th e Mor e Margina l Hig h an d the M0re-Tr0ndela g Faul t Comple x (Fig . 1) . The souther n boundar y o f th e provinc e i s th e area o f intersectio n betwee n th e NE-S W tren d of th e M0re-Tr0ndelag Faul t Comple x an d th e N-S t o NNE-SSW trends of the northern Nort h Sea a t c . 62°N (Fig . 1) .

The More Basin The Mor e Basin is bounded t o th e nort h b y the Jan Maye n Lineament , t o th e southeas t b y

Fig. 22 . Geoseismi c sectio n MM' . (Se e Fig . 2 for lin e location an d Fig . 4 fo r colou r legend. ) Modifie d fro m Blysta d et al. (1995).

Fig. 23 . Geoseismi c sectio n NN' . (Se e Fig . 2 for lin e locatio n an d Fig . 4 for colou r legend. ) Modifie d fro m Blysta d e t al. (1995) .

Fie 2 4 Seismi c line MB-06-9" > Th e overall downflexed natur e of the eastern flan k o f th e Mor e Basin . (Note the general eastward polarit y of the major fault s i n the flank.) The lin' e , s th e h'2 fo , easter n par , o f geoseismi c sectio n MM ' i n Fig . 2 2 (see Kig . 2 fo r lin e location) . Whit e arrow s poin t a t th e bas e Cretaceou s unconlorm.ty .

TECTONIC EVOLUTIO N O F CONTINENTA L MARGI N the M0re-Tr0ndela g Faul t Comple x an d t o th e west b y th e Faeroe-Shetlan d Escarpment . The basi n i s define d a t th e bas e Cretaceou s unconformity an d b y a greatl y expande d Cre taceous sequence . Th e basi n ha s a n overal l NE-SW trend an d in the axial parts of the basin the Cretaceous interva l may be up t o 6 km thic k and th e bas e o f th e Cretaceou s sequenc e lie s between 7 and 9 s tw t (Fig . 7). The lat e Mid-Jurassic-Earl y Cretaceou s rift ing wa s th e mai n tectoni c episod e i n th e M0r e Basin (Fig. 22) . The Cretaceous sediment s onlap the bas e Cretaceou s unconformit y al l alon g both th e eastern an d western margins, indicatin g that th e basi n wa s mainl y forme d b y down warping o f it s flank s subsequen t t o th e riftin g (Figs 2 2 an d 23) . Th e rif t fault s show polarit y away from the basin axis . The dip of the easter n flank i s a composit e o f th e downflexin g i n Early Cretaceous time s and th e Lat e Tertiar y uplif t o f the adjacen t mainland . Removin g th e effec t of th e Tertiar y uplif t leave s th e Cretaceou s More Basin as a perfectly symmetrical sag-basin. Influx o f coars e clasti c deposit s fro m th e west into a coastal o r shallow marine environmen t in the Slorebot n Sub-basin (Fig . 1 ) in Mid-Jurassic times (Jongpie r e t al. 1996 ) indicate s tha t th e deepest part s of the Cretaceous Mor e Basin were highly elevate d an d subjec t t o erosio n a t tha t time. Thi s i s i n agreemen t wit h th e observa tion tha t th e pre-Cretaceou s sequence s ar e thi n wherever recognize d withi n the Mor e Basin . The prominen t Vigr a High , clos e t o th e central axi s o f th e Mor e Basin , i s a large , NE-SW trendin g rotate d faul t bloc k onlappe d by the Lower Cretaceous sequenc e (Fig. 22). The bounding faul t alon g it s northeast margi n ha s a maximum thro w o f abou t 2500m . I t die s ou t towards th e south , an d end s a t th e Ja n Maye n Lineament to the north. The high was formed by large-scale faultin g i n th e lat e Mid-Jurassic Early Cretaceou s riftin g episode . Interna l reflec tors paralle l t o it s eastwar d slopin g summi t indicate th e existenc e of a thi n interva l of pre Cretaceous sediment s withi n the hig h (Fig . 24) . In sharp contras t t o the Voring Basin north of the Jan Mayen Lineament, the More Basin seems to hav e bee n tectonicall y quiet , an d experience d passive Cretaceou s subsidence subsequen t t o the late Mid-Jurassic-Earl y Cretaceou s rifting , an d no observabl e tectoni c activit y i n Lat e Cretac eous tim e an d onl y mino r activit y i n Tertiar y time. Th e Tertiar y tectonic s wer e mainl y asso ciated wit h reactivation s o f th e Ja n Maye n Lineament and minor faulting along the FaeroeShetland Escarpment . The thinnin g and onla p o f th e entir e Cretac eous sequenc e ont o th e slopin g basi n flank s

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strongly indicat e th e existenc e o f Cretaceou s bounding platform s an d sedimen t sourc e area s both acros s the now eroded mainlan d t o the east and across the M0re Marginal Hig h t o the west. Much o f th e wester n flank s o f th e Mor e Basi n are hidde n unde r th e Eocen e lava s s o tha t i t is no t possibl e t o produc e continuou s map s of detail s o f th e Cretaceou s features . However, in som e area s wher e th e lava s ar e reasonabl y seismically transparent , dat a sho w tha t th e Cre taceous sequenc e thin s b y onla p ont o th e west ern basi n flan k (Fig . 22) . Clos e t o th e presen t Faeroe-Shetland Escarpmen t th e Cretaceou s interval i s a s thi n a s o n th e uppe r part s o f the easter n basi n flan k (Fig . 22 ) an d o n th e edge o f th e Trondela g Platfor m (e.g . Fig s 6 and 11) . Thus, th e Faeroe-Shetland Escarpmen t is probabl y clos e t o th e Cretaceou s boundar y between th e Mor e Basi n an d th e wester n plat form (Fig . 22) .

The M0re-Trondelag Fault Complex The southeaster n margi n o f th e Mor e Basi n consists o f the M0re-Tr0ndelag Fault Complex , comprising a broa d NE-S W t o ENE-WS W trending syste m of fault-controlled ridges , highs and mino r basins . I n tw o o f thes e high s well s 6305/12-2 and 6306/10- 1 have proven crystalline basement overlai n b y a thi n Middl e Jurassi c sequence underneat h post-Albia n strat a (Jong pier e t al. 1996) . The Mane t Ridg e (south) , an d Gnausen, Giske , On a an d Goss a High s (north ) make u p a ro w o f such basement-core d high s that separat e th e mainl y Cretaceou s Magnu s Basin, Marul k Basi n an d Slorebot n Sub-basi n from th e main , deepe r par t o f th e Mor e Basi n (Brekke & Rii s 1987 ; Blystad e t al . 1995) . Thi s set of highs and mino r basins possibly originate d as a syste m o f horst s an d half-graben s alread y in Triassi c times , bu t wa s greatl y accentu ated durin g the lat e Mid-Jurassic-Earl y Cretac eous riftin g episode . Th e basement-core d Selj e High t o th e southeast i s also par t o f this system. The relativel y thi n Jurassi c an d Triassi c cove r to th e crystallin e basemen t alon g th e More Trondelag Faul t Complex , i s in agreemen t wit h the notio n o f a shallo w basemen t beneat h th e Cretaceous an d lates t Jurassi c basi n fil l i n the Mor e Basin (Fig . 22) . On th e mainland , th e faul t comple x ca n b e seen t o affec t th e Precambria n basemen t an d rocks o f Lower Palaeozoic , Devonia n an d Jur assic ages . Th e ENE-WS W structura l grai n i n northwestern Norwa y ca n b e related t o Caledo nian deformations , bu t th e faul t comple x ha s

Fig. 25 . Seismi c lin e MB-02-84 . Downfiexe d wester n slop e o f th e Froy a High . Whit e arrow s poin t a t th e bas e Cretaceou s unconformity . (Se e Fig . 2 fo r lin e location .

TECTONIC EVOLUTIO N O F CONTINENTA L MARGI N been reactivate d i n severa l episode s (Berin g 1992; Gr0nli e e t al 1994) .

The More Marginal High The margina l hig h i s situated t o th e west of the Faeroe-Shetland Escarpmen t an d bounde d t o the nort h b y th e Ja n Maye n Fractur e Zone . To th e south , th e M0r e Margina l Hig h merge s with th e Faero e Plateau , whic h include s th e Faeroe Islands . Th e hig h comprise s Tertiar y sediments o n to p o f thic k earl y Eocen e floo d basalts, and , lik e th e V0rin g Margina l High , probably contain s continenta l crus t tha t west wards become s increasingl y mor e intruded , faulted, thinne d an d stretche d (Fig . 22) . The substrat e t o th e floo d lava s in th e conti nental par t o f th e margina l hig h i s unknown . The westwar d thinnin g an d onla p o f th e Cre taceous unit s o f th e M0r e Basi n strongl y indicate tha t th e Cretaceou s sequenc e i s ver y thin o r missin g beneat h th e lava s o n th e high. The Paleocene sequenc e i n the M0re Basi n seems to thicken westwards towards the FaeroeShetland Escarpment , suggestin g earl y faul t activity alon g th e escarpmen t an d differentia l uplift an d erosio n o f th e margina l hig h a t tha t time. Th e fac t tha t th e earl y lava s flowe d fa r eastward o f th e presen t escarpmen t indicate s that th e are a wa s abov e se a leve l i n Lat e Paleocene-Early Eocene time . Subsequen t fault ing alon g th e escarpmen t probabl y submerge d and separate d th e 'inne r flows ' fro m th e marginal hig h (Fig . 20) . Th e presen t morphol ogy o f th e Faeroe-Shetlan d Escarpmen t seem s to b e mainl y th e resul t o f th e build-u p o f a volcanic fron t alon g tha t fault-relate d shoreline (see als o Smyth e e t al . 1983) . However , i n th e northern segment s i t i s enhance d b y mino r faulting i n Lat e Eocen e tim e (Fig. 22) . Probably becaus e o f a relativel y thi n cove r of lava in some places, it is possible t o identif y a deep-seated, faulte d angula r unconformit y within th e margina l high . A s i t i s juxtapose d by th e Cretaceou s sequenc e o f th e M0r e Basi n in th e hangin g wall , i t mus t b e olde r tha n the Cretaceou s sequence . I n Fig . 2 2 i t i s tenta tively suggeste d tha t i t ma y b e correlate d wit h the Uppe r Permia n unconformit y identifie d in th e Tr0ndela g Platform. The eastern basin flank s The easter n flan k o f th e M0r e Basi n obviously formed b y th e downflexin g o f th e rift-uncon formity o f lat e Mid-Jurassic-Earl y Cretaceou s time t o a wester n di p alon g th e basi n margin .

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A majorit y o f th e rif t fault s hav e a n easter n polarity an d thei r faul t plane s wer e rotate d t o a shallo w di p durin g th e flexurin g (Fig s 2 2 and 23) . I n th e presen t interpretatio n th e sam e flexure geometr y als o dominate s th e easter n flanks o f th e Ra s an d Traen a Basin s i n th e V0ring Basi n (Fig s 10 , 1 1 and 17) . This mean s that th e traditiona l map s showin g thes e basin s as the result of down-to-the-basin faulting alon g the Klakk , Revfalle t an d Ytreholme n Faul t Complexes ar e simplification s an d d o no t reveal th e tru e natur e o f th e crusta l mechanic s behind th e basi n formation . I n a n attemp t t o substantiate th e presen t interpretation , a serie s of seismic dip-lines are presented fro m the M0r e Basin i n th e sout h t o th e Traen a Basi n i n th e north (Fig s 24-30) . The dat a sho w th e genera l westwar d slop e o f the bas e Cretaceou s reflecto r an d tha t eve n th e apparently stee p western flank of the On a Hig h is a flexur e an d no t a faul t (seismi c dat a ar e plotted wit h a 2- 4 time s vertica l exaggeration ) (Fig. 24) . Along the Fr0y a Hig h th e basin flank is considerably steeper, but Fig . 2 5 indicates it to be the limb of a monoclinal flexure rather tha n a fault scarp . However , th e flexura l steepenin g o f the slop e i n some places ha s resulted i n a subse quent gravitationa l collaps e o f th e slope , givin g detached an d highl y rotated faul t block s (invol ving eve n th e basement ) hangin g i n th e slop e (Fig. 26) . Th e flan k o f th e Fr0y a Hig h i s par t of wha t i s terme d th e Klak k Faul t Complex , which als o run s alon g th e stee p flan k o f th e Halten Terrac e furthe r north . Th e dat a qualit y along the terrace flank is generally poor, bu t th e present autho r believe s th e flexurin g an d rota tion of early faults as shown in Figs 1 0 and 1 7 to be correct (compar e origina l data i n Figs 2 7 and 28, respectively) . The southern segmen t o f the Nordland Ridg e was strongl y reactivate d an d uplifte d durin g Late Cretaceou s an d Tertiary time , overprintin g much o f the late Mid-Jurassic-Early Cretaceou s tectonic features . Alon g thi s segmen t o f th e ridge, th e Lat e Cretaceou s basin-facin g fault s are clearl y a n importan t elemen t o f th e basi n flank (Fig . 5) . However , a clos e loo k a t th e original dat a reveal s tha t th e underlying , early structure o f th e Revfalle t Faul t Comple x i s a monoclinal flexur e (Fig . 29) . I n th e norther n segment o f the Nordlan d Ridg e thi s early (Lat e Jurassic-earliest Cretaceous ) flexurin g i s bette r preserved, bein g les s affecte d b y late r faultin g (Figs 1 1 and 30) . The characteristic back-rotate d Jurassic fault s facin g awa y fro m th e basi n ar e also clearl y observable. Accepting that th e underlying structure of the Late Jurassic-Earl y Cretaceou s easter n basi n

Fig. 26. Seismi c line FHM-91-102 . Gravitational collapse o f western slop e of the Froy a High . Whit e arrows point a t th e base Cretaceous unconformity. (Se c Fig . 2 for line location.) B y courtesy o f Nors k Hydr o a.s .

Fig. 27 . Seismi c line MB-04-84 . Th e overal l downflexed natur e o f th e wester n flan k o f th e F0y a High . (Not e th e indication s of Jurassi c fault s with a genera l eastwar d polarity in the upper part o f the flank.) The line is the basis for th e western part o f geoseismic section KK' in Fig. 1 0 (see Fig. 2 for line location). White arrow s poin t a t th e base Cretaceou s unconformity .

Fig. 28 . Seismi c line MB-6445-91, which is the basis for th e interpretation o f the overall downflexed natur e of the western flan k o f the Halte n Terrace in geoseismic sectio n JJ' i n Fig . 1 7 (see Fig . 2 fo r lin e location) . Whit e arrow s poin t a t th e bas e Cretaceou s unconformity .

Fig. 29 . Seismi c line NRGS-84-429, reprocesse d b y Sag a Petroleu m a.s , whic h illustrate s the underlying , Late Jurassi c monoclina l natur e o f th e wester n flan k o f th e Nordland Ridg e cut b y younger. Cretaceous faults . Th e lin e is the basis for the interoperation o f that part o f geoseismic section EE' in Fig. 5 (see Fig. 2 for line location). White arrow s poin t a t th e bas e Cretaceou s unconformity . By courtesy o f Sag a Petroleu m a.s .

Fig. 30. Seismi c line NRGS-84-462. The overall downflexed natur e of the western flank of the northern par t o f the Nordland Ridge . (Not e th e general eastward polarit y of the major Jurassic faults and the general lack of later. Cretaceous faults in this part o f the Nordland Ridge. ) White arrows poin t t o the base Cretaceou s unconformity . (See Fig. 2 fo r lin e location. ) B y courtesy o f Statoil .

TECTONIC EVOLUTIO N O F CONTINENTA L MARGI N flanks i s a flexur e implie s tha t bot h th e Mor e Basin an d th e V0rin g Basi n are a wer e symmetrical sag-lik e basi n area s i n Earl y Cretaceou s time (e.g . Fig s 6 an d 22) . Suc h a symmetry , combined wit h the grea t subsidenc e of the axial parts o f th e basi n area , strongl y indicate s tha t during th e lat e Mid-Jurassic-Earl y Cretaceou s rifting phas e the crus t of central parts of the rif t system was highly stretched an d attenuated. Th e accompanying hea t flo w mus t hav e bee n high , causing a significan t uplif t an d erosio n o f th e future basi n areas . A n earl y onse t o f increase d heat flow in Mid-Jurassic time would explain the enigmatic influ x o f san d fro m th e wes t i n th e western part s o f th e Halte n Terrac e an d Slorebot Sub-basi n durin g Jurassi c tim e (Gjel berg et al 1987 ; Jongpier et al 1996) . This could also explai n th e indication s o f a generall y thin pre-Cretaceous basemen t cove r all over both th e V0ring an d M0r e Basins . A s alread y argued , north o f the Sur t Lineamen t th e Cretaceou s (o r uppermost Jurassic ) units may even rest directly on th e basement . A possibl e fossi l opal-C T transition An anomalou s reflecto r o f possibl e diageneti c origin is observed in the Lower Miocene sequence throughout th e Vorin g Basi n wes t o f th e Fle s Fault Complex . Th e sam e reflecto r i s als o seen withi n th e Neogen e sequenc e o n th e V0r ing Margina l High , wher e i t wa s penetrate d b y well ODP-463 . According t o th e data fro m wel l ODP-643, opa l A i s converte d t o opal-C T at depth s o f 275-310 m (Roaldse t & H e 1995) . This dept h correlate s wel l wit h th e diageneti c seismic reflecto r (C . Magnus , pers . comm. ) (see also Fig . 6) . This reflecto r crosscuts th e beddin g reflectors where these are involve d in small-scal e compaction faultin g an d Lat e Miocen e doming . Th e reflected surfac e is not paralle l t o th e se a floor , nor t o th e Lat e Pliocene bedding , an d i s deeper and covere d b y considerably greater thicknesse s of Lat e Pliocen e sediment s t o th e eas t tha n t o the wes t (Fig . 6) . Th e transitio n surfac e i s parallel t o beddin g reflector s jus t belo w th e Upper Pliocen e unconformit y and i s situated a t a mor e o r les s constan t dept h o f 550 m belo w the Uppe r Pliocen e unconformity , bot h i n th e Voring Basin as well a s on th e Vorin g Marginal High. O n th e flank s o f th e larg e Lat e Miocen e arches an d dome s th e reflecto r clearly crosscuts the flexured bedding, but i s also seen to b e itself involved in the latest phase o f the flexuring. The same relationshi p i s observe d a t th e flexure d limbs of the V0ring Escarpment. These relationships indicate that th e phase transition surface is

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fossil an d wa s forme d probabl y i n lates t Miocene o r Earl y Pliocen e time . Th e fossi l nature o f th e phas e transitio n surfac e wa s als o suggested b y Roaldset & He (1995), on the basis of presen t temperatur e relationships . Th e esti mated transitio n temperatur e o f 55° C i n th e ODP-643 wel l (se e Fig . 6 ) correspond s t o a thermal gradien t o f 180°Ckm~ I , a s oppose d t o the presen t gradien t o f 50-65°C . Th e seismi c data (lin e NH-1-79: th e wester n en d o f geoseis mic section in Fig. 6 ) indicates a missing section of about 250 m as a result of slumping at the well location (OD P 643) , bu t allowin g fo r thi s wil l only reduce the calculated thermal gradient from 180 t o lOCTCknr 1 , stil l strongl y indicatin g a fossil process . Suc h fossi l opal- A t o opal-C T phase transitio n surface s are als o reporte d fro m the North Sea (Rundberg 1989 ) and th e Barents Sea (Roaldset & He 1995) . As th e opal- A t o opal-C T transitio n i s tem perature dependen t (Roaldse t & H e 1995) , th e fossil transitio n surface is a record of a peak heat flow at tha t tim e or a n even t of rapid reductio n of overburde n b y erosio n al l ove r th e wester n parts o f th e continenta l margin, o r a combina tion o f th e two. At wha t sedimen t depths th e phase transitio n took place is not known , as the depth o f erosion in Late Pliocene time is not known . Seismic data show that , i n a restricte d are a abov e th e Vigrid Syncline, a n extr a 15 0 m thic k sectio n o f Uppe r Miocene-Lower Pliocen e unit s i s preserve d a s an erosiona l remnan t i n th e shap e o f a large , gentle mound-lik e feature (lin e NPD-VB-17-90 : western end o f geoseismic section in Fig. 5) . The seismic correlation s indicat e tha t thi s 'mound ' contains th e younges t Uppe r Miocene-Lowe r Pliocene sediment s preserve d i n th e V0rin g Basin. Everywher e else in th e V0rin g Basi n this 150 m thick section seems to hav e been removed by erosion , givin g a minimu m estimat e o f missing sectio n beneat h th e Uppe r Pliocen e unconformity. However , a s th e proces s wa s submarine an d o f regiona l exten t thi s als o seems t o b e clos e t o a likel y maximu m o f eroded section . This implie s a maximu m burial depth o f abou t 700-80 0 m i n pre-erosion times . According t o Roaldse t & H e (1995) , th e expected transition temperature in siliceous sediments o f age s les s tha n 2 0 Ma i s abou t 40°C . This woul d impl y a Lat e Miocene-Earl y Plio cene therma l gradient o f a t leas t 50-60°Ckm~ 1 if th e phas e transitio n wa s du e t o highe r ther mal gradient s i n th e past . Th e existenc e o f a past therma l event i s also indicate d b y the hig h thermal gradien t a s calculate d fro m th e dat a i n ODP-643 wel l (Roaldse t & H e 1995) . If suc h a thermal event originated from mantl e processes ,

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H. BREKK E

Fig. 31 . Ma p o f the distributio n o f magmatic rock s i n the Norwegia n Se a continental shelf based o n seismi c interpretation. The distribution of the sill intrusions in the M0re and V0ring Basins is given in terms of both area and stratigraphy .

TECTONIC EVOLUTIO N O F CONTINENTA L MARGI N it mus t hav e bee n associate d wit h a phas e o f thermal uplif t o f th e whol e continental margin. This coul d fi t wit h th e 7 Ma perio d o f regiona l Miocene uplif t describe d earlier , bu t th e tim e since then (10-15 Ma) is far too short to cool the mantle an d crus t b y norma l conductio n t o it s present temperatur e gradien t an d dept h o f subsidence. Thi s call s fo r eithe r a n unknow n mechanism o f extremel y rapi d coolin g o f th e crust and underlying mantle or a mechanism for the heatin g o f th e uppe r part s o f th e crus t an d cover only , i n whic h cas e th e heatin g an d th e uplift ar e cause d b y separate processes . Alternatively, th e 'freezing ' o f th e transi tion surfac e ma y b e du e t o Pliocen e erosion . The cite d age-temperatur e relationshi p fo r th e opal-A-opal-CT transitio n give s a n expecte d burial dept h o f 1100-1200 m a t a norma l ther mal gradien t o f 35°Ckm~ 1 i n the V0rin g Basin, and c . 1000m a t th e presen t 55°Ckm~ l o n the Vorin g Margina l High . Thi s woul d impl y a regional submarin e erosio n o f th e orde r o f 700-500 m i n th e Vorin g Basi n an d 400-50 0 m on th e V0rin g Margina l High , i f th e presen t position o f th e transitio n surfac e i s du e t o Early Pliocene erosion . Rii s & Fjeldskaar (1992 ) suggested tha t th e mechanis m t o initiat e th e coincident majo r erosiona l even t onshore could be the change t o a colder climat e an d th e onse t of glaciation. The consequent increase in erosion rates woul d caus e uplif t o f th e mainlan d an d adjacent offshor e area s b y unloading. However , the mode l doe s no t indicat e tha t thi s uplif t mechanism woul d affec t th e oute r part s o f th e continental margi n a s fa r a s th e V0rin g Mar ginal High, an d the required erosion in that are a must hav e bee n purel y a n effec t o f a chang e i n the submarine currents. To the present author , it seems highl y unlikel y tha t submarin e current s have been responsible fo r removing a 500-700 m sediment column all over such a wide area as the Voring Basi n an d th e Vorin g Margina l High . One woul d expec t suc h submarin e erosio n t o be muc h mor e channelled . Thi s leave s u s with the optio n tha t th e fossi l opal-A-opal-C T surface record s a higher temperature gradien t in Late Miocene-Earl y Pliocen e tim e tha n a t present. Th e autho r find s i t mos t likel y tha t this wa s due t o a n even t of increased hea t flow in Mid-Miocen e times , thoug h cause d b y a n unknown mechanism .

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the sill s generally coincide s wit h th e Fle s Faul t Complex. However , alon g th e Ja n Maye n an d Bivrost Lineament s th e intrusion s ar e see n t o extend furthe r east , eve n int o th e flank s o f th e Tr0ndelag Platfor m (Fig. 31) . In the M0re Basin the genera l easter n limi t o f th e sill s coincide s with th e wester n fault s i n th e broa d M0re Tr0ndelag Faul t Comple x alon g th e easter n flanks of the basin . In bot h basin s the sill s are generally found a t their deepes t positio n an d lowes t stratigraphi c levels (Aptian? ) t o th e east . Fro m thi s positio n they typically step upward s toward s th e west t o their highes t level s (Maastrichtian) clos e t o th e front o f th e 'inne r flows ' (Fig . 31) ; thi s feature indicates tha t th e sill s originate d fro m feede r dykes i n th e centra l part s o f th e basins . I n th e Gjallar Ridg e in the western parts of the V0ring Basin, th e semi-horizonta l intrusion s cu t acros s the eastwar d tilte d beddin g (se e Figs 8 and 31) . These intrusion s in the Gjalla r Ridg e ma y hav e been fe d from dyke s in the Fenri s grabe n t o th e west or from th e base of the Gjallar Ridg e itself. Narrow, vertica l zones of deteriorated seismi c data, which are interpreted a s wrench fractures , are though t t o b e the loci of the feeder dykes to the sil l system . Thes e ar e distribute d ove r larg e parts o f the basin s an d i n man y places coincide with location s o f shift s i n stratigraphi c leve l of the sills, indicating that the y are linked to zone s of dilatio n accommodatin g th e vertica l step s i n the dyke-sil l system . Th e wrenc h movement s have resulte d i n 'bulges' , whic h appea r a s 'ey e structures' i n cross-section (Fig . 32) . The larges t eye structur e i s identifie d i n th e easter n par t o f the M0r e Basin , an d i s believed to b e th e loca tion o f a major feeder dyke to th e sills (Fig. 32) . A ro w o f stron g magneti c anomalie s coincide s with th e locatio n an d strik e of this feature. Such wrench-related eye structures (mainl y in the Paleocen e sequence ) occu r o n a variet y of scales an d th e transcurren t tectoni c movement s on thes e fracture s ar e probabl y minor , bu t clearly affec t al l strat a u p t o th e uppermos t Paleocene-lowermost Eocen e units . This would indicate a Paleocene-Earl y Eocen e ag e fo r th e magmatic activity . There ar e n o seismi c indications tha t th e magm a eve r reache d th e surfac e along these 'eye structures' in Paleocene-Eocen e time. Man y o f th e wrenc h zone s hav e bee n reactivated late r throug h Tertiar y tim e without accompanying magmatism.

Sill intrusions East o f the 'inner flows' the Cretaceous basi n fill is intruded b y numerous sill s of probably latest Paleocene-earliest Eocen e ag e (Fig s 5 , 6, 17 , 22 and 31) . In the V0ring Basin the eastern limit of

Discussion an d plate tectonic relationships The ver y thic k Cretaceou s sedimen t pil e an d the presen t grea t depth s t o th e bas e o f th e

Fig. 32 . Seismi c lin e MB-08-9 2 showing a larg e 'eye structure' typica l o f th e Mor e and Vorin g Basin s (central part o f figure) . Her e 'eye structure' a fleets mainly Paleocene and Eocen e strata , wit h some late r reactivation . Withi n th e 'eye ' th e bas e o f th e Tertiar y sequenc e i s downflcxed, wherea s th e Uppe r Paleocen e an d Eocen e strat a ar e upward bulging . The disturbe d /.one beneath 'ey e structure' i s interpreted t o b e major feede r dyk e t o th e sil l intrusion s seen t o protrud e fro m it . Thin, slanted arrow s poin t at nea r to p Eocene . Small , vertica l arrow s poin t a t nea r to p Paleocene . Thick , vertica l arrow s poin t a t bas e Tertiary . Open , wid e arrow s poin t a t sill s i n th e Cretaceou s section. Whit e triangle s point a t to p o f Eocen e lava s o f th e 'inne r Hows' . (Se e Eig . 3 1 fo r lin e location. )

TECTONIC EVOLUTIO N O F CONTINENTA L MARGI N Cretaceous sequenc e i n th e M0r e an d V0rin g Basins sho w tha t th e crus t beneat h th e basin s was stretche d an d thinne d significantl y ove r a n exceptionally wid e rif t are a i n th e lat e Mid Jurassic-Early Cretaceou s riftin g episode . Th e stretching wa s accompanie d b y norma l faultin g and fault-bloc k rotatio n acros s th e whol e basi n area, bu t withou t an y major, basinwar d bound ary faults. Instead, the basin margin s apparentl y formed b y symmetrica l tiltin g an d flexurin g during Earl y Cretaceou s time , indicatin g a les s brittle behaviou r o f th e crus t tha n seem s usua l for narrowe r rifts . Togethe r wit h th e indica tions o f a hig h elevatio n o f th e centra l basi n areas durin g Mid - an d Lat e Jurassi c time , thes e features poin t toward s a ver y hig h hea t flo w during th e stretchin g phase . N o attemp t ha s been mad e t o mode l ho w thi s mod e o f basi n formation woul d fi t wit h th e model s o f crusta l behaviour base d o n th e recent concep t o f 'depth of necking ' (Cloeting h e t al. 1985 ; Brau n & Beaumont 1989 ; Kooi & Cloetingh 1992) . How ever, suc h model s predic t case s wher e post rifting compressio n woul d ad d t o the subsidence of a basi n dependin g o n th e 'dept h o f necking ' (Kooi & Cloeting h 1992) . The post-rif t Cretac eous histor y o f th e V0rin g an d Mor e Basin s involves subsidenc e t o grea t depth s o f th e bas e of th e Cretaceou s sequence , th e developmen t o f relatively steep , polyphasall y tilte d (flexured ) basin flanks , an d interna l folding , al l o f whic h indicate tha t intermitten t phase s o f intraplat e compression hav e played a role in the subsidence history o f th e tw o basins . Thi s i s discusse d i n the following . The tectoni c histor y o f th e Cretaceou s an d Tertiary period s seem s t o b e closel y linke d t o the plat e tectoni c evolutio n o f Europ e an d th e North Atlanti c (Tabl e 1) . I n Cretaceou s time s the Mor e an d V0rin g Basin s wer e par t o f a system o f continenta l rif t basin s betwee n Nor way an d Greenlan d (Fig . 33) . Befor e break-u p the Ja n Maye n Lineamen t line d u p wit h Kon g Oscar Fjor d i n Greenland , an d th e souther n part o f th e Cretaceou s platfor m o f th e Vorin g Marginal Hig h wa s close to Liverpoo l Lan d an d Jameson Lan d (Fig . 33) . I n th e are a betwee n Liverpool Lan d an d th e Mor e Margina l High , which wa s late r rifte d of f t o becom e th e Jan Maye n Microcontinent , ther e wa s roo m for a Cretaceou s basin . However , recen t dat a from th e Ja n Maye n Microcontinen t indicat e only a moderatel y thic k Mesozoi c sectio n ther e (Kuvaas & Kodair a 1997) . Thes e reconstruc tions are taken to imply the existence of a central platform are a betwee n Greenlan d an d Norwa y throughout th e Cretaceou s perio d (Brekk e et al. 1999). Svalbard was situated nort h o f Greenlan d

369

and th e proto-Senj a Fractur e Zon e wa s prob ably alread y establishe d a s a crusta l tectoni c lineament (Brekk e & Rii s 1987). The tectoni c pulse s i n Lat e Cretaceou s time s (post-Cenomanian-Paleocene) wer e probabl y coupled t o th e openin g o f th e Labrado r Se a and th e associate d anti-clockwis e rotatio n o f Greenland tha t wa s initiate d i n Cenomania n time (Srivastav a & Tapscott , 1986 ) (Tabl e 1) . In th e V0rin g Basi n thi s wa s reflecte d a s tilting of basi n margins , increase d subsidence , faultin g and regiona l folding . Th e differen t structura l expressions indicat e tha t th e intra-plat e stres s regime, se t u p b y th e rotatio n o f Greenland , changed throug h tim e relativ e t o th e Vorin g Basin. I t i s possible tha t th e end-Cenomania n tilting o f th e basi n margin s an d th e increas e i n basin subsidenc e wer e due t o a n initia l phase o f weak intraplat e compression , inducin g furthe r crustal flexurin g (Cloeting h e t al . 1985 ; Kooi & Cloetingh 1992) . Faulting o n th e basi n margi n (Revfalle t an d Bremstein Faul t Complexes ) an d associate d growth o f th e Campania n sequenc e indicat e a change t o a n extensiona l regime that laste d int o latest Cretaceou s (Maastrichtian? ) tim e (Fig s 1 , 5, 6 and 17) . The foldin g o f th e Vorin g Basi n i n lates t Cretaceous tim e involve d fold s o f wavelength s much les s tha n th e widt h o f basin , an d apparently di d no t contribut e significantl y t o the overal l deepening o f the basin. However , th e large Vigri d an d Nagrin d Syncline s constitute d local depocentre s an d site s of furthe r deepenin g by downwar p durin g thi s phase . Th e chang e i n the scal e o f foldin g ma y hav e bee n du e t o th e reactivation o f the Fle s Faul t Comple x creatin g a regional anisotropy alon g the basin axis, which controlled th e distributio n o f folds . Excep t fo r the tiltin g an d steepenin g o f th e basi n flank s there ar e n o sign s o f thi s Lat e Cretaceou s tectonics i n th e Mor e Basin , which implie s tha t the Jan Maye n Lineamen t mus t hav e acte d a s a regional transfe r zon e durin g th e whol e o f Lat e Cretaceous time . Th e Maastrichtia n fol d axe s in th e V0rin g Basi n ar e generall y NE-SW , indicating a n orientatio n o f th e principa l com pressional stres s axi s norma l t o th e basi n axis . This implie s tha t th e Ja n Maye n Lineamen t a t that tim e mus t hav e acte d a s a dextra l transfer zone, wherea s th e southern , N- S trendin g seg ment o f th e Fle s Faul t Comple x acte d a s a conjugate sinistra l zone . Th e norther n segmen t of the Fle s Fault Comple x i s parallel t o th e fol d axes, implyin g compression an d revers e move ments o n reactivation . Tectonic activit y an d uplif t ar e als o eviden t north o f th e Bivros t Lineamen t an d i n th e

Table 1 . A regional comparison o f plate tectonics with relevance t o th e evolution o f th e Møre an d Vørin g Basins MORE BASI N LANDWARD UPLIFT AND WESTWARD TILTING

5 RAPID SUBSIDENCE

5 MIO

UPLIFT, EROSIO N ALONG FLANKS ONLY

6

7

GLACIATIONS, UPLIFT AND EROSION OFCIRCUM-BASINALLAN D AREAS SUBSIDENCE O F NORWEGIAN SEA AND ICELAND-F/EROE RIDGE PRONOUNCED CHANG E IN OCEANIC CIRCULATIO N ANDCLAIMATE •ACTIVE SPREADIN G O N PRESEN T KOLBEINSE Y RIDGE ESTABLISHED JAN MAYE N RIDGE SEPARATES FROM GREENLAND
TRANSPRESSION FORMING DOMES AND FLEXURES (ALONG JAN MAYEN LINEAM, ONLY)

36

TRANSPRESSION REACTIVATING V0RING ESCARPMENT FLES FAULTZONE, JAN MAYEN LINEAMENT AND NORDLAND RIDGE. GROWTH OF DOMES AND FLEXURES

EO

FAULTING ALONG F/EROE-SHETLAND ESCARPMENT

55

VOLCANISM ON M0RE MARGINAL HIGH

PAL

66 MAAST

FAULTING ALONG V0RING ESCARPMENT AND NYK HIG H VOLCANISM ON V0RING MARGINAL HIGH

SVALBARD SEPARATES FROM GREENLAND. EXTINCTIO N OF SEA-FLOOR SPREADING IN LABRADOR SEA AND BAFFIN BAY . GREENLAND PART OF NORTH AMERICA. SEA-FLOO R SPREADING CONTINUE S IN NORWEGIAN-GREENLAN D SE A

FORMATION OF ICELAND-F/EROE RIDGE, STRIKE-SLIP BETWEE N SVALBARD AND GREENLAND. CHANGE IN DIRCTION OF MOTION OF GREENLAND. (7POSSIBLE INITIATION OF SPREADING TO THE WEST OF THE XEGI R RIDGE? ) ACTIVE SEA-FLOOR SPREADING IN NORWEGIAN-GREENLAND SE A AND DAVIS STRAIT. CHANGE IN DIRECTIONOF MOTION OF GREENLAND. STRIKE-SLIP BETWEEN SVALBARD AND GREENLAND

MINOR FAULTING

POSSIBLE GRABEN FORMATION ALONG V0RING ESCARPMENT ELSEWHERE ONLY MINOR FAULTING

VOLCANISM IN DAVIS STRAIT, EASTERN GREENLAND, V0RING PLATEAU AND F/EROE ISLANDS. CHANGE IN DIRECTION OF MOTION BETWEEN GREENLAND AND NORTH AMERICA. STRIKESLIP BETWEEN SVALBARD AND GREENLAND

STRIKE-SLIP REACTIVATION (ALONG JAN MAYEN LINEAM. ONLY)

UPLIFT AND EROSION OF WHOLE BASIN AREA SYNCLINES AND ANTICLINES FORMED DUE TO TRANSPRESSION ALONG MAIN STRUCTURE S

SEAFLOOR SPREADIN G I N NORTHER N LABRADO R SE A RIFTING I N NORWEGIAN-GREENLAND SE A

71

MAIN COMPRESSION IN JURA MOUNTAINS

"NEOGENEOROGENr MAIN COMPRESSIONAL PHAS E RELATED TO TERMINATION OF NORTHERN SUBDUCTIO N (HELVETIC NAPPES)

83'

MAIN CRETACEOUS RIFTING BLOCK FAULTING ALONG NYK HIGH, GJALLAR ' RIDGE,REACTIVATION OF FLES FAULTZONE

SflNTON TILTING OFTR0NDELAG PLATFORM AND GJALLAR RIDGE

CONIAC

91

TURON

CEN

"PALEOGENE OROGENY" MAIN COMPRESSIONAL PHASE RELATED TO TERMINATION OF SOUTHERN SUBDUCTIO N (WESTERN, CENTRAL & EASTERN ALPS)

COMPRESSION (ULTRAHELVETIA , VALAIS, BRIANC.)

