Exhumation of the North Atlantic Margin: Timing, Mechanisms and Implications for Petroleum Exploration (Geological Society Special Publication, No. 196)

Exhumation of the North Atlantic Margin : Timing, Mechanisms an d Implications fo r Petroleum Exploratio n

Geological Society Special Publication s Society Book Editors A. J . FLEE T (CHIE F EDITOR ) P. DOYL E F. J . GREGOR Y J. S . GRIFFITH S A. J . HARTLEY R. E . HOLDSWORT H

A. C . MORTO N N. S . ROBIN S M. S . STOKE R J. P . TURNE R

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It is recommended that referenc e to all or part of this boo k shoul d be made i n one of the followin g ways: DORE, A . G. , CARTWRIGHT , J . A. , STOKER , M . S. , TURNER , J . P . & WHITE , N . (eds ) 2002 . Exhumation of th e North Atlantic Margin: Timing, Mechanisms an d Implications for Petroleum Exploration. Geo logical Society , London , Specia l Publications , 196 . BLUNDELL, D . J . 2002 . Cenzoi c inversio n an d uplif t o f souther n Britain. In: DORE , A . G., CARTWRIGHT , J. A. , STOKER , M . S. , TURNER , J . P . & WHITE , N . (eds ) Exhumation o f th e North Atlantic Margin: Timing, Mechanisms an d Implications fo r Petroleum Exploration. Geologica l Society , London , Specia l Publications, 196 , 85-101 .

GEOLOGICAL SOCIET Y SPECIA L PUBLICATION No. 196

Exhumation o f the North Atlantic Margin: Timing, Mechanisms and Implications fo r Petroleu m Exploration EDITED B Y

A. G. DORE Statoil, UK

J. A. CARTWRIGHT Cardiff University , UK

M. S. STOKE R

British Geological Survey, Edinburgh, UK

J. P. TURNE R

University of Birmingham, UK and

N. WHIT E Bullard Laboratories, Cambridge, UK

2002 Published by The Geological Societ y London

THE GEOLOGICAL SOCIETY

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Contents DORE, A. G. , CARTWRIGHT , J. A. , STOKER , M . S. , TURNER , J . P . & WHITE , N. J . Exhumatio n of 1 the North Atlanti c margin : introductio n an d background Mechanisms JONES, S . M. , WHITE , N. , CLARKE , B . J. , ROWLEY , E . & GALLAGHER, K . Presen t an d pas t 1 influence o f the Iceland Plum e on sedimentation

3

ROHRMAN, M. , VA N DE R BEEK , P . A. , VA N DER HILST , R . D . & REEMST , P . Timin g an d 2 mechanisms of North Atlantic Cenozoic uplift : evidenc e fo r mantle upwelling

7

NIELSEN, S . B. , PAULSEN , G . E. , HANSEN , D . L. , GEMMER , L. , CLAUSEN , O . R. , JACOBSEN , B . 4 H., BALLING , N. , HUUSE , M . & GALLAGHER , K . Paleocen e initiatio n o f Cenozoi c uplif t i n Norway

5

GRAVERSEN, O . A structura l transec t betwee n th e centra l Nort h Se a Dom e an d th e Sout h 6 Swedish Dome : Middl e Jurassic-Quaternar y uplift-subsidenc e reversa l an d exhumatio n acros s the eastern Nort h Se a Basin

7

BLUNDELL, D. J. Cenozoic inversio n an d uplif t o f souther n Britain 8

5

Scandinavia, Greenland an d adjacent margi n LlDMAR-BERGSTROM, K. & NASLUND , J. O . Landforms an d uplif t i n Scandinavi a 10

3

HENDRIKS, B . W . H . & ANDRIESSEN , P . A . M . Patter n an d timin g o f th e post-Caledonia n 11 denudation of northern Scandinavi a constraine d by apatite fission-track thermochronology

7

EVANS, D. , McGiVERON , S. , HARRISON , Z. , BRYN , P. & BERG , K . Along-slop e variatio n i n th e 13 late Neogene evolutio n of the mid-Norwegian margin in response to uplift an d tectonism

9

STROEVEN, A . P. , FABEL , D. , HARBOR , J. , HATTESTRAND , C . & KLEMAN , J . Reconstructin g th e 15 erosion histor y o f glaciate d passiv e margins : application s o f i n situ produce d cosmogeni c nuclide technique s

3

CEDERBOM, C . Th e thermotectoni c developmen t o f souther n Swede n durin g Mesozoi c an d 16 Cenozoic tim e

9

JAPSEN, P. , BIDSTRUP , T . & LIDMAR-BERGSTROM , K . Neogen e uplif t an d erosio n o f souther n 18 Scandinavia induced by the rise of the Sout h Swedish Dom e

3

HUUSE, M . Cenozoi c uplif t an d denudatio n o f souther n Norway : insight s fro m th e Nort h Se a 20 Basin

9

FALEIDE, J . L , KYRKJEB0, R., KJENNERUD , T., GABRIELSEN , R. H. , JORDT, H. , FANAVOLL, S. & 23 BJERKE, M . D . Tectoni c impac t o n sedimentar y processe s durin g Cenozoi c evolutio n o f th e northern Nort h Se a and surrounding area s

5

UK, Ireland and adjacent margin HALL, A . & BISHOP , P . Scotland' s denudationa l history : a n integrate d vie w o f erosio n an d 27 sedimentation a t an uplifted passiv e margin

1

ANDERSEN, M . S. , S0RENSEN , A. B. , BOLDREEL, L . O . & NIELSEN, T . Cenozoi c evolutio n o f th e 29 Faroe Platform , comparing denudatio n and depositio n

1

STOKER, M. S. Late Neogen e developmen t o f the UK Atlantic margi n 31

3

GREEN, P . F. , DUDDY , I . R . & HEGARTY , K . A . Quantifyin g exhumatio n fro m apatit e fission - 33 1 track analysi s an d vitrinit e reflectanc e data : precision , accurac y an d lates t result s fro m th e Atlantic margin of NW Europe

WARE, P. D. & TURNER, J . P. Sonic velocity analysi s o f the Tertiar y denudatio n o f the Iris h Se a 35 basin

5

ALLEN, P . A. , BENNETT , S . D. , CUNNINGHAM , M . J . M. , CARTER , A. , GALLAGHER , K. , 37 LAZZARETTI, E. , GALEWSKY , J. , DENSMORE , A . L. , PHILLIPS , W . E . A. , NAYLOR , D . & HACH , C. S. The post-Varisca n thermal an d denudational histor y o f Ireland

1

Implications for petroleum exploration DORE, A . G. , CORCORAN , D . V . & SCOTCHMAN , I . C . Predictio n o f th e hydrocarbo n syste m i n 40 exhumed basins , an d applicatio n t o the NW European margi n

1

PRICE, L . C . Geologica l an d geochemica l consequence s o f basi n exhumation , and commercia l 43 implications

1

PARNELL, J. Diagenesis an d fluid flow in response t o uplift an d exhumation 43

3

CRAMER, B. , SCHLOMER , S . & POELCHAU , H . S . Uplift-relate d hydrocarbo n accumulations : th e 44 release o f natural gas from groundwate r

7

CORCORAN, D . V . & DORE , A . G . Depressurizatio n o f hydrocarbon-bearin g reservoir s i n 45 exhumed basin settings : evidenc e fro m Atlanti c margin and borderland basin s

7

Index 48

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Exhumation o f the North Atlanti c margin : introduction and background A. G. DORE 1, J. A. CARTWRIGHT 2, M. S. STOKER 3, J. P. TURNER4 & N. J. WHITE 5 l Statoil (UK) Ltd, lla Regent Street, London SW1Y 4ST, UK (e-mail: agdo@ statoil.com) Department of Earth Sciences, Cardiff University, PO Box 914, Cardiff CF10 BYE, UK ^British Geological Survey, Murchison House, West Mains Road, Edinburgh EH9 SLA, UK ^University of Birmingham, School of Earth Sciences, Edgbaston, Birmingham B15 2TT, UK 5 Bullard Laboratories, Madingley Rise, Madingley Road, Cambridge CB3 OEZ, UK

Since consolidatio n durin g th e Caledonia n an d Variscan orogenies , N W Europ e ha s undergon e repeated episode s o f exhumatio n (th e exposur e of formerl y burie d rocks ) a s a resul t o f suc h factors a s post-orogenic unroofing , rift-shoulde r uplift, hotspo t activity , compressiv e tectonics , eustatic sea-leve l change , glaciatio n an d iso static readjustment . Modern measuremen t tech niques, suc h a s apatit e fission-trac k analysis , have helpe d t o establis h usefu l denudatio n chronologies fo r thi s entir e tim e span . How ever, th e mai n observationa l legac y o f exhumation aroun d th e Nort h Atlanti c i s preserve d i n the comparativel y youn g (Mesozoi c an d Cenozoic) geologica l recor d o f thi s region . This i s clearl y reflecte d b y th e unifyin g them e of thi s volume , whic h document s evidenc e fo r the widesprea d uplif t an d emergenc e o f larg e sections o f th e Nort h Atlanti c margi n i n Cenozoic time . All student s o f N W Europea n geolog y ar e aware o f th e compellin g palaeogeographica l evidence fo r th e transitio n a t th e en d o f th e Cretaceous fro m shel f sea s an d low-relie f landmasses t o a n are a dominate d b y highland s and newl y emergen t landmasses , flanke d b y shelves dominate d b y rejuvenate d clasti c depo sition. Similarly, it is also widely known that the highlands o f Norwa y an d Scotlan d d o no t represent th e origina l Caledonia n mountai n range bu t mus t b e instea d a produc t o f lat e emergence o r uplift . The Cenozoi c uplif t o f Fennoscandi a i n particular ha s a lon g histor y o f study . I t i s arguably one of the oldest debate s i n the history of systemati c geolog y an d feature d prominently in Lyell' s Principles o f Geology (Lyel l 1830-1875). Al l o f thi s earl y wor k was , o f course, base d o n onshor e observations. B y th e late 19t h century , i t wa s realize d tha t Norwa y

was essentially a tableland, a plain that had been reduced t o som e bas e leve l an d subsequentl y uplifted (e.g . Beete-Juke s 1872) . Th e ancien t land surface , no w considerabl y modifie d an d incised b y recent glacia l an d fluvial erosion, was termed th e Paleic Surfac e by Reusch (1901) and subsequently describe d i n detai l b y Gjessin g (1967). Base d solel y o n regiona l evidence , principally th e Alpine-relate d uplif t o f larg e parts of central Europe, it was inferred that such a surface mus t have been formed in late Mesozoi c or early Cenozoic time , and that uplift mus t have taken place a t some later stag e of Cenozoic time (see, e.g . Gregory 1913) . Overlapping wit h this work, similar planatio n surfaces an d episode s o f Cenozoi c uplif t wer e inferred i n Scotlan d (se e e.g . Godar d 1962 ; George 1966 ; Hal l 1991) , an d th e Cenozoi c emergence o f souther n Britai n wa s obvious , based o n widesprea d outcrop s o f Jurassic , Cretaceous and Eocene marin e rocks . Holtedahl (1953 ) made the critical observatio n that som e o f th e highland s bordering th e Nort h Atlantic wer e probabl y complementar y t o area s of downwar p an d depositio n o n th e adjacen t shelves. Althoug h b y n o mean s obviou s a t th e time, the hypothesi s was quickl y teste d by the explosion i n offshor e hydrocarbo n exploration , which confirme d tha t mos t o f th e surroundin g shelves wer e characterize d b y Mesozoic-Cen ozoic sedimentar y basins . Consequently , a n attempt coul d b e mad e t o matc h th e suppose d Cenozoic evolutio n o f th e lan d areas , includin g denudation i n response t o uplift , t o th e offshor e sedimentary response . Furthermore , recognitio n of th e importanc e o f th e Cenozoi c evolutio n in the formation of the offshore hydrocarbo n riches (e.g. Parke r 1975 ) provide d a commercial , a s well as academic , motivatio n for continue d research.

From: DORE , A.G. , CARTWRIGHT , J.A., STOKER , M.S. , TURNER , J.R & WHITE , N . 2002 . Exhumation o f th e North Atlantic Margin: Timing, Mechanisms an d Implications for Petroleum Exploration. Geologica l Society , London, Specia l Publications , 196 , 1-12 . 0305-8719/027 $ 15.00 © The Geological Societ y o f London 2002 .

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A. G . DORE ETAL.