MINOR COMPRESSION ? (VALAIS, BRIANC.)

COMPRESSION (EASTERN & CENTRAL & ALPS)

SUBSIDENCE CAMP

34

ALPS

i SPREADING IN NORWEGIAN SEA (/EGIR RIDGE) TERMINATES

21

31

RAPID SUBSIDENCE AN D WESTWARD TILTING . MINO R EROSION OF HIGHS TRANSPRESSIONAL DOMING ALONG JAN MAYEN AND SURT LINEAMENT, FLES FAULT ZONE, V0RING ESCARPMENT AND NORDLAN D RIDG E UPLIFT AND EROSION

OL

17

24

LANDWARD UPLIFT AND WESTWARD TILTING MINOR FAULTING ON NYK HIGH

NORWEGIAN-GREENLAND SEA AND RELATED AREAS

24

10 13

TRANSPRESSIONAL DOMING ALONG JAN MAYEN LINEAMENT

VORING BASI N

ACTIVE SEA-FLOOR SPREADIN G I N SOUTHERN LABRADO R SE A RIFTING I N NORTHERN LABRADOR SEA AND NORWEGIANGREENLAND SEA . COUNTER-CLOCKWISE ROTATIO N O F GREENLAND MAIN CRETACEOUS COMPRESSIONAL PHASE (EASTERN & CENTRAL ALPS) RIFTING BETWEE N GREENLAND AND LABRADO R AN D RELATED VOLCANISM

97 ALB SUBSIDENCE

109

COMPRESSION (EASTERN ALPS, VALAIS)

APT

40

MINOR FAULTING ASSOCIATED WITH FLES FAULT ZONE?

MINOR FAULTING ON FLES FAULTZONE? SUBSIDENCE

INITIAL, RIFT RELATED VOLCANISM ON LABRADOR SHELF RIFTING I N ROCKALL TROUGH

BAR 122-

GRAVITATIONAL SLIDING ON MARGINS

140

LATE JURASSIC

MAIN RIFTING PHASE, BLOCK FAULTING "BASIN MARGIN FLEXURE

GRAVITATIONAL SLIDING ON MARGINS MAIN RIFTING PHASE. BLOCK FAULTING BASIN MARGIN FLEXURE

Based o n o f Trumpy (1973) , Taiwan! & Eldholm (1977), Eldhol m & Thiede (1980), Srivastava & Tapscott (1986) , and Dewe y & Windley (1988).

TECTONIC EVOLUTIO N O F CONTINENTA L MARGI N

371

Fig. 33 . Palinspasti c reconstruction o f the Norwegian Se a continental shel f at anomaly 24 time (Ypresian). Plat e reconstruction take n fro m Srivastav a & Tapscott (1986) . Positions o f th e futur e intermediat e spreadin g ridge s compiled fro m Talwan i & Eldholm (1977 ) an d Nunn s (1983) . Other detail s o f map base d o n author' s ow n mapping an d Talwan i & Eldholm (1977) , Larse n (1984) , Stewar t e t al. (1992) and Whit e (1992) .

372

H. BREKK E

TECTONIC EVOLUTIO N O F CONTINENTA L MARGI N Barents Se a (Brekke & Riis 1987) . This indicate s that th e Lat e Cretaceou s stres s regim e reacti vated a syste m controlle d b y th e Ja n Maye n Lineament, th e Fle s Faul t Comple x an d th e proto-Senja Fractur e Zone . Associate d struc tures suc h a s th e Bivros t an d Sur t Lineament s were integra l part s o f th e reactivatio n system . To lin k thi s syste m t o th e openin g o f th e Labrador Sea , an d t o accommodat e th e differ ence betwee n th e quie t M0r e Basi n an d th e tectonically very active V0ring Basin, there must have bee n a tectonicall y active bel t wes t o f th e More High , perhap s alon g th e futur e break-u p line an d futur e JEgir Ridg e (Fig s 3 3 and 34) . The regiona l uplif t an d erosio n o f th e continental margi n a t th e beginnin g o f Tertiar y time wa s du e t o th e increase d hea t flo w o f th e area associate d wit h th e intens e riftin g centre d west o f th e margina l highs , whic h le d t o th e continental separatio n betwee n Norwa y an d Greenland i n lates t Paleocene-earlies t Eocen e time (Skogsei d & Eldhol m 1989) . Th e sever e extension an d riftin g als o induce d widesprea d intrusion activit y alon g th e axi s an d i n west ern part s o f th e V0rin g an d Mor e Basins . Although associated wit h th e ol d zone s o f weakness, th e intrusio n activit y wa s accompa nied b y minor, probabl y transtensional, tectonic movements, whic h wer e widesprea d an d evenl y distributed throughou t th e basin s wes t o f th e M0re-Tr0ndelag an d Fle s Faul t Complexes . This styl e o f tectoni c deformatio n an d th e volume o f intrusiv e rock s indicat e tha t a t thi s stage th e crus t beneat h th e basin s was generally weakened throug h heating . However, i t was no t sufficiently weakene d t o b e th e locatio n o f th e continental break-up . Thi s ma y indicat e tha t the Jurassic-Cretaceou s basin-formin g proces s itself ha d change d th e physica l properties o f th e underlying crus t t o mak e i t mor e resistan t t o rupture tha n th e crust beneat h th e western basi n bounding platforms (i.e . the present Vorin g an d M0re Margina l Highs) . The continenta l break-u p wa s accompanie d by enormou s volume s o f floo d basalts , an d th e JEgir an d Mohn s Spreadin g Ridge s wer e estab lished (Talwan i & Eldhol m 1977 ; Whit e 1992 ) (Table 1 Fig. 33) . Th e regiona l uplif t i n Paleo cene t o earlies t Eocen e times , jus t befor e th e break-up, affecte d area s eve n fa r awa y (900 1000km) fro m th e centr e o f th e upwellin g ho t

373

mantle material underneat h th e future spreadin g ridges. Thes e distan t area s (e.g . th e Tr0ndela g Platform an d th e Nort h Se a an d surroundin g mainland) experienced a n extremel y rapi d sub sidence in early Eocene times . Nadin & Kusznir (1995) explained thi s b y a dynami c uplif t o f th e crust b y latera l flui d flo w i n th e col d astheno sphere, which was induced by the dynamic flowfield set up by the upwelling hot mantl e material in th e distan t upwellin g centre befor e break-up . Subsequent t o break-u p th e mantl e flo w wa s focused o n th e ridg e axis , divertin g col d asthenosphere bac k fro m th e dista l part s o f th e uplifted region , an d thu s explainin g th e rapi d subsidence i n thos e areas . Up t o Lat e Eocen e time , there seems to b e no obvious lin k betwee n th e tectonics o f th e Norwegian Se a an d th e compressiona l phase s in the Alps (Table 1) . But from Lat e Eocene time onwards th e mai n orogeni c phase s i n th e Alp s coincide wit h the timing of events in the Norwe gian Sea . Thi s ma y b e du e t o th e fac t tha t th e Norwegian Se a was now established a s the plate boundary betwee n th e America n an d Eurasia n plates, an d tha t th e relative motions betwee n th e two wer e change d becaus e o f th e continuin g collision between the Eurasian and Gondwania n plates. I n particular , th e subduction-relate d col lisions o f the Alp s i n Lat e Eocen e an d Miocen e time cause d change s i n th e relativ e plat e motions, whic h wer e expresse d a s tectoni c adjustments a t th e ne w plat e boundary . A lin k between thes e event s an d th e Alpin e Orogen y and associate d plat e reorganizatio n ha s bee n suggested als o b y Cloeting h e t al. (1990 ) an d Gradstein & Backstrom (1996) . A s documente d by Gradstei n & Backstro m (1996) , th e Lat e Eocene-Early Oligocen e phas e o f flexurin g an d doming nort h o f th e Ja n Maye n Lineamen t coincides wit h a regiona l hiatu s o n th e easter n basin margin s acros s majo r part s o f th e Nort h Sea an d o n th e Halte n Terrace . Thi s Lat e Eocene Alpine orogenic phase may be correlated with th e en d o f th e sea-floo r spreadin g i n th e Labrador Se a and th e final passage o f Svalbar d past norther n Greenland, event s that lef t Green land a s a n integra l par t o f th e America n plat e (Eldholm & Thiede 1980) . This als o correspond s in tim e t o th e chang e i n azimut h o f th e Ja n Mayen Fractur e Zon e fro m NN W (easter n branch) t o N W (wester n branch ) (Fig . 34) ,

Fig. 34 . Palinspasti c reconstruction o f the Norwegia n Se a continental shel f a t anomal y 5 time (Late Miocene) . Plate reconstruction taken from Srivastav a & Tapscott (1986). Positions of the intermediate spreading ridges and the Ja n Maye n Microcontinen t compile d fro m Talwan i & Eldholm (1977 ) an d Nunn s (1983) . Tertiar y dome s around Faero e Islands based o n anticlina l axe s of Boldree l & Andersen (1993) . Other details of map base d o n author's ow n mappin g and Talwan i & Eldholm (1977) , Larse n (1984) , Stewar t e t al . (1992) an d Whit e (1992) .

374

H. BREKK E

implying a shif t i n th e relativ e plat e motion s (Talwani & Eldholm 1977) . At th e en d o f Oligocen e tim e th e sea-floo r spreading alon g th e JEgir Ridg e slowe d dow n and stoppe d (betwee n anomalie s 2 0 an d 7) . According t o Eldhol m & Thied e (1980) , th e present Kolbeinse y Ridg e wa s establishe d onl y just befor e anomal y 5 (1 0 Ma). I n th e perio d between anomal y 7 (26 Ma) an d anomal y 5 (i.e. Early t o lat e Mid-Miocen e time ) ther e wer e n o distinct spreadin g axi s sout h o f th e Ja n Maye n Fracture Zon e t o replac e th e ^Egi r Ridg e (Talwani & Eldhol m 1977 ; Eldhol m & Thied e 1980). However, Talwani & Eldholm (1977 ) suggested the existence of an intermediate spreading axis formin g a loca l oceani c basi n S W o f Ja n Mayen i n th e interva l betwee n anomalie s 6 A (21 Ma) an d 5 D (1 8 Ma), brackete d b y inter mittent period s o f n o spreading . Vog t e t al. (1980) postulate d spreadin g o n th e Kolbeinse y Ridge sinc e anomal y 6 C (23. 5 Ma), an d Nunn s (1983) proposed a model wit h conjugate wedges of sea-floo r spreadin g o n bot h side s o f th e Ja n Mayen Microcontinen t i n th e interva l betwee n anomalies 20 and 7 . However, the magnetic data shown b y Nunn s (1983 ) d o no t documen t an y well-defined symmetr y alon g th e Kolbeinse y Ridge befor e anomal y 5 B (1 5 Ma), no r an y anomaly olde r tha n anomal y 7 (2 6 Ma) wes t of the Jan Maye n Microcontinent. Betwee n what is interpreted b y Nunns a s anomalies 5 B and 7 the pattern seem s t o b e mor e i n lin e wit h a n intermediate axi s SS W o f Ja n Maye n a s proposed b y Talwan i & Eldhol m (1977) . I n thei r model th e intermediat e ridg e becam e extinc t along thei r anomal y 5D , whic h correspond s t o anomaly 6 of Nunns (1983) . The implicatio n o f this model is a local basin with a set of anomalies with a symmetrica l age distribution aroun d tha t central anomaly. Th e identification of anomalie s 6 an d 6 B by Nunn s (1983 ) further SW , close t o the present mainland of Greenland, may then be explained b y th e postulatio n (b y th e author ) of another , coinciden t intermediat e spreadin g ridge segmen t i n tha t area . Thi s mode l implie s the existenc e o f tw o e n echelo n loca l oceani c spreading basin s wes t o f th e Ja n Maye n Ridg e in th e perio d betwee n anomalie s 6 A an d 5 D (as correlate d b y Talwan i & Eldhol m (1977) ) (Fig. 34) . The offse t betwee n the two segments of the intermediat e ridge was taken u p b y a transverse zon e paralle l t o th e Ja n Maye n Fractur e Zone (Fig . 34) . Thi s offse t wa s late r cu t b y th e Kolbeinsey Ridge , whic h wa s establishe d b y a new ridg e shift sometim e between anomalie s 5 B and 5 (Fig. 34). The fan-shape d anomal y patter n aroun d th e Agir Ridge wa s explaine d b y Nunn s (1983 ) by

an anti-clockwis e rotatio n o f th e Ja n Maye n Microcontinent. Thi s seem s to b e impossible, a s the microcontinen t wa s separate d fro m Green land onl y subsequen t t o th e extinctio n o f th e ^Egir Ridge . Suc h a rotation woul d have to tak e place between anomalie s 2 2 and 7 to explain th e fan. A more likely explanation is that tha t was a period o f riftin g an d sever e stretchin g an d extension of the continental crus t in the souther n part o f th e futur e Ja n Maye n Microcontinent , which compensate d fo r th e fa n geometr y i n th e ocean crust . Th e mode l wit h tw o e n echelo n intermediate spreadin g ridge s implie s a substan tial siz e o f th e souther n par t o f th e Ja n Maye n Microcontinent (Fig . 34) , whic h woul d fit with such a n earl y stretchin g episode . Thi s is also i n agreement wit h the mappin g o f th e microconti nent based on seismic reflection dat a (Gudlaugs son et al. 1988) : the present wate r dept h an d th e intensity o f past block-faultin g increas e consid erably southward s o n th e microcontinent , indi cating increase d thinnin g an d stretchin g o f the crus t i n tha t direction . Anothe r implicatio n of suc h a mode l i s tha t Icelan d ma y b e under lain b y continenta l crus t o r fragment s o f suc h crust (Fig . 34) . From thi s discussion of the literatur e it seems safe t o conclud e tha t th e interva l betwee n anomalies 7 an d 5 B wa s on e o f slo w o r n o spreading a t th e ^gi r Ridg e an d a comple x pattern o f temporar y spreadin g centre s wes t o f Jan Mayen . Th e pictur e i s complicate d b y th e uncertain timin g of the riftin g eas t o f Liverpoo l Land an d th e subsequen t separatio n o f th e Ja n Mayen Microcontinen t (Gudlaugsso n e t al . 1988). I t ma y b e possibl e tha t th e riftin g i n th e continental crust compensated fo r periods o f n o spreading i n th e interva l between the extinction of the JEgir Ridg e and th e final establishment of the Kolbeinse y Ridge . This reorganizatio n o f th e spreadin g syste m coincided wit h th e onse t o f th e Neogen e Orog eny o f th e Alp s (Tabl e 1 ; Trumpy 1973 ; Dewey 6 Windley 1988) and the Mid-Miocene phase of compression i n th e V0rin g Basin . This indicate that th e Alpin e Neogen e Orogen y probabl y affected th e relativ e plat e motion s betwee n th e Eurasian an d America n plates . Thi s even t wa s also accompanie d wit h th e earlie r mentione d 7 Ma perio d o f regiona l uplift , erosion , nondeposition and subsequen t rapid subsidence seen all ove r th e Norwegia n Continenta l Shel f an d adjacent mainland . I n th e absenc e o f evidenc e of a regiona l hea t pulse , th e regiona l rapi d uplift an d subsidenc e ma y b e explaine d b y th e dynamic flui d mantl e flo w mode l o f Nadi n & Kusznir (1995) . After th e extinction of the ^gir Ridge, th e centr e o f mantl e upwellin g shifte d

TECTONIC EVOLUTIO N O F CONTINENTA L MARGI N

to the area beneat h the future Kolbeinse y Ridge. The spreadin g activit y betwee n th e Denmar k Strait Fractur e Zon e an d Ja n Maye n Fractur e Zone i n th e intermitten t perio d (Earl y t o Lat e Miocene time ) wa s slo w an d discontinuou s (see above), an d may have induce d centra l uplift just a s i n th e situatio n immediatel y befor e the continenta l break-u p i n Eocen e time . Thi s may i n tur n hav e induce d a dynami c latera l flow i n th e col d asthenospher e outward s fro m the upwellin g centre , resultin g i n uplif t i n th e distant flanks : th e V0rin g an d M0r e Basin s and th e margina l highs . Th e subsequen t rapi d subsidence followe d fro m th e fina l establish ment o f th e Kolbeinse y Ridg e a t th e en d o f th e overall plat e motio n adjustmen t i n lat e Mid Miocene time . A phas e o f stron g an d widesprea d Miocen e compression wa s als o documente d aroun d th e Faeroe Islands , b y Boldree l & Andersen (1993 , 1994) an d Anderse n & Boldree l (1995) . Thos e workers have als o identifie d phases o f compression i n Early Eocen e an d Oligocen e time , which fit wit h th e timin g o f tectoni c activit y i n th e V0ring Basin. As there is no evidence of Tertiary compressional tectonic s i n th e M0r e Basin , th e Jan Maye n Fractur e Zon e an d th e Ja n Maye n Lineament mus t hav e acte d a s a transfe r zon e between the Faeroe Islands are a an d th e V0ring Basin durin g Tertiar y tim e (as also inferre d for Late Cretaceou s time) . After lat e Mid-Miocene time the whole region again starte d t o subsid e an d th e tectonicall y created ocean-botto m topograph y wa s filled in until the period o f non-deposition i n late Earlyearly Lat e Pliocen e time . Th e regiona l Lat e Pliocene hiatu s o r unconformit y ha s bee n wel l documented b y Eidvin & Riis (1989, 1991, 1992 ) and Eidvi n et al. (1993). A phase o f considerable uplift an d dee p erosio n o f the mainlan d starte d in Late Pliocene time, resulting in the accumulation an d progradatio n o f a hug e sedimentar y wedge al l alon g th e subsidin g continenta l shel f (Riis & Fjeldskaar 1992) . However, in the mea n time th e therma l gradien t i n th e shel f ha s continued t o decrease, and the overburden accumulated sinc e Lat e Pliocen e tim e ha s stil l no t brought th e fossi l opa l transitio n leve l bac k down t o th e temperatur e o f its formation. The mechanis m behin d th e stron g Pliocen e uplift o f th e mainlan d an d th e symmetrica l subsidence of the continenta l shelf is still poorly understood (se e discussio n b y Va n de r Bee k (1995) an d Rii s (1996)) . The autho r i s gratefu l t o Nors k Hydro , Saga Petroleum, Statoi l and th e Norwegia n Petroleum Directorate for granting permission to publis h critical seismi c

375

data. L . O. Boldree l an d T . Torske are acknowledge d for thei r thorough review of the manuscript.

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ONGSTAD, P . & SPINNAGR , A . (eds ) Petroleum Geology o f th e North European Margin. Graha m and Trotman , London , 303-315. TALWANI, M . & ELDHOLM , O . 1972 . The continenta l margin of f Norway : a geophysica l study . Geological Society o f America Bulletin, 83 , 3375-3608 . & 1977 . Evolutio n o f th e Norwegian Greenland Sea . Geological Society o f America Bulletin, 88, 969-999. , MUTTER , J . C . & HINZ , K . 1983 . Ocean continent boundar y unde r th e Norwegia n con tinental margin . In: BOTT , M. H. P., SAXOV , S. , TALWANI, M . & THIEDE , J . (eds ) Structure an d Development of the Greenland-Scotland Ridge New Methods and Concepts . Plenum , New York, 121-131. , UDINTSEV , G . e t al . 1976 . Initial Reports o f th e Deep Se a Drilling Project, 38 , U S Governmen t Printing Office , Washington , DC . TORSKE, T . & PRESTVIK , T . 1991 . Mesozoi c detach ment faultin g betwee n Greenlan d an d Norway : inferences fro m Ja n Maye n Fracture Zon e syste m and associate d alkalic volcanic rocks. Geology, 19, 481-484. TRUMPY, R . 1973 . The timin g of orogenic events in th e Central Alps . In : D E JONG , K . A . & SCHOLTEN , R. S . (eds ) Gravity an d Tectonics. Wiley , Ne w York, 229-252 . VAN DE R BEEK , P . A . 1995 . Mechanism s o f Neo gene tectoni c uplif t i n souther n Norway . In : Tectonic evolution of continental rifts - inferences from numerical modelling and fission track thermochronology. Ph D thesis , Researc h Schoo l o f Sedimentary Geolog y (NSG) , Amsterdam . VOGT, P . R. , JOHNSON , G . L . & KRISTJANSSON , L . 1980. Morpholog y and magneti c anomalies nort h of Iceland . Journal o f Geophysics, 47, 67-80 . WHITE, R . S . 1992 . Crusta l structur e an d magma tism o f Nort h Atlanti c continenta l margins . Journal o f th e Geological Society, London, 149 , 841-854.

Late Cretaceou s an d Tertiary structura l evolutio n of th e northeastern part o f the Vering Basin , Norwegia n Se a TOMMY EGEBJER G MOGENSEN , RUN E NYBY , RIDVA N KARPU Z & PA L HAREM O

Norsk Hydro a.s. Exploration, Harstad, P.O. Box 31, N-9401 Harstad, Norway (e-mail: Tommy.Mogensen.Egebjerg @hydro.com) Abstract: Th e Late Cretaceous-Tertiary structural evolution of the northeastern par t of the V0ring Basin , mid-Norway , i s highl y complicated . Althoug h tectoni c activit y occurre d throughout Cretaceou s tim e in much o f the V0ring Basin , including the Gjalla r Ridg e an d along th e Fle s Faul t Zone , i n th e Vem a Dome-Ny k Hig h are a evidence o f suc h activit y is not observe d unti l the latest Maastrichtian time . In the Vema Dome-Nyk High area, several faults wit h both a NW-SE orientatio n an d a NE-SW orientatio n hav e experience d latera l movements. Th e NE-SW orientatio n i s the ol d Caledonia n trend . Th e comple x structura l evolution o f th e Vem a Dome-Ny k Hig h are a i s probabl y relate d t o th e existenc e o f continental weaknes s zones , whic h ar e th e onshor e extensio n o f know n oceani c fractur e zones. Th e two lineaments tha t delineate thes e weakness zones , and that delineate the Vema Dome-Nyk Hig h area , namel y the Bivros t and Sur t Lineaments , diverg e slightly from on e another towar d th e SE . This comple x structura l framework of NW-SE oriented lineaments and NE-SW oriented deep-seated basemen t structure s of Caledonian compressiona l and/or Mesozoic rif t origin , facilitate d mino r clockwis e rotatio n o f th e are a betwee n th e tw o lineaments during the extensional regime before the break-up o f the North Atlantic , as well as during the compressiona l regim e that ha s bee n propose d afte r th e break-up .

Several regiona l tectoni c studie s hav e bee n car ried ou t aroun d th e V0rin g Basi n o n th e mid Norway continenta l margi n (Fig . 1 ) (Brekke & Riis 1987; Dore& Gage, 1987 ; Planked al. 1991; Skogseid e t al . 1992 ; Hinz e t al . 1993 ; Skogsei d 1994). Mor e detaile d loca l structura l studie s o f the are a hav e bee n hampere d b y th e lac k o f closely space d reflectio n seismi c surveys. Befor e the Norwegia n 15t h licensin g round , a larg e number o f reflectio n seismi c data wer e acquired in mos t o f th e V0rin g Basin , whic h mad e i t possible t o carr y ou t a mor e detaile d structura l analysis in several key tectonic areas, such as the Vema Dome-Ny k Hig h are a (Fig . 2 ) (Blystad et al . 1995) . Several paper s hav e deal t wit h th e structural evolution of this area (Bj0rnset h et al. 1997; Robert s e t al . 1997 ; Walker e t al . 1997 ) and som e studies , suc h a s tha t b y Lundi n & Dore (1997 ) have focuse d on th e structura l role of the lineaments seen in the Norwegian Sea. All studies agre e o n a complicate d tectoni c evolu tion of the Vema Dome-Nyk High area and that the structura l evolutio n varie s i n severa l way s from tha t o f th e res t o f th e V0rin g Basin . Thi s paper addresse s th e structural setting before and after break-u p i n and aroun d th e Vema DomeNyk High , wit h it s basi s i n depocentr e config uration and fault orientations , and the structural influence b y th e Sur t an d Bivros t Lineaments.

Late Cretaceous-Paleocen e structura l evolution (pre-break-up ) Most o f th e structura l element s i n th e V0rin g Basin, e.g . th e Gjalla r Ridg e an d Fle s Faul t Zone, seem to be controlled by fault systems that have been technicall y reactivate d throughou t most of Cretaceous tim e (Fig. 3) . During most of Cretaceous time, however, the Vema Dome-Nyk High are a wer e i n a basin-floo r position , a s shown by the thick parallel Cretaceous sequence s across th e are a (Fig . 4) , simila r t o th e deposi tional settin g o f th e nearb y Vigri d Synclin e (Fig. 3) . Minor faulting and block rotation could have taken plac e i n the Vem a Dome-Nyk Hig h area in Early Cretaceous time, but i t is difficult t o differentiate th e deepes t sequence s becaus e o f limited seismic resolution at depth. The first clear evidence o f faul t activit y i n th e Vem a Dome Nyk Hig h are a i s recorded i n th e lates t par t o f Maastrichtian o r Earl y Paleocen e time , whe n a fault-related roll-ove r structur e wa s erode d a t the crest of the roll-over (Fig. 5 at A). From that time, thi s slowl y emerge d a s a resul t o f tec tonic activity, especially in the are a betwee n the present Vem a Dom e an d Ny k Hig h (Fig . 6) . Most o f th e Ny k Hig h are a continue d t o b e a site of deposition throughou t mos t o f the Paleo cene perio d (Fig . 6) , an d onl y i n th e lates t

From: N0TTVEDT , A . e t al . (eds ) Dynamics o f th e Norwegian Margin. Geologica l Society , London , Specia l Publications, 167 , 379-396 . l-86239-056-8/00/$15.0 0 © Th e Geologica l Societ y o f London 2000 .

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Fig. 1 . Locatio n o f th e Sur t an d Bivors t Lineament s i n th e Norther n par t o f th e V0rin g Basin , offshor e Mid Norway .

Paleocene o r earl y Eocen e tim e di d thi s are a become structurall y hig h a s a resul t o f tectoni c activity (Fig . 7 at A) . The Bas e Cretaceous is deeply burie d i n mos t of th e V0rin g Basin , an d faul t activit y i n pre Cretaceous tim e i s difficul t t o identify . Close r to th e Norwegia n mainland , o n th e Tr0ndela g Platform, wher e th e Jurassi c sequenc e i s wel l imaged, th e mai n Lat e Jurassi c fault orientatio n as wel l a s th e ol d Caledonia n structura l grai n is oriente d i n a NE-S W direction . W e there fore assum e tha t th e observe d NE-S W oriente d

extensional faultin g i n Lat e Maastrichtian Paleocene tim e too k plac e b y reactivatio n o f older NE-SW oriented Mesozoi c and/or Palaeo zoic faults , becaus e o f increase d extensiona l forces befor e th e Nort h Atlanti c break-u p i n this are a (se e als o Dor e e l al. 1997) . From Intra Paleocen e t o the Near Top Paleo cene tim e faultin g wa s intensifie d an d severa l transtensional features evolved (Fig. 8). Evidence of transtensio n i s the apparen t lac k o f extension at Nea r To p Cenomania n level , wherea s ther e has been clear extension at shallowe r Cretaceous

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Fig. 2. Structura l element map o f the northern part of the V0ring Basin, showing location of seismic sections and geosections, i n Fig s 3 , 5 , 8 , 1 0 and 19 .

levels (Fig . 8 a t A) . Depocentr e configura tion alon g fault s tha t hav e experience d latera l movements ca n b e use d a s a n indicato r o f th e orientation o f th e latera l movement , wit h increasing subsidenc e i n releasin g bend s an d less subsidenc e o r erosio n i n restrainin g bend s (Mogensen 1994) . Usin g thi s metho d th e depo centre configuratio n along th e NE-SW trending transtensional feature s in th e Vema Dome-Ny k High are a indicate s a sinistra l faul t movement , from Intr a Paleocene t o the Near To p Paleocen e time (Fig. 7 at B). Another indicatio n o f sinistral movement i s the configuratio n o f th e isochron s across a stee p fault , whic h w e assume migh t b e lateral (Fig . 6 at A). The non-eroded par t o f the uppermost Maastrichtia n t o Lowe r Paleocen e

sequence show s clearl y a n 'apparent ' sinistra l movement afte r th e depositio n o f th e sequenc e of the orde r o f c. 2km. There i s no structura l o r stratigraphic indication that thi s fault wa s active in Lates t Maastrichtia n o r Earl y Paleocen e times, s o th e 2k m o f sinistra l movemen t pos t Intra Paleocen e migh t b e a goo d estimat e o f sense an d amoun t o f displacement . Contemporaneous wit h the lateral movements along th e NE-S W trendin g faul t system , a relatively thick depocentre evolved in the western part o f th e Vem a Dom e alon g th e NW-S E oriented Ry m Faul t Zone-Sur t Lineament , where this structura l element meets the NE-SW oriented faul t syste m (Fig . 7 a t C , an d Fig . 9 at A) . Th e positio n an d configuratio n o f thi s

Fig. 3 . Geosectio n acros s the norther n par t o f th e V0ring Basi n fro m th e V0rin g Escarpmen t t o th e D0n n Terrace, showin g repeate d tectoni c activit y o n th e Gjalla r Ridg e an d the Fles Fault Zon e (thinning) throughout mos t of Cretaceous time. The tota l gravit y as well as the residual gravity field for the section i s also presented. (Fo r location , se e Fig. 2. )

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Fig. 4. Thic k parallel Upper Cretaceous sequences on top o f the Nyk High before Lat e Maastrichtian time, indicating that this area was in a basin-floor position in early Late Cretaceous time. This section is the same as in Fig. 8 flattened on th e Near Top Santonia n surface. Location , as for Fig . 8 is shown in Fig. 2.

depocentre compare d wit h th e configuratio n of the NW-SE trendin g faul t indicate s a releasin g bend depocentr e an d sinistra l faul t movement . The fault s transectin g th e forme r Vem a Dom e with an E-W orientatio n (Fig. 9) are interpreted as transfe r fault s betwee n th e NE-S W an d th e NW-SE faul t system . Influence o f bounding crusta l weakness zones The differen t structura l evolutio n o f th e Vem a Dome-Nyk Hig h compare d wit h mos t o f th e rest of the V0ring Basin seems to be governed by two lineaments or zones that delineate the Vema Dome-Nyk Hig h area , th e Sur t an d Bivros t Lineaments (Blysta d e t al 1995 ; Dor e e t al 1997) (Fig s 2 an d lO a an d b) . Th e existenc e of these lineaments or zone s has hithert o not bee n well constrained, bu t i t is thought thei r orienta tion i s NW-SE t o NNW-SSE . Ou r interpreta tion suggest s tha t th e orientation s o f th e tw o continental lineaments or zones diverge slightly, with th e Sur t Lineamen t havin g a NNW-SS E orientation wherea s th e Bivros t Lineamen t i s oriented NW-S E (Fig . 2) . Th e Bivros t Linea ment could more correctly be named the Bivrost Zone, a s there is no obvious single lineament, as for th e Sur t Lineament , bu t probabl y a series of lineaments actin g a s a hing e zon e toward s th e elevated areas of the Utr0st Hig h (Fig. 2). This is clearly see n o n th e Bas e Cretaceou s tim e ma p (Fig. 11) , which here is shown only with the 5. 5 s

twt contou r t o highligh t the structure. Also, th e total magneti c fiel d (Fig . 12) , th e tota l gravit y anomaly fiel d (Fig . 13 ) and th e bes t structura l image (filtere d fre e air ) (Fig . 14 ) identif y th e location o f th e lineamen t o r zone . Th e Bas e Cretaceous 5.5 s twt map i s superimposed o n all the potential field maps. The continental Bivrost Zone an d Sur t Lineamen t lin k u p wit h th e oceanic Lofote n Fractur e Zon e an d a majo r bend i n th e magneti c anomalie s 24 A an d 24 B (Fig. 15) . The earliest opening of the Norwegian Sea, north of the Jan Mayen Fracture Zone , was probably affecte d b y these weakness zones in the continental crust . Th e magneti c anomalie s 24B and 24 A betwee n th e Lofote n Fractur e Zon e and th e kin k i n the magnetic anomalies , ocean ward o f th e Sur t Lineament , hav e a n almos t E-W orientatio n (Fig s 1 2 an d 15 ) (Hagevan g et al. 1983; Karpuz et al. in press a, b}, indicating that th e initial opening within this compartmen t might hav e bee n N-S . Th e orientatio n o f Lat e Cretaceous fault s o n th e continenta l sid e als o seems t o reflec t a loca l N- S extensiona l fiel d within thi s segment , wit h E- W trendin g individual fault s (Fig . 11) . W e sugges t tha t suc h a local N- S extensiona l field in lat e Lat e Cretac eous tim e could reactivat e ol d crustal weaknes s zones, suc h a s bot h th e NNW-SS E oriente d Surt Lineamen t an d NW-S E oriente d Bivros t Zone, wit h a sinistra l component , settin g u p a clockwise rotatio n o f th e crusta l bloc k deli neated by thes e lineament s or zone s and the Vema Dome-Ny k Hig h faul t syste m (Fig . 9) .

Fig. 5 . Lin e drawing fro m seismi c section . (Not e erosio n a t A indicatin g that th e roll-ove r structur e wa s eroded a t thi s level, which i s in th e lates t par t o f Maastrichtia n time.) (Fo r location , se e Fig . 2. )

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Fig. 6 . Tim e difference ma p for the Base Tertiary to the Intra Paleocene sequence . Nyk High was at this time still in a depositiona l setting . Indication o f 2 km lef t latera l movement i n a pull-apart grabe n a t A . Wavy line shows the post-Bas e Tertiar y erosiona l limit . (See also Fig . 7.)

Fig. 7. Tim e difference ma p for the Intra Paleocene t o Near To p Paleocene sequence . Nyk High evolved as a high in thi s period, a t A . Depocentr e configuratio n with transtensiona l (releasing ) and transpressiona l (restraining) bends indicat e lef t latera l movement o n th e faul t a t B . (Note th e loca l depocentre a t C. )

Such a clockwis e bloc k rotatio n woul d explai n the observed sinistra l movements along bot h th e NE-SW trendin g Vem a Dome-Ny k Hig h an d the NNW-SS E trendin g Sur t Lineament . Although th e amoun t o f rotatio n i s relativel y small (sinistra l movement s o f the order of 2 km), it woul d caus e varyin g type s o f deformatio n that coul d explai n th e observe d comple x faul t patterns (Fig s 8 an d 9) , especiall y i n th e Vem a Dome area .

Latest Late Cretaceous structura l deformatio n seems t o diminis h awa y fro m th e oceanic continental transitio n withi n th e crusta l com partment betwee n th e Sur t Lineamen t an d th e Bivrost Zone, and is less at the Utgard Hig h tha n on th e Vem a Dome-Ny k High . Thi s indicate s that the Vema Dome-Nyk High fault syste m has taken u p mos t o f th e lates t Lat e Cretaceou s structural deformation , bot h di p sli p an d strik e slip, an d i t als o implie s tha t th e crusta l bloc k

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Fig. 8. Lin e drawing of seismic section. Extension on the Near Top Santonia n level, but no apparent extension on the Nea r Top Cenomania n leve l is a goo d indicatio n o f lateral movements o n th e Nyk High . (For location, see Fig. 2.)

rotation betwee n thes e tw o crusta l weaknes s zones primarily is restricted t o th e oute r part s of the continental crust , closes t t o the initial breakup (show n schematicall y i n Fig . 16) , which i s in good agreemen t wit h th e stretchin g configura tion o f Robert s e t al. (1997, fig. 6). Eocene t o presen t structura l evolutio n (post-break-up) The clockwis e rotation tha t seem s t o hav e bee n effective durin g th e pre-break-u p extensiona l regime migh t hav e continue d durin g th e post break u p compressiona l regime . With th e break-u p o f th e Nort h Atlanti c i n Eocene tim e a differen t tectoni c settin g wa s introduced t o th e V0rin g Basi n an d th e Vem a Dome-Nyk Hig h area . Th e whol e are a becam e exposed t o compressiona l forces , an d severa l phases o f inversio n hav e bee n identifie d fro m Eocene t o th e Lat e Miocen e time . Earthquak e focal mechanisms indicate that th e Vema DomeNyk Hig h are a i s toda y unde r compressiona l forces (Bungu m e t al . 1991) , bu t th e las t majo r inversion tectoni c activit y took plac e i n Mid - t o Late Miocen e time . The orientatio n o f th e compressiona l force s today i s NW-S E base d o n earthquak e foca l mechanisms (Bungu m e t al . 1991) , wherea s re activated fault s i n th e souther n par t o f th e V0ring Basi n an d o n th e Tr0ndela g Platfor m indicate tha t th e maximu m compressiona l forc e

during th e firs t compressiona l phase s i n Paleo gene tim e migh t hav e ha d a mor e WNW-ES E orientation. Th e origi n o f th e compressiona l forces i s stil l unde r debate , bu t migh t b e a combination o f several opposin g plat e tectoni c events (Lundi n & Dor e 1997) . Whatever origi n these compressiona l force s had , the y see m t o have reactivate d ol d lineaments , an d als o th e Surt an d Bivros t Lineament s o r Zones , onc e again wit h a sinistra l component (Fig . 17) . It i s difficult t o identif y an y latera l movemen t i n th e reactivation o f thes e zones , bu t lef t latera l movements ca n explai n th e configuratio n an d pattern o f th e inversio n structure s i n th e are a (Figs 1 7 an d 18) . Sinistra l movement s woul d again induc e a clockwis e rotation o f th e crusta l block betwee n th e Sur t an d Bivros t Lineament s or Zones , bu t thi s tim e i n a compressiona l setting. Th e Ny k Hig h an d Vem a Dom e woul d therefore becom e structure s tha t evolve d i n th e corners o f thi s crustal bloc k (Fig . 17) . Depend ing o n th e configuratio n of th e underlyin g faul t system thes e tw o structure s would evolv e differ ently. Th e Vem a Dom e woul d evolv e a s a normal inversio n structure , wherea s th e Ny k High, bein g a former rotate d footwall , was overrotated an d finall y locke d whe n th e mai n fault s became to o stee p (Fig . 19) . Thi s coul d als o explain th e inversio n structur e forme d i n th e former Nagrin d Synclin e area (Fig s lOa , 1 7 an d 19), indicatin g tha t th e whol e compartmen t between th e Sur t Lineamen t an d Bivros t Zon e

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Fig. 9 . Structura l developmen t in Late Maastrichtian-Paleocen e time . As a resul t o f local extensional force s a clockwise rotation an d sinistra l fault movement s of the area betwee n the Sur t Lineamen t an d th e Bivros t Lineament o r Zon e ca n explain the depocentre configuratio n and th e complex structura l evolutio n o f the area. Depocentre A unde r influenc e o f sinistral shear.

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Fig. 10 . (a ) Section across th e Surt Lineament and Bivros t Zone sout h of the Vema Dome-Nyk High , (b) Section across th e Sur t Lineamen t and Bivros t Zone nort h o f th e Vem a Dome-Ny k High . (Fo r location , se e Fig. 2. )

SE o f th e Vem a Dome-Ny k Hig h are a becam e locked. The primar y structura l evolutio n durin g th e compressional stage s i n Tertiar y tim e (Dor e & Lundin 1996 ) took par t i n the outermost par t o f the continental crust, as during the pre-break-u p structural evolution , probabl y becaus e o f th e weakened natur e of the continental crust close to the oceanic-continenta l transition . Since Lat e Miocen e time , larg e Plio-Pleisto cene clinoform s ar e evidenc e o f progradatio n from th e Norwegia n mainland , graduall y bury -

ing th e inversio n structures , an d onl y th e mos t oceanward inversio n structure s suc h a s the Ny k High ar e image d toda y o n th e se a bottom. Th e hiatus on th e apex o f the Nyk Hig h includes the Campanian t o Quaternar y period , s o tha t w e have n o exac t informatio n abou t whe n th e Ny k High submerged . The Nyk Hig h has clearly been exposed t o erosiona l force s fro m a t leas t Maastrichtian tim e unti l possibl y Eocen e tim e (Figs 7 an d 19) , bu t whethe r th e Ny k Hig h submerged a s early as Lat e Eocen e o r a s lat e as Late Miocen e tim e is still a n unsolve d question .

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Fig. 11. A Base Cretaceous tw t map wit h only the 5.5 s twt contour, showin g the Surt Lineamen t an d th e Bivrost Zone.

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Fig. 12 . Tota l magneti c field in th e norther n par t o f the V0rin g Basin with th e 5.5 s tw t contou r fro m th e Base Cretaceou s a s an overlay . Both th e Sur t Lineamen t an d th e Bivros t Zone ca n b e picked out . Locatio n of Fig. 3 is indicated .

Fig. 13 . Tota l free-ai r gravity field in th e norther n par t o f the V0rin g Basin with th e 5.5 s tw t contour fro m th e Base Cretaceou s a s an overlay . Bot h th e Sur t Lineamen t an d th e Bivros t Zone ca n b e picked out . Locatio n o f Fig. 3 is indicated .

Fig. 14 . Bes t structual Imag e (filtere d free ai r gravity) in the northern par t o f the V0ring Basin with the 5. 5 s twt contour fro m th e Bas e Cretaceou s a s an overlay .

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Fig. 15 . Locatio n o f th e Sur t Lineamen t an d Bivros t Zone , an d th e relationshi p t o th e magneti c oceani c spreading anomalie s (not e th e kin k o f anomal y 24 B and 24 A in prolongation o f the Sur t Lineament. )

Fig. 16 . Schemati c illustratio n o f the structura l evolutio n i n th e outermos t par t o f th e are a betwee n th e Sur t Lineament an d th e Bivros t Zone . Loca l extensiona l force s giv e ris e t o basemen t bloc k rotation .

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Fig. 17 . Structura l developmen t i n Eocene-Miocene time . Because o f th e regiona l compressional forces , clockwise rotatio n o f the are a betwee n th e Sur t Lineamen t an d th e Bivros t Lineament o r Zon e onc e agai n ca n explain th e structural evolutio n o f the area. The inversion structures , th e Vema Dome and the Nyk High , evolv e in th e corners o f the rotatin g basemen t block .

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Fig. 18 . Th e tim e difference ma p fro m th e se a bottom t o th e Nea r To p Paleocen e positio n i n th e Vorin g Basi n outlines the western and northernmost inversio n structures related to the limited post-Miocene deposition in these areas.

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Fig. 19 . Over-steepenin g o f the norma l fault s o n th e Nyk Hig h durin g the compression i n Tertiary tim e could lock th e faults . This would giv e ris e t o a transfe r o f the compressional forces t o th e whole basemen t bloc k southeast o f the Nyk High , which could explai n the inversion of the former Nagrind Syncline . (For location , see Fig. 2. ) The author s wis h t o than k B . T0rudbakken , R . Scrutton an d 0 . Birkelan d fo r thei r constructiv e comments durin g th e revie w process , an d E . M . Olsen fo r he r excellen t draftin g work . W e als o wan t to thank Nors k Hydr o a.s.a . for permission t o publish these studies .

References BJ0RNSETH, H . M. , GRANT , S . M. , HANSEN , E . K. , HOSSACK, J . R. , ROBERTS , D . G . & THOMSON, M . 1997. Structura l evolutio n o f th e V0rin g basin , Norway, durin g th e Lat e Cretaceou s an d Paleo gene. Journal o f th e Geological Society, London, 154, 545-550 . BLYSTAD, P. , BREKK E , H., F^ERSETH , R . B. , LARSEN , B. T. , SKOGSEID , J . & TORUDBAKKEN , B . 1995 . Structural Elements of the Norwegian Continental Shelf. Norwegia n Petroleu m Directorat e Bulletin , 8, 1535-1541 . BREKKE, H . & Rns, F . 1987 . Tectonic an d basi n evo lution o f the Norwegia n Shel f betwee n 62°N an d 72°N. Norsk Geologisk Tidsskrift, 67 , 295-321. BUNGUM, H. , ALSAKER , A. , KVAMME , L . B . & HAN SEN, R . A . 1991 . Seismicit y an d seismotectonic s of Norwa y an d nearb y continenta l shel f areas . Journal o f Geophysical Research, 24, 8249-8265 . DORE, A. G . & GAGE, M. S . 1987 . Crustal alignment s and sedimentar y domain s i n th e evolutio n o f th e North Sea , North-eas t Atlanti c margi n an d Barents Shelf . In: BROOKS , J . & GLENNIE , K . (eds) Petroleum Geology o f North West Europe. Graham an d Trotman , London , 1131-1148 . & LUNDIN , E . R . 1996 . Cenozoi c compressiona l structures o n th e N E Atlanti c margin : nature ,

origin an d potentia l significanc e fo r hydrocarbo n exploration. Petroleum Geoscience, 2, 299-311. -, FICHLER , C. & OLESEN, O. 1997 . Patterns of basemen t structur e an d reactivatio n alon g th e NE Atlanti c margin . Journal o f th e Geological Society, London, 154 , 85-92 . HAGEVANG, T., ELDHOLM, O. & AALSTAD, I. 1983 . Pre23 magneti c anomalie s betwee n Ja n Maye n an d Greenland-Senja Fracture Zones in the Norwegian Sea. Marine Geophysical Research, 5, 345-363. HINZ, K. , ELDHOLM , O. , BLOCK , M . & SKOGSEID , J . 1993. Evolutio n o f Nort h Atlanti c volcani c continental margins . In : PARKER , J . R . (ed. ) Petroleum Geology o f North West Europe. Geolo gical Society , London , 901-913 . KARPUZ, M . R. , FLECHE , H . & FARELLY , B . 2000^ . Magnetic tota l fiel d anomalies . Norsk Geologisk Tidsskrift, i n press . ,& 2000/7 . Fre e air (offshore)-Bougue r (land) anomalies . Norsk Geologisk Tidsskrift, i n press. LUNDIN, E . R . & DORE, A. G . 1997 . A tectonic mode l for th e Norwegia n passiv e margi n wit h implica tions fo r th e N E Atlantic : Earl y Cretaceou s t o break-up. Journal o f th e Geological Society, London, 154 , 545-550 . MOGENSEN, T . E . 1994 . Paleozoic structura l develop ment alon g th e Tornquis t Zone , Kattega t area , Denmark. Tectonophysics, 240, 191-214 . PLANKE, S. , SKOGSEID , J . & ELDHOLM , O . 1991 . Crustal structur e of f Norway , 62°-70°N . Tectonophysics, 189 , 91-10 7 ROBERTS, A . M. , LUNDIN , E . R . & KUSZNIR , N . J . 1997. Subsidence o f th e V0rin g Basi n an d th e influence o f th e Atlanti c continenta l margin . Journal o f th e Geological Society, London, 154 , 551-557.

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. E . MOGENSE N E T A

SKOGSEID, J . 1994 . Dimension s o f th e Lat e Cretac - o f Continental Break-up. Geologica l Society . eous-Paleocene Northeas t Atlanti c rif t derive d London , Specia l Publications , 68 , 305-320. from Cenozoi c subsidence . Tectonophysics. 240 , WALKER , I . M. , BERRY , K . A. , BRUCE , J . R. . 225-247. BYST0L , L. & SNOW. J. H . 1997 . Structural model, PEDERSEN , T. , ELDHOLM , O . & LARSEN , B . T . lin g of regional depth profile s i n the Vorin g Basin: 1992. Tectonis m an d magmatis m durin g N E implication s fo r th e structur e an d stratigraphi c Atlantic continenta l break-up : th e Vorin g developmen t o f th e Norwegia n passiv e margin . margin. In : STOREY , B . C. , ALABASTER , T . & Journal o f th e Geological Society, London. 154 . PANKHURST, R. J. (eds) Magmatism an d the Causes 537-544 .

Norwegian-Greenland Sea therma l field EIRIK SUNDVOR, 1 OLA V ELDHOLM, 2 TADEUS Z P GLADCZENKO 2 & SVERR E PLANKE 2 1

Institute of Solid Earth Physics, University of Bergen, Allegt. 41, N-5007, Bergen, Norway (e-mail: [email protected]) 2 Department of Geology, University of Oslo, Pb. 1047 Blindern, N-0316 Oslo, Norway Abstract: Therma l gradien t an d therma l conductivit y measurement s carrie d ou t i n th e Norwegian-Greenland Sea and in the western Eurasi a Basin from 196 7 to 199 6 have yielded 436 heat-flo w value s whic h hav e bee n store d o n a UNIX-platfor m database , HEAT , t o facilitate retrieval, documentation an d display . HEA T contains al l public domain dat a fo r th e region, an d th e databas e an d applicatio n program s ar e freel y available . Th e hea t flo w o n oceanic crus t reveal s a clear , first-orde r heat-flow-crustal-ag e relationship . Th e therma l conductivity i s relativel y unifor m an d consistent , wit h value s in th e 0.85-1.1 5 Wrrr1 K"1 range. Th e lowe r value s ar e fro m th e ocea n basin s an d th e highe r value s fro m th e conti nental margins. Local , very high heat flow along the present spreadin g axi s may imply activ e volcanism o r venting . We also documen t loca l continental slop e maxim a o n th e M0r e an d western Barent s Se a margins, contrastin g greatl y with th e typica l lo w hea t flow, generally <75 mWm~2 , on old oceanic and thinned continental crust . The abnormally high heat flow, exceeding lOOOmWrrr 2 , o n th e 'Hako n Mosby ' gas-seepin g mu d volcan o o n th e Barent s Sea margin suggest s that als o othe r loca l slop e maxim a ma y b e related t o activ e seeps .

Most o f th e earl y heat-flo w measurement s i n the Norwegian-Greenlan d Se a wer e collecte d by Lamont-Dohert y Geologica l Observator y (LDGO) during corin g operation s i n the perio d 1967-1973. Despite an uneve n geographica l dis tribution, thes e measurements outline d the firstorder therma l propertie s o f th e oceani c crus t (Langseth & Zielinski 1974) . Few additional measurement s wer e made until the acquisitio n o f a Norwegia n heat-flo w prob e dedicated t o high-latitud e work i n 1983 . At th e same time , petroleu m industr y interes t i n th e sedimentary basin s alon g th e Norwegia n con tinental margi n le d t o a French-Norwegia n programme, fro m 198 3 t o 1986 , t o asses s th e thermal and buria l history of sediment basins on the continenta l margi n b y mappin g th e crusta l structure and the therma l field alon g selecte d margin transects . Th e joint programme, betwee n Elf Petroleu m Norg e an d th e universitie s o f Bergen an d Oslo , include d seismi c wide-angl e measurements b y th e expande d sprea d profilin g technique (NORMA R Project) , an d heat-flo w measurements (FLUNORG E Project ) i n co operation wit h Institu t Frangai s d e Recherch e pour 1'Exploitatio n d e l a Me r (IFREMER) . A tota l o f 16 1 heat-flow value s wer e acquire d during FLUNORG E (Sundvo r e t al 1989 ; Sundvor & Eldhol m 1992) . These data , a s wel l as measurement s durin g man y Norwegia n

cruises, hav e greatl y expande d ou r understand ing o f th e therma l propertie s o f th e Norway Svalbard continenta l margin and the Jan Maye n Ridge. In addition, a number of values have been measured o n oceani c crus t i n th e Norwegian Greenland Se a and th e Eurasi a Basin . In thi s study , w e presen t th e heat-flo w an d thermal conductivity data fro m th e NorwegianGreenland Sea , western Eurasia Basi n and north ern Svalbar d margi n (Fig s 1 an d 2 ; Tabl e 1) , and discuss the regional character o f the thermal field ove r oceani c crus t an d alon g transect s across th e Eurasia n margin . Data an d database The heat-flo w survey s within the stud y are a ar e listed i n Tabl e 1 . The majorit y o f temperatur e gradients hav e bee n measure d b y thre e instru ment packages . Th e LDG O cruise s applie d th e Ewing thermogra d consistin g o f u p t o fiv e thermistors mounte d o n th e Ewin g piston core r (Gerard e t a l 1962 ; Haene l 1979) . Th e Norwe gian LDGO-manufacture d instrument , designed for fas t operations , employ s thermisto r probe s mounted o n outrigger s several centimetres away from a soli d lanc e tha t i s attache d t o a core head weight . I t also monitor s th e bottom water temperatur e and th e til t o f the instrument

From'. N0TTVEDT , A . e t al . (eds ) Dynamics o f th e Norwegian Margin. Geologica l Society , London , Specia l Publications, 167 , 397-410 . 1-86239-056-8/OO/ S 15.00 © Th e Geologica l Societ y o f Londo n 2000 .

398

E. SUNDVO R E T AL.

Fig. 1 . Heat-flo w measurements and selecte d heat-flo w contours . Simplifie d bathymetry in m. Mai n structural features an d sea-floo r spreading anomalie s fro m Talwan i & Eldholm (1977) , Skogseid & Eldholm (1987 ) an d Eldholm et al. (1990). Icelan d heat-flo w contours fro m Floven z & Saemundsson (1993) . RR, Reykjane s Ridge; KR, Kolbeinse y Ridge; AR, Aegir Ridge ; FIR , Faeroe-Icelan d Ridge ; JMR, Ja n Maye n Ridge; HR , Hovgaar d Ridge; JMFZ, Jan Maye n Fractur e Zone (WJMFZ and EJMFZ) ; SFZ, Senj a Fractur e Zone; GFZ , Greenlan d Fracture Zone; SPFZ, Spitsbergen Fracture Zone; HFZ, Hornsun d Faul t Zone; IP, Iceland Plateau; NB, Norway Basin; LB , Lofote n Basin ; GB, Greenlan d Basin ; BB, Boreas Basin ; F-SE , Faeroe-Shetlan d Escarpment ; VE, V0rin g Escarpment; GE , Greenlan d Escarpment ; YP , Yermak Plateau .

NORWEGIAN-GREENLAND SE A THERMA L FIEL D

399

Fig. 2 . Distributio n o f thermal conductivity measured in nea r sea-floo r sediments, and averag e values for separate physiographic province s (Tabl e 2) . Gridded, 1.5 ° x 0.5°, sediment thicknes s contours i n s twt fro m Johansen (1989) . Other reference s an d acronym s in Fig . 1.

package. The IFREMER heat-flow instrument is based on a modified Aanderaa thermisto r string, refurbished fo r enhanced resolution, attached t o the outer part of the piston cor e barrel . Thermal conductivitie s (Tabl e 2 ) hav e bee n measured o n sedimen t core s usin g th e needle -

probe metho d (vo n Herze n & Maxwel l 1959 ; Haenel 1979) . Durin g th e FLUNORG E an d Norwegian cruises a Fenwal probe consistin g of thermistor an d heatin g wir e wa s used, an d th e thermal conductivit y wa s measure d a t 10c m intervals along the cores and correcte d t o i n situ

Table 1 . Marine heat-fiow surveys i n Norwegian Greenland Se a an d Eurasia Basin Area

Survey

Number o f measurement s Tgrad.

s

Institution

Year

USGS; ON R LOGO Shirshov Inst. , Mosco w BGR

1965 1966 197 3 1971 1969, 197 2

Denmark Strai t Norwegian-Greenland Se a Norwegian -Greenland Se a Voring Plateau ; Lofote n Basi n

10 56 11 7

11 48 11 7

10 56 11 7

1 2 3 4

LOGO, UiB , UiO

1980

16

15

16

5

CGC, ONR , UiB , Ui O Elf, IFREMER , UiB , Ui O UiB

1981 1983 1984

4 71 96

4 18 2

4 71 96

6 7-8 9

Ice island Fra m II I FLUNORGE RV Hakon Mosbv

Elf, IFREMER , UiB , Ui O Elf, IFREMER , UiB , Ui O IKU UiB, Ui O OOP AWI, Geomar , Ui B OOP

1985 1986 1985 1986 1985 1987 1994

65 26 4 23 1 17 6

16 11 4 7 1 17 7

65 26 4 23 1 17 6

8 10 11 1 3 10 14 15 16

FLUNORGE FLUNORGE Shallow drilling RV Hakon Mosbv Leg 10 4 RV Polarstern Leg 15 1

NRL, UiB , Ui O OOP NRL, UiB , UiO, VNII O

1995 1995 1996

Norwegian-Greenland Sea , Yermak Platea u Yermak Platea u Norway-Svalbard Margi n Svalbard Margin , Greenlan d Sea , Yermak Platea u M0re Margi n Jan Maye n Ridg e Barents Sea Norwegian- Greenland Se a Norwegian- Greenland Sea Eurasia Basi n Norwegian Greenlan d Sea , Yermak Platea u Svalbard margi n Svalbard, Greenlan d margin s Norway Svalbar d margi n

Ice island Arli s I I RV Vema 23 , 27 , 28 , 29 , 3 0 RV Akademik Kurchatov RV Planet 1969 , RV Meteor 197 2 Ymer-80

25 2 16

77 2 13

25

17 18 17

RV Hakon Mosbv Leg 16 2 RV Profesor Logaehev

456

271

456

Total

Cond.

References Comment

Heat flow

9

16

AWI, Alfre d Wegene r Institut e fo r Pola r Research ; BGR , Bundesanstal t fur Geowissenschafte n un d Rohstoflfe ; Elf , El f Petroleu m Norge ; GSC , Geologica l Surve y o f Canada; I K U , IKU Petroleu m Research ; LDGO , Lamen t Dohert y Geologica l Observatory ; NRL , Nava l Researc h Laboratory ; ONR , Offic e o f Nava l Research ; UiB , University o f Bergen ; UiO, Universit y o f Oslo ; USGS , U S Geologica l Survey ; VNIIO, VNII-okeangeologia Institut e S t Petersburg . 1, Lachenbruc h & Marshal l (1968); 2, Langset h & Zielinsk i (1974); 3 , Udintse v & Lubimov a (1973) ; 4 , Haene l (1974) ; 5 , Crane e t al. (1982) ; 6, Jackso n e t al. (1984) ; 7, Cran e e t al . (1988) ; 8 , Sundvo r et al . (1989) ; 9 , Sundvo r (1986) ; 10 , Sundvor & Myhr e (1987) ; 11 , Siette m (1988) ; 12 , Zielinski et al . (1986) ; 13 , Loset h e t al . (1992) ; 14, Eldhol m e t al . (1989) ; 15 , Sundvo r & Tor p (1987) ; 16 , Myrh e et al . (1995) ; 17 , Eldhol m e t al . (1999) ; 18 , Janse n e t al . (1996) .

NORWEGIAN-GREENLAND SE A THERMA L FIEL D

401

Table 2 . Thermal conductivities measured i n near-surface sediment Physiographic provinc e

Number Averag of values (Wm"

e conductivit y Standar K~') deviatio

North Svalbar d margi n Yermak Platea u Svalbard margi n (74-80°N ) Barents Se a margin (70-74°N ) Norwegian margin (62-70°N ) Margin tota l

5 11 25 36 22 105

1.15 1.23 110 1.13 1.04 .111

0.33 0.23 0.26 0.15 0.28 0.24

Eurasia Basi n Greenland Se a Greenland Basi n Lofoten Basi n Iceland Platea u Norway Basi n Reykjanes Ridg e (nort h o f 61°N ) Basins tota l

17 25 11 27 36 0.8 20 0.8 27 0.8 163 0.9

1.11 .116 .014 .081 8 7 5 9

0.23 0.16 0.13 0.09 0.16 0.12 0.20 0.20

Hovgaard Ridg e regio n NE Norwa y Basi n Overall tota l

5 1.2 7 1.2 280 1.0

6 9 5

0.04 0.27 0.23

1

d n

Stations, Y l l an d Y1 8 (Cran e e t al. 1982 ) an d 15F8 3 (Cran e e t 0/.1988) , interprete d a s linearly increasin g conductivit y wit h depth , yiel d averag e value s >2.0Wm~ 1 K" 1 . Thes e values contras t greatl y wit h thos e o f adjacen t station s an d hav e bee n exclude d i n th e averages.

conditions (Ratcliff e 1960) . Th e LDG O an d IFREMER cor e barre l metho d allow s conduc tivity measurement s o n th e recovere d core , whereas th e Norwegia n lanc e prob e dat a wer e complemented by gravity cores at selected probe locations. A t location s withou t sedimen t cores , conductivities fro m neighbourin g station s hav e been used . Normally, th e averag e cor e conductivit y i s used t o calculat e th e hea t flow . However , increased conductivit y wit h dept h measure d i n some core s o n th e Svalbar d margi n mad e Crane e t al . (1982 , 1988 ) introduc e a linea r conductivity-depth relationshi p a t thes e sta tions. T o obtai n comparabl e heat-flo w values, we have calulated the averag e conductivity over the length of core, and recalculated the heat flow. Heat flo w ha s als o bee n determine d a t scien tific dril l site s (Fig s 1 an d 2 ; Tabl e 1) . Near continuous therma l conductivit y profiles exist at most Ocean Drillin g Program (ODP) sites, and a limited numbe r o f temperatur e gradient s hav e been measure d a t Site s 64 2 (Eldhol m e t al . 19876), 907-91 2 (Myhr e e t al . 1995) , an d 981 , 984, 98 6 and 98 7 (Jansen et al. 1996) . In general, the temperatur e gradient s ar e no t considere d well established , althoug h ther e i s a first-orde r consistency between heat flow in ODP hole s and in cor e o r prob e location s (A . M . Myhre , pers . comm.). In addition, a few heat-flow values have been publishe d fro m shallo w exploratio n hole s

in th e Barent s Se a (Zielinski et al . 1986 ; Saettem 1988; L0set h e t al . 1992). Marine heat-flo w dat a ar e affecte d b y instrumental error s an d bottom-wate r instability . Moreover, geolog y (sedimentation , hea t refrac tion, sea-floo r topography , uplif t an d erosion , etc.) ma y introduc e loca l anomalie s affectin g data reliabilit y an d interpretatio n potentia l (Langseth & Zielinsk i 1974 ; Haene l 1979) . Ideally, studie s o f lithospheri c properties base d on hea t flo w requir e a numbe r o f corrections . None th e less , th e commo n absenc e o f suc h corrections i n regional studie s ma y explai n par t of th e relativel y larg e standar d deviation s i n Table 3 (below). All heat-flo w data availabl e t o u s hav e bee n included i n a database , HEAT , t o facilitat e retrieval, documentatio n an d display . HEA T also include s som e survey s an d measurement s outside th e regio n show n i n Fig . 1 . The data base an d applicatio n program s t o mak e listings, maps an d section s hav e bee n documente d b y Planke (1989) . The databas e accepts , an d con verts between , c.g.s . an d S I hea t flo w units . Here, w e us e S I units , i.e . therma l gradien t i n mKm" 1 , therma l conductivit y of sediment s in Wm" 1 K"1, and sea-floor heat flow in mWm~ 2 . The softwar e is designed fo r a UNIX-platfor m and ma y b e linked with th e marin e geophysical GMT syste m (Wessel & Smith 1991) , as wel l a s with othe r displa y and interpretatio n packages .

E. SUNDVO R E T AL .

402

Table 3. Table 1. Average heat flow for various crustal provinces compared with th e averages of Haenel (1979) Region

Number Mea of value s valu

n Standar e deviatio

Earth All ocean s All continent s Northeast Atlanti c Ocea n Mid-Atlantic Ridg e Non-ridge area s Norwegian-Greenland Se a Norwegian Sea * Oceanic basin s >15Ma * Oceanic crus t <10Ma * Near-spreading axi s < 2 Maf Norwegian margi n M0re-V0ring M0re V0ring Jan Maye n Ridg e

5417 3718 1699 124 64 54 64 94 72 18 19

74.3 79.8 62.3 82.8 100.0 52.2 96.4 93.4 73.1 165.8 234.7

90.9 105.9 40.1 73.7 90.0 22.4 48.3 54.6 23.9 79.0 93.3

43 29 14 11

56.6 51.5 67.3 75.2

15.6 13.9 12.4 10.5

d Referenc n

e

Haenel (1979 ) Haenel (1979) Haenel (1979) Haenel (1979) Haenel (1979) Haenel (1979) Haenel (1979) This study This stud y This stud y This study This stud y This study This study This stud y

Distribution o f measurement s i n Fig . 1. * Includes only values on identified oceanic crust , i.e. south of Greenland-Senja Fractur e Zone . f Includes Knipovic h an d Mollo y ridges . The databas e an d th e HEA T softwar e ar e freel y available fro m Departmen t o f Geology , Uni versity o f Oslo .

Regional setting The early Tertiary break-up between Eurasia and Greenland durin g Chro n 24 r wa s accompanie d by transient , large-scale , subaeria l break-u p volcanism (Eldholm e t al. 1989). The subsequen t plate tectoni c evolutio n nort h o f th e Faeroe Iceland Ridg e may spatiall y be divided into four main region s separate d b y th e Ja n Mayen , Greenland-Senja an d Spitsberge n fractur e zon e systems (Fig. 1) ; and temporall y by the change in relative plat e motio n durin g Chro n 1 3 tim e (Talwani & Eldhol m 1977) . I n particular , th e Norway Basin-Icelan d Platea u an d Greenlan d Sea reveal a complex sea-floo r spreading history with unstable plate boundaries an d formatio n of the Ja n Maye n an d Hovgaar d microcontinents , respectively. Th e plat e geometr y delaye d crusta l extension i n th e norther n Greenlan d Se a unti l early Oligocene times , and generatio n o f oceani c crust was probably initiated as late as mid- to late Miocene tim e (Myhre & Eldholm 1988 ; Eldholm et al . 1994) . Th e plat e boundar y configuratio n through tim e also develope d a numbe r o f rifte d and sheare d continenta l margi n segments . The sedimen t distributio n o n oceani c crus t (Fig. 2 ) reflect s th e plat e tectoni c framewor k (Eldholm & Windisc h 1974 ; Johanse n 1989) .

However, lat e Neogen e continenta l uplif t an d the subsequen t Norther n Hemispher e glaciation led t o erosio n o f th e inne r shel f an d adjacen t landmass, an d t o depositio n o f thick , terrige nous sedimen t wedge s along th e easter n margi n (Eldholm e t al . 1994) . I n som e areas , hug e fa n complexes hav e bee n constructe d i n fron t o f prominent continenta l drainag e systems . Th e distal fan s commonl y extend onto oceanic crust, exerting a significant , exces s sedimentar y loa d (Myhre e t al . 1992 ; Faleide e t al . 1996) .

Thermal conductivit y The 27 1 therma l conductivit y measurement s show unifor m an d consisten t value s mostl y i n the 0.85-1.15 Wm-1 K~! range . Except fo r local areas, i t is not possibl e to make meaningfu l con tours. Therefore, we have illustrated the regional features b y calculatin g averag e conductivit y values fo r differen t physiographi c province s (Fig. 2), showing that th e conductivity is slightly higher o n th e margin s than i n th e ocea n basins . Moreover, th e conductivit y ma p outline s loca l areas o f relativel y hig h conductivity : th e Eas t Norway Basi n an d M0r e margin , a regio n tha t has been affecte d b y huge Quaternary submarine slides fro m th e uppe r slop e (Bugg e et al . 1987) ; the Nort h Greenlan d Se a i n th e vicinit y o f the Hovgaar d microcontinent ; the Barent s Sea Svalbard shel f an d uppe r slope ; th e wester n Yermak Plateau .

NORWEGIAN-GREENLAND SE A THERMA L FIEL D With th e exceptio n o f tw o core s samplin g Eocene mu d diapir s o n th e V0ring Plateau , th e cores sample d th e Quaternar y sequence . Th e thermal conductivity of marine sediments corre lates with bulk density and inversel y with water content (Hamilto n 1974) , and w e note tha t th e ODP core s o n th e V0ring margin revealed high densities an d lo w wate r content s i n th e dark , glacial sample s compare d wit h othe r glacia l o r interglacial sample s (Eldhol m e t al. 19876) . Therefore, th e variabilit y i n therma l conduc tivity i s probabl y relate d t o a dominanc e o f pelagic components i n the ocean basins, whereas low-conductivity sediment s deposite d durin g the presen t interglacia l ma y locall y hav e bee n removed o n th e continenta l margin . Thus , th e local area s o f hig h conductivit y o n th e margi n may b e indicator s o f presen t erosio n o r non deposition. Heat flow The regiona l distributio n o f th e 45 6 heat-flow stations nort h o f th e Faeroe-Icelan d Ridg e varies considerably . Mos t station s ar e locate d along transects across th e eastern margin, in the North Greenland Se a and in the North Norwa y Basin-Jan Maye n Ridg e region . Non e th e less , we have been abl e t o contour th e regional heat flow field (Fig. 1) . We als o illustrat e the spatia l and tempora l heat-flo w distributio n b y a serie s of regiona l transect s (Fig s 3-5) . Th e transect s are locate d i n region s o f relativel y dens e heat flow coverage , an d adjacen t station s ar e pro jected ont o th e transects . Th e minimu m water depth an d maximu m projectio n distanc e ar e annotated o n eac h transect . Mos t transect s approximate plat e tectoni c flowline s o r ru n perpendicular t o mai n structura l trends.

Oceanic crust Both the contour ma p and the transects reveal a clear first-order correspondence o f heat flow and age o f th e oceani c crus t (Fig s 1 an d 4) . Th e mid-oceanic ridge s provinc e an d adjacen t crus t is characterize d b y hea t flo w >100mWm~ 2 , whereas mos t station s o n crus t olde r tha n c. 30 Ma yiel d value s i n th e 50-7 5 mWm~2 range. Th e marin e dat a correspon d fairl y wel l to onshor e measurement s i n Iceland (Floven z & Ssemundsson 1993) . On the other hand, th e data coverage resolve s neithe r th e therma l signatur e of th e extinc t Aegi r spreadin g axi s i n th e Norway Basi n no r th e questio n o f a n extinc t axis west of the th e Ja n Maye n Ridg e (Talwan i

403

& Eldhol m 1977) . Previously , heat-flow-ag e relationships fo r variou s region s hav e bee n dis cussed base d o n sub-set s o f th e dat a i n Fig . 1 related t o lithospheric coolin g models (Langset h & Zielinski 1974; Zielinski 1979; Langseth e t at . 1990). W e restric t ou r analysi s t o station s o n reasonably well-identifie d oceani c crus t (Tal wani & Eldholm'1977 ; Hagevan g e t al . 1983 ; Nunns 1983 ; Skogsei d & Eldhol m 1987) , i.e . the crust sout h of the Greenland-Senja Fractur e Zone. Th e empirica l function , hea t flo w = 420 age"1/2, appear s t o fi t th e observe d dat a (Fig. 3). For referenc e t o th e globa l an d northeas t Atlantic mea n hea t flo w o f Haene l (1979) , w e have calculated averag e values. We attribute th e higher average s an d lowe r standar d deviation s in Tabl e 3 t o th e relativel y youn g ag e o f th e ocean, a hotte r tha n norma l uppe r mantle , an d data o f generally good quality . The heat flow along the present spreadin g axis south of the Spitsbergen Fracture Zon e is shown in Tabl e 3 fo r crust s les s the n c . 2 an d 1 0 Ma. We als o observ e som e local , ver y hig h near axis heat-flo w values . I n fact , th e 461mWm~ 2 value o n th e Knipovic h Ridg e (Langset h & Zielinski 1974 ) an d th e 85 5 an d 1164mWm- 2 values o n th e Nanse n Ridg e (Sundvo r & Torp 1987) ar e minim a becaus e th e measure d tem peratures exceede d th e calibrate d thermisto r ranges. W e sugges t tha t thes e extrema l value s reflect activ e volcanism and/or ventin g along the plate boundary .

Microcontinents The Ja n Maye n Ridg e microcontinen t wa s separated fro m Eas t Greenlan d b y riftin g an d sea-floor spreadin g sinc e fro m c . 44 Ma, an d complete separatio n wa s achieve d a t c . 24 Ma (Myhre e t al . 1984) . Fou r transect s cros s th e main ridg e block , whic h i s fla t toppe d i n th e north an d narrow s an d plunge s southward . The ridg e ha s a relativel y stee p slop e toward s the Icelan d Platea u an d a mor e gentl e slop e towards th e Norwa y Basi n (Fig . 1) . The tran sects sho w tha t th e ridg e ha s n o heat-flo w signature an d tha t ther e i s a regional , gentl y decreasing heat-flo w tren d fro m th e younge r Iceland Plateau toward s the older Norway Basin reflecting th e crusta l age s (Fig . 3) . Non e th e less, th e averag e hea t flo w o n th e ridg e bloc k proper, 75mWm~ 2 , i s greate r tha n tha t mea sured o n extende d continental crus t of f Norwa y (Table 3) . W e attribut e th e relativel y hig h heat flo w t o th e smal l siz e o f th e continen tal fragmen t an d th e fac t tha t it s crus t ha s

404

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Fig. 3 . Top : locatio n o f heat-flo w transect s (Fig s 3-5) , an d hea t flow v. age of oceani c crust fo r area s o f wellidentified crus t sout h o f the Greenland-Senj a Fractur e Zone . Bottom : Norwa y margi n heat-flo w transects. Red shadin g refer s t o loca l continental slop e maxima . WD, wate r depth ; D . distanc e norma l t o transect. 20 magnetic anomaly .

NORWEGIAN-GREENLAND SE A THERMA L FIEL D

405

Fig. 4 . Heat-flo w transects over predominantly oceanic crust and microcontinent s (Fig. 3). Red shading refers t o local continenta l slop e maxima .

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Fig. 5. Barent s Sea and Svalbard margi n heat-flow transects (Fig. 3). Red shading refers to local continental slope maxima

NORWEGIAN-GREENLAND SE A THERMA L FIEL D been thinned , an d intrude d b y two separat e rif t episodes (Gudlaugsso n e t al . 1988 ; Johanse n et a l 1988) . The Hovgaar d Ridg e microcontinen t (Fig . 1 ) separated fro m th e Svalbard-Barent s Se a mar gin by the readjustment of the plate boundary in early Oligocen e tim e (Eldhol m e t al . 1987a) . Recently, a continuation of the continental crust into th e Hovgaar d Basi n northwest o f the ridg e proper ha s bee n suggeste d b y A . M . Myhr e (pers. comm.) . The hea t flow ove r the inter preted continenta l fragmen t i s agai n simila r t o that i n th e adjacen t basin s (transect s 14-15 , Fig. 5) , an d ther e i s excellen t correspondenc e of th e shallo w measurement s an d thos e i n ODP hole s o n to p o f th e ridg e an d i n th e Hovgaard Basin . The averag e ridge heat flow is 91mWm~ 2 , increasin g t o 117mWirr 2 o n th e thinner crus t i n the Hovgaar d Basin .