Historical description s o f th e numerou s strands o f subsequen t investigatio n hav e bee n given b y Gabrielse n & Dor e (1995) , Stuevol d & Eldhol m (1996 ) an d Japse n & Chalmer s (2000). A summar y o f th e ke y finding s i s a s follows. • A s well a s in mainland Norwa y an d Britain, Cenozoic uplif t and/o r emergenc e too k plac e in Spitsberge n (Harlan d 1969) , Swede n an d part o f Denmar k (e.g . Japse n & Chalmer s 2000), Irelan d (e.g . Naylo r 1992) , Eas t Greenland (e.g . Johnso n & Gallaghe r 2000 ) and West Greenland (e.g . Mathiesen 1998) . In other words , i t i s a circum-Nort h Atlanti c phenomenon. • Cenozoi c exhumatio n als o too k plac e i n basins periphera l t o th e landmasses . I t wa s recognized earl y that the huge expanse of shelf forming th e Barent s Se a wa s expose d subaerially an d erode d durin g lat e Cenozoi c time (Nanse n 1904 ; Harlan d 1969) . T o thi s have subsequentl y bee n adde d th e Hord a Platform, Stor d Basi n an d Farsun d Basi n o f offshore Norwa y (e.g . Ghaz i 1992 ; Jense n & Schmidt 1993) , th e Wes t Shetlan d Inne r Moray Firt h an d East Iris h Se a Basins i n UK waters (described by, for example, Lewis el al (1992), Parnel l e t a l (1999) , Hilli s e t a l (1994) an d Rowley & White (1998) , respect ively), th e Slyne-Erri s an d Nort h Celti c Se a Basins of f th e Iris h coas t (describe d b y Scotchman & Thoma s (1995 ) an d Murdoc h et a l (1995) , respectively ) an d numerou s others. Earl y Cenozoi c uplif t an d subaeria l exposure als o too k plac e alon g th e volcani c highs margina l t o th e newl y developin g North Atlantic , area s no w submerge d t o depths o f a kilometre o r mor e (e.g . Eldhol m et a l 1989) . • Fro m the numerous studie s now carried out , both loca l an d regional , i t i s clea r tha t th e circum-North Atlanti c uplif t an d erosio n wa s variable in magnitude, locatio n an d timing. I t could thu s perhap s b e argue d tha t th e phenomenon represent s a patchwork o f effect s deriving fro m man y unrelate d causes . Never theless, a s integratio n betwee n th e variou s studies improves , i t seem s tha t a t leas t tw o events ha d regiona l significance : (1 ) a Paleocene episod e o f widesprea d emergenc e in NW Europe coinciden t wit h North Atlantic opening an d th e initia l effect s o f th e Icelan d Plume (e.g. White 1988;Brodie&White 1995); (2) a Neogen e (mainl y Plio-Pleistocene ) episode wit h n o obviou s tectoni c cause , emphasized b y rapi d glacia l erosio n an d isostatic adjustmen t o f landmasse s an d

bordering shelve s (e.g . Solhei m e t a l 1996 ; Japsen & Chalmer s 2000) , togethe r wit h redeposition an d a widesprea d chang e i n th e deep-water circulatio n pattern i n the adjacen t basins (Stratage m Partner s 2002) . Th e Paleo cene even t appear s t o hav e bee n particularl y significant in the British Isles, whereas the late Neogene even t seem s t o hav e extende d fro m Scandinavia t o the Atlanti c margin of Britai n and /Ireland. Additionally , compressiona l upli/t (inversion) associated with Alpine stress and/or ridge-pus h fro m Atlanti c spreadin g i s common throughou t muc h o f th e area , although localize d an d ver y variabl e i n effec t (e.g. Murdoc h e t a l 1995 ; Dor e & Lundi n 1996). • Criticall y fo r the petroleum industry , severa l North Atlanti c basin s containin g commer cially significan t hydrocarbo n resource s wer e both uplifte d an d exhume d durin g Cenozoi c time, wit h profoun d implication s fo r th e quantity an d natur e o f th e hydrocarbon s discovered. Thes e effect s hav e bee n system atically studie d i n th e Barent s Se a (Nylan d et al 1992 ) and to some extent in the east Irish Sea (Cowa n e t a l 1999) , bu t i n term s o f overall applicabilit y t o uplifte d terrane s ar e still underestimated. Despite thi s rapi d increas e i n th e under standing o f th e exhumatio n o f th e Nort h Atlantic borderlands , ther e are stil l man y unknowns. Th e relativ e intensit y o f th e variou s phases, an d thei r variatio n i n importanc e geographically, ar e stil l onl y understoo d i n a very genera l sense . Althoug h ther e i s n o shortage o f postulate d uplif t mechanisms , ther e is stil l a scarcit y o f observationa l evidenc e an d modelling studie s t o establis h beyon d reason able doub t whic h ar e th e primar y cause s o f exhumation, an d ho w thes e ma y var y fro m area t o area . Tie d t o thes e problem s i s th e larger-scale questio n o f whethe r th e circum North Atlanti c i s uniqu e o r whethe r it s behaviour i s typica l fo r passiv e margins . There hav e bee n severa l attempt s i n recen t years t o brin g togethe r researcher s t o addres s these questions , an d compilation s hav e bee n published tha t ar e th e direc t antecedent s o f thi s book (se e particularl y Jense n e t a l (1992) , Solheim e t al (1996 ) and Chalmers & Cloetingh (2000)). Thes e proceedings , however , hav e tended t o focu s o n on e particula r geographica l area or one particular exhumation phase. There is an acknowledge d nee d t o brin g togethe r disciplines tha t hav e traditionall y remaine d apart (e.g . geomorphologist s an d offshor e seismic interpreters ; Paleogen e an d Neogen e

INTRODUCTION

3

Fig. 1 . Topographi c an d bathymetri c ma p o f th e easter n Nort h Atlantic , showin g th e locatio n o f studie s represented i n this volume. The papers are numbered as follows: 1, Jones et al.\ 2, Rohrman et al.;3, Nielsen et al.\ 4, Graversen ; 5 , Blundell ; 6 , Lidmar-Bergstro m & Naslund ; 7 , Hendrick s & Andriessen ; 8 , Evan s e t al.\ 9 , Stroeven e t al.\ 10 , Cederbom; 11 , Japsen et aL\ 12 , Huuse; 13 , Faleide et aL\ 14 , Bishop & Hall; 15 , Andersen et al\ 16 , Stoker; 17 , Green et al; 18 , Ware & Turner; 19 , Allen et al\ 20, Dore et al\ 21, Price; 22, Parnell; 23, Cramer e t al.; 24 , Corcora n & Dore . Fo r paper s tha t attemp t onshore-of f shore even t correlation , th e are a o f interest i s represented b y the sam e numbe r onshore an d offshore .

workers; Scandinavia n an d British-Iris h research schools ) befor e a n integrate d stor y ca n emerge. B y providin g a n interdisciplinar y se t of studies ove r a wide latitudina l rang e o f th e N W European margin (Fig. 1) , this volume represents an initia l ste p i n this direction .

Exhumation and other terms: some definitions Terms suc h a s exhumation, erosio n an d denuda tion hav e a varied usag e in the literature an d are often use d interchangeably . Numerou s attempt s

4

A. G . DORE ETAL.

have been made at deriving forma l definitions for these terms , an d th e subtl e difference s betwee n these definition s ca n rende r communicatio n difficult. Furthermore , th e wa y suc h term s ar e defined may depend on the concerns an d research orientations o f th e users . So , fo r example , scientists primaril y concerne d wit h th e litho spheric unloadin g o f orogeni c belt s an d cor e complexes (e.g . Ring et al. 1999) may emphasize different factors , an d hav e differen t definitions , from thos e concerne d wit h th e behaviou r o f passive margin s (e.g . thos e contributin g t o thi s volume). We hav e no t attempte d t o impos e a rigi d standardization on the papers in this volume, and have deliberatel y use d th e descriptiv e an d les s rigorously defined ter m 'exhumation ' in the title. Nevertheless, i t i s wort h whil e examinin g som e alternative use s t o determin e whethe r an y consensus ca n b e draw n an d a s a backgroun d for th e papers that follow .

Uplift Uplift i s a term in common usag e and appeals t o standardization ar e likel y t o fail . I t is , however , useful t o examin e som e way s i n whic h th e concept i s formall y used . Englan d & Molna r (1990), Summerfiel d (1991 ) an d Rii s & Jense n (1992) all pointed to the existence of two types of uplift. Thes e ar e surface uplift, referrin g t o th e upward movemen t o f th e Earth' s surfac e (th e land o r se a bottom ) wit h respec t t o a specifi c datum, usuall y mean se a level o r the geoid, an d crustal uplift, referrin g t o upwar d movemen t o f the rock column with respect t o a similar datum. The relationshi p betwee n surfac e uplif t an d crustal uplif t depend s upo n th e amoun t o f denudation o r deposition . I f there i s n o denudation or deposition, the n surface uplift an d crustal uplift wil l b e equal . If , a s i s likely , denudation occurs the n crusta l uplif t wil l b e greate r tha n surface uplift , an d i f denudatio n i s greate r tha n crustal uplif t surfac e elevatio n wil l b e reduced . Surface uplif t i s mostl y relate d t o activ e tectonics, bu t crusta l uplif t ca n occu r simpl y as the isostatic response t o denudation. Th e term net uplift ha s been used by several of the 'Norwegia n school' o f worker s t o indicat e th e presen t elevation o f a marke r be d abov e it s maximu m burial dept h (Nylan d et al. 1992 ; Rii s & Jensen 1992; Dore & Jensen 1996) . It is a measure of the vertical distance a rock has been uplifted through a therma l fram e o f reference , an d i s thu s particularly usefu l i n studie s o f diagenesi s an d source roc k maturation.

Erosion Ring et al. (1999) defined erosion as 'the surficial removal o f mas s a t a spatia l poin t i n th e landscape b y bot h mechanica l an d chemica l processes'. Biologica l processe s coul d als o b e added to this definition. Leeder (1999 ) provided a similar definition , but pointe d ou t tha t erosio n must b e define d wit h referenc e t o co-ordinate s fixed beneath the surface, as the surface potential elevation ma y itself chang e becaus e o f tectonics. As indicate d b y Rii s & Jense n (1992) , erosio n may b e subaeria l o r submarine , thu s als o emphasizing th e nee d fo r referenc e t o fixe d subsurface coordinates . Erosion rates ma y refe r to loca l measure d absolut e values or to regional values integrate d ove r a wide r area . Th e difference betwee n erosio n an d denudatio n i s considered below . Denudation Denudation has been defined by Leeder (1999 ) as the los s o f materia l fro m bot h surfac e an d subsurface part s o f a drainage basi n o r regiona l landscape b y al l type s o f weathering , physica l and chemical. Leeder indicate d that, like erosion, denudation shoul d b e define d wit h referenc e t o fixed interna l co-ordinates . So , accordin g t o Leeder, the important difference between erosion and denudation is that erosion is measured a s an effect o n th e surface , wherea s denudatio n ca n include subsurfac e processe s suc h a s chemica l dissolution, an d i s no t alway s accompanie d b y erosion. I t shoul d b e note d tha t i f th e Leede r definition i s used , i t follow s tha t denudatio n should b e characterize d a s los s o f mas s rathe r than thickness . Rin g e t al . (1999 ) provide d a n important modifie r t o thes e definition s o f denudation b y includin g the los s o f materia l by tectonic processe s (e.g . thin-skinned extensional faulting i n core complexes) . We advocat e thi s consensual view o f erosio n and denudation , althoug h los s o f materia l b y tectonic unloadin g doe s no t pla y a n importan t part in this volume. It is, however, worth pointing out tha t man y othe r usage s exist . Fo r example , Summerfield (1991 ) fro m a geomorphologica l perspective, preferre d t o defin e erosio n a s mechanical denudation o r th e remova l o f soli d particles, as opposed to chemical denudation (i.e. the remova l o f dissolve d material) . A ver y different emphasi s wa s provide d b y Brow n (1991) an d Rii s & Jense n (1992) , arguin g fro m a subsurfac e petroleum exploratio n perspective . They define d denudatio n a s erosio n (i n thei r terms, decreasin g thicknes s of overburden) seen in a thermal frame o f reference and quantified b y

INTRODUCTION

such measures as vitrinite reflectance and apatite fission-track trend s (se e als o 'ne t uplift') . Thi s definition is, however, specialized compare d with most others in the literature. It is encumbered by the potential for thermal regimes to vary through time, thereb y givin g false value s for denudation or erosio n i f a fixed thermal regim e i s assume d (see discussio n give n i n thi s volum e b y Gree n et aL). Exhumation Exhumation i s th e mos t loosel y define d o f th e terms used in this volume, but precisely becaus e of its descriptiv e natur e it is useful a s an overal l shorthand to describe th e removal of material by any means from a basin or other terrane such that previously burie d rock s ar e exposed . I t ca n b e characterized a s a process o r a history (thus, this definition fro m Ring et aL (1999): 'the unroofing history o f a rock , a s cause d b y tectoni c and/o r surficial processes'). It is seldom characterized as a measure , unlik e erosio n o r denudation , which are ofte n describe d i n terms o f loss o f thickness mass or volume. Inversion Compressive reactivation with reversal of slip on formerly extensiona l faul t system s i s terme d basin inversio n and , ultimately , i t lead s t o folding, thrustin g an d expulsio n o f th e synrif t fill (see, for example, Coope r & Williams (1989 ) and paper s therein , an d Buchana n & Buchanan (1995) an d paper s therein) . Becaus e inversio n often cause s th e processe s describe d abov e (uplift, erosion , denudation , exhumation ) i t i s frequently use d interchangeabl y wit h thes e terms, especially i n NW Europe where compres sive reactivatio n an d genera l regiona l exhuma tion ofte n coexist . Thi s usag e is misleading an d should be avoided unless the underlying cause of the regiona l exhumatio n i s know n t o b e compression. Towards an understanding of North Atlantic exhumation The studies presented her e are based on a variety of techniques that have been employed to address the main concerns o f North Atlantic exhumation history, includin g timing , mechanism s an d th e sedimentary respons e o f the continenta l margin . The 2 4 paper s presente d i n thi s volum e hav e accordingly bee n arrange d i n fou r section s t o reflect the highly varied approac h to this subject, and th e commercia l implications . Par t 1 i s

5

mainly concerned wit h exhumation mechanisms; parts 2 an d 3 presen t curren t researc h o n th e continental margi n record o f offshore Scandina via an d Britain , Irelan d an d th e Faeroes , respectively; par t 4 cover s th e implication s o f exhumation for hydrocarbon-bearing basins . Key aspects o f thi s interdisciplinar y approac h ar e summarized below . Th e geographica l sprea d o f the papers i s shown in Fig. 1 .