Nor way-Svalbard continental margin The heat-flo w ma p reveal s loca l high s o n th e continental slop e of f Norwa y an d th e Barent s Sea (Fig . 1) . W e discus s th e slop e maxim a separately, an d thes e heat-flo w value s ar e no t included i n th e regiona l statistic s (Table 3) . The measurement s o n th e Mor e an d V0rin g margin cove r bot h th e margina l high s an d th e continental crust , whic h underwen t extensio n before th e earl y Tertiar y break-up . Transect s 6-9 exhibi t a very small regional decrease in heat flow toward th e continent; however , neithe r th e marginal high s no r th e Faeroe-Shetlan d an d V0ring marginal escarpment s hav e any appreci able heat-flo w signatur e (Fig . 3) . Modellin g of th e V0rin g margi n b y Zielinski (1979) shows that th e data fit a continent-ocean boundar y in the vicinity of the V0ring Escarpment , an d tha t effects o f latera l hea t flo w ma y occu r i n thi s setting. Th e averag e hea t flo w landwar d o f th e marginal escarpment s i s 57mWm~ 2 , bu t th e heat flo w i s distinctl y highe r o n th e V0rin g margin tha n farthe r south, i.e. 67 v. 52mWirT 2 (Table 3) . As first noted by Langseth & Zielinski (1974), th e lo w M0r e margi n hea t flo w i s surprising in view o f the conductivit y maximum in th e norther n par t (Fig . 3) . Ou r expande d database confirm s th e abnormal , an d stil l enigmatic, hea t flo w an d conductivit y i n thi s region. W e als o not e tha t transec t 9 whic h strongly affect s th e V0rin g average , crosse s th e Vema Dome , whic h underwen t bot h Paleocen e and Lat e Oligocene-earl y Pliocen e deforma tion, a s wel l a s th e activ e Vem a mu d diapi r field initiate d i n lat e Pliocen e tim e (Hjelstue n et al . 1997) .

407

The measurements o n oceani c crus t just wes t of the Senja Fracture Zone are internally consistent, exhibiting a very gentle increase wit h crustal age fro m 5 3 t o 75mWm~ 2 alon g th e north trending transec t 10 , whereas th e therma l fiel d is stable , averagin g 63.5±5.3mWm~ 2 , alon g transect 11 , which runs approximately alon g th e magnetic anomaly 1 3 lineation (Fig. 5) . The few measurements eas t o f th e fractur e zon e ar e no t statistically significant, but several values are less than 50mWm~ 2 . The Senj a Fractur e Zone , a well-define d structural boundar y demarcating a ver y narrow continent-ocean transitio n zon e (Faleid e e t al . 1993), doe s not exhibi t an appreciabl e heat flow signatur e i n transec t 10 ; whereas reliabl e values o f 121-122mWm~ 2, decreasin g t o <40mWm~ 2 c . 20km farthe r east , appea r i n transect 1 1 (Fig . 5) . I n vie w o f severa l othe r local high s in this region, the anomal y may not relate t o th e fracture zone and w e classify i t as a slope maximum. Both transec t 12 , o n th e post-lat e Pliocen e Storfjord Fa n (Hjelstue n e t al . 1996) , an d transects 13-1 6 betwee n 7 8 an d 79° N ar e located wes t of th e Hornsun d Faul t Zon e (Fig . 5). Thi s structur e o n th e oute r shel f sout h o f 79°N demarcates th e westernmos t extent o f th e continental crust (Myhr e & Eldholm 1988) . The region between the fault zone and the Knipovich Ridge provinc e i s characterize d b y thic k sedi ments o f whic h th e mai n par t comprise s prograded lat e Pliocene and Quaternar y glacial fans and wedges (Faleide e t al. 1996). In general, there is a larger scatte r in heat flow than farther south, as well as a high regional level outside the ridge province . W e attribut e th e forme r t o th e larger numbe r o f station s withi n th e partl y sediment-covered ridg e province, an d th e latte r to th e young crust, particularl y i n th e norther n Greenland Sea . The thermal evolution of the Svalbard margi n and th e norther n Greenlan d Sea-Fra m Strai t region has been discussed and modelled by Crane et al. (1982, 1988 , 1991) . The results are consist ent wit h th e regiona l plat e tectoni c framework of Eldhol m e t al. (19870), although a propagat ing rif t an d a migratin g plat e boundar y hav e been advocated fo r the northern Greenlan d Sea . North o f Svalbard , th e crusta l natur e o f th e margin i s stil l largel y unresolve d (Sundvo r & Austegard 1990) . Th e regiona l heat-flo w leve l west o f 35° E is generall y less tha n 75mWm~ 2 ; however, th e wester n par t o f th e Yermak Pla teau constitute s a n are a o f anomalousl y hig h heat flo w an d therma l conductivit y (Fig s 1 and 2) . Th e hig h hea t flow , an d th e lat e Mio cene volcanis m an d presen t therma l activit y i n

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northern Svalbard , hav e bee n relate d t o crusta l dyke injection from 1 6 Ma t o th e present (Cran e et a l 1982 ; Okay & Crane 1993) . O n th e othe r hand, ther e i s no direc t spatia l correspondenc e between th e onshor e therma l activit y and he offshore heat-flo w anomaly .

Slope maxima Sundvor e t aL (1989 ) pointed ou t severa l heatflow maxim a o n th e continenta l slop e o n th e M0re an d Barent s Se a margins . Suc h highs , shown in red in Figs 3-5 and as elongate maxima in Fig . 1 ar e characterize d b y local , greatl y enhanced hea t flo w wit h respec t t o th e surrounding regional field. For example , in transect 7 (Fig . 3 ) w e observ e maxim a o f 322mWm~ 2 and 118-172mWm- 2 a t depth s o f 69 5 an d 1100-1350 m respectively . Subsequently , Nava l Research Laborator y SeaMAR C I I sidesca n sonar an d swat h bathymetr y cruise s (Cran e e t al. 1995 ) revealed tw o circula r back-scatte r objects a t o f 1 km diamete r c . 1250m dept h o n the Barents Sea margin just sout h o f transect 11 . A cor e fro m th e flan k o f th e mai n anomal y contained ga s an d ga s hydrate , an d th e hea t flow was calculated t o b e 377mWm~ 2 (Vog t & Sundvor 1996) , wherea s Eldhol m e t al . (1999) , using a lowe r therma l conductivity , arrive d a t 298mWm~ 2 . Thi s discover y le d t o a detaile d survey i n 199 6 confirmin g that th e mai n struc ture, name d th e Hako n Mosb y Mu d Volcano , oozed mu d an d seepe d gas , an d tha t i t i s prob ably cappe d b y ga s hydrat e a t th e se a floo r (Vogt e t al . 1997) . Furthermore , th e heat-flo w maximum measure d o n th e structur e exceed s lOOOmWirr 2 (Eldhol m e t al . 1999). The mu d volcan o i s locate d withi n a larg e slide sca r o n th e post-lat e Pliocen e Bea r Islan d Fan, an d i s underlai n b y c . 33 Ma ol d oceani c crust (Hjelstue n e t al . 1999) . Farther south , th e heat flow maximum in transec t 7 and th e oute r maximum i n transec t 8 (Fig . 3) ar e measure d within th e Storegg a slid e (Bugg e e t al . 1987) . Although th e mechanis m causin g th e heat flow anomalie s i s stil l discusse d (e.g . Vogt & Sundvor 1996 ; Eldhol m e t al . 1999) , a relationship wit h large-scale , rapi d mas s movemen t inducing change s i n sedimen t temperatur e an d pressure a s wel l a s providin g new pathway s for fluids and ga s need s t o b e explored . Non e th e less, th e discover y o f th e Hako n Mosb y Mu d Volcano and its associated gas seep suggests that the other loca l slop e anomalie s o n the NorwaySvalbard margi n ma y als o b e cause d seepin g mud, flui d an d gas . However, detaile d loca l surveys are require d t o confir m this inference.

Summary All publishe d therma l gradient , therma l conductivity an d heat-flo w values fro m th e Norwe gian-Greenland Se a and i n the wester n Eurasi a Basin hav e bee n store d o n a freel y availabl e UNIX-platform database . Th e compilatio n o f heat-flow an d therma l conductivit y dat a show s the following . There i s a clear , first-orde r heat flow-crustal ag e relationshi p o n date d oceani c crust. Therma l conductivitie s ar e relativel y uniform an d consistent in the 0.85-1.15 Wm"1 K"1 range; lo w value s ar e measure d i n th e ocea n basin an d highe r value s on th e continental margin. Local area s o f high conductivity may reflect present erosio n o r non-deposition . Loca l ver y high hea t flo w alon g th e presen t spreadin g axi s implies probabl e activ e volcanis m o r venting . Continental slop e maxim a o n th e M0r e an d western Barent s Se a margin s contras t greatl y with th e typica l lo w hea t flow , generall y les s than 75 mW m~2 , on old oceanic and thinned continental crust . Abnormall y hig h hea t flow , ex ceeding lOOOmWrcr 2 , has been measured o n the Hakon Mosb y gas-seepin g mu d volcan o o n the Barents Sea margin. Continental slop e maxima elsewhere may also be related to active seeps. We ar e particularl y gratefu l t o G . L . Johnson , fo r supporting th e acquisitio n i n 198 3 o f a Norwegia n heat-flow prob e dedicated t o high-latitude studies, and to El f Petroleu m Norg e fo r initiatin g and supporting the French-Norwegian FLUNORG E Project . We also thank P . R . Vog t an d K . Cran e fo r thei r effort s an d advice during a lon g period o f active collaboration.

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& TORP , J . E . 1987 . Measurements o f Heat Flow and Thermal Conductivity in the Nansen Basin, Arctic Ocean. Universit y of Bergen, Seismologica l Observatory: Cruis e Report . , MYHRE , A . M . & ELDHOLM, O . 1989 . Heat Flow Measurements on Norwegian Continental Margin during th e Flunorge Project. Universit y o f Bergen , Seismological Observatory , Seismo-series , 27 . S^TTEM, J . 1988 . Varmestromsmalinge r i Barent shavet. 18. Nordiske Geologiske Vintermode, Geological Survey o f Denmark (extende d abstract) . 406-408. TALWANI, M . & ELDHOLM , O . 1977 . Evolution o f th e Norwegian-Greenland Sea . Geological Society o f America Bulletin, 88 , 969-999. UDINTSEV, G . B . & LUBIMOVA , E. A . 1973 . Teplovy e potoki vbliz i Islandii . Izvestiya Akademii Nauk SSSR, Fizika Zemli, 11. VOGT, P. R. & SUNDVOR, E. 1996. Heat flow highs in the Norwegian-Barents-Svalbard continenta l slope ; deep crusta l fractures , dewatering , o r 'memor y in th e mud \ Geophysical Research Letters, 23 , 3571-3574. et al . 1997 . Haako n Mosb y mu d volcan o provides unusua l exampl e o f venting . EO S Transactions, American Geophysical Union, 78 , 549-557. VON HERZEN , R . P . & MAXWELL , A . E . 1959 . Th e measurements of thermal conductivity of deep-se a sediments b y a needl e prob e method . Journal o f Geophysical Research, 64 , 1557-1563 . WESSEL, P . & SMITH , W . H . F 1991 . Fre e softwar e helps ma p an d displa y data . EO S Transactions, American Geophysical Union, 72, 441-446. ZIELINSKI, G . W . 1979 . O n th e therma l evolutio n of passiv e continenta l margins , therma l dept h anomalies an d th e Norwegian-Greenlan d Sea . Journal o f Geophysical Research, 84 , 7577-7588 . , GUNLEIKSRUD , T. , S^TTEM , J., ZUIDBERG , H . M .

& GEISE , J . M . 1986 . Dee p hea t flo w measure ments in Quaternary sediments on th e Norwegian continental shelf . Offshore Techologv Conference Houston, T X 1986, 277-282 .

Atlantic volcani c margins: a comparative stud y O. ELDHOLM , T . P . GLADCZENKO , J . SKOGSEI D & S . PLANK E Department of Geology, University of Oslo, P.O. Box 1047 Blindern, N-0316 Oslo, Norway (e-mail: [email protected]) Abstract: Volcani c margin s i n th e Atlanti c Ocea n revea l a serie s o f commo n crusta l unit s and structura l feature s developed durin g continenta l extension an d break-up . W e sugges t that fou r mai n crustal zones can be recognized o n volcanic margins. This tectono-magmati c zonation implie s a histor y o f developmen t wher e tectoni c an d magmati c style s an d dimensions depen d o n th e interactio n o f lithospheri c an d asthenospheri c propertie s an d dynamics. Th e amoun t o f exces s igneou s activit y depend s o n th e temperatur e an d flui d content o f the asthenosphere alon g th e incipient plate boundary and th e dynamic history of the lithospher e durin g th e rif t phase . A n adequat e understandin g o f th e margi n histor y requires studie s o f th e entir e rift , i.e . th e conjugat e margins . W e als o not e tha t th e spectacular wedge s o f seaward-dippin g reflector s observe d alon g man y rifte d margin s ar e only one of many igneous features originating during the process o f break-up and initial seafloor spreading. Probably , mos t passiv e rifte d margin s represent intermediat e cases relativ e to th e volcanic an d non-volcani c end-members. A mantl e plume impinging on lithospher e already unde r extensio n emplacin g Larg e Igneou s Province-typ e initia l oceani c crust , including an extensive extrusive cover, is considered th e most likel y explanation for volcanic margins. Hydrocarbo n resourc e evaluation s o f volcani c margin s hav e t o includ e thei r characteristic tectono-magmati c feature s an d thei r consequence s fo r vertica l motion , erosion, sedimentation , therma l and buria l histories, and maturation . The mai n Norwegia n contributio n t o th e Eur opean Union-Joul e I I researc h projec t 'Inte grated Basi n Studies ' comprise s th e modul e 'Dynamics o f th e Norwegia n Margin ' (IBS DNM), focusing o n the development an d evolu tion o f sediment basin s i n intraplate an d passiv e margin setting s (N0ttvedt e t al. 2000). The North Atlantic an d Nort h Se a Mesozoi c rif t system s merge o n th e mid-Norwa y continenta l margi n where th e subsequen t Lat e Cretaceous-Paleo cene rif t episod e le d t o sea-floo r spreading , accompanied b y massiv e igneou s activity , nea r the Paleocene-Eocene transition . Evidence o f massive, transient igneou s activity during th e fina l break-u p o f continent s an d th e initial perio d o f sea-floo r spreadin g exist s fro m many othe r o f th e world' s passiv e continenta l margins (Fig . 1) . To evalut e th e processe s tha t govern th e inceptio n and evolutio n of such margins i t i s necessary to compar e crusta l structure, tectono-magmatic relation s an d th e histor y o f vertical motion . Consequently , th e them e 'com parative volcani c margi n studies ' becam e par t of IBS-DNM . Here , w e presen t result s mainl y from Atlanti c margin s wit h a n emphasi s o n crustal structur e an d tectono-magmati c styl e and dimensions . We hav e compile d a globa l volcani c margi n database fro m th e scientific literature and studies

of selecte d margin s suc h a s th e Nort h Atlanti c conjugate margin s nort h o f Charlie Gibb s Fracture Zone, th e North Namibi a margi n an d othe r South Atlanti c margi n segment s sout h o f th e Abutment an d Sa o Paulo plateaux , th e U S Eas t Coast margin, th e western margi n of f India, an d the wester n Australi a margi n (se e Fig . 1) . Globally, th e Nort h Atlanti c volcani c margin s (Fig. 2 ) ar e th e bes t explore d bot h b y geophy sical survey s an d drillin g (se e Eldhol m e t al . 1995; Skogsei d e t al . 2000) , an d w e us e thi s region t o defin e typica l volcanic margin tectono magmatic features .

Volcanic margin s The massive extrusive complexes along the rifted margin segment s of f Norwa y (Fig . 2 ) wer e firs t recognized b y a n exceptionall y smooth acousti c basement surfac e nea r th e continent-ocea n boundary (COB ) (Talwan i & Eldhol m 1972 , 1977). Later , multichanne l seismi c (MCS ) line s imaged wedge s o f seaward-dipping , intrabase ment reflector s (Hin z & Webe r 1976 ; Hin z & Schliiter 1978 ; Talwan i 1978 ; Eldhol m e t al . 1979). Simila r wedge s wer e als o recognize d elsewhere (e.g . Hin z 1981) , an d th e numbe r o f observations ha s steadil y increase d wit h th e

From: NOTTVEDT , A . e t al . (eds ) Dynamics o f th e Norwegian Margin. Geologica l Society , London , Specia l Publications, 167 , 411^428. l-86239-056-8/00/$15.0 0 © Th e Geologica l Society o f Londo n 2000.

412

O. ELDHOL M E T AL .

Fig. 1 . Top : passive continental margin classification s (Eldholm e l al. 1995) . Middle an d bottom : distribution of volcanic margin s base d o n reporte d intrabasemen t an d seaward-dippin g reflectors. acquisition o f high-qualit y seismi c line s o n th e outer margin s (Fig . 1) . Outsid e th e Nort h Atlantic, extensiv e an d voluminou s extrusiv e units exist along the US East Coas t (e.g. Talwani et al . 1995 ) and i n th e Sout h Atlanti c sout h o f the Abutmen t an d Sa o Paul o plateau s (Hin z

et al . 1995 ; Gladczenk o e t al . 19976) . Further more, wide-angle seismic experiments commonl y reveal a high-velocit y lower-crusta l bod y (LCB ) beneath th e extrusiv e cover . Durin g th e pas t decade thes e margin s hav e bee n classifie d a s volcanic margin s (e.g. Eldholm e t al . 1995) .

ATLANTIC VOLCANI C MARGIN S

413

Fig. 2. Nort h Atlantic volcanic margins with distribution of flood basalts (updated fro m Eldhol m & Grue 1994) , and location s of DSDP and OD P dril l site s sampling igneous basemen t rock s (see Table 1) . SDW, main wedge s of seaward-dippin g reflectors . Volcanic margins , continenta l floo d basalt s (CFB), oceani c plateau x an d ocea n basi n floo d basalts constitut e th e mai n categorie s o f tran sient larg e igneou s province s (LIP ) compose d of voluminou s construction s o f predominantl y mafic igneou s rock s tha t hav e no t bee n emplaced b y norma l sea-floo r spreading . Th e transient, large-scal e volcanis m i s commonl y attributed t o mantl e plume s (e.g . Whit e & McKenzie 1989 ; Dunca n & Richard s 1991 ; Larson 1991) . The dimensions and emplacemen t rates fo r som e volcani c margin s sho w tha t the y contribute significantl y t o th e globa l crusta l production budge t an d tha t the y ma y induc e environmental chang e (Coffi n & Eldholm 1994) .

North Atlanti c conjugat e rifte d margin s The Earl y Tertiar y continenta l break-u p an d onset o f sea-floo r spreadin g betwee n Eurasi a and Greenland , c . 55 Ma, wa s accompanie d b y massive transien t volcanis m emplacin g onshor e flood basalt s (Dicki n 1988 ) an d massiv e coeva l

extrusive an d intrusiv e roc k complexe s o n th e rifted margin s (Fig . 2) . The break-u p volcanis m took place , i n part subaerially , alon g mor e tha n 2600km o f th e earl y Tertiar y plat e boundary , with mos t o f th e lava s extrude d durin g Chro n 24r. The transient even t contrasts wit h persistent subaerial volcanis m fo r c . 60 Ma alon g th e Iceland plum e trai l leavin g th e Greenland-Ice land-Faeroe ridge between the conjugate Faeroe and Eas t Greenlan d CFB s (e.g . Eldhol m e t al. 1989) (Fig . 2) . Several scientifi c dril l hole s hav e recovere d rocks fro m th e seaward-dippin g wedge s o n th e Hatton Bank , V0rin g an d S E Greenlan d mar gins (Fig . 2 , Tabl e 1) . Th e wedge s consis t o f mainly tholeiiti c basal t an d thi n interbedde d sediments reflectin g a subaeria l and/o r shallow water constructiona l environment . Th e max imum penetratio n wa s achieve d a t OD P Sit e 642, whic h drille d throug h c . 800m o f basalt s and c . 130m int o underlyin g dacitic-andesiti c lavas an d interbedde d sediments . Although th e seaward-dippin g wedge s yiel d a characteristic seismi c image, which is commonly

414

O. ELDHOL M E T AL . Table 1 . Summary o f scientific volcanic margin drill sites sampling basement rocks i n th e North Atlantic (Fig. 2) Site

Water depth (m )

Sediment thickness (m )

Basement penetration (m )

Reference

DSDP 338 DSDP 342 DSDP 553 DSDP 553 DSDP 554 DSDP 555 ODP 64 2 ODP 64 3 ODP 91 3 ODP 91 5 ODP 91 7 ODP 91 8 ODP 98 8 ODP 98 9 ODP 99 0

1297.0 1303.0 2301 2329 2574 1659 1286 2753 3318.6 533.1 508.1 1868.2 262.6 554.6 541.5

400.8 153.2 282.7 499.35 126.6 927.32 315.2 565.2+ 770+ 196.8 41.9 1189.4 10 4 211.9

0.95 17.3 31.3 183 82.4 37 914.2 12.6 833.0 15.0 22.0 80.2 130.8

Talwani e t al. (1976 ) Talwani e t al . (1976 ) Roberts e t al . (1984 ) Roberts e t al . (1984 ) Roberts e t al . (1984 ) Roberts e t al . (1984 ) Eldholm e t al . (1987 ) Eldholm e t al . (1987 ) Myhre e t al . (1995 ) Larsen e t al . (1994 ) Larsen e t al . (1994 ) Larsen e t al . (1994 ) Duncan e t al . (1996 ) Duncan e t al . (1996 ) Duncan e t al . (1996 )

ODP site s 64 3 an d 913 , which terminate d jus t abov e basement , ar e include d becaus e the y provide dat a o n volcani c margi n subsidence . taken a s a criterion for volcani c margins, severa l other igneou s feature s relat e t o th e transien t break-up even t (Tabl e 2) . I n particular , th e basaltic lava s ma y exten d fo r larg e distance s landward o f th e wedges , an d ther e i s an appar ent spatia l correlation betwee n th e LC B an d th e most voluminou s extrusive rocks. Th e 10-2 0 km thick initia l oceanic crus t wes t o f th e CO B thin s to norma l oceani c thicknes s ove r a distanc e o f 50-150 km (Fig . 3) . Moreover , sill s an d dyke s intrude th e pre-Eocen e continenta l crust . Thes e observations documen t tha t th e volcanic margi n encompasses intrusiv e and extrusiv e features far beyond th e dippin g wedges (Eldholm et al. 1989; Skogseid & Eldhol m 1989) . Thus , th e Nort h Atlantic LI P include s flood basalts bot h onshor e and o n th e rifte d margin ; i n fact , th e volcani c margin constitute s it s majo r component . Th e

LCB an d crusta l intrusiv e companions t o flood basalt volcanis m mak e als o significan t contribu tions t o th e crusta l volume . Minimu m extrusive and tota l igneou s crusta l volume s ar e estimate d to b e 1. 8 x 10 6 km 3 an d 6. 6 x 10 6 km 3 . respec tively (Eldhol m & Gru e 1994 ) (Tabl e 3) . The tectoni c dimensions , base d o n structura l features, subsidenc e modellin g and crusta l thickness variations, have been discusse d b y Skogsei d et al . (19920) and Skogsei d (1994) . The y sugges t that Nort h Atlanti c volcanic margi n formatio n was precede d b y a rif t phas e lastin g fo r abou t 18-20 Ma befor e break-u p (Tabl e 3) . The litho spheric extension , whic h affecte d a 300k m wid e region, separate d Eurasi a an d Greenlan d b y about 140k m (Skogsei d 1994) .

Table 2 . Geological features associated with transient, igneous activity during break-up in the North Atlantic LIP (Skogseid & Eldholm 1995)

The Sout h Atlanti c break-u p occurre d a t c. 130 Ma of f Sout h Afric a an d progresse d northward. Th e oldes t identifie d anomal y of f Namibia i s M4 . c . 127 Ma (Rabinowit z & LaBrecque 1979) . Th e Paran a an d Etendek a CFBs (Fig . 4 ) hav e bee n date d t o 137-12 7 Ma (Turner e t al . 1994 ) an d 130-12 5 Ma (Erlan k et al. 1984 ; Milner et al. 1992) , respectively. They are linked by the Walvis Ridge-Rio Grande Rise along th e Trista n plum e trai l (O'Conno r & Duncan 1990) . Larg e extrusiv e construction s exist o n bot h conjugat e margin s sout h o f th e plume trai l (Tabl e 3 ) (e.g . Hin z e t al . 1995 : Abreu e t al. 1996 ; Condi e t al. 1996 ; Gladczenk o et al . 19976) .

Continental floo d basalt s associated intrusiv e rock s Extrusive complexe s alon g continent-ocean transition seaward-dipping wedge s an d sub-horizonta l units associated introsiv e rock s Sills an d low-angl e dyke s Volcanic vent s Regional tephr a horizon s High-velocity lower-crusta l bodie s (LCB ) Thick initia l oceani c crust

Namibia an d other South Atlantic margin s

ATLANTIC VOLCANI C MARGIN S

415

Fig. 3 . Simplifie d crustal margin transects . SD\V , main wedge s o f seaward-dipping reflectors ; LCB , lower-crusta l high-velocity, high-density body. The continent-ocean boundary is placed a t th e seaward terminatio n o f the bas e reflector belo w th e inne r dipping wedge . Transec t location s i n Fig s 2 and 4-6 .

O. ELDHOL M E T AL .

416

Figure 3 . (continued} Table 3 . Tectono-magmatic dimensions for volcanic margin LIPs Volcanic margi n o r LI P

North Atlantic* South Atlantic ! US Eas t Coas t Margin !

Tectonic dimension s

Magmatic dimension s

Rift width (km)

Rift duration (Ma)

Length (km)

Area (xl06knr)

Extrusive volume (x!06km3)

Total crusta l volume (x!06km3)

300 280 200

18-20 25 70

2600 2400 1000

1.3

2.0 2.4 0.19

6.6

0.72

*Eldholm & Gru e (1994) . jGladczenko e t al. (1997/7) . tGladczenko e t al . (1994) .

We hav e interprete d a gri d o f commercia l MCS line s on th e oute r Nort h Namibi a margin , which revea l prominent seaward-dippin g wedge s and othe r coeva l igneou s features . Th e mar gin i s divided int o fou r tectono-magmati c zone s (Figs 3 an d 5) : (1 ) oceani c crust ; (2 ) thickene d oceanic crus t covere d b y seaward-dippin g wedges; (3 ) a c . 150km wid e break-u p relate d Late Jurassic-Earl y Cretaceous rif t (BR , Fig . 3), partly covere d b y th e dippin g wedge s i n th e

west an d lav a flow s an d intrusion s t o th e east ; (4) thicker continenta l crust . Th e crus t i n zone s (3) an d (4 ) ha s undergon e earlier . Lat e Palaeo zoic extension . Faulting in zone (3) sediments records th e Late Jurassic-Early Cretaceou s riftin g whic h culmi nated wit h break-up. Centra l rif t uplif t le d to a n erosional rif t unconformity . Th e subsequen t break-up volcanis m caused initial , subaerial sea floor spreading , seaward-dippin g wedges , as well

Fig. 4 . Mai n Sout h Atlanti c structura l feature s (Cand e e t al 1989) , Etendek a an d Paran a CFB s (Milne r e t al. 1992 ; Turne r et al. 1994) , an d distributio n o f seaward dipping reflectors . Bathymetr y i n m (ETOPO- 5 1988) .

418

O. ELDHOL M E T AL .

as lavas , abundan t sills , an d low-angl e dyke s east o f th e COB . Moreover , th e Earl y Cretac eous sequenc e ma y als o contai n lava s fro m th e Etendeka CFB. Th e most voluminous volcanis m took plac e o n th e Abutmen t Plateau , reflectin g the proximit y o f th e plum e an d th e persisten t volcanism alon g th e plum e trail. The CO B i s placed a t th e wester n termination of th e rift ; w e not e tha t th e rif t unconformit y defines a bas e o f th e innermos t dippin g wedge s and i s absent farthe r west. Thus, the boundary i s landward o f magnetic anomal y G of Rabinowit z & LaBrecqu e (1979) , whic h lie s ove r th e mai n wedge (Fig . 5) . A dee p continuou s intracrusta l reflector tha t ma y imag e th e to p o f a n LC B (Fig. 3 ) i s observe d belo w th e Lat e Jurassic Early Cretaceou s rif t zone . Newl y acquire d refraction seismi c dat a sho w a c. 5 km thic k 7.1-7.5 km s"1 LC B belo w th e oute r margi n (Bauer & Schulz e 1996) .

The volcani c margi n of f North Namibi a con tinues t o th e sout h (Fig s 4 an d 5) . Seaward dipping reflector s hav e bee n reporte d betwee n Walvis Ba y an d Liiderit z (Austi n & Uchup i 1982), an d of f th e Orang e Rive r (Gerrar d & Smith 1983 ) and Capetow n (Hin z 1981) . Thus , the entir e >2400k m lon g easter n margi n ha s a volcanic signature. Hinz el al. (1995) showed tha t the conjugat e margi n of f Sout h America , fro m the Sa o Paul o Platea u t o th e Falklan d Escarp ment, ha s a simila r character. I n particular , th e Uruguay an d Argentin e margin s (Fig . 3 ) have a tectono-magmatic zonatio n simila r t o tha t o f Namibia, an d th e cross-sectiona l dimension s of th e dippin g wedge s ar e als o similar. Assuming tha t th e Nort h Namibi a margi n transect i n Fig . 3 is representative, th e extrusive volume i s c . 0.2 x 10 6 km 3 fo r th e margi n seg ment i n Fig . 5 . Volum e estimate s farthe r sout h are uncertain , bu t appea r smalle r pe r lengt h

Fig. 5 . Namibi a margi n tectono-magmati c zonation . MC S profile s fro m Intera-ECL8 9 9 1 and PG S Nope c surveys. Magneti c anomalie s fro m Rabinowit z & LaBrecqu e (1979) . Bathymetr y in metre s (GEBC O 1994) . BR. break-u p relate d rif t zone ; COB . continent-ocean boundary.

ATLANTIC VOLCANI C MARGIN S unit. W e estimat e a volum e o f c . 0.58 x 10 6 km3 for th e entir e margin , an d c.O.S x 10 6 km 3 fo r its conjugate . Th e entir e Sout h Atlanti c LIP , including th e Parana-Etendek a CFB s (Milne r el al 1992 ; Peate e t al . 1992) , has a n extrusive volume o f a t leas t 2.3 5 x 10 6 km3 (Tabl e 3) . The onse t o f riftin g leadin g t o break-u p i s proposed a t c . 160 Ma (Ulian a et al. 1989; Niirnberg & Muller 1991) , an d dynami c modelling o f lithospheric extensio n i n th e Parana-Etendek a region suggests a rift duratio n o f c. 25Ma (Harr y & Sawye r 1992) . Th e latte r perio d i s consisten t with seismi c dat a o n th e Namibia n shel f (Ligh t et al . 1993) . We estimat e rif t width s o f 12 0 and

419

150km of T Namibia an d Argentina , respectively; and tha t th e u p t o 300k m an d 2400k m lon g rift underwen t extensio n fo r c . 25 Ma befor e break-up.

US Eas t Coas t margi n The US East Coas t margin (Fig . 6 ) was initiated by break-u p o f Nort h Americ a an d Afric a following a Lat e Triassic-Earl y Jurassi c rif t episode (Klitgor d e t al . 1988) . I t i s covere d b y very thic k sediment s limitin g seismi c resolutio n in the deep basins an d th e underlying crust . Th e

Fig. 6 . Distibutio n of seaward-dippin g wedges on th e th e U S Eas t Coas t margi n (O h e t al . 1995 ; Talwani et al . 1995) wit h selected seismi c profiles use d fo r volum e estimates in Tabl e 3 . GEBCO (1994 ) bathmetry i n metres, East Coas t Magneti c Anomaly (ECMA) fro m Talwan i e t al . (1995), an d fractur e zone s and sea-floo r spreading anomalies fro m Klitgor d e t al . (1988). SMV, submarin e volcanic rocks interprete d b y Austi n et al . (1990).

420

O. ELDHOL M E T AL .

on- and offshor e rift basin s exten d ove r a 200 km wide zone acros s Chesapeake Ba y into th e Baltimore Troug h (Benso n & Doyl e 1988) . Distinc t magnetic an d gravit y anomal y belt s delineat e crustal feature s o n th e margi n (e.g . Rabinowit z 1974; Also p & Talwani 1984) . Seaward-dipping reflector s wer e image d b y Klitgord e t al. (1988 ) an d Austi n e t al (1990) , and a 7.2-7. 5 km s"1 LC B was mapped b y wideangle profile s i n th e Baltimor e Canyo n (LAS E Study Group 1986 ) and Carolina trough s (Trehu et al. 1989) . Recent survey s have led to improve d mapping o f geometries an d distributio n of these rock complexe s (Holbroo k & Keleme n 1993 ; Sheridan e t al . 1993 ; Holbroo k e t al . 1994
underlying crus t i s uncertai n (O h e t al . 1995) . Holbrook & Kelemen (1993 ) calculated th e tota l igneous crus t emplace d durin g break-u p t o b e 1.6 x 10 6 km3 . On th e other hand, i f we apply th e same criteri a a s i n th e Nort h Atlanti c (Eldhol m & Gru e 1994) , we arrive a t a mor e conservativ e estimate o f 0.7 2 x 10 6 km 3 . Although n o seaward-dippin g wedge s hav e been reporte d o n th e conjugat e margin , extru sive rock s ma y explai n th e linea r magneti c anomaly of f Morocc o (Steine r & Roese r 1996) . Furthermore, a 7.1-7. 4 km s"1 LC B appear s t o replace typica l Laye r 3 velocitie s i n th e oldes t oceanic crus t (Holi k e t al . 1991) .

Discussion The volcani c margi n histor y depend s o n litho spheric an d asthenospheri c propertie s before , during an d afte r continenta l break-up . There fore, on e ha s t o stud y th e entir e tectono magmatic break-u p history , i.e . conside r th e lithospheric settin g befor e th e onse t o f conti nental extension , th e histor y o f magmatis m an d tectonism durin g riftin g an d break-up , an d th e subsequent margi n subsidence . Thi s implie s consideration o f th e entir e crus t a t conjugat e margins; however , th e databas e t o achiev e this goa l i s a s ye t meagre , eve n a t th e bes t explored margins . We observe changes in tectono-magmatic style and dimension s alon g singl e margi n segment s and amon g differen t margins . These are ascribe d primarily to the lithospheric configuration befor e rifting, mod e o f rifting , magnitud e of the mantle thermal anomaly , an d distanc e fro m th e plume . None th e less , w e not e gros s similaritie s i n tectono-magmatic styl e and dimensions , an d i n main crustal units (Fig. 3. Table 3) . In particular, the continental crust undergoes extensio n before break-up, formin g a wid e rif t zon e (BR . Fig . 3) . Hence, w e apply a crusta l zonatio n comprising : (1) norma l oceani c crust ; (2 ) expande d oceani c crust; (3 ) pre - an d syn-rif t sediment s an d con tinental basemen t rock s tha t ar e extended , intruded an d locall y covere d b y floo d basalt ; (4) norma l continenta l crust . Zone s (2 ) an d (3 ) are underlai n b y a high-velocity LCB. an d zone s (3) an d (4 ) ma y hav e undergon e previou s tectonic events. The CO B i s place d a t th e zon e (2)-(3 ) boundary, seawar d o f whic h ther e i s no bas e t o the dippin g wedge . Whethe r a distinc t CO B exists o n rifte d margin s i s debatable. W e infe r a narrow CO B o n volcani c margin s correspond ing t o rapid , latera l compositiona l change s i n

ATLANTIC VOLCANI C MARGIN S

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Fig. 7. Velocity-dept h functio n fo r expande d volcani c margi n oceani c crus t i n th e Nort h Atlanti c (lin e A ) (Eldholm & Grue 1994 ) an d of f th e U S Eas t Coas t (lin e B ) (Holbrook e t al. \994b) compare d wit h norma l continental (lin e C ) (Christense n & Mooney 1995 ) an d oceani c (lin e D ) (Whit e e t al . 1992 ) crusts. Suggeste d compositional colum n o n right . the uppermos t crystallin e crust belo w th e extrusive rocks . Thus , th e CO B (Fig . 3 ) separate s intruded, thinne d continenta l crus t fro m rock s emplaced entirel y afte r break-up . Hence , mag netic profiles may delineat e the CO B i f break-up occurred durin g period s o f frequen t reversals . On th e othe r hand , i f the entire crust i s included, zone (2 ) and th e mos t intrude d par t o f zon e (3 ) may b e considere d a transitio n zone . The velocit y distribution , th e thic k extrusiv e cover largel y emplaced subaerially , and th e LC B distinguish zon e (2 ) from norma l oceani c crust . The three-laye r Nort h Atlanti c zone (2 ) crust o f Eldholm & Gru e (1994 ) consist s o f a n uppe r extrusive layer , a mid-crustal layer an d th e LCB (Fig. 7) . This typ e of crust i s similar to th e giant Ontong Jav a Platea u LI P (Gladczenk o e t al . 19970) an d t o othe r oceani c LIP s (Coffi n & Eldholm 1994) . Th e consisten t LI P velocit y structure ma y sugges t commo n emplacemen t and compositiona l elements .