Techniques All o f th e mai n technique s fo r determinin g amount an d timin g o f uplift , erosio n an d denudation ar e represente d i n thi s selectio n o f studies. They ar e summarize d i n Table 1 , which also indicate s th e advantage s an d disadvantage s of eac h metho d an d where it is discussed i n this volume. Th e extrem e end s o f th e spectru m ar e those studie s tha t infe r exhumatio n chronolog y from eithe r exclusivel y onshore o r offshor e studies: th e forme r utilizin g moder n technique s of geomorphologica l analysi s (e.g . Lidmar Bergstrom & Naslund , Hal l & Bishop) , th e latter focusing on detailed seismic analysis of the offshore erosio n product s (e.g . Evan s et aL, Faleide et aL, Stoker). There i s no w ful l awarenes s o f th e nee d t o integrate onshore and offshore studie s to provide a mor e complet e pictur e o f uplift , erosio n an d sedimentary response . Availabl e approache s include simpl e graphica l reconstruction , tha t is, the projectio n o f offshor e stratigraphi c break s into onshor e surface s (e.g . Graversen , Japse n et aL} o r analysi s o f th e mas s balanc e betwee n basins an d hinterland s (Jone s et aL, Anderse n et a/.) . Thes e methods , b y thei r nature , provide only a coarse chronolog y an d are most effectiv e if combine d wit h physico-chemica l measure ment technique s fo r reconstructin g exhumatio n history. The apatite fission-track technique is by far the main 'growt h area ' an d i s represente d i n thi s volume b y severa l regiona l studie s (Hendrick s & Andriessen , Cederbom , Gree n et aL, Alle n et aL} Fission-trac k wor k i s provin g highl y effective i n establishing th e coolin g histor y of a rock, bu t wit h bot h thi s an d th e vitrinit e reflectance metho d (e.g . Japse n et aL, Gree n et aL} th e potentia l exist s fo r interpretin g anomalously hig h palaeo-hea t flow s a s highe r burial depth s i n th e past , an d vic e versa . A s indicated b y severa l workers , fission-trac k dat a are unreliable a t low temperatures (<60°C ) and can produc e spuriou s result s fo r geologicall y recent exhumation . Compaction studie s (Japse n et aL, War e & Turner ) benefi t fro m plentifu l

Table 1 . Summary o f techniques for measuring amount and timing o f uplift an d erosion-denudation, with attractions, limitations and where they ar e represented i n this volume Paper in this volume

Measurement technique

Attractions

Problems, limitation s

Geomorphological analysis

Field-based, plentifu l data , relate s landfor m developmen t t o geological histor y

Difficult t o obtain quantitativ e measures o f erosion -denudation, and to constrain timing ; most effectiv e when use d togethe r wit h offshore analysi s

6,9,10,14

Graphical reconstruction

Simple technique to estimate erode d thickness, correlate surface s and deriv e onshor e -off shore relationship s

Relies o n long-distanc e extrapolatio n an d assumptio n of thickness

4, 6 , 1 1

Offshore sedimentary response

Provides indicato r of denudation chronolog y o f sourc e areas ; seismic dat a allo w ful l sedimentar y sectio n t o be observe d

Erosion an d redepositio n o f offshore succession s complicat e correlation wit h onshore denudation chronology ; mos t effectiv e when use d togethe r wit h onshore analysi s

8, 12 , 13 , 16

Mass balanc e

Directly an d quantitatively correlates denudation wit h offshore sedimentation

Loss of mas s i n solution ; difficult y i n assignin g sedimen t t o correct catchmen t are a

1, 1 5

Vitrinite Preserve reflectance fissio

Apatite fissio track onshor

n Ca

Cosmogenic Use nuclide developmen

n trac

s recor d o f higher temperatures-burial depth s than k

n establis h detailed exhumatio n and buria l chronology bot h e an d offshor e d t o obtain detail on geologicall y recen t landform t not obtainabl e fro m fissio n trac k

Compaction, sonic Plentifu l sourc e o f wel l data (soni c logs) ; easier t o distinguish velocity change s i n basal hea t flo w fro m exhumatio n compared wit h thermal method s For th e numberin g o f th e paper s se e Fig . 1 .

Vitrinite absent fro m man y sediments and al l basement; vitrinit e 1 1 , 1 may b e misidentified ; difficul t t o separat e change s i n basal an d transient hea t flo w fro m exhumation ; reworked an d oxidized vitrinite a problem; onl y a crude timin g indicator Unreliable i n establishing recent exhumatio n events (coolin g 7 below annealin g temperature); difficul t t o separat e change s i n 1 basal an d transien t heat flo w fro m exhumatio n

7

, 10 , 17 , 9

Onshore only , mainly limited to chronology o f presen t lan d 9 surface Unreliability o f baseline compaction trend s fo r a basin o r a 1 1 , 1 lithology; difficult y i n identifyin g 'typical ' lithologie s for analysis; mechanica l compactio n retarde d b y overpressure , leading t o erosion underestimates ; only a crude timin g indicator

8

INTRODUCTION

data and , give n th e righ t stratigraphica l infor mation an d lithologies , ca n fill in som e o f thes e gaps. Th e mor e recen t onshor e denudatio n chronology is partially catered for by the rapidly developing cosmogeni c nuclid e techniqu e (Stroeven et a/.) , althoug h thi s metho d i s no t capable o f resolvin g larg e (kilometre-scale ) amounts of denudation. The essential conclusion, clear fro m Tabl e 1 , is tha t th e mor e technique s that can be brought to bear, the fuller th e picture that is likely t o emerge. Mechanisms Views o n th e natur e an d caus e o f th e Nort h Atlantic Cenozoi c exhumatio n hav e tende d t o depend o n wher e th e studie s hav e bee n carrie d out. Thu s th e olde r 'Scandinavia n school' , impressed b y th e evidenc e fo r rapi d Neogen e outbuilding an d th e concentri c natur e o f th e Cenozoic outcrops around the coasts, has tended to emphasiz e Neogen e effect s includin g glacio isostasy. Th e 'Britis h school', on the other hand, presented wit h evidenc e fo r hig h Paleocen e sedimentation rate s an d contemporaneou s vol canism aroun d th e islands , ha s mainl y stresse d Paleogene hotspo t model s t o explai n th e uplift , although George (1966) is a notable exception to this i n favourin g a substantia l amoun t o f Neogene reshapin g o f th e Britis h an d Iris h 'Massifs'. Thi s volum e i s a n initia l attemp t a t bringing thes e school s together , a s a mean s o f examining whethe r relationship s exis t an d ho w the differen t mechanism s ma y var y i n import ance i n tim e an d space . A summar y o f th e postulated mechanisms , togethe r wit h thei r attractions and limitations, is provided in Table 2. Part 1 o f th e boo k focuse s primaril y o n mechanisms related to the Iceland Plume. Jones et al. provid e mas s balanc e evidenc e tha t th e Iceland Plum e cause d significan t exhumatio n in Early Cenozoi c time , an d equall y impressiv e gravimetric evidenc e tha t th e plum e cause s a significant deflectio n o f th e geoi d ove r a ver y wide area today . It is widely believed tha t much of th e Paleogen e emergenc e ca n be attribute d to underplating o f th e lithospher e fro m th e plum e (Brodie & Whit e 1995) , bu t ther e i s stil l a requirement t o explai n th e apparen t absenc e o f (or lac k o f evidenc e for ) underplatin g i n som e areas, an d th e puzzlin g absenc e o f igneou s activity in classic uplifted area s such as Norway. Nevertheless, ther e i s soli d evidenc e fro m Bouguer gravit y an d tomograph y tha t eve n th e most recen t uplif t o f Norwa y occurre d wit h mantle involvement , perhap s upwellin g o r asthenospheric flo w withou t th e direc t impingement of a hotspot (Rohrma n et al.).

7

A variatio n o n thi s them e b y Nielse n et al. suggests tha t th e impingemen t o f th e Icelan d Plume hea d cause d delaminatio n o f th e lowe r lithosphere i n areas wit h deep crustal roots suc h as mainland Norway, leading to isostatic uplift. A central tene t o f th e Nielsen et al. mode l i s that, after initia l Paleocen e plume-relate d uplift , n o further activ e tectonic mechanis m i s required t o explain eithe r th e denudatio n chronology o r the present-day elevation of Norway. Post-Paleocene pulses o f erosio n an d depositio n ar e explaine d solely i n term s o f eustati c base-leve l fall , glaciation an d isostatic response t o these events . This view is emphasized in a supporting paper on the offshor e sedimentar y respons e (Huuse ) bu t rejected b y Japse n et al., wh o firmly believe i n the necessity fo r a Neogene tectonic uplift event . Detailed seismi c analysi s o f th e Cenozoi c sequences i n the norther n North Se a by Faleide et al. als o seem s t o suppor t a tectonic caus e fo r the late Neogene exhumation. Resolution of this fascinating debat e require s considerabl y mor e modelling an d researc h o n detaile d chronolog y and mas s balance . Graverse n provide s a controversial vie w (agai n involvin g mantl e upwelling) speculativel y linkin g domin g i n southern Scandinavi a wit h earlie r (Mesozoic ) doming in the central Nort h Sea . To complet e th e sectio n o n mechanisms , Blundell describe s th e rol e o f Alpin e stres s i n the Cenozoi c inversio n an d uplif t o f souther n Britain, supporte d i n thi s pape r b y thermo mechanical modellin g o f th e lithosphere . Blundell remarks , however , tha t apparen t lat e Neogene uplif t i n wester n part s o f th e are a cannot b e attribute d t o inversion . War e & T\irner, workin g i n th e Eas t Iris h Se a Basin , show how localization o f inversion-related strain superimposed o n a 'background ' epeirogeni c erosion signatur e cause d rapi d change s i n th e degree o f exhumation within a single basin. Continental margin record The paper s presente d i n part s 2 (Scandinavia ) and 3 (the Faeroes, th e British Isles an d Ireland) illustrate th e communicatio n tha t i s no w occurring betwee n the two regiona l researc h schools an d th e acknowledgemen t o f a multi phase Cenozoi c denudatio n chronology fo r both areas. Potentia l no w exist s fo r a compariso n between thes e area s t o determin e whic h event s are regionally correlated, an d which events are of local significanc e only . Fo r example , th e majo r geomorphological review s o f Scandinavi a (Lidmar-Bergstro m & Naslund) an d northern Scotland (Hal l & Bishop ) appea r t o detai l a similar event chronology, although late Neogene

Table 2 . Summary o f potential mechanisms for exhumation o f th e North Atlantic margin, with attractions, limitations an d where they are represented i n this volume Paper i n this volume

General grouping and mechanism

Attractions

Problems limitation s

Associated with mantle plume Dynamic uplif t Underplating

Regionally applicable ; supporte d b y geomorphological , palaeogeographical, mas s balance an d regional gravit y studies

Distribution o f underplatin g not full y understood ; absenc e o f igneous activit y in som e uplifte d areas , e.g . Fennoscandi a

1,3,12,13

Supported by Bougue r gravity and tomography , mainly southern Norway Supported by Bougue r gravity, mainly Norway

Causal agenc y poorl y understoo d

2,4

New mode l require s mor e testing ; connection betwee n plum e and delaminatio n beneath craton s no t clea r

3

Inversion tectonics widesprea d o n Atlantic margin; can explain rapid loca l variatio n i n exhumation

Difficult t o separate local from mor e regiona l backgroun d mechanisms; significan t shortenin g occur s before 'classic ' inversion geometrie s develop , henc e difficul t t o quantify bul k strain; cannot explain uplift o f craton s

5, 15 , 1 8

Strong correlation i n some area s between glaciatio n an d offshore sedimen t flux Could eliminat e need fo r enigmatic Neogene tectoni c uplif t mechanism

Emphasizes pre-existin g topography , bu t does not explain i t

12, 13 , 1 6

More studie s correlating denudation , sediment flux and glacio eustatic lowstand s required to remove necessit y fo r Neogen e tectonic uplif t

3, 1 2

Mantle upwelling without magmatism Ray leigh- Taylor instability Compressive Inversion Intraplate stres s Technically 'passive ' Glacio-isostatic Base-level change

For th e numberin g o f th e paper s se e Fig . 1 .

INTRODUCTION

doming i s considere d mor e prominen t i n th e former, wherea s th e latte r emphasize s th e particular importanc e o f Paleocene uplift . The fission-trac k coolin g chronologie s estab lished fo r Irelan d (Alle n et a/.), th e UK (Gree n et a/.) , Swede n (Cederbom ) an d norther n Norway (Hendrick s & Andriessen ) als o bea r comparison. Alle n e t a/.'s useful compilatio n o f fission-track dat a show s that , althoug h for any given tim e perio d Irelan d wa s a patchwor k o f differing exhumatio n profiles , som e genera l conclusions o n intensit y throug h tim e ca n b e drawn. Cederbom' s wor k o n souther n Swede n reaches th e sam e conclusion . Followin g thi s logic, a synthesis of the latest fission-trac k wor k over al l o f N W Europ e would provid e valuabl e insights int o th e regionall y significan t exhuma tion signatures , especiall y i f integrate d wit h methods capabl e o f resolvin g th e mos t recen t events suc h a s (U-Th)/H e thermochronolog y and perhaps even in situ and detrital cosmogeni c nuclide techniques. The integration of these three techniques is the key aim of the recently initiated CRUST projec t (Constrainin g Regiona l Uplift , Sedimentation & Thermochronology) , a colla borative projec t betwee n th e Universitie s o f Glasgow, Edinburg h an d Aberdeen , an d th e Scottish Universities ' Environmenta l Researc h Centre, Eas t Kilbride . As a final example, a similarly instructive view of regionally important events would be obtained by correlatin g th e Cenozoi c seismi c sequenc e chronologies reporte d fro m th e Mid-Norwegia n shelf (Evan s et al.\ Nort h Se a (Faleide et a/.), Faeroes shel f (Anderse n et al.) an d UK Atlantic margin (Stoker) . To this end, a unified, regiona l Neogene (Miocene-Holocene ) stratigraph y ha s recently bee n establishe d fo r th e entir e Atlanti c continental margi n between th e Lofote n Island s (off North Norway) and the Porcupine region (off southern Ireland), which includes adjacent deep water basin s suc h a s th e Norwegia n Basin , Faeroe-Shetland Channel , Rockal l Troug h an d Porcupine Basin (Stratagem Partners 2002). This study demonstrate s tha t th e entir e N E Atlanti c margin covere d b y thi s volum e (Fig . 1 ) wa s affected b y a n early-'mid'-Pliocen e even t tha t resulted, mos t significantly , i n th e initiatio n o f the Plio-Pleistocene prograding wedges . Implications for petroleum exploration Exhumation lead s t o coolin g an d lithostati c pressure decrease, an d these factors add a degree of difficult y t o petroleu m exploration . Th e following effect s ar e among the most important : (1) sealin g horizon s ar e remove d and/o r thei r effectiveness i s severel y reduced ; (2 ) fault s ar e

9

often reactivated , causin g the m t o becom e conduits fo r hydrocarbon leakag e to the surface ; (3) sourc e rock s wil l b e a t a highe r degre e o f maturation tha n expecte d fro m thei r presen t depth and will cease generation upon cooling; (4) potentially attractiv e reservoirs may , likewise, b e overcompacted an d downgraded ; (5 ) pressur e reduction during exhumation causes oil accumulations an d formatio n wate r t o exsolv e gas , causing ga s flushin g an d th e spillag e o f oi l accumulations; (6 ) regiona l tiltin g durin g uplif t results in changes to trap configurations and fluid migration directions . In spit e o f this , man y N W Europea n sedimentary basin s tha t underwen t sever e lat e Mesozoic an d Cenozoi c exhumatio n hav e retained prospectivity. These includ e the Barents Sea of f northern Norwa y and th e Eas t Iris h Se a Basin, whil e i n th e Wesse x basi n o f souther n England, Europe' s larges t onshor e oilfiel d a t Wytch Farm is located i n the footwall to a major compressionally reactivate d faul t zone . In par t 4 o f th e volume , five paper s describ e the significan t change s t o th e hydrocarbo n systems tha t occu r i n exhume d basins . Dor e et al. catalogu e these effects i n term s o f two of the standar d procedure s carrie d ou t i n th e oi l industry, namely , petroleum resourc e evaluatio n and ris k analysis . Thi s pape r document s bot h positive and negative implications, including the tendency toward s gas-dominate d systems , and references thes e t o a selectio n o f exhume d basins o n th e Nort h Atlanti c seaboard . Pric e considers th e sam e phenomen a i n som e o f th e well-documented uplifte d basin s i n wester n North America . Thi s pape r i s presente d i n abstract for m onl y becaus e o f th e untimel y death o f th e autho r durin g th e preparatio n o f these proceedings , a n even t tha t robbe d u s o f an importan t an d entertainin g contributo r to th e uplift debate . The three final papers consider specific aspect s of exhume d hydrocarbo n systems . A revie w b y Parnell examine s th e migratio n o f fluids , bot h hydrocarbons an d formatio n water , an d th e diagenetic change s tha t occu r durin g uplift o f a basin. Cramer & Poelchau draw attention to the potential fo r liberatio n o f hug e amount s o f methane fro m formatio n wate r durin g pressur e and temperatur e decrease , a proces s tha t ha s contributed significantly t o the giant gas fields of Western Siberia , an d b y implicatio n t o man y other gas accumulations worldwide. Complete or partial failur e o f seal s a s a result o f brittleness , stress-induced fracturin g and hydrofracturin g is shown b y Corcora n & Dor e t o b e a frequen t outcome o f exhumation , leadin g t o th e predic tion o f underfille d trap s an d near-hydrostati c

10

A. G . DORE ETAL.

gradients. Thes e inference s ar e agai n illustrate d by cas e studie s o n the North Atlantic margin. Concluding remarks The Nort h Atlanti c margi n describe d i n thi s volume is , o f course , par t o f a mor e globa l system o f margina l uplift s borderin g man y oceanic basins . Thes e includ e th e almos t classical are a fo r geomorphologist s o f th e southern Africa n margina l escarpments , alon g with th e margina l uplifte d massif s o f Easter n Australia an d Antarctica . Althoug h thes e othe r marginal uplift s wer e beyon d th e scop e o f thi s volume, man y o f th e technique s develope d an d refined fo r th e Nort h Atlanti c margin s ar e applicable mor e generally , an d ma y hel p t o broaden an d deepe n th e investigatio n o f thi s global tectoni c process . Perhaps th e ke y t o reall y advancin g ou r understanding o f th e processe s involve d i n oceanic margi n uplift s lie s o n a globa l scale , that is, to identify thei r spatial distribution within a high-resolutio n chronostratigraphi c frame work. Ther e i s thu s a prim e nee d fo r bette r resolution o f th e timin g an d magnitud e o f al l these margina l uplifts , a s par t o f an y attemp t t o synthesize their distributio n globally. For th e Nort h Atlanti c margins , significan t progress ha s bee n mad e alon g thes e lines , a s shown b y man y contribution s i n thi s volume . Additional effort i s required to fully integrat e the work of the 'Scandinavian ' an d 'British' schools, and als o t o forg e close r link s acros s disciplines . The sequenc e stratigrapher s workin g o n th e offshore seismi c recor d nee d greate r awarenes s of th e man y contribution s mad e b y onshor e geomorphologists an d vic e versa . Bette r resol ution of climate variation through Cenozoic tim e is o f paramoun t importanc e i n constrainin g th e long- an d short-ter m rate s o f erosio n an d sediment transport . I n thi s way , depositiona l systems analysi s offshor e ca n b e linke d mor e effectively t o surfac e uplif t onshore , an d lag s between uplif t an d sedimen t flu x fro m th e uplifted region s ca n b e identified . Thi s researc h problem i s trul y demandin g an d multidisciplinary, an d althoug h much is still to be done , much has alread y bee n achieve d toward s a bette r process understanding. References BEETE-JUKES, J . 1872 . Th e Student's Manual o f Geology. 3r d Edition. A . & C. Black, Edinburgh . BRODIE, 1 & WHITE , N.J . 1995 . Th e lin k betwee n sedimentary basi n inversio n an d igneou s under plating. In: BUCHANAN , J.G. & BUCHANAN , P.O.

(eds) Basin Inversion. Geological Society , London , Special Publications, 88 , 21-38. BROWN, R.W. 1991 . Backstacking apatite fission-trac k 'stratigraphy': a method fo r resolving the erosional and isostatic rebound components o f tectonic uplif t histories. Geology, 19 , 74-77. BUCHANAN, J.G. & BUCHANAN, P.G. (eds) 1995 . Basin Inversion. Geologica l Society , London , Specia l Publications, 88 . CHALMERS, J.A . & CLOETINGH , S . (eds ) 2000 . Neogene Uplift and Tectonics around the North Atlantic. Global and Planetary Change, Special Issue, 34 (3-4). COOPER, M.A . & WILLIAMS , G.D . (eds ) 1989 . Inversion Tectonics. Geologica l Society . London , Special Publications , 44. COWAN, G. , BURLEY , S.D. , HOEY , A.N . & 5 OTHER S 1999. Oi l an d ga s migratio n i n th e Sherwoo d Sandstone o f th e Eas t Iris h Se a Basin . In : Fleet , A.J. & Boldy , S.A.R . (eds ) Petroleum Geology o f Northwest Europe: Proceedings of the 5th Conference. Geologica l Society , London , 1383-1398. DORE, A.G . & JENSEN , L.N. 1996 . Th e impac t o f lat e Cenozoic uplif t an d erosio n o n hydrocarbo n exploration: offshor e Norwa y an d som e othe r uplifted basins . Global an d Planetary Change. 12 , 415-436. DORE, A.G . & LUNDIN , E.R . 1996 . Cenozoi c compressional structure s o n th e N E Atlanti c margin: nature , origi n an d potentia l significanc e for hydrocarbo n exploration . Petroleum Geoscience, 2, 299-311. ELDHOLM, O. , THEIDE , J . & TAYLOR , E . 1989 . Evolution o f th e V0rin g volcani c margin . In : ELDHOLM, O. , THEIDE , J . & TAYLOR . E . (eds ) Proceedings of the Ocean Drilling Program, Scientific Results, 104. Ocea n Drillin g Program , College Station , TX, 1033-1065 . ENGLAND, P. & MOLNAR . P. 1990. Surface uplift, uplif t of rocks , an d exhumatio n o f rocks . Geology. 18 . 1173-1177. GABRIELSEN, R.H . & DORE , A.G . 1995 . Histor y o f tectonic model s o n th e Norwegia n continenta l shelf. In : HANSLIEN , S . (ed. ) Petroleum Exploration an d Exploitation i n Nonvay. Norwegia n Petroleum Societ y (NPF ) Specia l Publication , 4 . 333-368. GEORGE, T.N . 1966 . Geomorphi c evolutio n i n Hebridean Scotland . Scottish Journal o f Geology. 2, 1-34 . GHAZI, S.A . 1992 . Cenozoi c uplif t i n th e Stor d Basi n area an d it s consequence s fo r exploration . In : JENSEN, L.N . & Rus , F . (eds ) Post-Cretaceous Uplift and Sedimentation along the Western Fennoscandian Shield. Norsk Geologisk Tidsskrift. vo!72, 285-290. GJESSING, J.P . 1967 . Norway' s palei c surface . Norsk Geografisk Tidsskrift. 21 , 69-132. GODARD, A . 1962 . Essai s d e correlatio n entr e Taltitudes de s relief s e t le s caractere s petro graphique de s roche s dan s le s socle s d e TEcoss e du nord . Comptes Rendus d e rAcademic de s Sciences, 255. 139-141 .

INTRODUCTION GREGORY, J.W. 1913 . Th e Nature and Origin of Fjords. John Murray, London . HALL, A.M. 1991 . Pre-Quaternary landscape evolutio n in th e Scottis h Highlands . Transactions o f th e Royal Society o f Edinburgh: Earth Sciences, 82 , 1-26. HARLAND, W.B . 1969 . Mantl e change s beneat h th e Barents Shelf . Transactions o f th e Ne w York Academy o f Sciences, Series 2 , 31, 25-41. HILLIS, R.R. , THOMSON, K . & UNDERBILL , J.R. 1994 . Quantification o f Tertiar y erosio n i n th e Inne r Moray Firt h usin g soni c velocit y dat a fro m th e Chalk an d th e Kimmeridge Clay . Marine an d Petroleum Geology, 11 , 283-293. HOLTEDAHL, O . 1953 . O n th e obliqu e uplif t o f som e northern lands . Norsk Geografisk Tidsskrift, 14 , 132-139. JAPSEN, P . & CHALMERS , J.A . 2000 . Neogen e uplif t and tectonic s aroun d th e Nort h Atlantic : over view. In : Chalmers , J.A . & Cloetingh , S . (eds ) Neogene Uplift and Tectonics around the North Atlantic. Global an d Planetary Change, 2 4 (3-4), 165-174 . JENSEN, L.N . & SCHMIDT , BJ . 1993 . Neogen e uplif t and erosion offshore sout h Norway: magnitude and consequences fo r hydrocarbo n exploratio n i n th e Farsund Basin . In : SPENCER , A.M . (ed. ) Generation, Accumulation and Production of Europe's Hydrocarbons, HI . Specia l Publicatio n o f th e European Associatio n o f Petroleu m Geoscientists , 3, 79-88. JENSEN, L.N. , Rns , F . & BOYD , R . (eds ) 1992 . Post Cretaceous Uplif t an d Sedimentatio n alon g th e Western Fennoscandia n Shield . Norsk Geologisk Tidsskrift, Symposium Issue, 3 , 72 . JOHNSON, C . & GALLAGHER , K . 2000 . A preliminar y Mesozoic an d Cenozoi c denudatio n histor y o f th e North East Greenlan d onshor e margin . Global and Planetary Change, 24 , 261-274. LEEDER, M.R . 1999 . Sedimentology an d Sedimentary Basins. Blackwell Science , Oxford . LYELL, C . 1830-1875 . Principles o f Geology, 1 2 editions. Joh n Murray, London. MATHIESEN, A. 1998 . Modelling o f Uplift History from Maturity and Fission-track Data, Nuussauq, West Greenland. Danmark s o g Gr0nland s Geologisk e Unders0gelse Rapport , 87 . MURDOCH, L.M. , MUSGROVE , F.W . & PERRY , J.S . 1995. Tertiar y uplif t an d inversio n histor y i n the Nort h Celti c Se a Basi n an d it s influenc e o n source roc k maturity . In : CROKER , PF . & SHANNON, P.M . (eds ) Th e Petroleum Geology of Ireland's Offshore Basins. Geologica l Society, London , Specia l Publications , 93 , 297-319. NANSEN, F . 1904 . Th e bathymetrica l feature s o f th e North Pola r seas , wit h discussio n o f th e continental shelve s an d previou s oscillation s o f the shoreline . In : NANSEN , F . (ed. ) Th e Norwegian North Polar Expedition 1893-1896, Scientific Results. Jaco b Bydwad , Christiani a (Oslo). NAYLOR, D . 1992 . Th e post-Varisca n histor y o f Ireland. In: PARNELL, J. (ed.) Basins on the Atlantic

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Seaboard: Petroleum Geology, Sedimentology and Basin Evolution. Geologica l Society , London , Special Publications, 62, 255-275. NYLAND, B. , JENSEN , L.N. , SKAGEN , J. , SKARPNES , O. & VORREN , T . 1992 . Tertiar y uplif t an d erosion i n th e Barent s Sea ; magnitude , timin g and consequences . In: LARSEN , R.M. , BREKKE , H., LARSEN , B.T . & TALLERAAS , E . (eds ) Structural and Tectonic Modelling and its Application t o Petroleum Geology. Norwegia n Petroleum Societ y (NPF ) Specia l Publication , 1 , 153-162. PARKER, J.R. 1975 . Lowe r Tertiar y san d developmen t in th e centra l Nort h Sea . In : WOODLAND , A.W . (ed.) Petroleum and the Continental Shelf o f NorthWest Europe. Applie d Science , Barking , UK , 447-454. PARNELL, J., CAREY , P.P. , GREEN , PF . & DUNCAN , W . 1999. Hydrocarbo n migratio n history , wes t o f Shetland: integrate d flui d inclusio n an d fissio n track studies . In : FLEET , A.J . & BOLDY , S.A.R . (eds) Petroleum Geology o f Northwest Europe: Proceedings o f th e 5t h Conference. Geologica l Society, London , 613-626 . REUSCH, H . 1901 . Nogl e bidra g til l forstaaelse n a f hvorledes Norge s dal e o g fjeld e e r blevn e til . Norges Geologiske Unders0gelse, Aarbog, 14 , 96-102. Rns, F . & JENSEN, L.N. 1992. Introduction: measuring uplift an d erosion—proposal for a terminology. In: Jensen, L.N. & Riis, F. (eds) Post-Cretaceous Uplift and Sedimentation along the Western Fennoscandian Shield. Norsk Geologisk Tidsskrift, 72 , 223-228. RING, U. , BRANDON , M.T. , LISTER , G.S . & WILLETT, S.D . 1999 . Exhumatio n processes . In : RING, U. , BRANDON , M.T. , LISTER , G.S . & WILLETT, S.D . (eds ) Exhumation Processes: Normal Faulting, Ductile Flow and Erosion. Geological Society , London , Specia l Publications , 154, 1-27 . ROWLEY, E . & WHITE , N . 1998 . Invers e modellin g o f extension and denudation in the East Irish Sea and surrounding areas . Earth an d Planetary Science Letters, 161 , 57-71. SCOTCHMAN, I.C . & THOMAS , J.R.W . 1995 . Tertiar y uplift an d inversion history in the North Celtic Se a Basin and its influence o n source rock maturity. In: CROKER, PF . & SHANNON , P.M . (eds ) Th e Petroleum Geology of Ireland's Offshore Basins. Geological Society , London, Specia l Publications , 93,385-411. SOLHEIM, A. , Rns , E , ELVERHOI , A. , FALEIDE , J.I. , JENSEN, L.N . & CLOETINGH , S . 1996 . Impac t o f glaciations o n basi n evolution : dat a an d model s from th e Norwegia n margi n an d adjacen t areas — introduction an d summary . Global an d Planetary Change, 12 , 1-9 . STEUVOLD, L.M . & ELDHOLM , O . 1996 . Cenozoi c uplift o f Fennoscandia inferre d from a study of th e mid-Norwegian margin . Global an d Planetary Change, 1 2 (1-4), 359-386. STRATAGEM PARTNERS , 2002. The Neogen e stratigra phy o f th e Europea n margi n between Lofote n an d

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. G. DORE ETAL.