Extrusive cover The dippin g wedge s i n th e Nort h Atlanti c consist o f u p t o 6k m thic k floo d basal t an d

very thi n interbedde d sediments . Th e velocit y increases rapidl y fro m c . 3.5 t o >5.0kms~ 1 i n the uppermos t par t wit h a gentle r velocit y gradient a t dept h (Fig . 7) . Velocitie s o f 6.0 6.5kms -1 nea r th e bas e o f th e thickes t dippin g wedges may sugges t a n increasin g proportion o f dykes wit h depth . Integratio n o f log , cor e an d seismic dat a fro m OD P Site s 64 2 (Planke 1994 ; Planke & Eldhol m 1994 ) an d 91 7 (Plank e & Cambray 1998 ) (Fig. 2 ) yields rock propertie s o f lavas an d interbedde d sediments . Plank e (1994 ) determined 0.6-18.5 m flo w thicknesse s a t Sit e 642, wherea s mos t sedimen t layer s ar e < l m . Physical propertie s a t flo w an d composite-flo w scale, an d seismi c modellin g sho w tha t mos t lavas ar e to o thi n an d withou t th e physica l property distributio n required t o produc e reflec tors resolve d b y standar d MC S surveys . Th e dipping reflector s appea r t o originat e fro m extensive, thic k individua l flows and fro m seis mic interference . Therefore, MC S dat a ar e no t suited fo r interpretin g th e detaile d interna l stratigraphy o f th e seaward-dippin g wedges . We als o not e tha t Site s 64 2 an d 91 7 a s wel l as vertical seismic profiling (VSP) experiments in Iceland sugges t tha t lav a velocitie s recorded o n

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the surfac e may b e 10-20 % to o hig h a s a resul t of transvers e isotropi c propertie s o f th e lava s (Planke & Eldhol m 1994 ; Plank e & Floven z 1994). an d tha t availabl e seismic profile s recor d 2D images whereas volcano an d fissure eruptions construct 3 D feature s (Eldholm et al. 1995). Sites 642 and 91 7 were drilled near th e feather edge o f a dippin g wedge , i.e . landwar d o f th e most typica l dipping reflectors . Though ther e i s seismic continuit y fro m th e site s t o th e mai n wedges, ther e ar e fe w distinct , extensiv e reflec tors a t th e site s proper . Therefore , thic k lav a series ma y exis t als o withou t distinc t intrabase ment reflectors , i.e. flood basalts ma y exten d fa r beyond th e prominen t wedges . I n fact , th e dipping wedg e i s onl y on e o f severa l igneou s features cause d b y th e break-u p even t (Tabl e 2 ) (e.g. Anderse n 1988 ; Wood e t al . 1988 ; Eldholm et al . 1989) . an d rifte d margin s ma y compris e extrusive construction s no t image d b y th e seismic record . Th e variet y i n seismi c styl e i s related t o volum e an d rat e o f magm a produc tion, constructiona l environment , an d syn - an d post-constructional deformatio n an d subsi dence. Fo r example , Plank e e t al . (1999 ) hav e proposed tha t th e varying seismic characteristics reflect a chang e fro m subaeria l floo d basal t through shallow-wate r hyaloclasti c mound s t o deep-water flows.

Middle and lower crust Zone 2 middle crust (Fig . 7 ) has a 6.5-6.7 km s^ 1 velocity at the top an d a gentle velocity gradient, resembling a thickene d oceani c laye r 3 A (Ewing & Hout z 1979 ; Whit e e t al . 1992) . I t probabl y consists of dykes at th e transition wit h the extrusive cove r an d gabbr o belo w (Zehnde r e t al . 1990). Th e 7 + kms- 1 LC B velocit y i s no t typical fo r norma l oceani c o r continenta l crust s (Meissner 1986 ; Christense n & Moone y 1995) , but i s characteristic o f LIP s (Coffi n & Eldhol m 1994). There is still uncertainty i n LCB geometr y and velocity , partl y becaus e o f dat a quality , acquisition an d interpretatio n techniques . Thus , one ha s t o b e carefu l i n usin g seismi c velocit y alone t o distinguis h crusta l typ e an d composi tion. Moreover , linearl y scale d model s o f 'normal' oceani c (Zehnde r e t al . 1990 ; Mutte r & Mutte r 1993 ) an d continenta l crust s hav e obvious geneti c implication s whic h ma y no t b e valid i n vie w o f mel t volum e an d emplacemen t setting fo r th e initia l oceani c crust . Ther e i s similarity o f th e uppe r an d middl e crus t i n zone (2) with Icelandic crus t (Mutte r et al. 1984) , thus the term Icelandi c oceani c crust has been applie d (e.g. Eldhol m e t al . 1989 ; Hin z e t al . 1993) .

We relat e zon e (2 ) t o LIP-typ e crusta l emplacement (Coffi n & Eldhol m 1994) . char acterized b y increase d decompressiona l partia l melting during break-up emplacing the LCB. th e middle crust i n zon e (2 ) and th e extrusiv e cover. High-quality expande d sprea d profil e (ESP ) and ocean botto m seismograp h (OBS ) experiments (e.g. Eldhol m & Mutte r 1986 ; Hin z e t al . 1987 ; Fowler e t al . 1989 ; Olafsso n e t al . 1992 ; Mjelde et al . 1993 ; Holbrook e t al . 19940 ) yiel d a rang e of velocities , 7.1-7.7kms~ 1 , fo r th e LCB . Th e fact tha t increase d Mg O conten t i n ponde d decompressional basalti c melt s a t th e bas e o f the crus t yield s onl y 7.1-7. 2 km s"1 velocitie s (White & McKenzie 1989 ) suggests that the LC B velocity rang e relate s t o a varying degree of melt fractionation. Th e uppe r LC B ma y represen t a transitio n fro m gabbr o t o olivin e cumulates derived fro m picriti c melts (Fig. 7). On th e othe r hand, th e 7 + kms" 1 velocit y ma y als o repre sent a secondary , metamorphi c facie s boundar y (Eldholm & Gru e 1994 ; Eldhol m e t al. 1995) . probably th e gabbro-garnet-granulit e transi tion. However , this process require s the presenc e of substantia l amount s o f fluid s shortl y afte r emplacement, fo r whic h a viabl e sourc e i s no t obvious (Gladczenk o e t al . \991a). LCB s ar e commonly describe d a s magmati c underplatin g (e.g. LAS E Stud y Grou p 1986 ; Whit e e t a l 1987), a proces s tha t refer s t o accumulatio n o f mantle-derived materia l below continenta l crust requiring a melt-crus t densit y contras t (Herz berg e t al . 1983 ; Furlon g & Fountai n 1986) . Because a densit y filte r i s no t applicabl e during oceanic crus t formatio n onl y the LC B belo w th e extended continenta l crus t i n zon e (3 ) i s trul y underplated (Fig . 3).

Tectono-magmatic dimensions For LIP s globally, our database allow s only firstorder volum e estimate s of the offshor e extrusive component an d o f th e tota l igneou s crust . Most volume s ar e considere d minimu m value s (Table 3) . I t i s notabl e tha t th e mai n contribu tion t o th e igneou s volume s a t volcani c margi n LIPs wit h coeva l CFBs , suc h a s th e Nort h an d South Atlanti c (Tabl e 3 ) an d th e Deccan Seychelles, i s foun d offshore , showin g thes e margins contribut e significantl y t o th e globa l LIP inventory . At margin s associate d wit h mantl e plume s there i s som e evidenc e o f narrowing , an d les s voluminous wedge s awa y fro m th e plume . Th e wedge is thickest, c. 15 km, of f the U S Eas t Coast (Holbrook e t al . 19940) , wher e i t extend s dow n to th e LCB . Individua l reflectors have als o bee n

ATLANTIC VOLCANI C MARGIN S interpreted t o thi s leve l o n th e Hatto n Ban k margin (Spence e t al. 1989). These reflectors ma y originate withi n gabbroi c rock s an d no t fro m extrusive rock . The margin s i n Fig . 3 sho w breaku p relate d rift zone s wit h 150-20 0 km half-widths , an d rifting appear s t o hav e laste d fo r 50-2 0 Ma before break-u p (Tabl e 3) . Th e Namibi a an d Voring margin s experience d on e o r mor e rif t episodes pre-datin g th e break-up rift . Hence , th e continental crystalline crust ma y be thinned ove r a wid e region , wherea s th e break-u p relate d rif t is les s tha t 350k m wide . Flo w o f lava s ont o continental crus t an d pervasiv e intrusion s inhi bit seismi c resolutio n an d commonl y hid e rif t structures. Th e apparen t lac k o f extensiona l features ha s le d t o model s o f ver y rapi d break up o f the continental lithospher e withou t signifi cant riftin g (Mutte r e t al . 1984 ; Larse n 1990 ; Hopper et al. 1992) . I n contrast, w e show that a protracted rif t phas e i s compatibl e wit h dat a from man y margins . Thus , a separat e tectoni c framework fo r volcani c margins is not required . The similarit y in structural style and dimensions of volcani c margins , non-volcani c margin s an d continental rift s make s u s sugges t tha t th e principal differenc e betwee n volcani c an d non volcanic margin s i s derive d fro m th e mel t potential o f th e asthenospher e durin g riftin g and break-up .

Margin asymmetry and rifting style There is magmatic and/o r tectonic asymmetry on many conjugate volcanic margins. The magmati c asymmetry, expresse d b y th e on - an d off-shor e extrusive an d LC B distributio n an d volume , may exis t alon g an d acros s th e initia l plat e boundary. I n th e Nort h Atlantic , th e are a an d volume o f basalt s o n continenta l crus t ar e greatest sout h o f Iceland , becomin g smalle r t o the nort h (Eldhol m & Grue 1994) . Thi s config uration ma y reflec t stepwis e propagation o f th e plate boundar y durin g break-u p resultin g i n diminished mel t potentia l northward . Extrusive across-plate-boundar y asymmetry , commonly show n b y distributio n o f CFB s an d dipping wedges , ma y reflec t th e positio n o f a mantle plum e wit h respec t t o th e lin e o f break up. The prominent wedges off the US East Coast without obviou s equivalent s o n th e conjugat e Morocco margi n ma y b e anothe r example . Th e present distributio n o f basalti c lava s ha s bee n related t o th e combine d effect s o f variabl e mel t production, vulnerabilit y of the continental crus t to mel t penetration , multipl e transien t feeders , lateral mel t migration , constructiona l environ ment an d erosio n (e.g . Eldhol m e t al . 1995) .

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The tectoni c style of the margi n is determined by th e pre-rif t lithospheri c settin g an d th e styl e of th e rif t deformatio n outline d b y faul t an d detachment distribution s and geometries, and by conjugate transfe r system s (Liste r e t al . 1991) . Crustal break-u p awa y fro m th e rif t axi s (Kee n 1987) wil l creat e asymmetri c margi n structures , as doe s simple-shea r extension proposed fo r th e US Eas t Coas t margi n (Benso n & Doyl e 1988 ; Klitgord e t al . 1988) . Simple-shea r extensio n and associate d syn-constructiona l listri c fault s may als o explai n th e abrup t seawar d termina tion sometimes observed a t large dipping wedge s (Eldholm e t al . 1989) .

Volcanic margins and mantle plumes A relationshi p between most LI P emplacement s and mantl e plumes , recognize d b y hotspots , i s well documente d (e.g . White & McKenzie 1989 ; Duncan & Richard s 1991) . Th e igneou s activ ity relate d t o th e transien t break-u p even t i s caused b y decompressiona l melting . I f a plum e reaches th e bas e o f th e lithospher e i n a regio n under extension, or i n a region with pre-existing thinned lithosphere , meltin g wil l b e amplifie d and th e exces s melt s ma y resul t i n a volcani c margin. Th e variabilit y in bot h extrusiv e cover and tota l igneou s crustal volum e emplaced dur ing break-u p lea d t o th e inferenc e that volcani c margins ar e expression s o f asthenospheri c mel t anomalies o f differen t magnitudes . Simila r rela tions appl y t o transien t LIP s i n genera l (Coffi n & Eldhol m 1994) . Notin g th e rang e i n size , Eldholm e t al . (1995 ) pointe d ou t tha t som e observations ma y challeng e th e plum e mode l a s the onl y mechanis m fo r volcani c margi n initia tion. Fo r example , th e length s o f th e volcani c rifted margin s in the North Atlantic (Fig . 2) and in th e Sout h Atlanti c (Fig . 4 ) require ver y larg e diameters fo r collapse d plum e heads . There ar e n o obviou s plume s t o explai n th e US Eas t Coas t (Holbroo k & Keleme n 1993 ; Talwani et al. 1995 ) and Wes t Australia margins (Mutter e t al . 1984 ; Hopper e t al . 1992 ; Colwell et al . 1994) , although a plum e relationshi p wa s inferred b y Wilson (1997) for the U S East Coas t and th e conjugat e Wes t Afric a margins . Th e inferred volcani c margin-plum e relationshi p i s commonly base d o n les s excessive , persisten t volcanism cause d b y th e tai l o f th e plume , an d recognized b y a submarin e ridg e o r seamoun t chain suc h a s th e Icelan d an d Trista n plum e trails expresse d b y th e Greenland-Iceland Faeroe ridg e (Fig . 3 ) and th e Walvis Ridge-Ri o Grande Rise , respectivel y (Fig. 4) . On th e othe r hand, th e plum e concep t ma y b e retaine d i f the

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plume sourc e generate s a mantle k blotT (Griffith s & Campbel l 1991 ) rathe r tha n a persisten t plume, o r i f th e plum e i s locate d beneat h a plate remainin g relativel y stationar y wit h respect to th e asthenosphere . Thus , a volcani c margi n may stil l hav e a dee p mantl e therma l sourc e without being located i n the vicinity of a hotspot . In vie w o f th e variet y i n siz e an d distributio n of igneou s volume s o n volcani c margins , w e prefer t o trea t th e mantl e plum e a s a sufficient , but no t necessary , conditio n fo r exces s igneou s activity durin g complet e plat e separation . Th e asthenospheric mel t potentia l i s determine d by th e therma l stat e an d flui d conten t i n th e asthenosphere and th e dynamic state of the lithosphere. i.e . magnitud e an d duratio n o f rifting . Consequently, th e combine d effec t o f smal l regional asthenospheri c temperatur e an d flui d content anomalie s i n th e asthenospher e an d lithospheric extension may induce exces s melting during break-up . Thi s concep t doe s no t depen d on a plum e o r a specifi c mantl e circulatio n model, althoug h th e existenc e o f a plum e wil l greatly facilitat e large-scal e melting, particularly if i t impinge s o n lithospher e tha t i s alread y under extension or ha s bee n thinne d by previous rift episode s (Thompso n & Gibso n 1991) . Th e similarity i n tectoni c styl e o f volcani c o r non volcanic margin s suggest s tha t th e classificatio n in Fig . 1 refers t o end-membe r type s an d tha t most margin s ar e intermediat e types . Th e mar gins of f W Australi a tha t hav e seaward-dippin g wedges and LCB s o f relatively small volumes are examples o f intermediate margin s (Mutte r e t al. 1989. Hoppe r e t al . 1992 ; Planke e t al . 1996) .

Implications fo r resourc e evaluation Large-scale transien t geologica l event s influenc e the palaeoenvironmen t b y changin g oceano graphic an d atmospheri c circulatio n pattern s and compositions . I n particular , th e syn-rif t uplift an d subsequen t massive , transien t volcan ism durin g break-u p affect s environment s o n local, regiona l an d possibl y globa l scale s b y modification o f basi n geometries , ne w deposi tional an d erosiona l environments , an d change s in th e compositio n o f th e biosphere . Th e effect s have been discusse d b y Coffi n & Eldholm (1994 ) for LIP s i n general, an d b y Eldholm & Thomas (1993) an d Eldhol m e t al . (1995 ) fo r volcani c margins i n particular . The potentiall y globa l environmenta l impac t of volcanic margin formatio n ha s been suggeste d for th e Nort h Atlantic , wher e sediment s sho w that th e floo d basal t emplacemen t nea r th e Paleocene-Eocene boundar y wa s accompanie d

Table 4. Tectono-magmatic an d depositional effects of volcanic margin formation having potential resource implications Pre- amd syn-rift (pre-opening > basins • Syn-rif t uplif t erosion redeposition restricted basin s • Therma l imprin t • Faultin g • Intrusiv e activity Post-opening (early opening) basins • Along - and across-margin barrier s restricted basi n high biogeni c productivity • Centra l sedimen t sourc e • Therma l imprin t • Floo d basalt s • Margi n subsidence LCB influenc e Primarily base d o n studie s of th e margi n off Norway (Fig . 2).

by regiona l ashfalls . Ther e i s als o a n apparen t temporal correspondenc e betwee n this boundary event an d th e globa l Paleocene-Eocen e extinc tion event . Subsequently , th e Eart h entere d th e early Eocen e greenhouse , th e warmes t perio d over th e pas t 7 0 Ma (Eldhol m & Thomas 1993) . In terms of hydrocarbon exploration , the crustal movement s an d therma l regim e associate d with volcani c margi n formatio n influenc e th e resource potentia l o f th e pre-openin g sedimen tary basins , i. e pre - an d syn-rif t basins , a s wel l as th e post-openin g margi n basin s (Tabl e 4) . In particular , the combination o f transfer faults , fracture zone s an d centra l rif t uplif t durin g th e syn-rift an d earl y post-rif t period s ma y for m across- an d along-margi n barrier s an d thu s develop a serie s o f restricte d basins , whic h in som e case s existe d ten s o f million s o f year s after break-up . Th e correspondence o f restricted basins an d period s o f globa l warmin g may . i n fact, induc e favourabl e condition s fo r sourc e rock formation . Few studie s hav e ye t addresse d thes e ques tions, except o n th e margin of f Norway, where it has bee n show n tha t th e LC B cause s a sig nificant, quantifiabl e reductio n i n the subsidence of th e oute r margi n (Skogsei d 1994) . Thus , th e LCB mus t b e include d durin g modellin g o f relative vertica l motio n an d subsidence-derive d lithospheric extension (Skogseid et al. 2000). The uplifted centra l region wa s eroded an d acte d a s a main sourc e o f Paleocene an d Eocen e sediment s into th e regiona l Vorin g an d Mor e basins , whereas sediment s fro m th e eas t firs t reache d

ATLANTIC VOLCANI C MARGIN S the high s i n mid-Eocen e time . Th e extru sive rock s becam e completel y sedimen t covere d as lat e a s mid-Oligocene t o earl y Miocen e time , i.e. c . 30 Ma afte r break-u p (Skogsei d & Eld holm 1989 ; Skogsei d e t al 19920,6) . Th e thermal imprin t i s almos t entirel y restricte d t o the part o f basins overlying the LCB. Here, there is u p t o 200 % increas e i n hea t flow , an d potential sourc e rock s reac h thei r maximu m temperature a fe w million year s afte r break-up , and retur n t o norma l therma l condition s 15 — 20 Ma late r (Pederse n e l al . 1996) . Moreover , modelling o f 100 m o r thicke r sill s show s considerable maturatio n increas e a t distance s 3-4 time s th e sil l thicknes s (Pederse n e t al . 1996). Thi s effect , documente d b y well s o n th e Exmouth Platea u o n th e wester n Australi a margin (Reeckman n & Mebberso n 1984) , wil l become eve n more importan t i f convective heat transport i s achieved. This stud y ha s benefite d fro m result s an d comment s from a number o f colleagues an d student s involve d i n continental margi n studie s a t th e Universit y o f Oslo . In additio n t o th e IB S institutiona l an d industr y partners, w e are gratefu l fo r advic e an d dat a suppor t from th e Australia n Geologica l Surve y Organization , Bundesanstalt fii r Geowissenschafte n un d Rohstoffe , Germany, an d PG S Nopec , Norway . Th e wor k ha s been supporte d b y the Researc h Counci l of Norway a s part o f th e IB S (Integrate d Basi n Studies ) projec t under th e JOUL E I I researc h programm e funde d b y the Commissio n o f Europea n Communitie s (Contrac t JOU2-CT 92-0110), an d i n par t b y th e ProPetr o research program. This paper i s an IB S Contribution .

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Crustal stress an d tectonics in Norwegian regions determined fro m earthquake foca l mechanism s CONRAD D . LINDHOLM, 1 HILMA R BUNGUM, 1 ERI K HICKS 1 & MARI O VILLAGRAN 2 1

2

NORSAR, Box 51, 2007 Kjeller, Norway Present address: NOC-SOREQ, Nuclear Research Centre of Israel, 81800 Yavne, Israel Abstract: A databas e o f 10 9 earthquake foca l mechanism s fo r Scandinavi a an d Svalbar d has bee n compiled , comprisin g 2 9 ne w solutions . Stres s direction s extracte d fro m thes e earthquake foca l mechanism s hav e bee n calculate d fo r southeaster n an d southwester n Norway, th e Norwegia n Se a an d th e Barents-Finnmar k region . Th e principa l horizonta l stress is c. N45°W, and indicate s a clockwise rotation a s one move s fro m sout h t o north , in accordance wit h stress trajectories tha t ca n b e expected fro m ridg e push. The dat a indicat e that i n souther n Norwa y norma l faultin g regimes ar e mor e dominan t i n onshor e regions , whereas thrus t faultin g dominate s th e offshor e seismi c activity . Offshor e mid-Norwa y a preference fo r norma l faultin g is also found, whereas the shallo w earthquakes i n Finnmark , northern Norway , ar e al l reverse . Foca l dept h distribution s wer e studie d i n th e norther n North Se a region, revealin g substantia l difference s with depths in the rang e 5-3 0 km .

In th e las t par t o f th e 1980 s th e Worl d Stres s Map (WSM ) project ha s made large amount s of crustal stres s dat a availabl e an d accessibl e (Zoback 1992) . O f al l th e dat a i n th e WS M database mor e tha n 50 % ste m fro m earth quakes, indicatin g ho w importan t tha t sourc e of informatio n i s o n a globa l scale . Th e WS M data hav e no w convincingl y demonstrate d tha t stresses, no t only at the plate margins, but also in intraplate areas , ar e closel y connected t o globa l tectonics i n genera l an d t o plat e motion s i n particular. Th e Europea n par t o f the WSM ha s been discusse d i n detai l b y Miille r e t al. (1992) and fo r Fennoscandi a b y Gregerse n (1992) . For th e norther n Nort h Se a are a Lindhol m et al. (1995) compared shallo w borehole inferre d stress direction s wit h direction s extracte d fro m earthquakes. Stresses i n th e crus t ar e describe d b y a stres s tensor wher e th e thre e principa l stresse s a\ (maximum) <7 2 (intermediate) and cr 3 (minimum) form a n orthogona l set . Th e stres s tenso r describes completel y th e stresse s i n a bod y b y magnitude an d directio n o f principa l axe s (e.g . Aki & Richard s 1980 ; Turcott e & Schuber t 1982). Th e magnitud e o f th e principa l stresse s can b e inferred b y in situ methods such as hydrofracturing an d overcorin g measurement s (Has t 1958, 1969 ; Stephanson et al. 1987) . In this study we hav e limite d th e descriptio n o f th e stres s tensor t o th e directio n of th e principa l axes.

The understandin g o f th e present-da y stres s patterns and the sources and mechanisms behind these have a direct bearing on the understanding of th e present-da y tectonic s (Bungu m e t al . 1991) a s wel l as practica l application s i n petro leum exploitation (Lindhol m e t al . 1995) . The present-da y crusta l deformatio n (th e brittle part ) i s reflecte d i n th e seismi c activity , which i s show n i n Fig . 1 . I t ca n b e see n tha t Fennoscandia ma y b e divide d i n thre e seismo tectonic activ e regions: th e western Norwa y an d northern Nort h Se a region; th e Norwegia n Se a with coasta l areas ; th e weakl y indicate d zon e from th e Oslofjord an d Skagerra k area s towards the northeas t alon g th e Swedis h coas t o f th e Gulf o f Bothni a an d furthe r northward s fro m the Bothnia n Bay . Th e larges t know n earth quake i n Fennoscandi a wa s the 3 1 Augus t 181 9 earthquake i n th e Ran a area , wit h an estimate d magnitude o f 5. 8 (Muir Woo d 1989) . The purpos e o f thi s stud y ha s bee n t o utiliz e earthquakes t o compil e a databas e o f ol d an d new foca l mechanisms , and t o synthesiz e regional stres s direction s fro m tha t database . Th e criteria use d fo r th e definitio n o f principa l horizontal compressio n (<J H) ar e th e sam e a s used i n th e WS M databas e (Zobac k 1992) . Th e study als o include s a n analysi s of foca l depth s for norther n Nort h Se a earthquakes. Wave-form dat a hav e bee n use d i n som e o f the investigations , togethe r wit h reading s an d

From: NOTTVEDT , A . e t al . (eds ) Dynamics o f th e Norwegian Margin. Geologica l Society , London , Specia l Publications, 167 , 429-439 . 1-86239-056-8/OO/ S 15.00 © Th e Geologica l Societ y o f Londo n 2000 .

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Fig. 1 . Earthquake s i n Scandinavi a fro m 180 0 to 1991 . Dat a fro m FENCA T (Ahjos & Usk i 1992).

locations o f event s i n th e norther n Nort h Sea . The data use d ste m fro m the Norwegian Seismi c Network (NSN ) an d fro m NORSA R arrays .

Sources o f data A foca l mechanism s databas e ha s bee n estab lished tha t currentl y comprise s 10 9 focal mech anism solution s fo r th e Fennoscandia n region . For Swedis h territor y onl y th e larges t earth quakes wer e included , a s take n fro m Arvidsso n & Kulhane k (1994) . Th e bul k o f th e dat a hav e otherwise bee n publishe d earlie r b y Bungu m et al. (1991), an d ar e also mostly include d i n th e WSM database . Othe r source s tha t hav e bee n used ar e Lindhol m e t al . (1995) , Bungu m & Lmdholm (1997 ) an d Hick s (1996) . Th e las t study include d 2 9 ne w solutions , whic h hav e been detaile d an d discusse d togethe r wit h al l older foca l mechanisms b y Hicks et al. (in prep.), and onl y th e 1 4 that wer e compile d withi n th e frame o f th e IB S projec t ar e detaile d below .

Figure 2 give s in thi s respec t a n overvie w of al l the foca l mechanisms use d i n this study (see also Lindholm e t al . (i n prep.)) , an d th e 1 4 solutions are liste d in Table 1 (of which most ar e show n in Fig. 3) . Most o f th e foca l mechanis m solution s wer e obtained onl y b y mean s o f firs t motion polarity information, som e o f th e olde r one s wer e constrained b y amplitud e rati o information , and som e o f th e ne w solution s (Tabl e 1 ) hav e been constraine d b y wave-for m modelling . Al l of the solutions have been evaluate d with respect to qualit y on a scal e fro m A (good ) t o D (poor). The qualit y schem e suggeste d b y th e WS M project wa s applied, but ther e is also a significan t amount o f exper t judgement i n thi s evaluation .

Results From th e larg e databas e wit h 10 9 Fennoscan dian an d Svalbar d foca l mechanis m solutions . 77 were located withi n the fou r regions show n in

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Fig. 2 . Earthquak e foca l mechanisms i n Scandinavia use d in this study. The four regions discussed i n the text ar e indicated. Structura l element s fro m Blysta d e t al. (1989), Brekke et al. (1989) and Gabrielse n e t al. (1990). Table 1 . Recent earthquake focal mechanism solutions Date Date

Lat.

Lon.

D

M

Ti

T2

P\

P2

1990 226 1990 51 6 1991 9 1 1991 1231 1992 21 9 1992 41 4 1992 630 1993 12 0 1993 62 6 1993 91 3 19931227 1994 72 7 199411 19 1995 2 6

57.67 66.04 79.02 61.98 59.27 59.50 60.88 64.75 62.61 66.37 61.25 62.63 60.17 59.84

6.91 6.26 3.59 4.23 10.88 5.66 11.53 4.81 4.14 5.72 2.84 3.90 11.06 6.51

6.4 30 10 15 10 12 12 15 17 20 20 10 13 10

3.4 3.4 5.0 3.3 3.6 3.0 2.8 3.5 3.9 3.9 3.3 3.7 3.5 3.0

252 225 84 210 258 202 299 175 280 7 0 66 185 6

39 1 0 5 9 30 5 90 37 9 58

139 135 174 30 160 292 182 285 180 246 121 334 68 115

26 10 0 85 39 0 78 0 12 72 18 30 54 44

9

18 19

c c

c

c c c B B

C B B

C C B

C

Date, year , month , day ; Lat., latitude ; Lon. , longitude ; Z> , dept h (km) ; M , magnitude ; 7^ , azimuth o f T-axis ; T 2, di p o f T-axis ; P\, azimut h o f P-axis ; P 2, di p o f /^-axis ; Q, quality evaluation (A-D , wher e A i s best).

Fig. 2 . O f these , 3 8 an d 2 1 solution s wer e i n southwestern an d southeaster n Norway , respec tively, an d 1 3 and 5 solution s wer e i n th e tw o northern area s a s indicate d in Fig . 2 . For al l th e region s ros e diagram s o f th e directions o f th e principa l horizonta l compres -

sions (CT H) hav e bee n prepared . Th e stres s direction ha s bee n calculate d b y followin g th e rules use d i n th e WS M projec t (Zobac k 1992) . For eac h regio n a ros e diagra m o f
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C. D . LINDHOL M E T AL . From th e triangl e diagra m inse t i n Fig . 4 it is evident tha t norma l an d strike-sli p faultin g processes ar e relativel y more abundan t i n southeastern Norwa y tha n revers e faulting , an d tha t the maximu m principa l horizonta l stres s i s between N10 0 E an d N150 : E.

Southwestern Norway and Adjacent seas As pointe d ou t i n earlie r paper s (e.g . Huseby e et al . 1978 ; Bungu m e t al . 1991) . this regio n i s among th e seismicall y mos t activ e i n north western Europe , wit h frequent magnitude 4 an d stronger earthquakes . Th e largest recen t one was a magnitud e 5. 3 earthquak e tha t occurre d o n 23 Januar y 198 9 (Hanse n e t al . 1989) . Never theless, th e calculatio n o f foca l mechanism s i n this regio n ar e ofte n difficul t becaus e th e statio n coverage i s i n genera l limite d t o th e coasta l stations i n wester n Norway.

Fig. 3. Th e 2 3 southernmost o f th e 2 9 most recen t earthquake foca l mechanism s fo r souther n Norway .

grouped i n bin s o f 10" , wit h the numbe r o f dat a points withi n eac h grou p easil y countabl e (se e also Fejersko v e t al. in prep.) . To illustrat e th e typ e o f faultin g involve d w e have use d triangl e plot s a s introduce d b y Frohlich & Apperson (1992) .

Southeastern Norway Southeastern Norwa y experience d i n 190 4 th e most damagin g earthquak e i n thi s centur y i n Fennoscandia. Locate d i n th e Oslofjor d region , it i s thought t o hav e occurred o n th e flan k o f the Oslo Graben . Thre e earthquake s hav e bee n fel t in thi s region , o n 1 9 February 199 2 (M L 3. 3 i n Fredrikstad), o n 3 0 Jun e 199 2 (M L 2. 1 nea r Odalen), an d o n 1 9 Novembe r 199 4 (M L 3. 2 near K10fta) . Th e activit y rat e i n th e Osl o Gra ben are a is , however , lo w compare d wit h wes t coast seismicity , even thoug h thes e ne w event s substantiate th e earlie r observatio n o f Bungu m et al . (1991) that thi s area i s more activ e tha n it s surroundings. Twenty-on e foca l mechanism s (Fig. 4 ) hav e unti l no w bee n calculate d fo r thi s area, largely because the are a is particularly wel l covered wit h seismi c instruments .

Earthquake focal depths. Earthquak e foca l depth i s on e o f th e mos t difficul t factor s t o determine with sufficient precision . This is a wellrecognized seismological problem, an d i s valid as much fo r th e norther n Nort h Se a area a s it is for other part s o f th e world . Th e reaso n fo r thi s is tha t mos t o f th e station s ar e onshore , an d often a t distance s between 20 0 and 40 0 km fro m the epicentre , whereas accurate depth s generally require station s withi n distance s o f twic e th e focal depth . Som e attempt s hav e bee n mad e to stud y earthquak e dept h i n wester n Norwa y (e.g. Engell-Sorense n & Havsko v 1986 ; Han sen e t al . 1989 ; Bungu m e t al . 1991) , bu t th e depths foun d s o fa r hav e largel y bee n onl y indicative. I n southeaster n Norway , earthquak e focal depth s coul d i n man y case s b e determine d with accurac y t o th e lowe r crust . Thi s highe r location accurac y fo r th e smalle r earthquakes i n the NORSA R sitin g are a i s a resul t of th e dens e instrumental coverage . Since a reliabl e foca l dept h estimat e i s a prerequisite fo r a reliabl e foca l mechanis m solution, w e hav e i n thi s stud y analyse d 50 0 earthquakes wit h respec t t o foca l dept h i n th e northern Nort h Sea , coverin g th e perio d 1984 1994. Th e event s wer e binne d i n roughl y l O O k m x 100k m bin s a s show n i n Fig . 5 an d then locate d wit h fixe d dept h i n 1 km dept h intervals {usin g differen t locatio n method s an d crustal velocit y models) . Th e roo t mea n squar e (r.m.s.) value s o f th e trave l tim e residual s wer e then calculate d and summe d fo r eac h bin . Only first P and S arrival times (i.e. no guide d phases ) were used , an d becaus e o f th e station-hypocen tre distance s thes e d o i n almos t al l case s correspond t o Moh o refracte d phases .

Fig. 4 . Epicentre s wit h focal mechanisms i n southeastern Norway . Th e ros e diagra m show s th e directions o f the principal horizonta l compressions .

Fig. 5 . Dept h distributio n fo r earthquake s i n th e norther n Nort h se a calculated b y minimizatio n o f trave l tim e residuals. Curve s sho w th e trave l tim e residual s fo r tw o differen t velocit y models, an d th e number s indicat e number o f event s behin d th e calculation s i n eac h bi n (2 ° x 1°) . Horizontal scale s indicates : foca l dept h (km) ; vertical scal e indicate s r.m.s . o f trave l residuals.

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C. D . LINDHOL M E T AL .

Fig. 6. Foca l mechanism s i n the norther n Nort h Sea . The ros e diagra m show s th e direction s o f th e principa l horizontal compresions , an d th e triangl e plot th e sens e o f faultin g involved . (See also Fig . 3 for detaile d presentation o f the foca l mechanis m beachballs.) The result s are show n i n Fig. 5 where fo r eac h bin the sum of the r.m.s. travel time residuals are plotted a s function o f foca l depth . I t i s seen tha t the best fitting earthquake depth s (lowest su m of r.m.s.) within each bi n ar e clearl y different, wit h shallow depth s (5-1 5 km) o n th e Hord a Plat form an d eastwar d int o th e coasta l areas , an d with deepe r event s (20-3 0 km) a t th e wester n flank o f th e Vikin g Grabe n an d northwar d towards th e Tampe n area . Th e summe d r.m.s . values v . depth s i n Fig . 5 ar e average d an d weighted number s fro m larg e populations , an d therefore d o no t exclud e th e possibilit y o f dee p events i n a shallo w zon e an d vic e versa . Whe n evaluating Fig . 5 it i s also importan t t o bea r i n mind tha t th e dept h resolutio n (an d locatio n precision i n general) is higher close r t o th e coas t (where th e station s are ) tha n furthe r offshore . Earthquake focal mechanisms. Figur e 6 show s the foca l mechanism s wit h a triangl e plo t an d a stress ros e diagra m a s insets . Th e figur e show s a ver y distinc t geographica l distribution , illus trating th e Stord-Hardangerfjord cluste r (which

has lon g been recognized a s an area with few 7 bu t notable earthquakes) , an d th e are a northwes t o f the Sognefjor d mouth , whic h i s possibl y th e most seismicall y activ e regio n i n Fennoscandi a at present , wit h earthquake s greate r tha n o r equal t o magnitud e 4. 0 nearl y ever y year . From th e inse t ros e diagra m (Fig . 6 ) i t i s clearly see n tha t th e principa l horizonta l stres s direction i s WNW-ESE , whic h compare s wel l with th e expecte d directio n i f ridg e pus h i s assumed a s th e principa l stress-generatin g mechanism. Lindhol m e t al. (1995 ) investigated this are a i n mor e detail , als o includin g bore hole breakou t measurements , an d conclude d that ther e i s a wea k anomalou s tren d i n th e Tampen-Sogn Grabe n area , wher e CT H seem s to b e rotate d 9 0 norma l t o th e regiona l tren d shown i n Fig . 6 . Thi s stres s rotatio n wa s originally note d b y Bungu m e l al . (1991) . The typ e o f faultin g involve d i n thi s area ca n be see n fro m th e triangl e inset in Fig . 6 . Reverse and obliqu e strike-sli p faultin g i s clearl y domi nant, als o indicatin g a n overal l compressiv e stress regime . However , thi s conclusion require s

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Fig. 7. Firs t motio n distributio n fro m a selecte d grou p o f 1 2 earthquakes i n th e Sog n Grabe n area .

a qualification , a s bot h Bungu m e t al. (1991 ) and Lindhol m e t al . (1995 ) showe d tha t ther e are indication s tha t activit y i n shallo w Stord Hardanger i s dominate d mor e b y norma l faulting processe s tha n i n th e offshor e region , implying th e possibilit y o f greate r glacia l rebound effect s i n thi s region . A stres s mode l supporting thi s wa s presente d b y Stei n e t al . (1989), an d ha s bee n furthe r discusse d b y Fejer skov & Lindhol m (thi s volume). Earthquake firs t motion distribution. A furthe r attempt t o ma p stres s direction s wa s don e b y selecting earthquake s fro m a smal l regio n a s shown i n Fig . 7 an d rea d th e polarit y o f firs t P-wave motion s fro m al l th e station s wheneve r possible (Hick s 1996) . Th e quadran t i n Fig . 7 only show s th e expecte d directio n fo r com pressions an d dilatation s fro m a hypothesize d strike-slip foca l mechanis m wit h th e principa l horizontal compressio n i n a NW-S E direction . As ca n b e see n fro m Fig . 7, th e relativ e dominance o f observe d dilatation s support s thi s assumption i n general , bu t wit h th e FO O (station a t Floro , Fig . 7) bein g anomalous . Thi s i s

expected i f th e earthquake s i n thi s regio n ar e sub-Moho events , a s evidence d als o fo r th e Ms 5. 1 earthquak e o f 23 January 1989 , a t 62°N , 4.4°E, dept h 26.2k m (Hanse n e t al . 1989) . Eve n though individua l focal mechanism s canno t b e calculated fo r mos t o f these events , th e statistic s of th e firs t motion s ar e indicativ e o f compres sion i n th e uppe r lef t quadrant , i.e . as expecte d from ridg e push . Ther e ar e reason s t o believ e that man y differen t type s of faultin g d o occu r in the Sog n Grabe n region , bu t th e dat a i n Fig . 7 indicate tha t ther e i s a dominanc e o f tectoni c activity leading to horizonta l compression i n the NW quadrant .