Porcupine. Stoker , M.S . (compiler) . A produc t o f WHITE , R.S . 1988 . A hot-spot model for Early Tertiary the EC-funde d STRATAGE M project . Online . volcanis m in the North Atlantic. In: MORTON, A.C. Available a t http://www.stratagem-europe.org. & PARSONS, L.M. (eds) Early Tertiary Volcanism in SUMMERFIELD, M.A . 1991 . Global Geomorphology. th e North Atlantic. Geologica l Society , London , Longman, Harlow , UK. Specia l Publications , 39, 241-252.

Present an d past influence of the Iceland Plum e o n sedimentatio n STEPHEN M . JONES 1, NICKY WHITE 1, BENJAMIN J. CLARKE 1'2, ELEANOR ROWLEY 1'3 & KERRY GALLAGHER 1'4 l Bullard Laboratories, Madingley Rise, Madingley Road, Cambridge CB3 OEZ, UK (e-mail: jones@ esc.cam.ac.uk) 2 Present address: Department of Geology and Geophysics, University of Edinburgh, King's Building, West Mains Road, Edinburgh EH9 3JW, UK 3 Present address: Shell Eygpt, 6 Hassan El-Sherley Street, PO Box 2681 El Horreya, Heliopolis, Cairo, Eygpt 4 TH Huxley School of Environment, Earth Science, and Engineering, Imperial College of Science, Technology and Medicine, South Kensington, London SW7 2AS, UK Abstract: Th e Cenozoic developmen t o f the North Atlantic provinc e ha s been dramaticall y influenced b y the behaviour of the Iceland Plume , whose striking dominance i s manifest by long-wavelength free-air gravity anomalies and by oceanic bathymetri c anomalies. Here, we use these anomalies to estimate the amplitude an d wavelength o f present-day dynami c uplif t associated wit h this plume. Maximum dynamic support in the North Atlantic is 1.5-2 km at Iceland itself . Mos t o f Greenlan d i s currently experiencing dynami c suppor t o f 0.5-1 km, whereas th e N W Europea n shel f i s generall y supporte d b y < 0.5 km. Th e proto-Icelan d Plume had a n equally dramatic effec t o n the Early Cenozoi c palaeogeograph y o f the North Atlantic margins , a s w e illustrat e wit h a stud y o f plume-relate d uplift , denudatio n an d sedimentation o n th e continenta l shel f encompassin g Britai n an d Ireland . W e infe r tha t during Paleocen e tim e a hot subvertica l shee t o f asthenospher e welle d u p beneath a n axi s running from the Faroes through the Irish Sea towards Lundy, generating a welt of magmatic underplating of the crust which is known to exist beneath thi s axis. Transient an d permanen t uplift associate d wit h this magmatic injection caused regional denudation, and consequently large amount s of clastic sedimen t hav e been she d into surrounding basins during Cenozoi c time. Mass balance calculations indicat e agreemen t betwee n the volume of denuded materia l and the volume of Cenozoic sediment s deposited offshor e in the northern Nort h Se a Basin and th e Rockal l Trough . Th e volum e o f materia l denude d fro m Britai n an d Irelan d i s probably insufficien t to account fo r the sediment i n the Faroe-Shetland Basin and an excess of sediment has been supplied to the Porcupine Basin. We emphasize the value of combining observations fro m bot h oceani c an d continenta l realm s t o elucidat e th e evolutio n o f th e Iceland Plume throug h spac e an d time.

The free-ai r gravit y fiel d ove r th e norther n estimate s of dynami c support derived from th e hemisphere i s dominate d by a long-wavelengt h bathymetr y an d gravit y fiel d o f th e Nort h high centre d o n Iceland , stretchin g fro m th e Atlanti c oceanic real m t o constrai n the present Azores t o Siberi a an d fro m Baffi n Islan d t o magnitud e of dynami c suppor t of bot h oceanic Denmark (Fig . 1) . Bot h theoretica l consider - crus t and the adjacent continenta l margins, ations and observations from th e world's oceans Next , we consider the region that was affecte d suggest that long-wavelength ( > 1000km ) posi- b y th e plum e in th e past . Temporal and spatial tive anomalie s ar e generall y associate d wit h variatio n of plume-relate d uplift ha s playe d an mantle upwellin g an d dynami c uplif t (Sclate r importan t role in the evolutio n of both margins, et al 1975 ; McKenzie 1994). If these inferences especiall y i n controllin g th e generatio n an d also hold in the North Atlantic province, the areal distributio n o f clasti c sediments . Here , w e extent o f th e gravit y hig h suggest s tha t th e conside r the specifi c exampl e of sedimen t mass continental margin s of N W Europ e and easter n balanc e around Britain and Ireland. Geochemical Greenland are dynamically supported at present, evidenc e has proved that this region experienced In th e firs t par t o f thi s paper , w e compar e magmati c underplatin g o f th e crus t durin g From: DORE , A.G., CARTWRIGHT, J.A. , STOKER , M.S. , TURNER, J.P . & WHITE , N. 2002 . Exhumation of th e North Atlantic Margin: Timing, Mechanisms an d Implications for Petroleum Exploration. Geologica l Society , London, Special Publications, 196, 13-25 . 0305-8719/02/$15.00 © The Geological Society of London 2002.

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Fig. 1 . Free-air gravity anomaly ove r part of the northern hemisphere , displaye d using a Lambert azimutha l equal area projectio n centre d o n Iceland . Th e long-wavelengt h gravit y hig h centre d o n Iceland , whic h extend s fro m south o f the Azore s t o Spitzberge n an d fro m Baffi n Islan d t o Scandinavia , shoul d b e noted .

Paleocene time , leadin g t o permanen t uplift , which drov e denudatio n (Thompso n 1974 ; Brodie & Whit e 1995) . Mas s balanc e provide s a mean s o f comparin g an d verifyin g comp lementary measure s o f Earl y Cenozoi c denuda tion, estimate d fro m th e onshor e recor d o f denudation an d th e offshor e record o f sedimen t accumulation. Two forms of surface uplift are associated wit h mantle plumes . Dynami c suppor t alway s results when abnormall y ho t mantl e i s emplace d beneath th e lithosphere. Thi s support i s transient and disappear s whe n th e therma l anomalie s i n the asthenospher e an d lithospher e dissipat e b y convection an d conduction . Th e Nort h Atlanti c region i s dynamicall y supporte d toda y an d i t must als o hav e bee n dynamicall y supporte d during Earl y Cenozoi c time , a s th e volcani c record attest s t o abnormall y ho t mantl e a t tha t time. Permanen t uplif t ma y als o occu r i f mantle thermal anomalies induce melting and the melt is injected int o or just beneath the crust. Permanen t uplift affecte d Britai n an d Irelan d durin g

Paleocene time , an d th e region s flankin g th e line o f continental separation a t the Paleocene Eocene boundary. Comparison o f the two themes we presen t i n thi s paper therefor e demonstrate s the relativ e spatia l exten t an d magnitud e o f transient an d permanen t uplift . Th e ter m epeirogenic uplif t refer s t o uplif t tha t coul d b e permanent and/o r transient . I n general , epeiro genic uplif t refer s t o uplif t tha t is no t generate d by horizonta l plat e motions , an d th e ter m nee d not impl y an y particula r mechanism . However , the presen t an d pas t epeirogeni c uplif t o f th e North Atlanti c region w e discuss in this study is generated b y the Iceland Plume. Present-day dynamic support The Nort h Atlanti c Ocea n i s anomalousl y shallow i n th e vicinit y o f Iceland . Anomalou s topography culminate s a t Icelan d itself , whic h rises t o c . 2 km abov e sea-level , o r c . 4. 5 km above the average depth of the global mid-ocea n ridge system . Tw o methods ca n b e employe d t o

ICELAND PLUM E PAST AND PRESEN T

investigate dynami c suppor t o f th e regio n o f oceanic crus t roun d Iceland . Th e firs t metho d exploits th e fac t tha t th e dept h o f oceani c crus t away from th e influence of mantle plumes varies with ag e i n a well-understoo d manner , b y comparing th e present-da y bathymetr y aroun d Iceland wit h thi s referenc e depth . Th e secon d method involve s establishin g a lin k betwee n dynamic suppor t an d th e long-wavelengt h free air gravit y anomaly . Analysi s o f gravit y anomalies should always be treated with caution, as an y gravit y fiel d ca n b e explaine d b y a n infinite numbe r o f densit y distributions , eac h with differen t implication s for dynamic support. In this section, we first establish that estimates of dynamic suppor t derive d fro m bathymetr y an d gravity ar e i n genera l agreemen t ove r oceani c crust aroun d Iceland . Thi s resul t the n give s u s confidence i n estimatin g dynami c suppor t o f the adjacen t continenta l margin s usin g gravity alone.

15

Estimates from bathymetry Figure 2 is a plot of anomalous topography in the North Atlantic , calculate d b y subtractin g th e well-known age-dept h mode l o f Parson s & Sclater (1977 ) fro m th e bathymetry . Unfortu nately, th e anomalou s topograph y i n Fig . 2 cannot be interpreted solel y i n terms o f present day dynami c suppor t because i t als o include s a component o f permanent topograph y caused b y spatial variation s i n th e thicknes s o f oceani c crust. However, the magnitude of this permanent topography ca n b e estimate d give n determi nations o f oceani c crusta l thicknes s fro m wide angle seismi c experiments . I n th e ocean s surrounding Icelan d bu t awa y fro m th e con tinental margin s an d the Greenland-Iceland Faroes Ridge , anomalousl y ho t mantl e ha s generated crus t 7-10k m thic k (Whit e 1997) . Isostatic balancing shows that this variation of up to 3k m greate r tha n th e thicknes s o f standar d

Fig. 2 . Anomalou s topograph y o f th e Nort h Atlanti c Ocean , calculate d b y subtractin g th e age-dept h coolin g relationship of Parsons & Sclater (1977) from th e ETOPO5 bathymetr y grid. The age of oceanic lithosphere was taken from Miiller et al (1997) ; the Greenland-Iceland-Faroes Ridge was excluded from the calculation because its age is not well known. Anomalous topography is corrected for sediment loading of oceanic basement using the method of Le Douaran & Parsons (1982) and the sediment thickness map of Laske & Masters (1997). It should be noted that the anomalous topography displayed here contains both a component of present-day dynamic support and a permanent component cause d by crustal thickness variations.

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oceanic crus t (7 km, White et al. 1992 ) generates permanent topograph y o f 0-0.5 km. Thi s valu e of 0. 5 km i s smalle r tha n th e tota l variatio n i n anomalous topograph y acros s thes e regions , suggesting tha t most o f th e anomalou s topogra phy i s generate d b y present-da y dynami c support. However , clos e t o th e continenta l margins an d th e Greenland-Iceland-Faroe s Ridge, th e thicknes s o f oceani c crus t i s 10-15 km, varyin g ove r distance s o f <50k m (Barton & White 1997; Wei r et al. 2001). In these regions, permanen t topograph y probabl y accounts fo r mos t o f th e anomalou s topography in Fig. 2 . To summariz e th e discussio n above , anom alous topography o f those regions away from the continental margin s an d th e Greenland Iceland-Faroes Ridg e ca n b e interprete d a s a n estimate o f dynamic suppor t with an error rang e of 0.5 km. However, short-wavelengt h variation s in anomalou s topograph y associate d wit h con tinental margin s an d th e Greenland-Iceland Faroes Ridge are most likely to reflect permanent topography associate d wit h crusta l thicknes s variations. Bearin g i n min d thes e caveats , thre e important result s concernin g present-da y dynamic suppor t ca n b e gleane d fro m Fig . 2 . First, the amplitude o f dynamic suppor t clos e to Iceland itsel f i s 1.5- 2 km. Secondly , th e amplitude o f dynami c suppor t i s partiall y controlled b y the active mid-ocean ridg e system . This fac t i s clearl y see n bot h t o th e nort h o f Iceland, wher e th e activ e Kolbeinse y Ridg e i s more anomalousl y elevate d tha n th e extinc t Aegir Ridge, and to the south of Iceland, where a tongue o f anomalou s uplif t extend s alon g th e Reykjanes Ridge . Thirdly , th e continenta l margins o f easter n Greenlan d an d N W Europ e are currentl y experiencin g a t leas t 0.5k m o f dynamic support , implyin g tha t th e adjacen t continental shelve s ar e als o experiencin g significant dynami c suppor t a t present. Estimates from gravity The free-air gravity field over a convecting layer results fro m tw o competin g effects : densit y variations withi n th e laye r an d surfac e defor mation induced by the convective circulation . In the cas e o f upwellin g plumes , ho t mantl e ha s a lower density than the surrounding mantle, which acts to reduce the gravity. However, a n upwelling plume deform s th e Earth' s surfac e upwards , which acts to increase the gravity. The sign of the total free-ai r gravit y anomal y depend s o n th e relative magnitude of thes e effects . A t Rayleig h numbers appropriat e t o Earth , th e effec t o f surface deformatio n outweigh s the effec t o f th e

reduction i n density, so upwelling regions in the mantle ar e characterize d b y positiv e free-ai r gravity anomalie s a t the surface. The transfe r functio n betwee n th e free-ai r gravity anomaly g and topography h is called the admittance an d i s given by Admittance i s a functio n o f wavenumbe r k . A t wavelengths shorte r tha n c . 500km , admittance is controlled b y the mechanical propertie s o f the lithosphere. Admittanc e decrease s systemati cally with increasing wavelengt h (i.e . decreasing wavenumber), an d th e rat e o f decreas e i s dependent upo n th e effectiv e elasti c thicknes s of th e lithosphere . McKenzi e (1994 ) an d McKenzie & Fairhea d (1997 ) exploite d thi s behaviour t o estimat e th e effectiv e elasti c thickness o f th e lithospher e i n th e Pacifi c an d Indian Oceans fro m th e short-wavelengt h part of the gravit y an d topograph y fields . A t wave lengths abov e c . 500k m th e admittanc e calcu lated fro m th e observe d gravit y and topography diverges fro m tha t calculate d fro m theoretica l models assumin g a n elasti c plate . A t thes e lon g wavelengths, th e flexura l strengt h o f th e litho sphere play s n o part i n supportin g topography . Instead, topograph y i s supporte d b y stresse s exerted o n th e bas e o f th e lithospher e b y th e convecting mantl e an d b y long-wavelengt h variations i n th e densit y structur e o f th e lithosphere, which are isostatically compensated. Thus, dynami c suppor t produce s long-wave length anomalou s topography A/z conv tha t corre lates wit h th e long-wavelengt h free-ai r gravit y anomaly. Isostaticall y compensate d topography, such a s th e permanen t anomalou s topograph y that arises from crustal thickness variations in the North Atlantic , is no t correlate d wit h a measurable gravit y anomaly. The behaviou r o f th e admittanc e functio n within th e wavelengt h ban d 500-3000k m suggests a linea r relationshi p betwee n gravit y and topography . I f thi s simpl e relationship , observed i n th e Pacifi c an d India n Oceans, als o holds i n th e Nort h Atlanti c the n anomalou s topography cause d b y dynami c convectiv e support ca n b e calculate d fro m th e long wavelength free-ai r gravit y field using The value of Z to be used can b e constrained by observations. It has lon g been recognize d tha t Z ~35mGalkm~ 1 i s appropriat e fo r Earth' s oceans (Sclate r e t al . 1975) . Anothe r measur e of th e admittanc e t o b e use d t o calculat e dynamic support can be obtained from numerical

ICELAND PLUM E PAST AND PRESEN T

convection experiments . McKenzi e (1994) sum marized severa l numerica l model s o f axisym metric plumes, includin g the model of Watson & McKenzie (1991 ) tha t successfull y matched th e observed gravity , topograph y an d mel t pro duction o f th e Hawaiia n Plume . Thes e convec tion model s ar e characterized b y admittance s i n the rang e 34. 4 ± 2. 2 mGal km"l . Hence , observed an d theoretica l value s fo r th e admit tance betwee n topograph y an d gravit y o f th e oceans a t lon g wavelength s ar e i n goo d agreement. The free-ai r gravit y datase t use d her e i s a compilation o f poin t measurement s ove r land , together wit h th e satellit e gravit y datase t o f Sandwell & Smit h (1997 ) ove r th e oceans , a s described b y McKenzie & Fairhead (1997) . Th e short-wavelength part o f the gravit y field that is influenced b y th e flexura l strengt h o f th e lithosphere was removed usin g a low-pass filter. The regio n o f interes t ha s dimension s compar able with the radius of Earth, so low-pass filtering was carried out using a spherical harmonic model of th e gravit y dataset . Th e long-wavelengt h gravity mode l was generate d usin g spherica l harmonic coefficient s o f degrees 0 < / = m ^ 53 (equivalent to a low-pass filter of c . 750km ) with a suitabl e tape r a t th e uppe r cut-of f t o prevent ringing.

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Figure 3 show s th e dynami c componen t o f anomalous topography A/iconv estimated from the free-air gravity field using two different values of Z. These estimates are simply a scaled version of the long-wavelengt h gravit y fiel d an d resembl e the original gravit y field fairly closel y (compar e Figs 1 and 3). In the region surroundin g Iceland, estimates o f convective suppor t from th e gravity field agre e reasonabl y wel l wit h independen t estimates of convective support from topography (compare Fig s 2 an d 3) . I n particular , th e dynamic suppor t estimate s fro m gravit y reinforce th e thre e observation s derive d fro m Fig. 2 an d note d i n th e previou s section . First , peak dynami c suppor t i s c . 1.8k m a t Iceland . Secondly, th e activ e spreadin g axi s exert s a n important control on the long-wavelength gravity field an d o n th e magnitud e of dynami c support . Dynamic suppor t is centred o n the Mid-Atlantic Ridge, an d t o th e nort h o f Icelan d th e activ e Kolbeinsey Ridge is experiencing greater support than th e extinc t Aegi r Ridge . Thirdly , th e continental margin s ar e currentl y experiencin g significant dynami c suppor t of 0.5-1 km. It i s importan t t o not e tha t estimate s o f present-day dynami c suppor t calculate d directl y from long-wavelengt h gravity anomalies ar e not always i n agreemen t wit h estimate s calculate d from th e bathymetry . Figur e 3 suggest s tha t

Fig. 3 . Estimates of present-day dynami c suppor t in the North Atlanti c region , calculate d b y dividing the longwavelength free-air gravity field by a constant admittance, as discussed in the text. Bold continuous lines indicate continent-ocean boundaries , (a ) Dynami c suppor t predicte d usin g a n admittanc e o f Z = 3 5 mGal km"1 (appropriate for subaqueous regions); (b) dynamic support predicted usin g an admittance of Z = 5 0 mGal km"1 (equivalent value for subaeria l regions). K , Kangerlussuaq; S, Scoresby Sund.

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significant dynami c suppor t occur s from Iceland to th e Azores . However , Fig . 2 indicate s n o measurable dynami c suppor t jus t sout h o f th e Charlie Gibb s fractur e zone. Thus, the amplitud e of dynamic suppor t calculate d fro m gravity dat a is unlikel y t o be correc t i n detai l an d shoul d b e treated wit h som e caution . Nevertheless , th e agreement betwee n dynami c suppor t estimate s based o n bathymetr y an d gravit y i n th e vicinity of Icelan d suggest s tha t Fig . 3 provide s a reasonable estimat e o f dynami c suppor t o f th e Greenland-Iceland-Faroes Ridg e an d th e adjacent continenta l shelves . Free-air gravit y anomalie s ma y b e use d t o estimate dynamic suppor t of the continents in the same wa y a s fo r th e oceans . Wherea s i t i s appropriate t o emplo y a n admittanc e o f Z = 35mGalkm~ 1 to estimate dynami c suppor t of region s covere d b y water , a n equivalen t air loaded valu e o f Z = 50mGalkm~ 1 shoul d b e used fo r subaeria l regions . Dynami c suppor t estimates calculate d usin g bot h value s o f admittance ar e show n i n Fig . 3 . Greenlan d i s currently experiencin g dynami c suppor t o f 0.5-1km. Dynami c suppor t i s greates t o n th e east coast , notabl y i n th e regio n betwee n Kangerlussuaq an d Scoresb y Sund , adjacen t t o the Greenland-Iceland-Faroe s Ridge . I n con trast, the magnitud e of dynamic suppor t beneat h the N W Europea n shel f seldo m exceed s 0. 5 km. Dynamic suppor t increase s westward s fro m Scandinavia acros s th e Norwegia n continenta l shelf. Th e Nort h Se a an d souther n Englan d d o not appear to be dynamically supported . Cenozoic denudation of Britain and Ireland In thi s section , w e focu s o n on e aspec t o f th e problem o f determinin g th e Earl y Cenozoi c shape o f th e Icelan d Plume , namel y th e relationship between Cenozoic uplift , denudation and sedimentatio n o n th e continenta l shel f surrounding Britai n an d Ireland . Ou r long-ter m goal i s t o us e mas s balanc e calculation s i n conjunction wit h estimate s o f permanen t uplif t caused b y magmati c underplatin g t o determin e the behaviou r o f th e Icelan d Plum e throughou t Cenozoic time . We hav e chose n t o carr y ou t mas s balanc e calculations fo r th e Britis h Isle s fo r thre e important reasons . First , geochemica l evidenc e collected fro m rock s o f th e Britis h Cenozoi c Igneous Provinc e prove s tha t th e crus t beneat h much o f Britai n an d Irelan d ha s bee n thickene d by a substantia l amoun t o f igneou s materia l a s we explai n below , implyin g significan t perma nent uplif t durin g Paleocen e time . Secondly , a substantial bod y o f informatio n abou t th e

denudation o f Britai n an d Irelan d i s availabl e based o n modellin g o f subsidence , vitrinit e reflectance, apatit e fission-trac k an d soni c velocity datasets . Thirdly , th e product s o f Cenozoic denudatio n hav e bee n carefull y mapped i n al l o f th e surroundin g sedimentar y basins. Although detailed wor k ha s been carrie d out previousl y o n bot h onshor e denudatio n an d offshore deposition , n o attempt s hav e ye t bee n made t o check these estimate s b y constructing a mass balanc e o n a regional scale . The evidence for magmatic underplating of the crust beneat h Britai n an d Irelan d i s no t widel y recognized, despit e th e fac t tha t th e conse quences of suc h an igneous addition, in terms of uplift, denudatio n and sedimentation, can explain many feature s o f th e surfac e geolog y an d thu s have obviou s implication s fo r th e hydrocarbo n industry. For decades i t has been recognized that the compositio n o f floo d basalt s fro m th e Hebrides can be explaine d onl y by fractiona l crystallization o f u p t o 70 % o f thei r origina l liquid mas s (Thompso n 1974) . The crystallize d residuum must therefore remain at depth. From a surface processe s poin t o f vie w th e argumen t ends here , becaus e i f suc h materia l i s adde d anywhere within the upper half of the lithosphere its densit y wil l b e les s tha n tha t o f th e asthenosphere i t displaces , an d isostati c balan cing shows that permanent uplift wil l result. The depth a t whic h crystallizatio n occurred, an d a t which the crystallized residuu m remains , ca n be established usin g geobarometr y techniques , which determin e th e pressur e a t whic h th e major-element composition s o f bot h floo d basalts an d th e phenocryst s the y contai n ar e i n equilibrium. Pressures o f c . I GP a are estimated, equivalent t o a dept h o f aroun d 30km , i.e . th e depth o f th e Moh o (Thompso n 1974 ; Brodi e & White 1995) . This result is not surprising , as the fact tha t basalti c mel t i s les s dens e tha n th e mantle lithospher e but roughly the sam e density as the lower crust means that the melt should rise through th e mantl e an d pon d a t th e bas e o f th e crust, givin g ris e t o th e concep t o f magmati c underplating. Wide-angl e seismi c experiment s across Nort h Atlanti c volcani c continenta l margins hav e image d high-velocit y bodie s a t the bas e o f th e continenta l crust , whic h ar e usually interprete d a s pod s o f igneou s materia l underplated beneat h th e crus t a t th e tim e o f continental separatio n (Barto n & Whit e 1997) . Preliminary result s fro m modellin g o f a wide angle seismi c lin e spannin g th e Iris h Se a suggest tha t a high-velocit y zon e exist s nea r the Moho , whic h probabl y represent s th e igneous materia l tha t w e expec t t o se e underplated beneat h th e onshor e par t o f th e

ICELAND PLUM E PAS T AND PRESEN T British Cenozoi c Igneou s Provinc e (S . Al-Kindi, pers . comm.) . However , a s thi s discussion implies , seismi c evidenc e alon e ca n never directl y revea l th e ag e o r natur e o f thes e high-velocity layers , s o it is always necessary t o interpret seismi c evidenc e i n conjunctio n wit h geochemical evidence . T o conclude, i t should b e noted that crustal magmatic underplating is not a peculiar featur e o f th e Nort h Atlanti c bu t i s common to all continental flood basalt province s (Cox 1993) .

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Mass balance calculations The mas s balanc e calculatio n ha s thre e parts: th e exten t an d amoun t o f denudatio n of th e sedimen t sourc e are a o r catchmen t (Fig. 4) ; th e mas s o f sedimen t accumulate d through tim e i n th e sedimentar y basin s immediately adjacen t t o th e Britis h Isle s (Fig. 5) ; an d th e mas s o f materia l los t fro m the syste m b y solutio n an d b y escap e t o th e deep ocean .

Fig. 4. Estimate s o f Cenozoi c denudatio n o f Britain an d Ireland base d o n modelling subsidenc e historie s an d apatite fission-trac k lengt h distributions. +, locations of well sections in extensional sedimentary basins used for subsidence modelling. The amount of missing post-rift basin fill was predicted by fitting a theoretical subsidenc e curve to the remnant synrif t stratigraphy, assumin g the standard lithospheric stretching model. Thi s technique has been describe d full y b y Rowle y & White (1998) , an d results hav e been tabulate d by Hal l (1995 ) an d Rowley (1998). O , location s o f apatit e samples . Apatit e fission-trac k lengt h distribution s wer e modelle d t o fin d Mesozoic-Cenozoic therma l historie s b y K . Gallaghe r usin g th e Laslet t e t al. (1987 ) annealin g mode l fo r Durango apatite and the method described by Gallagher (1995); th e results were reported by Rowley (1998). The amount of Cenozoic coolin g wa s converted t o a range of denudation estimate s b y Monte Carlo modellin g usin g the range of geothermal gradients observed in the North Sea today (Rowley 1998). Data covering Ireland are from Allen e t al . (2002) . (a ) Minimu m denudation estimat e foun d b y contourin g result s derive d fro m subsidenc e analysis, th e lowe r bound s o f the Rowle y (1998 ) denudatio n estimate s derive d fro m modellin g apatit e fission track length distributions, and the minimum denudation estimate fo r Ireland o f Allen et al. (2002, fig . lOb) . (b ) Maximum denudation estimate foun d b y contouring the modes of the Rowley (1998) denudation estimate s an d the maximum denudation estimat e for Ireland o f Allen et al (2002 , fig. lOa). These contou r plots were generate d by taking the mean of all estimates within blocks of dimension 1° longitude X 307 latitude (c. 50 km X 50 km) an d gridding the resulting values using the continuous curvature spline method of Smith & Wessel (1990). It should be noted that a further se t of Cenozoic denudation estimates based on modelling of vitrinite reflectance profiles fro m a subset of the wells used for subsidence analysis yields a denudation estimate that lies between the minimum and maximum estimates illustrate d her e (Rowle y 1998) .