Mid-Norway Thirteen earthquakes , generall y alon g th e continental margin , are show n i n th e boxe d are a i n Fig. 8 . The principa l horizonta l stres s directio n is relativel y consisten t an d slightl y clockwis e rotated compare d wit h th e norther n Nort h Sea are a (Fig . 6 ) an d southeastern.-Norway . The bul k o f th e dat a follo w th e mai n tectoni c

436

C. D . LINDHOL M E T AL .

Fig. 8. Earthquak e foca l mechanism s i n th e mid-Norwa y region . Th e ros e diagra m show^ s th e direction s o f th e principal horizonta l compressions . structures alon g th e Norwegia n margi n (Fig . 2), and bot h norma l an d revers e faulting is observed in th e area . The combinatio n o f bot h revers e faultin g an d normal faultin g yielding th e sam e <J H directio n (NW-SE) is unique. A tentativ e explanatio n (a s pointed ou t b y Byrkjelan d et al. (in press)) ma y be connecte d t o th e fac t tha t a clea r correlation exists betwee n rapidl y deposite d post - Miocen e sediments an d th e seismicit y distribution o n th e mid-Norwegian shelf . A subsidin g basi n wil l cause gravit y fault s t o develo p a t basi n mar gins, an d i f thi s proces s i s place d i n a regiona l NW-SE compressiona l stres s regime, only those gravity fault s tha t hav e a strik e nea r NW-S E will develop .

Northern Norway and the Southern Barents Sea Region In th e Barent s Sea region th e seismicit y is generally ver y lo w excep t fo r th e activit y alon g the mai n fractur e zones; also , th e statio n cover age i s poo r i n thi s region , an d thes e factor s

together explai n the fe w (five ) foca l mechanisms in Fig . 9 . Also, th e solution s show n i n Fig . 9 are not a s wel l constraine d a s on e woul d normall y require (Bungu m & Lindhol m 1997) . Never theless, th e principa l horizonta l stres s directio n (weakly indicated) seem s to b e clockwise rotate d compared wit h Fig . 8 an d thre e o f th e fou r Finnmark event s tha t occurre d o n o r nea r th e Stuoragurra neotectoni c faul t ar e indicatin g reverse faultin g i n goo d agreemen t wit h surface expressions an d borehol e result s (Bungu m & Lindholm 1997) .

Discussion Even i f detaile d comparativ e studie s ar e now possible, usin g th e presen t database , w e ar e i n this stud y focusing only on th e regiona l picture. The ros e diagram s o f th e principa l horizonta l stress fo r eac h regio n ar e plotte d i n Fig . 1 0 together wit h stress trajectories generated b y th e ridge pus h force , assumin g a spreadin g pol e a t 47.3 : N an d 123.0 :E for the mid-Atlantic spread ing axis.

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Fig. 9 . Foca l mechanism s i n th e Norwegia n Sea. Th e ros e diagra m show s the direction s of th e principal horizontal compressions.

Fig. 10 . Stres s trajectories an d ros e diagrams fo r eac h o f the fou r region s discusse d i n the text . The smalle r ros e dirgam i s from nin e northern Swede n earthquake foca l mechanisms.

We kno w fro m th e dat a tha t considerabl e uncertainty i n th e individua l foca l mechanism s exists, bu t eve n the n th e extracte d principa l stress direction s ar e surprisingl y stable . More over, ther e seem s t o b e a systemati c clockwis e rotation whe n movin g fro m southwester n t o

northern Norway , whic h matche s th e rotatio n that woul d b e expected i f the ridg e push forc e i s the principa l stress-generatin g mechanis m (se e Bott 1991) . Figur e 1 0 demonstrate s tha t ridg e push should , als o i n terms o f stress direction, b e regarded a s a majo r (first-order ) sourc e fo r th e

438

C. D . LINDHOL M E T AL .

observed stresses , ye t withou t underestimatin g possible (an d likely ) contribution s fro m othe r (higher-order) stres s sources . An interestin g detail is shown i n Fig. 1 0 by th e rose diagra m o f stres s direction s extracte d fro m nine Swedis h earthquake s (Arvidsso n & Kulha nek 1994) . Th e E- W directio n foun d fo r thi s region i s nearl y perpendicula r t o th e Fennos candian uplif t centr e i n th e inne r Gul f o f Bothnia, indicatin g that thi s an d possibl y othe r more loca l source s o f stres s i n thi s cas e ar e obscuring th e mor e regiona l ridg e pus h correla tion generall y found elsewhere . With respec t t o th e faultin g process involve d in th e databas e a significan t differenc e betwee n the earthquake s offshor e an d onshor e souther n Norway i s found , wit h norma l faultin g bein g more abundan t i n th e onshor e area s (compris ing thos e i n southeaster n Norwa y an d thos e in th e Stor d area) . I n th e Stor d are a th e events see m t o occu r a t systematicall y shallower depths tha n offshore , an d ar e als o mor e shallo w than th e average dept h i n the Oslo Graben area . In th e mid-Norwegia n are a norma l an d revers e faulting i s foun d sid e b y side , bu t wit h a con sistent stres s NW-S E direction . I n th e offshor e areas (an d i n Finnmark) , revers e an d strike-sli p faulting i s th e mor e commo n process . The correlatio n betwee n seismi c activit y an d main fault s o n a regiona l scal e i s generall y high, an d ca n b e seen whe n comparing th e map s of Larse n e l al. (i n prep. ) an d Lindhol m e t al . (in prep.) .

Conclusions From th e curren t stud y th e followin g ca n b e concluded: • Th e databas e o f 10 9 Fennoscandian earth quake foca l mechanisms confirm s the overall stress pictur e tha t wa s establishe d throug h the WS M projec t (Gregerse n 1992 ; Mulle r et al . 1992) . • A clockwis e rotatio n o f th e stres s directio n when movin g fro m sout h t o nort h i s indicated b y th e data . Thi s ca n b e explaine d as a result of ridge push ; however , mor e dat a (especially i n norther n Norway ) ar e neede d before mor e firm conclusions in this direction can b e drawn . • Earthquak e foca l depth s i n th e norther n North Se a see m t o b e deepe r tha n thos e onshore, an d differenc e i n depth ma y b e cor related wit h a differenc e i n tectoni c regime . • Deepe r northern Nort h Se a events more often show revers e faulting , wherea s mor e shal low foc i onshor e (Stor d area ) earthquake s

indicate preferences fo r normal faulting . One geodynamic mode l explainin g thi s observa tion i s the glacia l reboun d o f Fennoscandia . • Earthquake s in southeastern Norwa y exhibit a preferenc e fo r strike-sli p an d norma l faulting regimes . This pape r ha s focuse d o n th e presentatio n o f data fro m earthquake s tha t ca n revea l th e crustal stat e o f stres s i n Norwegia n regions . Fejerskov & Lindhol m (thi s volume) have given a mor e detaile d discussio n o f possibl e geody namic models , wit h incorporatio n als o o f stres s data fro m borehol e breakouts . The author s ar e gratefu l t o B . T . Larse n an d M. Fejersko v fo r the many comments an d suggestions , and w e than k E . Bogoslowsk i fo r th e linguisti c comments. Thi s wor k wa s supporte d b y th e Researc h Council o f Norway , a s par t o f th e IB S (Integrate d Basin Studies ) projec t unde r th e Joul e I I researc h programme funde d b y th e Commissio n o f Europea n Communities (Contrac t JOU2-C T 92-0110).

References AHJOS. T . & USKI . M . 1992 . Earthquake s i n northern Europe i n 1375-1989 . Tectonophysics. 207 . 1-23. AKI. K . & RICHARDS . P . G . 1980 . Quantitative Seismology; Theory an d Methods. Freeman . Sa n Francisco. CA . ARVIDSSON. R . & KULHANEK . O . 1994 . Seismodynamics o f Swede n deducte d fro m earthquak e focal mechanisms . Geophysical Journal International 116 . 377-392. BLVSTAD, P. . F.^ERSETH . R. . LARSEN . B . T.. SKOGSEID . J. & TORUDBAKKEN . B . 1989 . Nomenclature o f tectonic unit s i n th e Norwegia n Sea betwee n 6 2 North an d 69. 5 North . In : Structural an d Tectonic Modelling and its Application to Petroleum Geologv. Norwegia n Petroleu m Directorat e Conference. 18-2 0 October 1989 . Stavanger . BOTT, M . H . P . 1991 . Ridg e push an d associate d plat e interior stres s i n norma l an d ho t spo t regions . Tectonophysics. 200 . 1 7 32 . BREKKE. H. . F/ERSETH . R. . GABRIELSEN . R.. COWERS . M. B . & PEGRUM . R . M . 1989 . Nomenclature o f tectonic unit s i n th e Norwegia n Nort h Se a sout h of 6 2 degree s north . In : Structural and TectonicModelling and its Application to Petroleum Geologr. Norwegia n Petroleu m Directorat e Confer ence. 18-2 0 Octobe r 1989 . Stavanger . BUNGUM. H . & LINDHOLM . C . D . 1997 . Seismo - an d neotectonics in Finnmark. Kola, an d i n the southern Barent s Sea . Par t 2 : Seismologica l analysi s and seismotectonics . Tectonophysics. 270 . 15-28. . ALSAKER . A. . KVAMME . L . B . & HANSEN . R . A . 1991. Seismicit y an d seismotectonic s o f Norwa y and nearb y continenta l shel f areas . Journal o f Geophysical Research. 96 . 2249-2265 . BYRKJELAND. U. . BUNGUM . H . & ELDHOLM . O . 2000 . Seismotectonics o f th e Norwegia n Continenta l Margin. Journal o f Geophysical Research. I n press .

CRUSTAL STRES S AND TECTONIC S IN NORWEGIA N REGION S ENGELL-S0RENSEN, L . & HAVSKOV , J . 1986 . Recen t North Se a seismicit y studies. Physics o f th e Earth and Planetary Interiors, 45 , 37-44. FEJERSKOV, M . & LINDHOLM , C . 2000 . Crusta l stress i n an d aroun d Norway : a n evaluatio n of stress-generatin g mechanisms. This volume. ,, LARSEN , B . T . & NOTTVEDT , A . 2000 . Stress Map . Nort h Atlantic Area. I n preparation . FROHLICH, C . & APPERSON , K . D . 1992 . Earthquak e focal mechanisms , momen t tensors , an d th e consistency o f seismi c activity near plat e bound aries. Tectonics, 11 , 279-296. GABRIELSEN, R. , F^ERSETH , R. , JENSEN , L . N. , KALHEIM, J. E . & Rns, F. 1990 . Structural Elements o f the Norwegian Continental Shelf. Part I: The Barents Se a Region. Norwegia n Petroleu m Direc torate Bulletin , 6. GREGERSEN, S . 1992 . Crustal stres s regim e i n Fennos candia fro m foca l mechanisms . Journal o f Geophysical Research, 97, 1 1 821-11 827. HANSEN, R . A. , BUNGUM , H . & ALSAKER , A . 1989 . Three recen t larg e earthquake s offshor e Norway. Terra Nova, 1 , 284-295. HAST, N . 1958 . Th e Measurement o f Rock Pressure in Mines. Sverige s Geologisk a Undersokning , Serie s C, Arsbok, 52(3) . 1969. Th e stat e o f stres s i n th e uppe r par t o f th e earth's crust . Tectonophysics,^, 169-211 . HICKS, E . 1996 . Crustal stresses i n Norway an d surrounding areas as derived from earthquake focal mechanism solutions and in situ stress measurements. Cand.Scient . thesis , Universit y of Oslo . , BUNGUM , H . & LINDHOLM , C . 1999 . Stres s inversion o f earthquak e foca l mechanis m solu tions fro m onshor e an d offshor e Norway . I n preparation. HUSEBYE, E . S. , BUNGUM , H., FYEN , J . & GJOYSTDAL , H. 1978 . Earthquak e activit y i n Fennoscandi a between 149 7 and 197 5 and intraplat e tectonics. Norsk Geologisk Tidsskrift, 58 , 51-68 .

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LARSEN, B . T., NOTTVEDT , A., VAGNES , E . e t al. 1999. Tectonic map . Nort h Atlanti c area. This volume. LINDHOLM, C . D. , BUNGUM , H. , BRATLI , R . K. , AADNOY, B . S. , DAHL , N. , TORUDBAKKEN , B . & ATAKAN, K . 1995 . Crustal stres s i n th e norther n North Se a a s inferre d fro m borehol e breakout s and earthquak e focal mechanisms. Terra Nova, 1, 51-59. ,, HICKS , E. , FEJERSKOV , M. , OLESEN , O. , LARSEN, B . T . & NOTTVEDT , A . 1999 . Seismicity, earthquake foca l mechanism s an d neotectonics , North Atlanti c area. I n preparatio n MUIR WOOD , R . 1989 . The Scandinavia n earthquake s of 2 2 Decembe r 175 9 an d 3 1 Augus t 1819 . Disasters, 12(3) , 223-236 . MULLER, B. , ZOBACK , M . L. , FUCHS , K . e t al . 1992 . Regional pattern s o f tectoni c stres s i n Europe . Journal Geophysical Research, 97 , 1 1 783-11 803. STEIN, S. , CLOETHINGH , S. , SLEEP , N . H . & WORTEL , R. 1989 . Passive margi n earthquakes , stresse s an d rheology. In : GREGERSEN , S . & BASHAM , P . W . (eds) Earthquakes a t North Atlantic Passive Margins: Neotectonics and Postglacial Rebound. Kluwer, Dordrecht , 231-259 . STEPHANSON, O. , DAHLSTROM , I. , BERGSTROM , K . et al . 1987 . FRSDB-Fennoscandian Rock Stress Database. Departmen t o f Mining , Norwegia n Institute o f Technology , Trondheim . TURCOTTE, D. , & SCHUBERT , G. 1982 . Geodynamics; Applications of Continuum Physics to Geological Problems, Wiley , New York . WESSEL, P . & SMITH , W . H . F . 1991 . Free softwar e helps ma p an d displa y data . EO S Transactions, American Geophysical Union, 72, 441 , 445-446. ZOBACK, M . L . 1992 . First an d secon d orde r pattern s of stress in the lithosphere : the Worl d Stres s Ma p Project. Journal o f Geophysical Research, 97 , 11703-11728.

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Crustal stress i n and around Norway: a compilation of i n situ stress observations MORTEN FEJERSKOV, 1 ' 3 CONRA D LINDHOLM, 2 ARN E MYRVANG & HILMA R BUNGUM 2

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Norwegian University of Science and Technology, 7034 Trondheim, Norway 2 NORSAR, 2007 Kjeller, Norway 3 Present address: Saga Petroleum ASA, P.O. Box 490, 1300 Sandvika, Norway (e-mail: [email protected]) Abstract: In-situ roc k stresse s yiel d informatio n abou t geodynami c processe s i n th e crust , and ar e importan t inpu t dat a for almos t al l kinds o f geomechanics work. Compilation s o f rock stres s measurements, suc h as the World Stres s Map, have been used to characterize th e regional stres s field . O n th e basi s o f ne w dat a fro m th e Norwegia n region , a ma p o f th e maximum horizonta l stres s (<J H) direction , fo r bot h on - an d offshor e Norway , ha s bee n established. Fou r main stres s province s wit h differen t regiona l stres s trend s ar e identified . The Barents Sea and norther n Norwa y province exhibits a very consistent N-S cr H direction, with hig h horizonta l stresses . I n th e Norwegia n Se a an d mid-Norwa y provinc e th e
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stress field. On th e basi s of Hast's measurements they conclude d tha t th e maximu m horizonta l stress directio n wa s E- W i n souther n an d N- S in th e norther n Fennoscandia . Overcorin g dat a (Myrvang 1976 ) an d earthquak e foca l mechan isms (Bungu m & Fye n 1979 ; Bungu m e t al. 1991) fro m Norwa y seeme d t o fi t wel l into thi s pattern. Klei n & Bar r (1986 ) use d stres s dat a from al l ove r Europ e t o identif y a regiona l NW-SE trend in western Europe , but character ized th e stres s fiel d i n Fennoscandi a a s compli cated. O n th e basi s o f overcorin g an d hydrauli c fracturing measurements , mainl y don e fo r engi neering purposes , a Fennoscandia n Roc k Stres s Database (FRSDB ) wa s compiled (Stephansso n el al . 1987) . Today , th e FRSD B contain s mor e than 50 0 measurement s fro m mor e tha n 12 0 locations in Finland, Norwa y and Sweden . Thes e data wer e availabl e fo r th e Worl d Stres s Ma p (WSM) projec t (Zobac k e t a l 1989 ; Zobac k 1992), which , whe n finishe d i n 1992 , containe d more tha n 730 0 stres s dat a fro m 1 8 countries. The primar y goa l o f th e WS M wa s t o establis h the global stres s field within the plates, excluding data though t t o b e influence d b y loca l condi tions. Fro m th e WSM dat a a well-define d NW SE tren d i n th e centra l part s o f th e Europea n plate was identified, wherea s this trend i s weaker and mor e disperse d i n souther n Fennoscan dia an d th e Nort h Se a (Miille r et a l 1992 ; Muller 1993) . On th e basi s o f dat a fro m th e FRSD B an d WSM, togethe r wit h ne w borehol e breakou t and foca l mechanis m dat a fro m th e Nort h Sea , a ne w stres s ma p ha s bee n establishe d (Fig . 1) . Four mai n province s wit h differen t regiona l stress trend s hav e bee n identified . These trend s have then been related to different possibl e stressgenerating mechanisms.

Stress province s alon g the coas t of Norway Northern Norway and the Barents Sea The Barent s Sea (Fig . 1) , extendin g between th e north o f Norwa y an d th e Svalbar d archipelago , has fo r th e las t decad e bee n subjecte d t o oi l and ga s exploration . Mor e tha n 5 5 exploratio n wells has bee n drille d by 1996 . Eighteen o f these have bee n analyse d fo r borehol e breakouts . Th e occurrence o f breakouts i s high an d occur s ove r depths o f 700-450 0 m. Th e maximu m horizon tal stres s (cr H ) directio n define d b y th e break out dat a i s consistentl y N-S , a tren d whic h i s confirmed b y nin e overcorin g measurement s i n northern Norwav .

Overcoring measurement s i n Norwa y ar e pri marily carrie d ou t i n connectio n wit h minin g and civi l engineerin g activities , an d fo r thi s purpose a modifie d 3 D Leema n cel l i s use d (Myrvang 1976) . For us e in this compilation th e measurements hav e bee n subjecte d t o a stric t quality control an d onl y data withou t significan t topographical effect s o r disturbance s fro m near by opening s ar e presented . I n Finnmark . over coring dat a revea l ver y hig h tectoni c stresse s (15-25MPa) a t a shallo w dept h (0-20 0 m) i n both th e N- S an d E- W direction s (Myrvang et al . 1993) . The Barent s Se a i s characterize d b y ver y lo w seismic activit y (Lindholm e t al . \995b) an d n o focal mechanism s hav e bee n derived . However , in Finnmar k som e seismi c activit y ha s bee n recorded, primaril y trace d bac k t o post-glacia l activity o n th e Stuoragurr a faul t zone . Thes e data represen t compressiona l stres s regime s an d reverse faultin g wit h th e maximu m horizonta l stress trendin g N- S o r E-W . Anothe r seismi cally activ e region, th e Senj a fractur e zone , als o exhibits revers e faultin g wit h maximu m hori zontal principa l stres s strikin g paralle l t o th e fracture zone .

Mid-Norway and the Norwegian Sea This are a comprise s dat a fro m th e Norwegia n Shelf and mainlan d Norwa y (62 :-70 N ) (Fig ^ 2). Oil an d ga s exploratio n activitie s ar e concen trated a t Haltenbanke n an d th e Trondelag Plat form. Mor e tha n 10 0 exploratio n well s hav e been drille d i n thi s area . Fiftee n well s have been analysed fo r borehol e breakouts , an d breakout s were encountere d fro m 110 0 t o 4800m . yieldin g a ver y consisten t NW-SE <J H direction. Overcoring measurement s hav e bee n per formed a t severa l location s alon g th e coast , and th e dominan t stres s directio n i s NW-SE . This confirm s th e direction s derived from bore hole breakouts . The measurement s indicate very high horizonta l stresses , occasionall y exceedin g 3dMPa (Hansse n & Myrvan g 1986 ; Myrvan g 1993) clos e t o th e surface . The Norwegia n margi n i s subject t o relativel y high seismi c activity, wit h tw o interestin g earthquake swarms in the coastal regio n at Melo y and Steigen (Bungu m & Fye n 1979 : Ataka n e t al . 1994). Thi s ha s mad e i t possibl e t o comput e ten earthquak e foca l mechanis m solution s fo r the area , whic h primaril y indicat e compressiv e reverse faultin g wit h a consisten t NW-S E a H direction. Tw o events , however , giv e norma l faulting. One . o n th e shel f edge, fits well int o the model o f passive continental margins (Stein et al.

Fig. 1 . Stres s data fro m norther n Norway and th e Barent s Sea. Ros e diagram s show th e compiled result s for eac h measurin g technique . The numbe r o f data points i s given below each ros e diagram .

Fig. 2 . Stres s dat a fro m mid-Norwa y and th e Norwegia n Sea . Ros e diagram s show th e compiled result s fo r eac h measurin g technique . The numbe r o f data points is given below each ros e diagram.

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1989) an d th e othe r i s connecte d t o a highl y elevated an d glaciate d onshor e regio n clos e t o Mel0y and Steigen .

Western Norway and the northern North Sea The major area o f investigation in this study has been th e norther n Nort h Sea , comprisin g th e Tampen Spu r area , th e Vikin g Graben an d th e Horda Platfor m (Fig . 3) . Th e are a ha s bee n subjected t o extensiv e oi l an d ga s exploration , and man y oi l an d ga s field s ar e a t presen t i n production. Th e numbe r o f exploratio n well s in th e Norwegia n secto r exceed s 370 . A tota l o f 125 well s hav e been examine d wit h respec t to borehol e breakouts , o f whic h 7 3 exhibite d breakouts. From th e UK secto r an additional 30 wells with breakouts ar e reported. Breakout s are encountered a t depth s fro m 150 0 t o 5000m . Breakout dat a fro m th e norther n Nort h Se a yield a more disperse d stres s orientatio n tha n i n the Barent s Sea and o n th e Norwegia n margin . However, eve n thoug h a blurre d NE-S W tren d is observed , th e majo r directio n is WNW-ESE. This bimoda l distributio n i s als o prominen t i n the othe r dataset s fro m thi s province . Overcoring measurement s i n thi s regio n ar e primarily mad e i n connection wit h roa d tunnel ling and constructio n of hydropower plants. The relief i s high, with deep fjords , an d th e measure ments thu s contai n larg e topographica l contri butions. Topograph y ofte n force s th e principa l stresses to alig n with the fjor d an d eac h measur ing sit e ha s t o b e examine d carefully . Eve n though overcorin g dat a tha t reflec t onl y loca l topographical effect s hav e not been included, the rose diagra m look s diffuse . However , a n overal l NW-SE trend i s obtained, an d b y correcting fo r the topographica l contributio n tectoni c stresse s of the orde r o f 10-1 5 MPa ar e calculated . The seismi c activity in thi s regio n i s relatively high an d mainl y concentrate d i n th e Tampe n Spur, th e Sog n Grabe n an d th e Stor d area . An indicative focal depth distribution fro m mor e than 50 0 earthquake s i n thi s regio n reveale d different characteristi c foca l depth s withi n sub regions (Lindhol m e t al. \995b). Earthquake s seem t o b e more shallo w (c. 5 km) i n th e coasta l region o f Stord , wherea s i n th e Tampe n Spu r (c. 15 km) an d th e wester n sid e o f th e Vikin g Graben (c . 30 km) the y appea r deepe r i n th e crust. Foca l mechanism s i n th e sam e regio n reveal consisten t trends , wher e th e shallowe r earthquakes i n the Stord area are mostly relate d to norma l faulting but i n the Tampen Spu r an d the Sog n Grabe n the y ar e relate d t o revers e

and strike-sli p faulting. Evaluatio n o f th e stres s direction derive d fro m th e foca l mechanism s shows tha t th e major tren d i s WNW-ESE, with a weaker NE-SW trend als o present . This correlates wel l with borehole breakou t data fro m th e same are a (Lindhol m et al . 19950) . To explai n th e bimoda l distributio n i t ha s been necessar y t o loo k a t som e o f th e i n situ measurements i n mor e detail . Thre e majo r oi l and ga s field s (Troll , Visun d an d Snorre ) hav e been chose n fo r detaile d breakou t studies , an d the result s revea l a rathe r unifor m pattern . Breakouts, appearin g i n well s situated centrally within th e faul t blocks , yiel d a relativel y con sistent WNW-ES E trend . Th e NE-S W stres s direction i s the n primaril y connected t o break outs encountere d i n well s clos e t o majo r faul t zones, an d thu s reflect s a stres s reorientation . where the maximum horizontal stress aligns with the tectoni c structures . Furthermore , borehol e breakout dat a indicat e a n increas e i n tectoni c stress with depth, which is believed to b e related to formatio n stiffness . Triassi c an d Jurassi c strata exhibi t near-lithostati c stresses , possibl y favouring a slight strike-slip faulting regime. The Cretaceous an d Tertiar y sediments , however , exhibit lo w tectoni c stress , yieldin g a norma l faulting regim e (Fejerskov 1996).

Southwestern Norway and the central North Sea The are a sout h o f 59° N i s referre d t o a s th e central Nort h Se a i n thi s paper . Th e majo r structural element s are th e Centra l Grabe n an d the Norwegian-Danis h Basi n (Fig . 4) . This was the first area wher e hydrocarbons were encountered, an d th e are a ha s sinc e bee n subjecte d t o extensive oil and ga s activities. In the Norwegian sector alone , mor e tha n 30 0 exploratio n well s have bee n drilled , an d Norway , th e UK , Den mark an d th e Netherland s have strong interests in thi s area . I n thi s compilation , 4 5 well s i n the Norwegia n secto r an d 3 8 well s i n th e U K sector exhibite d breakouts . Th e maximu m hor izontal stres s direction change s vigorously , both between well s a s wel l a s withi n singl e wells . However, a vagu e patter n ma y b e identified , a s wells situate d centrall y i n th e grabe n revea l a NW-SE stress orientatio n paralle l t o the graben system, an d well s o n th e flank s o r nea r larg e fault zone s mor e ofte n sho w a NE-S W stres s direction norma l t o th e grabe n structure . Still , breakouts d o no t see m t o appea r wit h the sam e consistency a s in the northern area . This may be explained b y th e extensiv e salt diapiris m i n thi s area, whic h ha s resulte d i n larg e anticline s an d

CRUSTAL STRES S OBSERVATION S I N AN D AROUN D NORWA Y

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Fig. 3 . Stres s data fro m wester n Norway an d th e northern North Sea . Ros e diagram s sho w the compiled result s for eac h measurin g technique . Th e numbe r o f data point s i s given belo w eac h ros e diagram .

Fig. 4 . Stres s dat a fro m southwester n Norwa y an d th e central Nort h Sea . Ros e diagram s sho w th e compile d results for eac h measurin g technique . Th e numbe r o f data point s i s given belo w eac h ros e diagram .

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doming structures . Near-equa l horizonta l stres ses may als o explai n th e observations . Very fe w overcoring measurement s have bee n carried ou t i n th e coasta l region s o f souther n Norway. Th e databas e i s s o spars e tha t n o overcoring dat a wer e use d fo r comparison . Seismic activit y i s als o lo w i n th e souther n North Sea , an d onl y thre e foca l mechanism s have bee n obtaine d fro m th e region .

Discussion When comparin g stres s dat a fro m th e area s i n question, it is particularly interesting to not e th e coinciding stres s direction s tha t ar e obtaine d from differen t stres s determinatio n method s (Fig. 5) . Althoug h th e method s cove r differen t depth interval s (0-1 k m fo r overcoring , 1- 5 km for borehol e breakout s an d 5-3 0 km fo r foca l mechanisms) the y yiel d approximatel y th e sam e stress orientation , wit h th e sam e regiona l an d internal variation . Thi s supports the assumptio n of a tectoni c stres s fiel d penetratin g larg e part s of th e crust . I t furthe r indicates a clos e connec tion betwee n stresse s in crystalline and sedimen tary rocks , fo r bot h on - an d offshor e Norway . This definitel y support s th e hypothesi s o f a n attached basi n (Bel l 1993) , i n contras t t o a detached basi n wher e independen t stres s fields are observe d abov e an d belo w a detachmen t zone. A t th e ver y least , th e consisten t resul t obtained b y th e variou s technique s increase s the confidenc e i n eac h o f th e stres s determina tion techniques . Most o f th e measurement s indicat e extensive compressive stresse s generall y exceedin g th e vertical stress . Earthquak e foca l mechanism s primarily indicat e a strike-sli p or thrus t faultin g stress regime , althoug h tw o mino r area s wit h normal faultin g ar e observed . Overcorin g mea surements als o yield strike-slip o r thrus t faultin g stress regime s wit h substantia l tectoni c stresses . The highes t tectoni c stres s i s measure d i n massive gneis s regions , wherea s th e mor e frac tured Caledonia n rock s exhibi t somewha t lowe r tectonic stress . Thi s i s primaril y believe d t o b e due t o th e stiffnes s o f th e roc k masse s an d thereby thei r abilit y t o conduc t tectoni c stress . Borehole breakout s yiel d n o direc t informatio n about stres s magnitudes , but , i n th e norther n Viking Graben , w rhere othe r wellbor e stres s information ha s bee n utilize d togethe r wit h th e breakout data , a possibl e strike-sli p stres s regime ma y b e inferre d i n Triassi c an d Jurassi c strata (Fejerskov 1996) . Cretaceous an d Tertiar y strata d o no t sho w extensiv e sign s o f tectoni c

stresses. Thi s i s primaril y believe d t o b e du e to th e sof t sediment s low abilit y of conduc t tec tonic stress . Although som e area s exhibi t scattere d data , the general trend indicate s a rotation o f the stress field from N- S i n th e Barent s Sea , t o NW-S E on th e Norwegian margin and t o WNW-ESE in the North Sea . This is supported b y all the stress measuring methods . T o identif y th e reaso n fo r this stres s rotation , severa l possibilitie s hav e t o be investigated . An observe d horizonta l stres s field may b e though t o f a s bein g compose d o f different components . Th e gravit y componen t yields isotropi c horizonta l stresses , wherea s tectonic component s ar e directiona l an d thu s form a n anisotropi c horizonta l stres s field, as is experienced i n bot h on - an d offshor e Norway . For convenience , tectoni c stresse s ar e ofte n fur ther divide d int o plate-wide , regiona l an d local , or first- , second - an d third-orde r stres s fields, according t o thei r are a o f influenc e (Zobac k 1992). Observe d tectoni c stresse s ma y the n be explaine d b y a plate-wid e stres s field , whic h is possibl y overprinte d b y regiona l an d loca l effects. Fo r a first-orde r approximatio n i t i s therefore essentia l t o examin e plate-wide stressgenerating mechanisms . Her e onl y a brie f discussion o f th e possibl e stres s generatin g mechanisms i s made . Fo r details , th e reade r i s referred t o th e followin g pape r o n stress-generating mechanisms (Fejerskov & Lindholm 1999) . The dominan t sourc e o f tectoni c stres s i n western Europ e i s a combinatio n o f collision related force s a t th e Euroasian-Africa n an d ridge pus h alon g th e Mid-Atlanti c an d Arcti c spreading ridg e (Miille r 1993) . Th e vicinit y t o the spreadin g ridge , an d th e fac t tha t th e ridg e push forc e ma y contribut e significantl y t o intraplate tectoni c stresse s (Forsyt h & Uyed a 1975; Dahle n 1981) . ha s mad e th e ridg e pus h effect a prim e suspec t fo r th e tectoni c stres s observed i n th e Norwegia n area . Using , a s a n approximation, th e pol e o f rotatio n a s a basi s for determinin g th e directio n o f ridg e push , traces indicatin g th e directio n o f plat e move ment hav e been superimpose d o n the stress map. Figure 5 shows tha t mos t o f th e stres s rotatio n can b e explaine d b y th e flowlines . A mor e precise wa y t o describ e th e ridg e pus h forc e i s to us e the ridg e geometry an d bathymetr y data. In the arctic region the Lomonosov Ridg e is very prominent, an d henc e coul d contribut e signifi cantly t o a rotation o f the ridg e push forc e fro m NW-SE i n th e Atlanti c t o mor e N- S i n th e Arctic. Therefore , fo r a mor e detaile d stud y o f the ridg e pus h forc e th e actua l 3 D geometr y (bathymetry, latera l variatio n an d plat e thick ness) ha s t o b e modelled . Althoug h th e ridg e

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Fig. 5 . Compilatio n o f Norwegian stres s data . Ros e diagrams indicat e maximum horizonta l stres s direction s from borehol e breakout s (B) , focal mechanisms (F) an d overcorin g (O) within each o f the four stress provinces. Geological feature s such as the spreading ridge and th e continental margin are included. Flowline s (dashed line), representing the directio n o f spreading and indicatin g the directio n of ridge push a s a first-order approximation correlate wel l wit h the stres s trends.

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push effec t ca n explai n most o f the stres s obser vations, othe r possibl e cause s fo r th e tectoni c stress, suc h a s densit y variations, sediment load and deglaciation , als o hav e t o b e considered . These effects , however , ar e o f muc h smalle r lateral extent, and wil l primarily affect th e stres s field only locally . The souther n areas , encompassin g th e Nort h Sea, yiel d mor e scattere d dat a tha n d o th e Barents Se a an d th e Norwegia n Sea . Thi s ma y be a result of the available number of data, but is rather anticipate d to b e caused b y variation s in the horizonta l stres s anisotropy . I f th e Nort h Sea exhibit s lowe r stres s anisotrop y tha n th e Barents Sea , measurement s i n Nort h Se a wil l more easil y be influence d by loca l tectoni c an d geological structures , yieldin g a mor e scattere d stress direction . The chang e in stress anisotropy may b e cause d b y change s i n th e magnitud e of the ridge push force, but can also be related t o the relativ e angl e betwee n th e continenta l margin an d th e ridg e pus h force . I n th e wester n Barents Se a th e ridg e pus h forc e act s nea r parallel t o th e continenta l margin , wherea s i n the Norwegia n Se a i t act s perpendicula r t o it . As the continental margi n i s known t o generat e deviatoric tension in the continental crust (Stein el al. 1989) , th e continenta l margi n effec t ma y counteract th e ridg e pus h forc e i n th e Nor wegian Se a simultaneousl y a s th e horizonta l stress anisotrop y i n th e wester n Barent s Se a is enhanced. Conclusion The differen t technique s for stress determinatio n applied o n the Norwegian continenta l shel f yield very consisten t results , although the y cove r different depths . Thi s increase s th e confidenc e i n each stres s determinatio n method , support s th e assumption of tectonic stresse s penetrating larg e parts o f the crust, an d indicate s a close connec tion betwee n stresse s i n bot h crystallin e an d sedimentary rocks, thereby supporting th e hypothesis o f an attache d basin . The first-orde r stres s patter n i n Norwa y an d adjacent offshor e region s seem s to b e characterized b y hig h compressiv e horizonta l stresses . A rotatio n o f th e maximu m horizonta l stress , from N- S i n the Barent s Sea, to NW-SE on the mid-Norwegian margi n an d t o WNW-ES E i n the Nort h Sea , i s observed. I t i s considered tha t most o f this rotatio n ca n generall y be explained by th e stresse s generated fro m ridge push, bu t i t is likely that othe r effect s ma y als o contribute . The stres s direction s ar e mor e scattere d i n the Nort h Se a compare d wit h th e well-define d

trends i n th e Barent s Se a an d th e Norwegia n Sea. Thi s i s believe d t o b e cause d b y change s in stres s anisotropy, indicating a stronge r stress anisotropy i n th e norther n areas . Chang e i n stress anisotrop y may b e du e t o change s i n th e magnitude o f th e ridg e push force , or relate d to a chang e i n angl e betwee n th e continenta l margin an d th e ridg e push force. This wor k wa s supporte d b y th e Researc h Counci l of Norway, as a part o f the IBS (Integrated Basin Studies) project unde r th e Joule I I researc h programm e funde d by th e Commissio n o f Europea n Communitie s (Con tract JOU2-CT-92-0110) . Dat a fo r th e stud y wer e supported b y Nors k Hydro . Sag a Petroleum . Statoi l and th e PROFIT-project. Th e authors especiall y wan t to than k M . D . Zoback . R . Gabrielsen . A . Mascle . B. Torudbakken . an d J . Palme r fo r valuabl e inpu t during th e writing .

References ATAKAN. K. . LINDHOLM , C . D . & HAVSKOV . J . 1994 . Earthquake swarm in Steigen. Northern Norway: an unusua l example of intraplate seismicity. Terra Nova, 6 , 180-194 . BELL, J . S . 1993 . Globa l sedimentar y stres s projec t o f the international lithospher e program. Poste r P71 . 55th Meetin g o f Associatio n o f Exploratio n an d Geophysics, Stavanger , 7-11 June . BLJNGUM, H . & EVEN , J . 1979 . Hypocentral distribution, foca l mechanisms, an d tectoni c implication s of Fennoscandia n earthquake s 1954-1978 . Geologiska Foreningens i Slokholm Forhandlingar. 101(4), 261-273. , ALSAKER , A.. KVAMME , L. B . & HANSEN . R . A . 1991. Seismicit y an d seismotectonic s o f Norwa y and nearb y continenta l shel f areas . Journal o f Geophysical Research. 96. 2249-2265 . DAHLEN, F . A . 1981 . Isostas y an d th e ambien t stat e of stres s i n th e oceani c lithosphere . Journal o f Geophysical Research, 86, 7801-7807 . FEJERSKOV, M 1996 . Determination o f i n sit u rock stresses related to petroleum activities on the Norwegian Continental Shelf. Dr . Ing . thesis . Norwegian Institut e of Scienc e an d Technology . Trondheim. & LINDHOLM . C . 1999 . Crusta l stres s i n an d around Norway : a n evaluatio n o f stres s generat ing mechanisms . This volume. FORSYTH, D . W . & UYEDA . S . 1975 . On th e relativ e importance o f th e drivin g forces of plat e motion. Geophysical Journal of the Roval Astronomical Society, 43 . 163-200 . HANSSEN. T . H . & MYRVANG . A . 1986 . Roc k stresse s and roc k stres s effect s i n th e Kobbel v area , northern Norway. Proceedings of th e International Symposium on Rock Stress and Rock Stress Measurements. Stockholm . 1- 3 September, 625-634 . HAST. N . 1958 . Th e Measurement o f Rock Pressure i n Mines. Swedis h Geologica l Survey . Series C. 560 .

CRUSTAL STRES S OBSERVATION S I N AN D AROUN D NORWA Y KLEIN, R . J . & BARR , M . V . 1986 . Regiona l stat e o f stress in western Europe . Proceedings of th e International Symposium on Rock Stress and Rock Stress Measurements, Stockholm , 1- 3 September , 21-32.

LlNDHOLM, C, BUNGUM , H. , BRATLI , R . K. , AADN0Y ,

B. S. , DAHL , N. , TORUDBAKKEN , B . & ATAKAN , K. 1995<3 . Crusta l stres s i n th e norther n Nort h Sea a s inferre d fro m borehol e breakout s an d earthquake foca l mechanisms . Terra Nova, 7, 51-59 ,, VILLAGRAN, M. & HICKS, E. 19956 . Crustal stress an d tectonic s i n Norwegia n region s deter mined fro m earthquak e foca l mechanisms . Proceedings from Workshop on Rock Stresses in the North Sea, Trondheim, 13-1 4 February, 77-91 . MULLER, B . 1993. Tectonic stress i n Europe - borehole guided waves affected by stress-induced anisotropy. PhD thesis , Universit y of Karlsruhe . , ZOBACK , M . L. , FUCHS , K. , MASTIN , L. , GREGERSEN, S. , PAVONI , N. , STEPHANSSON , O . & LJUNGGREN , C . 1992 . Regiona l pattern s o f tectonic stres s i n Europe . Journal o f Geophysical Research, 97(B8), 1 1 783-11 803. MYRVANG, A . 1976 . Practica l us e o f roc k stres s measurements i n Norway . Symposium o n Investigation of Stress in Rock - Advances in Stress Measurements, Sydney , 11-1 3 August, 92-99.

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1993. Roc k stres s an d rock stres s problem s i n Norway. In: HUDSON , J . A . (ed. ) Comprehensive Rock Engineering, Vol. 3 . Pergamon , Oxford , 461-471. , HANSEN, S . E. & SORENSEN , T. 1993 . Rock stress redistribution aroun d a n ope n pi t min e i n hard rock. International Journal o f Rock Mechanics, Mining Science, an d Geomechanics Abstracts, 30 , 1001-1004. RANALLI, G. & CHANDLER, T. E . 1975 . The stres s field in the upper crus t a s determined fro m i n situ measurements. Geologische Rundschau, 64 , 653-674 STEIN, S. , CLOETINGH , S. , SLEEP , N . & WORTEL , R . 1989. Passiv e margi n earthquakes , stresse s an d rheology. In : GREGERSEN , S . & BASHAM , P . (eds ) Earthquakes at North-Atlantic Passive Margins: Neotectonics and Postglacial Rebound, Nato A SI Series C . Kluwer, Boston , MA , 231-259 . STEPHANSSON, O. , DAHLSTROEM , L . O. , BERG STROEM, K. et al. 1987. Fennoscandian Rock Stress Data Base - FRSDB. Lule a University , Research report, Lule a 1987:06 . ZOBACK, M . L . 1992 . First- and second-orde r pattern s of stress in the lithosphere: th e World Stres s Ma p project. Journal o f Geophysical Research, 97(B8) , 11703-11728. et al . 1989 . Globa l pattern s o f tectoni c stress. Nature, 341 , 291-298.

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Crustal stress in an d around Norway : a n evaluation o f stress-generating mechanism s MORTEN FEJERSKOV 13 & CONRA D LINDHOLM 2 1

Norwegian University of Science and Technology, 7034 Trondheim, Norway 2 NORSAR, 2007 Kjeller, Norway ^Present address: Saga Petroleum ASA, P.O. Box 490, 1300 Sandvika, Norway (e-mail: [email protected]) Abstract: Recen t stres s observation s fro m i n situ measurement s an d earthquak e foca l mechanisms i n th e Norwegia n onshor e an d offshor e area s ar e evaluate d wit h th e ai m o f characterizing th e mos t importan t mechanism s fo r th e presen t stres s fiel d i n an d aroun d Norway. Th e evaluatio n i s base d o n a se t o f stres s indicator s tha t includ e both shallo w measurements (overcorin g an d borehol e breakouts ) an d dee p dat a (earthquak e foca l mechanisms). Compute r simulation s o f simpl e stress-generatio n model s wer e use d t o constrain th e relativ e importance o f different stress-generatin g mechanisms. Th e ridg e push force associate d wit h sea-floo r spreadin g i n th e Nort h Atlanti c i s considere d t o b e th e primary sourc e o f th e compressiona l stres s field observed i n Norway . Regiona l influences from th e continental margi n densit y contrast , topograph y an d flexure induced b y sedimen t loading ar e o f limite d latera l extent , bu t ar e importan t i n reorientin g th e stres s fiel d i n certain areas . Th e observed tectonic s an d stresse s are generally also in accord wit h tectonics expected fro m Fennoscandia n uplift .

The orientatio n o f th e stresse s i n th e crus t is amon g th e principa l indicator s o f curren t dynamic processes . Earthquak e foca l mech anisms togethe r wit h i n situ measurement s (overcoring, hydrauli c fracturin g an d borehol e breakout analysis ) are importan t source s o f in formation o n stres s orientatio n an d magnitude . The firs t i n situ stres s determinatio n method s were developed i n the 1950 s (see Hast 1958) , and were primarily relate d t o mining activity. Before that time , roc k stresse s ha d onl y bee n regarde d as gravit y induced , bu t i n th e 1950 s th e ne w measurements indicate d tha t horizonta l stresse s exceeded wha t coul d hav e bee n expecte d fro m only gravitationa l forces , an d sinc e th e 1960 s stress dat a wer e interprete d i n a wide r geophy sical context . Sinc e th e firs t measurement s wer e obtained i n Norwegia n areas , th e stres s data base ha s increase d dramatically , primarily as a result o f thre e scientifi c programmes : th e Fen noscandian Rock Stres s Data Bas e (Stephansson et al 1987) , the World Stres s Map (Zobac k e t al. 1989; Mulle r e t al . 1992 ; Zoback 1992) , and th e Dynamics o f the Norwegia n Margi n (Fejersko v et al . 1995 ; Lindholm e t al . 19956) . I n th e map ping o f crusta l stresse s i n northwester n Europ e and Norway , th e followin g contribution s can be considered a s milestones : Has t (1969) , Ranalli

& Chandle r (1975) , Bungu m & Fye n (1979) , Klein & Bar r (1986) , Claus s e t al . (1989) , Bungum e t al . (1991) , Mulle r e t al., (1992) an d Muller (1993). Parallel t o th e compilatio n o f ne w dat a o n observed stresses , analytica l an d numerica l models an d compute r simulation s were applie d to investigat e th e characteristic s o f differen t stress-generating mechanisms . Model s rangin g from globa l analysi s o f differen t plat e drivin g forces (Forsyt h & Uyed a 1975 ; Bott & Kuszni r 1984; Richardsso n 1992 ) t o regiona l an d loca l features suc h a s densit y inhomogeneities , an d flexural stresse s fro m sedimen t loading , glacia l rebound an d topograph y (Stephansso n 1988 ; Stein e t al . 1989 ; Span n e t al . 1991 ) wer e developed. Th e presen t pape r review s availabl e dat a and modellin g result s wit h th e goa l o f charac terizing th e mos t importan t stress-generatin g processes behin d th e horizonta l stres s fiel d i n Norway an d adjacen t offshor e areas .

Data Ranalli & Chandle r (1975 ) wer e th e firs t t o compile stres s dat a fro m Scandinavia . O n th e

From: NOTTVEDT , A . e t al . (eds ) Dynamics o f th e Norwegian Margin. Geologica l Society , London , Specia l Publications, 167 , 451-467 . 1-86239-056-8/OO/ S 15.00 © Th e Geologica l Societ y o f Londo n 2000 .

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basis o f Has f s measurement s (largel y overcor ing) the y conclude d tha t th e directio n o f th e principal horizonta l stresse s wa s E- W i n south ern Scandinavi a an d N-S i n northern Scandina via. Althoug h thi s pictur e ha s becom e mor e differentiated wit h th e acquisitio n o f additiona l data, th e main conclusion o f this early investigation i s stil l valid . Th e databas e no w a t han d has bee n compile d an d qualit y checked, an d ha s been presente d b y Fejersko v e t al. (1995 ) an d Lindholm e t al . (\995b). A synthesi s of al l th e data i s show n i n Fig s 1 an d 2 , togethe r wit h the regionalization used in this paper. Withi n the area define d in Fig . 1 altogether 35 1 data point s give th e azimut h o f th e larges t principa l hor izontal stres s component , an d th e observe d stress direction s are summarize d a s follows: • Th e maximu m horizonta l stres s directio n (<JH) rotate s fro m N- S i n norther n Norwa y and th e Barent s Se a t o WNW-ES E i n western Norway an d the northern Nort h Sea .

The stress direction s a t shallo w depths deter mined b y i n situ techniques , an d deepe r observations fro m th e middl e t o lowe r crus t based o n earthquak e foca l mechanisms , ar e similar an d indicat e that th e tectonic stresse s are homogeneou s i n directio n ove r a larg e depth range . In offshor e area s revers e an d strike-sli p faulting dominates , indicatin g compressiv e stress regimes. Shallow earthquakes , particularl y foun d in a n are a aroun d Stord , bu t als o i n th e Oslo regio n an d i n th e Melo y an d Steige n sequences (Bungum et al. 1979 ; Atakan e t al. 1994) indicat e tensiona l stres s regime s with normal faulting . No chang e i n horizonta l stres s directio n i s observed betwee n sedimentar y an d crystal line rocks . Overcoring measurement s show a more scattered stress direction compared wit h borehole breakouts an d earthquake foca l mechanisms.

Fig. 1 . Fiv e regions with rose diagrams indicating the maximum horizontal stress direction (<J H ) determined fro m focal mechanism s (F) , borehol e breakout s (B ) and overcorin g (C) . A detaile d descriptio n o f th e dat a use d ha s been give n by Fejerskov e t al. (1995) and Lindhol m e t al. (\995a.b). The shel f edge is shown with the dashed line. S, Stord area ; M, Mel0y ; St , Steigen ; V.B. , Vorin g Basin ; Sv , Svartisen ; L Jostedalsbreen ; F. Folgefonna ; M.B.. Mor e Basin ; Esc., Escarpment ; Gr. , Graben ; F.Z. , Fractur e Zone .

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Fig. 2. Histogram s o f borehole breakout , earthquak e focal mechanis m an d overcorin g data fo r eac h region .

They als o revea l hig h stres s magnitude s (10-20MPa tectoni c stress) , particularl y i n massive gneis s area s o f mid-Norway .

Regions Five mai n region s wit h differen t regiona l hor izontal stres s trend s hav e bee n identifie d i n Norway. Th e regio n boundarie s d o no t repre sent abrup t change s i n th e stres s field , bu t a continuous latera l variation .

Northern Norway and the Barents Sea As see n fro m Fig . 1 th e are a comprise s larg e offshore area s an d th e northernmos t par t o f Norway (Finnmark) . I n situ measurement s i n the Barent s Se a revea l a clea r N- S stres s direction (Dar t et al. 1995 ; Fejersko v e t al. 1995). Overcorin g measurement s fro m onshor e Finnmark are consistent wit h the offshore obser vations an d revea l ver y hig h tectoni c stresse s (15-25MPa) bot h i n N- S an d E- W direc tions a t shallowe r depth s (0-20 0 m) (Myrvan g et al . 1993) . I n contras t t o th e lo w seismi c activity i n th e Barent s Sea , neotectoni c move -

ments associate d wit h larg e prehistori c earth quakes hav e bee n mappe d i n Finnmar k an d northern Sweden . Displacement s o n NE-S W and NNW-SS E trendin g structure s indicat e a compressional stres s regim e wit h a NNW-SS E stress directio n (Olese n e t al . 1992 ; Bungum & Lindholm i n prep.).

Mid-Norway and the Norwegian Sea This area s comprise s a larg e continenta l plat form an d mid-Norway . I n situ stres s obser vations o n th e continenta l platfor m an d i n mid-Norway indicate a NW-SE stres s direction , which is consistent wit h earthquake foca l mech anisms. Th e seismi c activit y i n th e regio n i s generally high, with two recent interestin g earthquake swarm s recorde d i n th e coasta l regio n a t Mel0y and Steige n (Bungum et al. 1979 ; Atakan et al . 1994) . Earthquak e foca l mechanism s reveal a compressiona l stres s regim e offshore , whereas seismic onshore activities close to highly elevated an d glaciate d regions , suc h a s th e Mel0y-Steigen area, indicate a shallow tensiona l stress regime. Overcorin g measurement s i n midNorway revea l hig h tectoni c stresse s (u p t o 30MPa), especiall y i n massiv e gneis s region s (Hanssen & Myrvan g 1986).

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Western Norway and the northern North Sea In situ stres s observation s i n the norther n Nort h Sea yield a WNW-ESE regional stress direction , but a NNE-SS W directio n i s als o weakl y indicated bot h b y i n situ measurement s an d b y earthquake foca l mechanism s (Lindhol m e t al . \995a). Th e seismi c activit y offshor e i s high , especially centre d aroun d 61.5 C N an d 3.0°E , and a compressiona l stres s regim e i s indicated . Onshore, seismi c activit y i s concentrated o n th e seaward sid e o f a highl y elevate d an d glaciate d region jus t eas t o f th e islan d o f Stord . Foca l mechanisms indicat e a smal l are a wit h shallo w tensional stres s regim e i n thi s area . Stres s direc tion an d magnitude s fro m overcorin g measure ments i n wester n Norwa y ar e ofte n strongl y influenced b y topography , however , indicatin g high tectoni c stresse s (10-1 5 MPa).

Southwestern Norway and the central North Sea Borehole breakout s i n th e centra l Nort h Se a yield a comple x stres s pattern . Th e NW-S E European regiona l tren d i s observe d i n th e dataset, bu t als o othe r stres s direction s fre quently appear. The observations ar e of very low quality an d probabl y also largel y reflec t loca l features suc h a s fault s an d sal t diapir s (Cowgil l et al . 1994 ; As k 1996) . A s relativel y fe w foca l mechanisms an d overcorin g measurement s hav e been obtaine d fo r thi s area , i t ha s bee n difficul t to accomplis h a regiona l interpretation . Th e seismic activit y i n thi s regio n i s low .

Southeastern Norway This regio n i s centred aroun d th e Permia n Osl o Graben, an d becaus e o f the proximity t o seismi c monitoring station s foca l mechanism s coul d b e calculated fo r man y mino r events . Th e dat a indicate bot h relatively deep an d shallo w seismi c

activity wit h a dominanc e o f norma l faultin g in th e uppe r crust . Th e foca l mechanism s indicate a NW-SE compressiona l (NE-S W ten sional) stres s direction . Thi s i s onl y partl y confirmed b y overcorin g measurements , whic h generally yiel d a NE-SW tren d mor e parallel t o the grabe n structure , an d wit h stresse s u p t o 20 MPa a t shallowe r depths .

Stress-generating mechanism s Stresses in the Earth's crust may be distinguished with respec t t o thei r origi n an d latera l exten t (Engelder 1974 ; Sykes & Sbar 1994) , an d ca n b e divided into continental, regional and loca l stress fields (Tabl e 1) . Th e observe d stres s i s the n composed o f a plate-wid e continenta l stres s field overprinte d b y regiona l an d loca l effect s (Zoback 1992) . T o identif y th e stres s source s in th e Norwegia n areas , i t i s importan t t o examine an d identif y th e mos t importan t stress generating mechanism s an d t o quantif y th e stress magnitude , latera l extensio n an d dept h variation. Thi s ha s bee n achieve d primaril y b y analytical o r numerica l stres s modelling , wher e the plate has bee n regarde d a s pure elastic. Thi s is a coars e simplificatio n o f th e crust' s brittle ductile behaviour . I n th e simplifie d model s th e computed stres s direction s an d stres s regime s will b e clos e t o reality , wherea s absolut e stres s magnitudes ar e probabl y exaggerated . Henc e the model s ma y b e use d onl y qualitatively , and to deriv e appropriat e stres s magnitude s th e effects o f rheolog y hav e t o b e evaluate d (Cloe tingh & Buro v 1996) . I n case s wher e w e hav e not performe d th e modellin g a revie w o f othe r modelling result s an d a qualitativ e evaluation has bee n made . The generall y accepte d explanatio n fo r th e observed first-orde r intraplat e stresse s ar e plat e tectonic force s actin g a t convergen t an d diver gent plat e boundarie s (Zobac k 1992) . Severa l workers, who have modelled the significance an d relative magnitud e o f differen t plat e tectoni c

Table 1 . Stress-generating mechanisms Stress field

Ist-order, continenta l

2nd-order, regiona l

3rd-order, loca l

Lateral exten t

> 1000 km

100- 1000 km

< 100 km

Stress-generating mechanisms

Plate tectoni c forces : Ridge pus h Slab pul l Basal dra g

Large-scale densit y inhomogeneities: Continental margi n Flexural stresses : Deglaciation Sediment loading Wide topographica l load s

Topography: Fjords an d mountai n ranges Geological features : Faults Hard an d sof t inclusion s

CRUSTAL STRESS-GENERATIN G MECHANISMS forces, hav e concluded tha t ridg e push, sla b pull and collisiona l resistance forces are the principal contributors bot h t o plat e kinematic s an d fo r the globa l stres s fiel d (Forsyt h & Uyed a 1975 ; Chappie & Tullis, 1977 ; Richardsso n 1992) . Muller (1993 ) analyse d th e Europea n stres s field an d conclude d tha t th e dominan t sourc e of regional tectonic stress in western Europe i s a combination o f collision-relate d force s a t th e southern plat e boundar y an d ridg e pus h alon g the wester n an d norther n boundary . Thi s i s based o n a matc h betwee n ridg e pus h an d collisional torqu e direction s an d stres s direc tions (Richardsso n 1992) , an d a similarit y between stres s orientation s an d relativ e plat e motion directions betwee n th e African-Eurasia n and Nort h American-Eurasia n plates . Becaus e of th e Norwegia n region' s proxima l posi tion relativ e t o th e Mid-Atlanti c an d Arcti c spreading ridge , th e majo r first-orde r tectoni c stresses i n this regio n ar e attribute d t o th e ridg e push force .

455

Second-order stres s fields arise fro m regiona l density inhomogeneities , topographica l load s and plat e flexur e relate d t o deglaciatio n o r sediment loading . Som e loca l effects , a s indi cated i n Table 1 are belo w th e resolutio n o f ou r stress observation s an d ar e therefor e no t dis cussed i n detai l in thi s paper. In additio n t o stress-generatin g feature s th e paper als o focuse s o n th e effec t o f crusta l thickness a s a stress-modifyin g feature tha t ha s importance i n Norwegia n regions . A s tectoni c stresses are concentrated in the brittle part o f the crust, a loca l o r regiona l crusta l thinnin g wil l significantly enhanc e th e stres s magnitud e an d vice vers a (Hasegaw a e t al. 1985) .

The ridge push force Mid-ocean ridge s ar e area s o f shallo w bathy metry i n approximatel y isostati c equilibriu m because th e elevate d crus t i s compensate d a t depth b y hot , low-densit y material . A s th e

Fig. 3 . Schemati c cross-section o f th e spreadin g ridg e model (a). Variation in temperatur e (b) , density (c) and deviatoric stresses (d) as a function o f age and depth . Assumed temperature and densit y of the mantle are 1350° C and 330 0 kg m'3, respectively; a therma l diffusivit y o f 0.008cm 2 s"1 an d coefficien t o f thermal expansio n o f 3.2 x 10- 5oC-' ar e anticipated .

456

M. FEJERSKO V & C . D . LINDHOL M Table 2 . Stress magnitudes associated \\ith ridge push reported bv various researchers Reference

Stresses relate d t o ridg e pus h

Bott & Kuszni r (1984 ) Stein e t al (1989 ) Bott (1991 )

Average tectoni c stres s o f the orde r o f 20-30 MPa Average tectoni c stres s o f a fe w tens o f MP a Maximum deviatori c stress o f abou t 4 0 MPa i n th e upper lithospher e 200 km fro m th e cres t Maximum deviatori c stres s of 3 1 MPa a t th e surfac e of 6 0 Ma oceani c lithospher e Average tectoni c stres s of 2 5 MPa fo r a n 80k m thic k 60 Ma oceani c crus t

Dahlen (1981 ) Fleitout & Froidevau x (1983 )

oceanic lithospher e age s an d move s awa y fro m the mid-ocea n ridg e i t cool s an d subsides . Th e ridge elevatio n create s a n outwar d compres sional, gravity-generate d forc e perpendicula r t o the crest . A s a resul t o f th e rathe r simpl e ridg e geometry an d th e well-constraine d boundar y conditions th e ridg e pus h forc e i s quantitatively well understoo d an d ha s bee n analyse d bot h analytically an d numericall y (Liste r 1975 ; Par sons & Richte r 1980 ; Dahle n 1981) . By applyin g a coolin g half-spac e model , Dahlen (1981 ) wa s abl e t o estimat e th e stresse s as a functio n o f bot h dept h an d ag e o f th e oceanic lithosphere . an d o n th e basi s o f hi s work. Fig . 3 show s th e distributio n o f tempera ture, densit y an d deviatori c stresse s i n th e oceanic lithosphere . (Deviatori c stres s (S) i s the differenc e betwee n tota l stres s (a ) an d mea n Earth pressur e (P).

compression i n th e oceani c crust ) clos e t o th e margin slope , decreasin g wit h bot h dept h an d distance fro m th e margin . Th e effec t i s no t expected t o penetrat e mor e tha n abou t 100k m into th e continental plate . As lon g a s th e continenta l effec t i s evaluate d alone, stresse s wil l alway s ac t perpendicula r t o the margin , bu t Golk e e t al . (1995) showe d tha t stress directio n an d stres s anisotrop y nea r th e continental margi n ar e dependen t o n th e angl e between th e applie d ridg e pus h forc e an d th e continental margin. Figure 4 illustrates the effec t of a chang e i n orientatio n o f th e continenta l margin, where margins orientated norma l t o th e tectonic far-fiel d stres s directio n wer e foun d t o exhibit lowe r stress anisotropy (clos e t o zero ) in the continenta l crus t compare d wit h margin s orientated paralle l t o th e fa r field.

Glacial rebound and flexwal stresses related to deglaciation where P = (cry + (TH + cr//)/3, and S H = -Si an d SH = 0 is assumed. ) The ridg e pus h forc e (F R) wil l b e zer o a t th e ridge cres t an d increas e linearl y wit h age . The deviatori c stresse s increas e wit h age , bu t will decreas e wit h depth . Fo r a 6 0 Ma oceani c crust, a maximu m deviatoric stress o f 3 1 MPa i s obtained. Thi s correspond s wel l t o value s reported b y variou s worker s liste d i n Tabl e 2 .

Glaciers ca n for m a significant loa d o n th e lithosphere, an d deglaciation , o r remova l o f th e ic e load, introduce s stres s change s i n th e uppe r crust. Tw o models , bot h considerin g th e crus t as elastic , hav e bee n propose d t o estimat e horizontal stresse s cause d b y glacia l reboun d

Density contrast at the continental margin The transitio n from dense r an d thinne r oceani c crust t o lighte r and thicke r continental crus t wil l create a tensiona l stres s stat e i n th e continenta l crust wher e i t tend s t o 'sprea d out ' ove r th e oceanic lithospher e (Artyushko v 1973) , whereby tensional deviatori c stresse s wil l b e generate d i n the continenta l crus t an d compressiona l devia toric stresse s i n th e oceani c crust . Stei n e t al . (1989) computed maximum deviatori c stresse s o f 40-50 MPa (tensio n i n the continental crus t an d

Fig. 4 . Schemati c mode l o f a continenta l margi n perturbed b y a tectoni c far-fiel d stress . Tectoni c stresses applie d paralle l t o th e margi n yiel d a higher horizontal stres s anisotrop y i n th e continenta l crust compared wit h whe n th e margin strike s normal t o th e far-field stres s direction .

CRUSTAL STRESS-GENERATIN G MECHANISM S

457

Fig. 5. Tw o models for the estimation of stresses associated with post-glacial rebound. The model of Stephansson (1988) (a ) predicts compressional horizonta l stresse s o f the orde r o f 2-3 MPa beneat h a n ic e sheet o f 2k m thickness with 140 m o f uplif t remaining . The mode l of Stei n et al. (1989) (b) predicts tensional stresses of a few tens o f MP a fo r a n ic e load o f 2 km thickness . (Fig. 5) . Stephansso n (1988 ) assumed th e litho sphere t o b e i n equilibriu m befor e ic e loadin g and compute d compressiona l stresse s beneat h the ice , s o tha t whe n th e ic e i s remove d hori zontal stresse s slowl y diminis h unti l th e uplif t ceases an d equilibriu m i s reached . Stei n e t al . (1989) assume d th e plat e t o b e i n equilibriu m with th e ic e loa d applie d an d thu s derive d tensional stresse s associated wit h the uplift . The choic e o f mode l depend s o n glacia l history an d th e preferre d plat e relaxatio n time . If th e plat e relaxatio n tim e i s lon g compare d with th e glacia l history , th e plat e wil l no t hav e time t o adjus t it s equilibriu m and stresse s fro m deglaciation ar e bes t calculate d b y usin g a n undeformed plat e (Stephansson' s model) . Fo r short relaxatio n time s some viscou s deformatio n would appea r i n th e lithosphere , changin g th e state o f equilibrium , which makes th e mode l o f Stein e t al . more appropriate . Close t o th e forme r ic e margin th e mode l o f Stein e t al . (1989 ) compute d deviatori c stresse s in th e orde r o f 30 MPa fo r a 2 km thic k ice sheet (Fig. 6) . Th e mode l o f Stephansso n predict s much smalle r stres s magnitude s (3- 4 MPa) and a n opposit e stres s pattern . I n reality , th e result i s assumed t o li e somewhere betwee n th e two models , probabl y favourin g sligh t tensio n (P. Johnston, pers . comm.) .

Flexural stresses from sediment loading Subsiding sedimen t basin s ar e drive n b y a combination o f tectoni c crusta l stretchin g an d

thinning and th e increasing sediment load. A s in the cas e o f ic e loading , th e sedimen t loa d wil l introduce flexura l stresse s wit h compressiona l stresses i n th e uppe r crus t underneat h th e sedi ment basi n an d tensiona l stresse s outsid e th e basin (Fig . 7) . Stei n e t al . (1989 ) compute d stresses o f severa l hundre d MP a fo r a 10k m thick sedimentar y basin . Thi s i s a n obviou s artefact o f the elastic model, an d b y introducing viscoelastic an d brittle-ductil e plat e behaviou r stresses decreas e t o som e ten s o f MPa. B y varying lithospher e thicknes s an d sedimentatio n rates, Stei n e t al . furthe r conclude d tha t th e highest stresse s ar e relate d t o basin s wit h hig h sedimentation rate s underlai n b y a thi n litho sphere. Ver y fe w place s i n th e worl d sho w sufficiently hig h sedimentatio n rate s t o generat e strong seismi c activity , whic h lead s t o th e conclusion tha t sedimen t loadin g generall y ha s only a mino r influenc e o n th e regiona l stres s field. Recent observation s fro m th e Norwegia n area, wher e area s o f extraordinar y hig h sedi mentation rate s occur, however , show a correla tion betwee n seismicit y and sedimentar y basin s (Byrkjeland 1996) .

Topography Topography represent s a loa d o n th e litho sphere, whic h ha s t o b e supported b y stresses i n the crust . Thre e mai n type s o f topographica l surface loadin g situation s ma y b e identifie d (Bott 1971) , an d ar e qualitativel y illustrated i n

458

M. FEJERSKO V & C . D . LINDHOL M

Fig. 6. Deformatio n an d horizonta l stress magnitudes resulting from th e removal of 2000 m of ice from a n elastic crust floating on a fluid. Modelling parameter s use d are : flexural rigidity (D= xlO 2 -), flexural parameter (a = 205km) , restorin g densit y (p a = 2300 kg m~3 ), plat e thicknes s ( T — 110km). Fig. 8 : narrow load ; wid e uncompensated load ; wide compensate d load . For smal l surface loads les s than 5 0 km wide, the lithospher e wil l no t ben d significantly , an d stresses are adequately modelled b y applying the load o n a n elasti c half-spac e (Jaege r & Coo k 1969; Bott & Kusznir 1984) . Beneath th e applie d load deviatori c tensio n develops , wherea s beyond th e edg e o f the load , sligh t compressio n appears. Mor e genera l model s o f symmetrica l

Fig. 7. Schemati c representatio n o f th e bendin g stresses associated wit h sedimen t loading . Beneat h th e sediment basin compressional stresses appear, whereas tensional stresse s dominate on th e flanks.

ridges i n bot h tw o dimension s (Savag e e l al. 1985; Pan & Amadei 1993 ) and thre e dimensions (Liu & Zobac k 1992 ) arrive a t simila r conclu sions, bu t indicat e slightl y lowe r stresse s com pared wit h a constan t load . I n al l case s th e stresses ar e believe d t o b e to o smal l t o caus e tectonic activity , but ma y be of local importanc e and trigge r earthquake s i n area s wit h enhance d regional stres s or pre-existing zones of weakness. For surfac e load s significantl y wide r tha n 50km, whic h ar e no t compensate d a t depth , the sam e type of bendin g as caused b y sediment loading wil l appear . Th e flexure associated wit h bending creates compression on the concave side of th e plat e an d tensio n o n th e convex side , an d the stres s magnitude s ca n b e ver y hig h (several hundreds o f MPa ) i f elasti c model s ar e applie d (Bott 1971) . Brittle-ductil e behaviour , however , reduces th e stres s magnitude s significantl y an d entails stres s dissipation by transien t cree p over a relativ e shor t geologica l time scale. If th e surfac e loa d i s compensate d a t depth , the loa d wil l b e counterbalanced by a n upthrust and n o bendin g wil l occur . Th e vertica l stresse s

CRUSTAL STRESS-GENERATIN G MECHANISM S

459

Sources of observe d stresse s in the Norwegian provinces On th e basi s o f th e describe d source s o f crusta l stress an d th e observationa l dat a i t i s now tim e to lin k th e stress-generatin g mechanism s t o th e defined stres s provinces . Th e effec t o f eac h stress-generating mechanis m in eac h stres s pro vince is discussed in vie w of the stres s observa tions, an d th e sourc e model s ar e evaluate d i n order o f latera l extent (plate-wide to local ) an d ranked i n orde r o f importance . Figur e 9 provides a n overvie w of th e regiona l variation s and Tabl e 4 contains a summar y of th e results.

Ridge push Fig. 8 . Topographica l loa d model s investigated in this study, (a) Narrow topographica l load s caus e tensional deviatoric stresses beneath an d compressiona l stresses adjacent t o th e load. Th e maximum stress difference is 0.64 time s the loading pressur e an d define s a semicircular ar c containin g the end s of th e load , (b) Large compensated load s wil l also cause deviatoric tension underneat h th e load, bu t her e the maximum stress difference wil l be higher (approximately equal t o the loadin g pressure) , (c ) Uncompensated wid e loads result i n bending, wher e the uppe r par t o f th e lithosphere beneath th e loa d i s subjected t o compression an d th e flanks to tension . The opposit e situation appear s a t depth . Th e maximum bending stresses appea r fo r load s tha t ar e abou t 4. 4 times the flexural parameter an d wil l b e about si x times th e loading pressure . in th e crus t wil l the n increas e b y a n amoun t proportional t o th e exces s topography . Hori zontal stresse s ar e relativel y unaffecte d an d deviatoric tensiona l stres s wil l therefor e appea r in th e regio n betwee n loa d an d compensatio n and compressio n a t th e loa d edges . Deviatori c stresses clos e to th e weigh t of the exces s surface load ar e thu s calculated, which means that for a 1500m elevate d regio n deviatori c stresse s of th e order o f 40 MPa ma y b e assumed (Tabl e 3) .

Since lat e Paleocen e tim e (6 0 Ma) spreadin g has occurre d fro m bot h th e mid-Atlanti c ridge and it s northwar d continuation , th e Arcti c mid-ocean ridg e (Jackso n & Gunnarsson 1990) . In th e pola r regio n th e ridg e geometr y i s wel l defined an d th e Barent s Se a i s expecte d t o b e subjected t o a SSE-directe d ridg e push fro m th e Arctic mid-ocea n ridg e (Fig . 9a) . I n th e north ern Atlanti c Ocean , th e ridg e (Knipovitc h an d Molloy Ridges ) turns N-S an d i s cut b y several transform zones , and thes e zones complicate th e ridge geometry and reduc e the ridg e push effect . It shoul d therefor e b e anticipate d tha t th e Barents Sea is influenced b y NNW-SSE tectonic stresses, which , fo r a 6 0 Ma crust , generat e maximum deviatori c stresse s o f th e orde r o f 20-30 MPa. Th e observed stres s direction (N-S ) in th e Barent s Se a i s rotate d c . 20° clockwise from th e expecte d ridg e pus h directio n an d hence ridge push alone cannot explain the observations i n thi s region. Further south , th e Mohn s Ridge , a relatively straight ENE-WS W strikin g ridge bounde d b y the Jan Maye n an d Senj a fractur e zones , gener ates a clea r NNW-SS E tectoni c stress . Thi s i s consistent wit h th e observe d stres s directio n i n the Norwegia n Se a and norther n Norway , an d

Table 3. Predicted patterns an d magnitudes of stress associated with topographical loading Type o f loadin g Narrow loa d Wide load , uncompensated Wide load , compensated

Maximum stres s magnitude (MPa) 25

-240 40

Lateral variation

Variation with dept h

Zone o f influence

Compression outsid e loa d Tension underneath Tension outsid e loa d Compression underneath Compression outsid e loa d Tension underneath

Monotonic

Local

Reverses

Insignificant

Monotonic

Regional

All value s are compute d fo r a 1500 m elevation.

460

M. FEJERSKO V & C . D . LINDHOL M

Fig. 9 . Fou r importan t stress-generatin g mechanisms an d thei r regional effec t o n Norxva y an d adjacen t regions , (a) The mid-Atlanti c spreadin g ridg e induce s a NW-SE deviatoric stres s field of the orde r o f 20-30 MPa i n the oldest oceani c crust . Th e magnitud e i s lowered somewha t i n th e thic k continenta l crust, bu t migh t sho w loca l maxima i n areas with thinne d crust , (b ) The continenta l margin s ar e characterize d b y a deviatori c tensiona l stress field in the continental crust , an d deviatori c compressio n i n the oceani c crust , norma l t o th e margin , (c) Post-glacia l uplif t contour s indicat e substantia l vertica l movemen t o f Fennoscandi a sinc e deglaciation o f the region . Deviatori c stresses , tensio n beneat h th e forme r ice. and compressio n beyon d th e ic e edge, ar e associated wit h deglaciatio n i n accordanc e wit h th e mode l o f Stei n e l al. (1989). Maximu m stresse s occu r clos e to th e forme r ic e margin, and wil l reverse bot h wit h dept h an d laterally , (d ) Sediment loadin g cause s bendin g stresses t o develo p i n th e lithosphere , wit h compressio n beneat h th e loa d an d tensio n o n th e flanks wit h stres s reversal a t depth . Bendin g stresse s ar e dependen t o n a hig h sedimentatio n rate . modelling als o her e indicate s maximu m devia toric stresse s o f the orde r o f 20-30 MPa clos e t o the continenta l margin . South o f Jan Mayen , th e Iceland-Ja n Maye n Ridge define s th e mid-Atlanti c spreadin g ridge .