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Fig. 5 . Map of the NW European continenta l shel f showin g the Cenozoic soli d sedimen t thickness. This isopac h map wa s constructe d fro m a databas e o f 2 D an d 3 D seismi c reflectio n survey s calibrate d wit h well-lo g information. Soli d thicknes s wa s calculated fro m th e observe d thicknes s b y subtractin g the pore-space volum e predicted b y th e standar d exponentia l relationshi p betwee n fractiona l porosit y =
The best-constraine d elemen t i n th e mas s balance proble m i s th e mas s o f Cenozoi c sediment i n th e offshor e basins . Here , th e volume of soli d sedimen t i s used i n place o f the mass o f sediment . Th e soli d volum e o f th e offshore sedimen t pil e is calculated b y removin g the volum e accounte d fo r b y porosity , predicte d by the standard exponentia l porosit y model . The principal sourc e o f error i n calculatin g the soli d volume i s th e dept h conversio n procedure . However, the error introduced by this uncertainty can be realistically quantifie d by using a range of velocity-depth functions, and the resulting erro r estimates ar e illustrate d in Fig. 6 . Parameteriza tion o f th e porosity-dept h relationshi p use d i n the compactio n calculatio n i s relativel y wel l determined fo r th e basin s surroundin g Britai n and Ireland . Fo r example , usin g parameter s fo r the end-membe r lithologie s o f shal e an d san d determined by Sclate r & Christie (1980 ) for the North Se a yields a variation in solid volume that

is only 14 % of the error range associated wit h the depth conversio n calculatio n i n tha t region . Figure 5 show s th e soli d sedimen t thicknes s accumulated aroun d Britai n an d Irelan d durin g Cenozoic time . Durin g thi s tim e th e mai n sediment sink s wer e th e Porcupin e Basin , th e Rockall Trough , th e Faroe-Shetlan d Basi n an d the Nort h Se a Basin . Th e soli d volum e o f Cenozoic deposit s i n souther n Englan d an d offshore souther n Ireland i s negligible. The tota l Cenozoi c denudatio n i s show n i n Fig. 4 , base d o n thre e independen t line s o f evidence. Th e firs t lin e o f evidenc e depend s o n our knowledg e o f extensiona l sedimentar y basins. Man y extensiona l basin s aroun d Britai n and Irelan d contai n a fault-controlle d syn-rif t stratigraphy bu t th e anticipate d post-rif t strati graphy i s partiall y o r entirel y absent (Brodi e & White 1995) . Th e amoun t o f missin g post-rif t stratigraphy ca n b e predicted , base d o n ou r knowledge o f th e kinematic s o f extensiona l

ICELAND PLUME PAST AND PRESEN T

Fig. 6. Summary of mass balance calculations for NW European continenta l shelf . Ope n bar s represen t th e amount o f soli d sedimen t accumulate d offshore , calculated b y integratin g Fig . 5 ove r eac h offshor e depocentre. Error range s reflec t uncertaintie s i n dept h conversion. The soli d sedimen t volum e for th e North Sea plotted here has been halved to account for the fact that Scandinavia ha s als o supplie d sedimen t t o th e North Sea . Fille d bar s represen t volum e o f roc k denuded fro m onshor e catchments , calculate d b y integrating th e minimu m an d maximu m estimate s given in Fig. 4 over the regions marked in Fig. 5. The Porcupine sedimen t sin k i s supplie d b y tw o catch ments: SI , souther n Irelan d an d th e Iris h Sea ; WI, western Ireland . Th e othe r sedimen t sink s hav e on e catchment each .

basins (Rowle y & White 1998) . Th e othe r tw o methods o f estimatin g denudatio n rel y o n thermal indicator s withi n th e sedimen t pil e tha t retain a memor y o f thei r buria l history . Reflectance o f vitrinit e increase s wit h risin g temperature b y mean s o f a non-reversibl e reaction, s o vitrinit e retain s a memor y o f th e highest temperatur e i t has experienced. Denuda tion estimate s ca n b e obtaine d b y comparin g vitrinite reflectanc e profile s wit h a globa l reflectance datase t fro m non-inverte d basin s (Rowley 1998) . Analysi s o f fissio n track s i n apatite als o provide s temperatur e historie s tha t can be interpreted in terms of denudation through time. The most prominent featur e of Fig. 4 is the peak in denudation centred on NW England. This region suffere d denudatio n o f 1-2. 5 km during Cenozoic time . A peak in denudation centre d o n

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NW Englan d ha s als o bee n suggeste d b y previous studie s base d o n apatit e fission-trac k data alone (Lewis et al. 1992). The magnitude of Cenozoic denudatio n generall y decrease s towards th e presen t offshor e areas . Anothe r important messag e fro m Fig . 4 i s tha t th e erro r range on the denudatio n estimat e for any particular region is c. 1 km. It is difficult t o quantify los s of mass from th e system. Loss by solution depends strongl y on the rock typ e being eroded . I t woul d b e difficul t t o include th e amoun t o f los s b y solutio n i n th e mass balanc e calculation s becaus e o f th e uncertainty ove r th e origina l lithologie s bein g eroded an d the variety of lithologies withi n each catchment. Her e w e appl y n o correction s t o account for loss o f mass from th e system . When loss o f mas s fro m th e syste m i s neglected , th e solid volum e o f sedimen t measure d i n offshor e basins provides a lower bound on the denudation. If th e soli d sedimen t volum e measure d offshor e is foun d t o be greater tha n the onshor e estimat e of denudatio n the n a n additional, unidentifie d source of sediment to the offshore basin would be implied. O n the other hand , if the offshor e soli d sediment volume is found to be significantly les s than the onshore estimate o f denudation then the difference betwee n th e tw o estimate s ma y provide a n estimat e o f th e amoun t o f materia l lost from th e system. Mass balance results Figure 6 summarizes the result of balancing th e volume o f soli d sedimen t accumulate d offshor e with direc t estimate s o f denudatio n mad e onshore fo r th e Porcupine , Rockall , Faroe Shetland an d norther n Nort h Se a systems . Th e drainage catchment matche d wit h each offshor e sediment sink is shown in Fig. 5. The boundaries of these catchments are based on the present-day topography an d bedloa d transpor t patterns . Palaeotopographic reconstruction s o f Britai n and Irelan d durin g earlies t Eocen e time , a t th e time o f maximu m dynami c suppor t b y th e Iceland Plum e (se e discussio n later) , sugges t that th e north-sout h drainag e divid e runnin g through Scotlan d an d Englan d ha s remaine d relatively stati c sinc e Paleocen e tim e (Jone s 2000). Th e position s o f th e west-east-oriente d drainage divide s ar e more likel y t o have altere d position durin g Cenozoi c time , an d we conside r the effec t o f suc h migratio n i n th e followin g discussion. In the Rockall Troug h and northern North Sea systems, th e volum e o f Cenozoi c sedimen t accumulated offshor e i s th e sam e a s th e denudation measure d onshore , withi n error .

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However, i n th e Faroe-Shetlan d system , th e volume o f sedimen t accumulate d offshor e i s greater than the denudation measured onshore by at leas t 2 x 10 4km3. Thi s discrepanc y implie s that th e Faroe-Shetlan d Basi n ha s anothe r source o f Cenozoi c sedimen t i n additio n t o th e catchment i n th e regio n o f N W Scotland . Th e additional sedimen t sourc e ma y b e a catchmen t in th e regio n o f th e Faroe s o r Greenland . Alternatively, th e boundarie s o f th e drainag e catchment show n i n Fig . 5 coverin g N W Scotland ma y b e incorrect . Th e souther n boundary o f tha t catchmen t i s no t wel l constrained an d it i s possibl e tha t sedimen t wa s also source d fro m wester n Scotlan d an d chan nelled betwee n th e Oute r Hebride s an d th e Scottish mainlan d toward s th e Faroe-Shetlan d Basin. A thir d possibilit y i s tha t additiona l sediment wa s transporte d int o th e Faroe-Shet land Basi n b y current s runnin g alon g th e basi n axis. Stron g botto m current s flowin g south westwards alon g th e Faroe-Shetlan d Troug h were initiate d i n Oligocen e tim e an d persis t t o the present da y (Davie s e t al. 2001). Lack o f a significan t differenc e betwee n th e two estimate s o f denudatio n i n th e Rockal l an d northern Nort h Se a system s suggest s tha t littl e mass has been lost from thos e systems. However, the tota l sedimen t supplie d t o th e Porcupin e Basin b y bot h th e catchmen t coverin g wester n Ireland an d th e catchmen t coverin g souther n Ireland an d th e Iris h Se a i s greate r tha n th e material accumulate d offshor e b y a t leas t 4 x l 0 4 k m 3 . Thus , a significan t mas s o f sediment ha s bee n los t fro m thi s system . Th e magnitude o f this discrepancy show n in Fig. 6 is a lower bound because no account has been taken of sedimen t supplie d t o th e wester n sid e o f th e Porcupine Basi n b y th e Porcupin e Ban k an d Ridge. Thre e effect s hav e probabl y contribute d to th e los s o f mas s fro m th e Porcupin e system . First, Carboniferou s limeston e crop s ou t ove r a large area of the Irish catchments , an d erosion of this limeston e produce s negligibl e clasti c sedi ment. Secondly , i t i s likel y tha t mos t o f th e sediment supplie d t o th e souther n en d o f th e Porcupine Basin has escaped directly to the deepsea sediment fa n observed on the North Atlanti c oceanic abyssa l plai n vi a th e canyo n syste m i n the S E o f th e Porcupin e Basi n (Fig . 5). Thi s canyon syste m ha s bee n activ e sinc e Earl y Oligocene tim e (Jone s 2000) . Thirdly , a variety of evidenc e fro m bot h th e continenta l shel f surrounding Britai n an d Irelan d an d adjacen t oceanic crus t suggest s tha t th e hea d o f th e Iceland Plum e expande d rapidl y durin g earlies t Eocene time, causing dynamic uplift centred N W of Britai n an d Irelan d (se e discussion below) .

The resultin g southeasterl y til t o f Britai n an d Ireland probabl y cause d mos t o f th e materia l eroded fro m souther n Irelan d an d th e Iris h Se a during Eocen e tim e t o b e she d southward s into the Bay of Biscay, rather than westwards into the Porcupine Basin.

Cenozoic permanen t an d transient uplift In this section, we discuss evidence fo r tempora l and spatia l variatio n i n th e Icelan d Plum e throughout Cenozoi c time . Paleocene permanent uplift The histor y o f sedimen t flu x int o th e basin s surrounding Britai n an d Irelan d durin g Paleo cene tim e add s detai l t o th e denudatio n history established usin g th e mas s balance . Th e rat e o f sediment flux into an offshore basi n is related t o the siz e an d th e rat e o f denudatio n o f th e corresponding drainag e catchment. The lag time between a n increas e i n denudatio n rate an d th e corresponding increas e i n sediment flux offshore is likely t o be < lOOk a (Readin g 1991 ; Burgess & Hoviu s 1998) . Therefore , sedimen t flu x histories i n th e basin s surroundin g Britai n an d Ireland ca n b e directl y relate d t o uplif t o f thei r sediment sourc e regions . Calculatio n o f th e volumes o f Paleocen e an d Eocen e sedimen t sequences i n the northern North Se a and FaroeShetland Basin s ha s show n tha t sedimen t flu x into these basins grew through Paleocene tim e to a maximum at 59-58 Ma and then decreased into Eocene tim e (Reynold s 1994 ; Clark e 2002) . Thus w e ma y infe r tha t epeirogeni c uplif t o f Scotland an d norther n Englan d was initiate d i n Early Paleocen e tim e an d th e rat e o f uplif t peaked durin g mid-Lat e Paleocen e time . I n contrast, sedimen t flux into the Porcupine Basin remained lo w throughou t Paleocen e time , suggesting tha t uplif t rate s wer e lowe r awa y from th e regio n o f Scotlan d an d norther n England a t this time (Jones 2000) . The denudatio n histor y o f Scotlan d inferre d from th e histor y o f sedimen t flu x aroun d Scotland correlate s wit h th e histor y o f onshor e igneous activity, which also peaked a t 59-58 Ma (White & Lovel l 1997) . A s w e hav e seen , th e sediment flu x dat a sugges t tha t Paleocen e epeirogenic uplif t wa s confine d t o th e regio n close t o th e present-da y surfac e expressio n o f onshore igneou s activity . W e therefor e sugges t that th e Paleocen e igneou s activity , epeirogeni c uplift an d denudatio n wer e initiate d whe n a subvertical shee t o f unusually hot asthenospher e was injecte d beneat h th e presen t Faroes-Iris h Sea-Lundy axis . Th e surfac e expressio n o f

ICELAND PLUME PAST AND PRESENT

onshore igneou s activit y i s concentrate d i n western Scotlan d an d Ulster , bu t thi s igneou s activity i s offse t fro m th e locu s o f maximu m denudation, whic h i s centre d o n norther n England an d th e Iris h Sea . I f th e denudatio n shown i n Fig . 4 i s drive n mainl y b y permanen t uplift resultin g from magmati c underplating, the greatest amount of melt must have been added to the crus t beneat h thi s region . Preliminar y modelling o f a wide-angle seismi c lin e crossing Ireland, th e Iris h Se a an d norther n England has imaged a high-velocit y po d nea r th e Moho , which is thickest beneath the centre of the profile and thin s toward s wester n Irelan d an d towards the Nort h Se a (S . Al-Kindi, pers . comm.) . Thi s pod ma y represen t th e laye r o f magmati c underplating beneat h th e crus t tha t i s predicte d by bot h petrologica l an d sedimentologica l evidence. Paleocene-Eocene dynamic support The sedimentar y record s i n al l basins surrounding th e denude d are a o f Britai n an d Irelan d indicate a major regression-transgression cycle , with maximu m regressio n correspondin g t o th e upper Flett Formation and its lateral equivalents, which wer e deposite d durin g earlies t Eocen e time (e.g . Milton e t al 1990 ; Ebdo n et al 1995) . It is now generally believed that this regressiontransgression cycl e i s relate d t o a phas e o f transient dynami c uplif t tha t peake d in earlies t Eocene tim e (Nadi n e t al . 1997 ; Jone s e t al . 2001). The magnitude of dynamic support can be quantified withi n extensiona l sedimentar y basins, where we can isolate epeirogenic vertica l motions fro m tectoni c vertica l motions , whic h are ultimatel y cause d b y horizonta l plat e motions. Given the synrif t subsidenc e history of an extensiona l basin , th e post-rif t subsidenc e history ca n b e calculate d usin g th e well established lithospheri c stretchin g model . Post rift marke r horizon s wit h sedimentologicall y well-constrained water depths are then compared with the anticipated post-rif t subsidence curv e to reveal the magnitude of dynamic support through time. Th e mos t importan t o f thes e marke r horizons aroun d Britai n an d Irelan d ar e th e earliest Eocene delta-top coal s of the upper Flett Formation an d its lateral equivalents. The results of thi s subsidenc e analysi s sho w tha t pea k dynamic support was c. 0.5 km in the Porcupine and Faroe-Shetlan d Basins , an d dynami c sup port decreased in a southeasterly direction to zero across souther n Englan d (Nadi n e t al . 1997 ; Jones 2000; Jones e t al 2001) . Peak dynami c suppor t a t th e Paleocen e Eocene boundar y wa s coeva l wit h voluminous