Although th e ocean-floo r geometr y i n thi s region i s complicated b y th e Ja n Maye n micro continent, th e ^Egi r Ridg e an d th e oceani c swel l beneath Iceland , a ne t ridg e pus h forc e i s generated wit h a NW-S E direction . Th e stres s

CRUSTAL STRESS-GENERATIN G MECHANISM S direction ma y show a radial orientatio n aroun d Iceland, a s pointe d ou t b y Bot t (1991) , bu t th e first-order stres s patter n i s expecte d t o b e oriented norma l t o th e ridg e (NW-SE) . Th e stress magnitud e i s difficul t t o asses s fro m a simple coolin g model , bu t i t i s expecte d t o b e similar to or slightly less than that computed fo r the origina l 6 0 Ma oceani c crus t (20-30 MPa). The tectoni c stresse s se t up b y the ridg e pus h are know n t o penetrat e th e entir e nort h Eur opean plat e (Miille r e t al 1992 ; Miille r 1993 ) and wil l therefor e b e activ e bot h i n th e north ern an d souther n Nort h Sea , a s wel l a s i n southwestern Norway . I n th e norther n Nort h Sea and wester n Norwa y th e predicted NW-S E direction fit s th e observation s fairl y well . I n th e central Nort h Se a and southwester n Norwa y a complex stres s patter n appears , bu t als o her e a NW-SE direction i s clearly present i n th e data. The compressional stres s magnitude fro m th e ridge pus h i s affecte d b y crusta l thicknes s an d plate rheology . Becaus e th e crusta l thicknes s of the Balti c Shiel d i s significant , a forc e lik e th e ridge pus h wil l hav e t o b e distribute d ove r a wider dept h range . Thi s wil l eventuall y reduc e the tectoni c stres s regionall y i n centra l part s o f Norway an d Fennoscandia , an d enhance s th e possibilities fo r second-orde r stres s generator s to dominat e th e observe d stres s field . O n th e other hand , th e thinne d crus t underneat h off shore rifte d basin s (mid-Norwegia n margin , Viking an d Centra l Grabens ) enhance s tectoni c stresses an d explain s th e compressiona l foca l mechanisms an d hig h seismi c activit y observe d in thi s regions . In summary , th e ridg e pus h explain s mos t o f the stres s observation s i n th e Norwegia n are a and n o area seems to be outside th e influence of this stress . However , th e observe d stres s direc tion in the Barent s Sea, a s well as local onshor e tensional region s an d possibl e latera l variation s in horizonta l stres s anisotropy , impl y tha t second-order effect s ar e als o present . Othe r stress-generating mechanisms , suc h a s th e con tinental margin effec t an d flexura l stresse s associated wit h sedimen t loadin g o r deglaciation , therefore nee d t o b e taken into consideration.

Continental margin To th e west , th e Barent s Se a i s bounde d b y a N-S trendin g continental margin , whic h generates extensiona l horizonta l deviatori c stresse s perpendicular t o th e margi n i n th e continenta l crust (Fig. 9b) . When superimposing thi s margin effect o n the stresses from th e ridge push a stress rotation ca n appear . Dependin g o n th e rati o

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between th e stresse s se t u p b y th e margi n effec t and th e stresse s generate d fro m ridg e push , th e continental margi n ma y rotat e th e stresse s s o that th e stres s directio n wil l b e mor e o r les s parallel t o th e margin . Clos e t o th e margin , where the deviatoric stresse s from th e continen tal margin are at their maximum, the continental margin effec t wil l b e mos t significan t an d henc e may explai n th e clockwis e rotatio n fro m th e NNW-SSE ridge push directio n t o the observe d N-S direction . In th e Norwegia n Se a the continental margi n is defined b y the Mor e and V0rin g escarpment s and th e shel f edg e of f th e Lofote n islands . The margin here is approximately paralle l to the spreading ridge , which means that th e continental margi n effec t introduce s deviatori c tensiona l stresses i n th e continenta l crus t paralle l t o th e compressional stresse s caused b y the ridge push . The overal l effec t wil l b e a more isotropi c stres s state in th e continental crust nea r th e margin. Further sout h th e continenta l margi n bend s around th e Unite d Kingdo m an d i s relatively distant fro m th e regio n unde r focus . It s effec t i n the North Sea and southern Norway will thus be insignificant.

Sediment loading During Tertiar y an d Quaternar y tim e th e Barents Se a region wa s uplifted an d extensively eroded, an d th e eroded sediment s were redepos ited as a wedge of up to 3 km thickness covering the continental margi n (Reemst e t al. 1994). The sedimentation rat e in Pliocene time (1.6 mm a"1) was sufficientl y hig h t o se t up significan t devia toric stresse s i n th e underlyin g rock. Computa tions indicat e tha t fo r a sedimen t wedg e o f 2-3 km thicknes s coverin g th e continenta l mar gin, tensiona l deviatori c stresse s o f the orde r of 10-20 MPa can be set up in the continental crust close t o th e margin . Th e stresse s alig n wit h th e continental margin effec t an d ma y b e significan t both in reorienting the stress field and in increasing th e stres s anisotropy clos e to th e margi n in the Barents area. In th e mid-Norway and Norwegia n Sea province a stron g differentia l til t an d uplif t o f th e mainland initiate d a stron g erosion an d deposition o f a thic k progradin g sequenc e (abou t 1500m) ove r th e mid-Norwegia n Shel f i n lat e Neogene time . Th e hig h sedimentatio n rat e i n Pliocene time (up to 0.8 mm a"1) may have been large enoug h t o caus e flexura l stresse s o n th e Mid-Norwegian Shelf . Byrkjelan d (1996 ) also found a clea r correlatio n betwee n seismi c activity an d Pleistocen e an d post-Miocen e sedimen t

Table 4 . Stress-generating mechanisms an d their importance fo r th e stress field i n th e Norwegian area Mechanism

Theoretical effect s an d magnitude s

Ridge pus h 20-3 0 MP a fo r 6 0 Ma ol d oceani c crus t Th (continental) Norma l t o ridg e Fennoscandi Compression i n oceani c an d continenta l crust Monotonic wit h depth

Effects i n Fennoscandi a e basi c stres s generato r throughou t a

Crustal thicknes s Modifie s the presen t stres s fiel d Area s wher e th e crus t ha s bee n extensivel y (regional) A forc e wil l generat e highe r tectoni c stres s i n thinne d o r i n area s o f extraordinar y thick thin crus t compare d wit h thic k crus t crus t

Comments Yields a WNW ES E to NNW-SS E stress field, affecting th e whol e region

High compressional stresses i n the thinned crust under th e offshor e basins o f mid- and centra l Norway ma y b e reflecte d b y th e hig h seismi c activity The thic k Fennoscandia n crus t dampen s th e tectonic stresse s i n centra l parts o f Norwa y and Swede n

Continental 10-5 margin Norma (regional) Tensio

0 MPa l t o margi n n i n continental crus t an d compressio n in oceani c crus t Monotonic wit h depth

Probably importan t i n th e wester n Barent s Sea and o n th e Norwegia n continental margin

In th e wester n Barent s Sea , th e continenta l margin effec t supplement s th e tectoni c stres s from ridg e push ; a t th e Norwegia n continental margi n thi s mechanis m counteracts th e ridg e push ; thi s account s fo r the observe d reductio n i n stres s anisotrop y from nort h t o sout h

Sediment loadin g 16 (regional) wit

0 MPa fo r 2 k m o f sediment s (decrease s h time) Compression beneat h basin s an d tensio n at margin s Compression i s strongest i n th e centr e an d decreases outward s Stress directio n reverse s with dept h High sedimentatio n rate s require d

Western Barent s Se a an d tentativel y th e Norwegian margin , Vikin g an d Centra l Grabens

The Wester n Barent s Sea wa s th e locu s o f moderate sedimentatio n rate s i n Paleocen e time; bending stresses i n the Barent s Sea may reinforce th e tectoni c stresse s fro m ridg e push; o n th e Norwegia n margi n an d i n th e Viking an d Centra l Grabens th e sedimentation rate s are lowe r and th e effec t i s less clea r

Deglaciation 2 (regional) wit

0 3 0 MPa fo r 2 3 km ic e load (decrease s h time ) Tension beneat h uplifte d area s an d compression beyon d Most prominen t clos e t o th e forme r ice margi n Stress reverse s wit h dept h

Tentatively affect s whol e regio n

In Norwa y th e effec t wil l b e smal l compare d with othe r mechanisms ; observe d faul t type s in norther n Norwa y ma y b e explained by this mechanism, a s als o th e norma l faultin g mechanisms observed i n coastal area s i n midand wes t Norwa y

Topographical 4 0 MPa fo r 150 0 m elevate d regio n wide loa d Tensio n beneat h th e ic e load an d (regional) compressio n beyon d Monotonic wit h dept h

The mountai n ranges i n western Norwa y ma y support thi s mechanis m

Topography i s moderate (c . 1000m elevation) ; underneath th e elevated region s i n norther n and southwester n Norway tectoni c stresse s from ridg e pus h wil l b e reduced, wherea s i n the surrounding (bot h coastal an d landward ) areas th e tectoni c stresse s ar e expected t o increase slightly ; in offshor e region s t o th e northwest th e directio n o f compressio n wil l be similar to tha t o f ridg e push

Topographical 2 5 MPa fo r 1500 m elevate d regio n narrow loa d Tensio n beneat h th e elevate d region s (local) an d compressio n beyon d

Uncertain

Centres o f seismicit y with normal faultin g mechanisms ar e locate d clos e t o th e glacier s Svartisen and Folgefonn a (remnant s from the glaciation?); thi s may b e explained by loca l areas o f hig h topograph y an d possibl y strengthened b y loca l uplif t connecte d t o the glacier s

Surface relie f 10-3 (local) Compressiona

0 MPa

Throughout region , bu t particularl y in high relief area s o f norther n an d southwestern Norway

The larg e variatio n i n stres s directio n fro m shallow onshor e measurement s along fjord s and valley s in the coastal region s o f norther n and wester n Norway ma y b e explained by local, shallo w topographical effects ; depending o n th e orientatio n of fjords an d valleys, topographica l stres s ma y or ma y no t interact wit h the regiona l stress

Faults Modifie (local) Stres

s the presen t stres s field s rotate s norma l o r paralle l t o th e faults

Throughout th e region , but mos t clearly observed i n the Nort h Se a basin s

Large variatio n in stres s direction (sometimes rotation?) i n th e Nort h Se a basins may b e explained b y low horizontal stress difference s favouring loca l influenc e fro m fault s

Salt diapir s beneat h th e souther n Nort h Sea Gneiss window s in mid-Norwa y

The comple x stres s patter n i n the Centra l Graben ma y possibl y b e explained by low horizontal stres s difference s favourin g loca l influence fro m sal t an d othe r domin g structures The ver y hig h tectoni c stresses measured i n the gneissic areas compared wit h surrounding Caledonian area s may b e explained by variation s i n fractur e systems , an d stiffness contrast s

l stres s acts paralle l t o th e surface gradien t Only important ver y clos e t o th e surface

Inclusions Modifie (local) chang

s the presen t stres s field (up t o 50 % e i n stres s magnitude) Stress bend s aroun d sof t inclusion s such a s salt diapir s an d i s directed int o area s o f high stiffnes s suc h a s massiv e roc k masse s

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depocentres o n th e mid-Norwegia n shelf , sub stantiating th e ide a o f sedimentation-induce d stresses. Analysi s o f foca l mechanism s under neath th e loade d are a yield s a compressiona l regime wherea s onshor e mechanism s indicat e a tensional regime . I t is , however , difficul t t o assess th e quantitativ e importanc e o f th e sedi ment loadin g effect , a s it acts constructively an d in th e sam e directio n a s ridg e push . In Pleistocen e time , th e Vikin g an d Centra l Grabens subside d an d th e sedimentatio n rat e was relativ e low . Hence , bendin g stresse s ar e relatively small and th e sediment loadin g effec t i s expected t o b e negligibl e here .

Deglaciation The deglaciatio n model s indicat e maximu m deviatoric stres s concentration s clos e t o th e former ic e edge , wit h a shif t i n polarit y acros s the edg e (largel y along th e coast ) (Fig . 9c) . Thi s is als o t o som e exten t observe d i n th e region , with compressiona l offshor e foca l mechanism s and som e shallo w onshor e mechanism s (Stord , Oslo an d Mel0y-Steigen ) indicatin g a tensiona l stress regime . Th e observation s thu s suppor t the deglaciatio n mode l o f Stei n e t al. (1989 ) well. However , th e onshor e faultin g ma y als o be explaine d b y sedimen t loadin g an d topog raphy effects . Finnmark an d th e souther n Barent s Se a have also been covered by thick ice sheets (Kjemperud & Fjeldskaa r 1992) . possibl y inducin g tectoni c stresses in connection wit h deglaciation. Th e observed shallow reverse faulting in Finnmark, with a NNW-SS E compressio n reporte d b y Olese n et al. (1992) and Bungu m & Lindholm (i n prep.), may favou r Stephansson' s uplif t model , bu t n o lateral shif t i n stres s regim e i s identifie d fro m the availabl e data. Th e deglaciatio n effec t i n th e northern par t o f Norway i s therefore believe d t o be either negligibl e or maske d b y othe r regiona l effects. Th e observed compressive stress regime is more probabl y connecte d t o effect s suc h a s th e compression from ridg e push. If th e deglaciatio n effec t i s a n importan t source fo r th e observe d stres s fiel d i n Norwa y the observation s surel y favou r th e mode l o f Stein e t al . (1989).

Topography The bathymetr y of th e Barent s Se a i s relativel y shallow an d th e onshor e relie f i s lo w wit h gentle slopes . Th e topographica l effects , bot h on a regiona l an d o n a local scale , ar e therefor e not expecte d t o play an important rol e in north ern Norway.

Mid-Norway i s characterized b y locall y high altitude mountains , especiall y i n th e norther n part o f th e region , an d topograph y clearl y influences th e near-surfac e stresses . Th e moun tains ca n b e regarde d a s narro w load s an d modelling ha s show n tha t tensiona l deviatori c stresses exten d downwar d t o a dept h compar able wit h th e widt h o f th e load . I n th e Mel0y Steigen area s ther e ar e example s o f location s relatively nea r hig h mountai n complexes . Her e earthquake swarms , indicatin g inhomogeneou s stress conditions , ar e observed ; however , th e shallow norma l faultin g foca l mechanism s ar e contrary t o wha t shoul d hav e bee n expecte d from topographica l loadin g (Fig . 8) . Western Norwa y i s also characterize d b y high mountains cu t b y dee p fjord s an d valleys , an d hence, nea r i n situ measurement s ar e ver y much influenced b y loca l topographica l features . Th e mountain rang e als o contribut e o n a mor e regional scale, as the large topographical feature s act a s a wid e compensate d load , introducin g tensional deviatoric stresses underneat h elevate d onshore areas. The shallow tensional focal mechanisms observed i n southwestern Norwa y can be attributed t o topographica l effects . The relie f i n souther n an d easter n Norwa y is lower an d mor e gentl e than i n western Norway. Hence, topographica l effect s wil l no t b e o f th e same importanc e a s i n th e high-relie f an d high altitude area s an d ma y therefor e b e neglected .

Discussion The N- S stres s directio n observe d alon g th e western margi n o f th e Barent s Se a i s wel l explained b y th e ridg e push , whe n modifie d i n azimuth b y th e continenta l margi n effec t an d flexure associated wit h rapid sedimentation . The ridge pus h i s also believe d t o caus e th e shallo w reverse faultin g an d hig h horizonta l stres s magnitudes measure d i n norther n Norway . Th e consistent stres s direction i n this region indicate s a relativel y hig h stres s anisotropy . I n contrast , the lo w seismi c activit y shoul d indicat e lo w stress anisotropy . Thi s behaviou r i s difficul t t o explain, an d a s ye t i s not full y understood . The observe d regiona l stres s directio n o n th e Norwegian margi n i s consisten t wit h th e direc tion o f ridge push, an d th e high tectonic stresse s observed ar e als o attributabl e t o thi s source . Because o f th e loca l thinnin g of th e continental crust, the tectonic stresse s ma y be amplified, an d this ma y explai n th e revers e faultin g an d hig h seismic activit y observe d offshore , whic h i n a broad sens e i s concentrated i n area s o f thinne d crust. Sedimen t loadin g effect s als o induc e

CRUSTAL STRESS-GENERATIN G MECHANISMS offshore compressional deviatori c stresses underneath th e basins. The observed stres s anisotropy seems t o b e lowe r o n th e Norwegia n margi n than i n th e Barent s Sea, presumably because o f the continenta l margi n effec t counteractin g th e ridge pus h effect . A s thre e effect s (sedimen t loading, topograph y an d deglaciation ) produc e more or less the same stress pattern in this area it is difficul t t o evaluat e their relative importance. The regiona l stres s directio n i n wester n Norway an d th e norther n Nort h Se a correlate s well wit h th e directio n o f ridg e push . Loca l thinning o f th e crus t i s als o assume d t o b e a n important facto r i n amplifyin g tectonic stresses in th e continenta l crus t underneat h th e graben . This explain s th e larg e amoun t o f earthquake s (many wit h reverse foca l mechanisms ) observe d in thi s region . Sedimen t loadin g fits partly wit h the observe d stresses , an d ma y also possibl y b e an importan t facto r i n th e area s o f extrem e sedimentation rate s in recent time . The observed local stres s reorientatio n in th e norther n Nort h Sea may indicat e tha t stres s anisotrop y i s lower in this region compared wit h other areas, o r tha t local feature s ar e prominen t an d deflec t th e regional stres s field . Onshore , th e shallo w normal faultin g activity is attributed t o regional and loca l topography , bu t regiona l deglaciation effects ma y als o b e important . Possibl e impor tant loca l stres s modifier s ar e faults , a s wel l as loca l topographica l feature s suc h a s fjord s and valleys . The majo r stres s directio n i n southwester n Norway an d th e centra l Nort h Se a coincide s with th e 'Europea n trend ' (Miille r e t al 1992) , indicating that ridg e push is an important sourc e here a s well. None of the other stress-generatin g mechanisms examine d i n thi s pape r see m t o b e of great importanc e i n this region . The complex stress pattern observed from borehol e breakouts in th e Centra l Graben are a i s primarily believed to b e a resul t o f low-qualit y data , bu t ma y also tentativel y b e cause d b y lo w horizonta l stress anisotrop y an d significan t local influence by geologica l structure s suc h a s fault s an d salt diapirism . The NW-S E stres s directio n observe d i n southeastern Norwa y ma y b e attribute d t o th e ridge pus h effect , bu t a s crusta l thicknes s increases eastwar d th e regio n wil l probabl y b e subjected t o somewha t lowe r tectoni c stresses . Of th e othe r effect s analyse d i n thi s paper onl y the deglaciatio n effec t wa s foun d t o b e abl e t o influence th e stres s fiel d i n thi s region . Ver y tentative indicator s (reversa l o f stres s regim e with dept h indicate d b y foca l mechanism s an d N-S t o NE-S W stres s direction s fro m over coring) favou r th e deglaciatio n mode l Stei n

465

et al.lt shoul d b e note d tha t als o i n thi s regio n the zone s o f enhance d seismi c activit y coincid e with area s o f thinne d crus t (th e Oslo Graben ) and tha t near-surfac e overcorin g measurement s often yiel d directions paralle l to th e major faults in thi s region .

Conclusions All stress-generatin g mechanism s examine d i n this pape r ar e summarize d i n Tabl e 4 togethe r with a brie f descriptio n o f thei r importanc e fo r the Norwegia n area . Th e interna l rankin g ha s been base d o n latera l extensio n o r regiona l influence, a s mos t mechanism s ar e capabl e o f generating hig h deviatoric stresses. Som e o f th e mechanisms, know n a s non-renewable , reliev e with time , an d th e magnitude s reporte d ma y exaggerate thei r importance , a s the y ar e maximum value s compute d fo r a perfectl y elasti c crust. Other , mor e loca l mechanisms, whic h ar e not formall y stres s generators , bu t mor e pre cisely describe d a s modifiers, are als o include d in Tabl e 4 . The onl y continenta l stress-generatin g mech anism tha t ca n b e responsibl e fo r th e hig h compressional stresse s observed throughou t th e Norwegian regio n i s th e ridg e pus h effect . Modelling indicate s tha t horizonta l deviatori c stresses from ridge pus h i n the oceani c crus t ar e of th e orde r o f 20-3 0 MPa. I n th e continenta l crust th e change s i n tectoni c stresse s depen d on crusta l thicknes s an d abilit y o f th e crus t t o accommodate differentia l stresses . Region s wit h thinned crust , suc h a s th e Norwegia n margin , the Nort h Se a Grabe n syste m an d th e Osl o Graben, wil l hav e highe r tectoni c stresses . Thi s is also clearl y demonstrate d b y th e enhance d seismic activit y in thes e regions. The othe r stress-generatin g mechanisms analysed her e ar e foun d t o b e o f variabl e regional importance. The y modif y th e tectoni c stresse s generated b y th e ridg e pus h an d ar e respon sible fo r locall y observe d stres s reorientation , horizontal stres s anisotrop y an d variation s i n seismicity. Close t o th e wester n margi n o f th e Barent s Sea both th e continental margi n an d recen t hig h sedimentation rate s ac t constructivel y wit h th e ridge pus h an d pla y a n importan t rol e i n rota ting an d strengthenin g th e tectoni c stres s field. In th e Norwegia n Se a th e continenta l margi n acts agains t th e ridg e push , causin g a lowe r stress anisotrop y in thi s region. Along the coast of southern Norwa y local and regional topograph y play s a n importan t rol e i n modifying stres s magnitude s an d explainin g th e shallow norma l foca l mechanism s observed.

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The deglaciatio n o f Fennoscandi a ha s bee n regarded as an important stress-generating mechanism. Th e modellin g di d no t indicat e thi s a s a sourc e o f hig h stres s anisotropy , bu t al l along the coastal area s and i n southeastern Nor way th e observe d stres s direction s and tectonic s are i n reasonabl e accor d wit h prediction s fro m deglaciation. Stress modellin g base d o n a n elasti c plat e behaviour i s a coars e simplificatio n o f th e crust's brittl e ductil e behaviour ; nevertheless , the model s ca n b e applie d qualitatively . Fo r more accurat e determinatio n o f stres s magni tudes viscoelasti c modellin g i s required , wher e also th e yield strengt h i s accounte d fo r (Cloe tingh & Buro v 1996) . Furthe r researc h shoul d therefore b e base d o n mor e realisti c models . A quantitativ e evaluatio n o f th e contributio n from differen t stress-generatin g mechanisms can then b e conducte d unde r otherwis e identica l mechanical an d boundar y conditions . Thi s wil l shed more light on the mutual contribution fro m topography, sedimen t loadin g an d deglaciation , taking bot h loadin g histor y (duratio n an d tim e since maximu m loading ) an d latera l influenc e into consideration .

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Index

Page number s i n italics refer t o Figure s an d Table s jEgir Ridg e 37 4 Artinaskian-Ladinian, stretchin g 94-96 asthenosphere, anomalie s 42 4 Atlantic rif t zon e 239 , 242 , 29 5 geological framewor k 296-29 8 palaeogeography 307-31 6 volcanism 316-32 0 Baltic delta 22 0 Barents Se a 295-307, 313 crustal stres s 436 , 442 , 446-448 , 453-464 basalt 413-41 4 flood basal t 307 , 320 , 353 basement rock s 3 4 basin infil l 167-17 3 basin modellin g 15-16 , 84-100, 89, 273-292 basin overfil l 28 6 basin subsidence , Vikin g Graben 64-6 6 Bathonian-Berriasian, stretchin g 96-97 Bergen Arc s 3 1 biostratigraphy, Tertiar y 227-22 8 Bivrost Fractur e Zon e 328, 329 Bivrost Lineamen t 328, 334-337 , 348 , 353 , 379-39 5 Brent delt a 141-14 4 Brent Grou p 64 , 72 , 10 7 structure 61-63, 10 9 British Tertiar y Igneou s Projec t 31 6 Carboniferous sedimentary rock s 329 tectonic activit y 337 Cenozoic successio n 24 5 absolute ag e conversion 22 8 biostratigraphy 224-22 8 computer simulatio n 273-29 2 environmental influence s 24 5 geological evolutio n 219-242 mineralogical analysi s 223-224 , 246-247 mineralogical chang e 268-27 0 sediment compositio n 245 , 247-270 seismic stratigraph y 224-242 , 245-271 , 276-279 tectonic subsidenc e 286-29 2 see also Tertiary sediment s computer simulatio n basin modellin g 273-29 2 continental rifting , N E Atlanti c 296-321 Cretaceous Atlantic rif t zon e 299 , 308-31 2 basin formatio n 340 , 343 sediments 27 6 structural evolutio n 379-38 6 subsidence 275 tectonic activit y 329, 35 7 uplift 35 3 crustal fault s 55 crustal stres s 429 , 441-446, 451-452 in situ observation s 441-44 8 Norway 432-438 , 442-446, 447, 453-454 see also stress-generatin g mechanism s

crustal structur e 29-30, 36 , 46-49 data analysi s 19-2 9 crustal thicknes s estimate s 91-93 , 99 crustal thinnin g 34-36 /^-curves 3 6 estimates 298 see also Permo-Triassi c rifting , Jurassi c riftin g Deep Se a Drillin g Project 32 7 DEMOSTRAT 273 , 282-284 detachment fault s 105-12 8 modelling 125-12 7 Devonian sediment s 17 , 96, 99 , 11 7 domino fault s 112-113 , 117 , 124-12 5 D0nna Terrac e 34 2 Draupne Formatio n 191-199 , 200 earthquakes 432-43 8 focal mechanism s 429-438, 431 East Shetlan d Basi n 30 East Shetlan d Platfor m 17 , 219 Tertiary sedimen t depositio n 231 , 276 eclogite formatio n 3 1 Eocene flood basal t 353 , 35 7 stratigraphy 228-229 , 239 , 253 , 331 , 348 Erlend Platfor m 32 9 eustacy 144 , 239, 241-242, 245 , 284 exploration 32 7 extensional modellin g 45-4 6 Faeroe Platea u 35 9 Faeroe-Shetland Basi n 329 Faeroe-Shetland escarpmen t 357 , 359 fault geometr y 51-55 , 105 fault plan e reflectio n 50 , 51-5 3 Fjerritslev faul t zon e 21 9 Fles faul t comple x 332 , 340 , 342 , 34 8 reverse reactivatio n 34 8 flexural basin mode l 66-7 7 footwall uplif t 76 , 91, 96 , 202-206 forward modellin g 85 , 91, 273 Froan Basi n 337 Fr0ya Hig h 33 7 • Fulla Ridg e 340 Fur Formatio n 22 8 geochemistry, Cenozoi c sediment s 245-27 0 Gjallar Ridg e 340-34 2 glacial reboun d 456 , 464 graben system s 19 , 108 gravity modellin g 29 Greenland 296-29 7 NE Greenlan d margi n basi n 309-31 1 Greenland Se a see Norwegian-Greenland Se a Gullfaks detachmen t 113 , 12 6 Gullfaks faul t block 91 , 109-11 7 Gullfaks Fiel d 110-113 , 117 , 123-128 Hakon Mosb y Mu d Volcan o 408 Halten Terrac e 329 , 337-340 , 342

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INDEX

hanging wall subsidence 9 1 Hatton-Rockall margi n 295 , 304, 306-307 HEAT 40 1 heat flo w 320 , 397-408 surveys 40 0 Heather Formatio n 18 4 Hel Graben 340 , 34 2 Helland Hanse n Arc h 300 , 348 Horda Platfor m 17 , 31, 36 , 48, 9 1 Tertiary depositio n 228 , 236, 239 , 276-27 8 uplift 24 1 horst an d grabe n structur e 352, 353 Hutton alignmen t 96 , 109-11 0 Iceland mantl e plume 316 , 413 , 423 Integrated Basi n Studie s Projec t 1 database 3 results 4-10 inter-rift period s 14 1 intra-basement detachmen t 117-12 3 intra-mantle reflection s 49, 53, 55 isostacy 88 , 284 Jan Maye n Fractur e Zon e 328, 329 , 334 Jan Maye n Lineamen t 328, 332-337 , 348 , 35 3 Jan Maye n Microcontinen t 37 4 Jennegga Hig h 35 3 Jurassic riftin g modelling 96-100 sediments 14 1 structure 13 8 see also Viking Graben Klakk Faul t Comple x 35 9 Kolbeinsey Ridg e 374-37 5 Labrador Se a 297 , 31 1 large igneou s province s 413, 421-422 crustal emplacemen t 42 2 North Atlanti c 414 , 422-423 listric fault s 54-55, 105-10 6 lithosphere extension mode l 67-7 9 flexural strength 66-67 , 70-72, 75-77 , 99-100 lithostratigraphic nomenclature 137 Lofoten Ridg e 329 , 35 3 low angl e fault s se e detachment fault s magmatic rock s Norwegian Se a 366 see also sill s magmatism 31 6 magnetic anomalie s 295 , 334 , 420 magnetic modellin g 29 Magnus Basi n 3 0 mantle fault s 5 3 mantle plume s 316-319 , 413, 423 mantle structure s 48-9 , 53-4, see also mantle plume s Mesozoic sediment s 61 , 85 Mid Atlanti c Ridg e 417 middle Miocen e unconformit y 34 8 middle Triassi c subsidenc e 139-14 1 Miocene sequenc e 231-236 , 241 , 263 , 266 Modgunn Arc h 34 8

Moho 31-34.48 , 74 , 85, 92 M0re Basi n 328 faulting 35 9 formation 297 , 305 , 353-36 5 stratigraphy 330-33 2 tectonic framewor k 329 , 369-375 , 370 M0re Margi n 305-30 7 More Margina l Hig h 328, 357, 359 M0re-Tr0ndelag faul t comple x 35 7 Naglfar Dom e 348 Nagrind Synclin e 340 , 342 , 386 Namibia se e North Namibi a margi n necking dept h 68-69 , 78 , 88 Nordland Ridg e 342 , 35 9 North Namibi a margi n 414-419 , 423 North Se a basin formatio n 1 5 rift syste m 17 , 107-11 0 North Se a Basin, geologica l settin g 274-27 6 Northern Nort h Se a 134 basement rock s 3 4 basin infil l 167-17 3 basin model s 15-16 , 60-66, 84-100 , 273-29 2 Cenozoic successio n 24 5 crustal geometr y 46-55 crustal structur e 16 , 29-36, 91-93 gravity an d magneti c data 2 9 seismic data 19-2 9 faults 105-12 8 geological settin g 17-19 , 46, 274-276 lithostratigraphy 13 7 rift 108-110,306-30 7 sedimentation 90, 172-17 3 tectonic models 52-5 3 Norway Margin , hea t flow 407 Norwegian Petroleu m Directorat e 32 7 seismic dat a 329 Norwegian Se a exploration 32 7 structure 328, 329, 379-38 6 tectonic histor y 329 , 331, 366, 367-37 5 unconformities 331-332 , 334, 348 Norwegian stres s province s 447 Norwegian-Danish Basi n biostratigraphy 227 mineralogical analysi s 223-22 4 seismic interpretatio n 221-227 , 23 0 stratigraphy 222-223 Tertiary sequenc e 228-23 6 Norwegian-Greenland Se a heat flow measurements 397-408 , 398, 400 thermal conductivit y 399, 402-403 Nyk Hig h 340 , 342 , 34 8 structural evolutio n 379-388 oceanic crust , therma l propertie s 39 7 Oligocene sequenc e 231 , 239-241, 257 , 263 opal-CT transition 270 , 365 , 36 7 Oseberg formatio n 65 , 70 , 140 , 148-15 1 0ygarden faul t zon e 51-52 , 85 , 97, 109-110 , 140 . 219 Tertiary sediment s 228 , 231 , 236 , 27 6

INDEX palaeo-waterdepth 9 8 palaeoenvironment 42 4 Paleocene sequence 239 , 248-249, 331 Penguin half-grabe n 153-15 8 Permian, sedimentar y rock s 329 Permo-Triassic riftin g modelling 94-10 0 sediments 138-13 9 structure 13 8 see also Hord a Platfor m planar fault s 10 5 plate reconstructio n 307-31 6 plate tectonics , Norwegia n Se a 367-375 Pliocene sequenc e 236 , 241 , 266-26 8 post-rift sediment s 6 1 loading 8 8 modelling 279-290 uplift 7 1 see also syn-rif t sediments , Tertiar y sediments , Cretaceous sediment s progradation 222-223, 227-228, 231 , 236 , 239 Ras Basi n 330-331 , 340 reservoir facie s 179-18 1 resource evaluatio n 424 restored basi n geometr y 307-31 6 reverse modellin g 8 5 Revfallet Faul t Comple x 342 , 359 rheology, couple d v decoupled 54 , 66-67, 70-72, 75-77, 99-100 Ribban Basi n 35 3 ridge pus h forc e 455-456 , 459-461 rift evolutio n 73 , 140-14 1 rifting se e Permo-Triassic rifting , Jurassi c riftin g rifting M0re Basi n 357 Tr0ndelag Platfor m 337-34 0 V0ring Basi n are a 340-35 0 Rockall Troug h 297 , 306-307 , 311-31 2 rotational faulting 83 , 140 , 14 1 Rym Faul t Zon e 34 2 salt diapir s 31 3 sea floo r spreadin g 329 , 373-374, 413 sediment compaction 89 , 279-283 correction 63-6 4 composition, Cenozoi c 245, 247-27 0 transport coefficient s 283-28 4 sediment loadin g 457-459 , 461-464 seismic dat a 19-29 , 42-45, 85 , 105 seismic stratigraphy , Cenozoi c 224-242 , 245-271, 276-279 shear model s 15-16 , 45 Shetland Platfor m 96 , 99 sills 320 , 367 , 41 4 S10rebotn Sub-basi n 35 7 Snorre faul t bloc k geological settin g 18 4 syn-rift evolutio n 180-21 5 stratigraphy 188-19 1 Snorre half-grabe n 161-16 3

Snorre-H footwall uplif t 202-20 6 tectonostratigraphic evolutio n 206-21 0 Sorgenfrei-Tornquist zon e 219 South Atlantic , structur e 417 South Atlanti c margins 414 structural feature s 417 Statfjord-Statfjord Nort h are a 158-16 1 stratigraphic modelling, problems 97-9 9 stress-generating mechanisms 454-466 stretching distribution 93-97 sub-aerial depositio n 286-28 7 sub-aerial erosio n 287 , 290 submarine erosion , Miocen e 348 subsidence 91 Tertiary 219-220 , 245 , 284, 287 , 290-292 Triassic-Jurassic 139-14 1 see also therma l subsidenc e supra-basement detachmen t 110 , 123-12 8 Surt Lineamen t 342-343 , 379-39 5 Svalbard Margin , hea t flow 407 syn-rift evolutio n 199-20 6 syn-rift sediment s 61 , 90, 138-139 , 141-17 3 stratigraphy 8 8 see also post-rif t sediments , Snorre faul t bloc k Tampen Spu r 30, 11 0 lithostratigraphy 18 1 Tarbert formatio n 141 , 146 , 148 tectonic movements, se e also therma l subsidenc e tectonics subsidenc e 284-29 2 tectono-magmatic model s 319-32 0 Tertiary reactivation phase s 34 8 structural evolutio n 386-388 volcanism 413 Tertiary sediment s isopach map s 238 Norwegian-Danish Basi n 228-24 2 seismic data 221-22 3 sonic lo g velocity 221 , 23 2 thermal evolutio n 97 thermal plume s se e mantle plume s thermal subsidenc e 219 , 275, 284-292, 318 , 340 Traena Basi n 330-331, 340 Tristan plum e trai l 414, 423 Tromso Basi n 299-300, 307 , 313 Tr0ndelag Platform 328, 330 , 337-34 0 Unst Basi n 89 , 138 uplift 99 , 219 , 239, 241-242, 316 Cretaceous 35 3 Lofoten-Utr0st are a 35 3 Tertiary 35 7 V0ring Basi n 348 Upper Cretaceous , stratigraph y 329 upper Paleocen e sequenc e 228-229 US Eas t Coas t margi n 419-420, 423 Utr0st Ridg e 353 Vema Dom e 34 8 structural evolutio n 379-38 8 Vestfjorden Basi n 35 3 Vigra Hig h 35 7

471

472

Vigrid Synclin e 340 , 342 Viking Grabe n 51 , 9 1 basin modellin g 60-7 7 fault geometr y 48 , 105-128 formation 1 7 lithosphere extensio n 67-7 9 rifting 59-6 1 sediments 141-15 1 Tertiary depositio n 219, 228-231 , 276-279 structure 61-6 6 subsidence 64-66 , 139-14 0 Vingleia faul t comple x 33 7 Visund detachmen t 117 , 72 2 Visund faul t bloc k 11 3 Visund half-grabe n 163-16 4 Visund S0r0s t detachmen t 113-117 . 123 , 126 volcanic margi n formatio n 316 , 319-320, 411-42 4

INDEX Voring Basi n 32 8 folding 34 2 formation 297 , 316, 340-35 0 sediment 21 9 stratigraphy 329-33 2 structural evolutio n 379-38 6 tectonic framewor k 329 , 369-375. 37 0 Voring escarpmen t 346-348 , 351 Voring Margin , crusta l thinnin g 299-30 6 Voring Margina l Hig h 303 . 328, 350-353 water dept h 98 , 279. 286-287 West Afric a margi n 42 3 West Australi a margi n 42 3 Western Gneis s Regio n 3 1 World Stres s Map (WSM ) 42 9