23

intrusive an d extrusive igneou s activit y offshor e NW o f Britai n an d Irelan d tha t wa s associate d with break-u p o f Europ e an d Greenlan d abov e unusually ho t asthenosphere . Whit e (1997 ) collated oceani c crustal thickness measurement s and showe d tha t th e ho t asthenospher e o f th e Iceland Plum e hea d extende d a t leas t 1000k m along th e continent-ocea n boundarie s t o th e north an d sout h o f th e presen t Greenland Iceland-Faroes Ridge . Dynami c suppor t esti mates fro m subsidenc e analyse s aroun d Britain and Ireland suggest that the plume head extended a simila r distanc e inboar d o f the N W Europea n continental margin. We migh t expec t tha t a s maximu m transient dynamic suppor t o f th e basin s surroundin g Britain an d Irelan d occurre d durin g earlies t Eocene time , th e acm e o f denudatio n an d offshore sedimen t flux should also have occurred at thi s time . However , a s w e discusse d above , maximum sedimen t flu x int o th e Faroe-Shet land an d Nort h Se a Basin s actuall y occurre d 3-4 Ma previously, during Late Paleocene time . Only in the Porcupine Basin, west of Ireland, was the pea k sedimentatio n rat e coeva l wit h pea k dynamic suppor t i n earlies t Eocen e tim e (Jone s 2000). Thi s discrepanc y i n timin g o f severa l million years between the two different measures of epeirogenic uplif t i s an important observatio n that ha s ye t t o b e explained . On e possibl e explanation i s tha t ther e wer e tw o separat e phases o f mantl e plum e activity , th e firs t a n upwelling hot vertical shee t that led to magmatic underplating o f th e crus t beneat h th e Faroes Irish Sea-Lundy axis, and the second the growth of a mushroom-shape d plum e hea d beneat h a region ove r 1000k m in radius. Oligocene-Recent dynamic support Variations i n crusta l thicknes s an d structur e around Iceland record changes in the siz e of the thermal hea d o f the Icelan d Plume , measure d at the Mid-Atlantic Ridge , through Cenozoic time . As we have alread y discussed , th e radius o f the plume hea d wa s > 1000 km a t th e tim e o f continental separatio n betwee n Europ e an d Greenland i n earlies t Eocen e time . However , the distributio n o f normal-thicknes s oceani c crust sout h of Iceland suggest s that during Late Eocene times, the part of the plume head beneath the Mid-Atlantic Ridge extended < 300 km from the presen t centr e of Iceland (Whit e 1997) . Th e portion of the plume head beneath the ridge then increased t o it s presen t radiu s o f c . 1000k m between Oligocen e tim e an d the present (Fig s 2 and 3). Studies of plate motion with respect to the hotspot reference frame suggest that the centre of

24

S. M. JONES ETAL.

the Iceland Plum e lay beneath Greenlan d durin g mid-Cenozoic time an d ha s effectivel y move d eastwards toward s Europ e throug h tim e (Lawve r & Mulle r 1994) . W e sugges t tha t thi s relativ e motion betwee n th e Mid-Atlanti c Ridg e an d a plume o f relativel y constan t mas s flu x ma y account fo r th e apparen t increas e i n th e siz e o f the head of the Iceland Plum e between Oligocen e time and the present. Thi s history of motion als o agrees wit h ou r observatio n tha t Greenlan d i s currently experiencin g greate r dynami c suppor t than N W Europ e (Fig . 3) . Th e reaso n i s tha t a greater volum e o f ho t plume-hea d materia l accumulates beneat h th e plat e tha t th e plum e stem i s movin g awa y fro m tha n accumulate s beneath th e plat e tha t th e plum e ste m i s approaching (Rib e & Delattr e 1998) . A s Rohrman & van der Beek (1996 ) have suggested , relative movement o f the Iceland Plum e toward s Europe ma y als o provid e a n explanation fo r th e Miocene-Recent epeirogeni c uplif t an d conse quent increas e i n denudatio n tha t i s know n t o have affecte d Scandinavia .

Conclusions The Cenozoi c evolutio n o f th e Nort h Atlanti c province ha s bee n dominate d b y interactio n between th e Icelan d Plum e convectiv e syste m and sea-floo r spreadin g betwee n Europ e and Greenland o n a hierarchy o f spatial an d tempora l scales. Th e continenta l sedimentar y recor d seems to be influenced by the relative importance of dynami c support , whic h ha s varie d throug h through time , an d permanen t uplift , whic h wa s driven b y magmati c underplatin g o f th e crust . Further progres s i n understandin g thi s relation ship will depend upon measuring the distribution of Paleogen e magmati c underplatin g an d upo n an improve d quantitativ e understandin g o f Cenozoic denudation . Our conclusions concernin g th e present shap e and siz e o f th e Icelan d Plum e an d concernin g Cenozoic mas s balanc e are : (1) maximu m dynami c suppor t i n th e Nort h Atlantic is 1.5-2 km at Iceland itself . (2) Th e magnitud e o f dynamic suppor t i n the North Atlantic is influenced by the location of the active spreadin g centres . (3) Th e continent-ocea n boundar y o f N W Europe i s currently experiencing dynami c uplif t of c . 0.5km , whic h decrease s t o zer o acros s Scandinavia and the North Sea . Th e continent ocean boundary of eastern Greenland is currently experiencing dynami c uplif t o f c. 1 km. (4) Cenozoi c mas s balanc e validate s onshor e denudation estimate s b y showin g the y ar e

compatible wit h th e soli d volum e o f sedimen t accumulated offshore . S.MJ. wa s supporte d b y a n NER C studentshi p an d B.J.C. wa s supporte d b y a B P studentship . W e ar e indebted t o B . Mitchene r an d J . Perr y o f B P fo r providing dat a an d support . A . Carte r an d T . Hurfor d provided th e fission-track dat a tha t wer e modelle d t o produce Fig . 4 , an d P . Allen allowed u s t o us e results reported elsewher e i n thi s volume i n th e sam e figure . We thank H. Walford for help in producing Fig. 3 , and M. Shaw-Champion for help in producing Fig. 5. A.G. Dore, J.-I . Faleide , B . Lovel l an d M . Rohrma n provided helpfu l reviews . Thi s pape r i s Departmen t of Eart h Science s Contributio n ES.6677 .

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Timing and mechanisms o f North Atlantic Cenozoi c uplift : evidence fo r mantle upwellin g MAX ROHRMAN 1'2, PETER A . VAN DER BEEK 3, ROB D. VAN DER HILST 4 & PAUL REEMST 5 1 Landmark Graphics, Stavanger, Norway 2 Present address: Shell UK Exploration & Production, 1 Aliens Farm Rd, Nigg, Aberdeen AB12 3FY, UK (e-mail: max. rohrman @expro. shell, co. uk) 3 Laboratoire de Geodynamique des Chaines Alpines, Universite Joseph Fourier, Grenoble, France 4 Department of Earth, Atmospheric and Planetary Sciences, MIT, Cambridge, MA 02139, USA 5 Geologica AS, Stavanger, Norway (present address: NAM, Assen, The Netherlands) Abstract: Postrif t doma l uplif t pattern s ar e a distinc t featur e o f norther n Nort h Atlanti c margins. O n th e basi s o f apatit e fission-trac k data , offshor e seismi c stratigraphy , geomorphology, gravit y an d seismi c tomography , w e argu e tha t souther n Norwa y i s characterized by predominantly Neogen e domal uplift. Th e uplift is technically driven an d estimated at around 1. 5 km. Low flexural rigidity (c . 10 22 N m) and corresponding equivalen t elastic thickness T e (c. 15km ) value s fo r the souther n Norwegia n lithospher e indicate tha t the lithospher e i s relatively weak . Additionally , hig h temperatur e estimate s derive d fro m low-velocity mantle P - and S-wave seismi c tomography belo w th e dome sugges t a therma l anomaly a t depth. Therefore, th e observed topography i s most plausibly explaine d by mantle upwelling. Thi s woul d superced e othe r previousl y propose d primar y mechanism s suc h a s eustasy, isostati c readjustmen t t o glacia l erosion , magmati c underplatin g an d intraplat e compression. Currentl y availabl e dat a sugges t simila r processes fo r othe r uplifte d region s such a s Spitsbergen, norther n Norway , th e Britis h Isle s and parts o f East Greenland .

During recen t decade s i t ha s becom e apparen t Neogen e phase (Rohrman & van der Beek 1996 ; that th e Nort h Atlanti c margin s experience d Japse n & Chalmer s 2000) , whic h ha s a mor e substantial vertical movements in Cenozoic time enigmati c cause. Thi s even t i s characterize d b y (e.g. Whit e & Lovel l 1997 ; Dor e e t al 1999 ; substantia l uplif t o f Britain , part s o f Norway , Japsen & Chalmer s 2000 ) characterize d b y th e Spitsbergen , th e Faeroe s an d Greenland (Japsen coupled emergenc e o f rif t margin s an d anom - & Chalmers 2000) an d by associated subsidenc e alous subsidenc e o f th e nearb y basin s (e.g . o f nearb y basin s (e.g . Nort h Sea , M0r e Basin , Cloetingh et al. 1990) . However , the mechanism offshor e Greenland , S0rvestnaget Basin, Rockall as well a s its temporal and spatial resolution has Trough , Porcupin e Basin) . Becaus e o f th e been a matter of debate. Recently, some progress associatio n o f uplif t an d subsidence , th e has bee n mad e wit h respect t o timing , an d tw o mechanis m drivin g thes e processe s i s probabl y phases tha t hav e influence d th e whol e Nort h tectoni c (e.g . Dor e e t al . 1999 ; Japse n & Atlantic regio n hav e bee n identifie d (e.g . Rii s Chalmer s 2000) . Variou s mechanism s hav e 1996; Martinse n e t al . 1999) . Th e firs t i s a bee n propose d includin g intraplate compression Paleogene phas e primaril y associate d wit h (Cloeting h et al. 1990 ; Dore et al. 1999) , mantle Eocene riftin g an d contemporaneou s volcanis m phas e change s (Rii s & Fjeldskaa r 1992) , generated by the Iceland mantle plume (e.g. Dore magmati c underplatin g (Co x 1993 ; Brodi e & et al. 1999) . Ther e i s considerable evidenc e tha t Whit e 1995 ) an d small-scal e asthenospheri c this even t affecte d al l region s aroun d the Nort h convectio n (Vagne s & Amundse n 1993 ; Rohr Atlantic (Britain , Norwa y an d Greenland) , ma n & van der Beek 1996 ; Stuevol d & Eldholm although it s magnitud e differed fro m regio n t o 1996) , amon g others. region (e.g . Gree n e t al . 1993 ; Rii s 1996 ; Da m I n thi s paper , w e briefl y revie w propose d et al . 1998 ; Dor e e t al . 1999) . Th e secon d i s a mechanism s for uplift aroun d the North Atlantic. From: DORE , A.G., CARTWRIGHT , J.A., STOKER , M.S., TURNER , J.P . & WHITE , N . 2002 . Exhumation of th e North Atlantic Margin: Timing, Mechanisms and Implications for Petroleum Exploration. Geological Society , London, Special Publications , 196 , 27^3. 0305-8719/027$ 15.00 © The Geological Societ y of London 2002 .

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Subsequently, we focus on southern Norway, and review the evidence for the timing and magnitude of Neogen e uplif t an d denudation . W e calculate tectonic uplif t fro m th e elevatio n o f plateau surface remnant s an d a n estimat e o f flexura l rigidity fro m th e coherenc e betwee n Bougue r gravity an d topography . Thes e result s ar e the n used t o differentiat e betwee n mechanism s proposed. Additionally , w e presen t seismi c tomographic image s o f P-wav e velocitie s i n th e upper mantl e aroun d th e Nort h Atlanti c an d correlate thes e wit h othe r data , t o arriv e a t th e most plausibl e mechanis m operatin g durin g Neogene time .

Mechanisms proposed Early studie s wer e mainl y focuse d o n Norwa y and primaril y base d o n geomorphologica l observations (e.g . Holtedah l 1953 ; Torsk e 1972). Th e rugge d mountain s o f souther n Nor way, spectacula r fjord s an d preservatio n o f plateau surface s a t hig h altitud e (Gjessin g 1967; Peulvas t 1985 ; Riis 1996 ) strongly suggest a recen t upwar p o f th e souther n Norwegia n landmass. Intensiv e hydrocarbo n exploratio n o f offshore Norwa y ha s provide d furthe r evidenc e from observe d structura l basinwar d di p o f pre Neogene strat a an d build-u p o f larg e clasti c wedges (Jord t e t al. 1995) . Palaeogeographica l reconstructions hav e adde d additiona l evidenc e for a lat e emergenc e (Dor e 1992a , 1992b) . Apatite fission-trac k (AFT ) dat a constrai n denudation t o b e a mainl y Neogen e event , starting a t c . 3 0 Ma. Denudatio n occurre d i n a dome-like pattern wit h the amplitude of denudation decreasin g radiall y outwar d fro m a maxi mum o f 2. 0 ± 0. 5 km a t th e centr e (Rohrma n et al . 1995) . Th e timin g o f th e onse t o f denudation i s consisten t wit h offshor e strati graphic evidenc e (Rundber g & Smalle y 1989 ; Jordt e t al. 1995) . Any mechanis m proposed , t o b e successful , has to explain this timing of events (van der Beek & Rohrma n 1997) . Som e worker s favou r a Paleogene onse t o f uplif t fo r souther n Norway , but mos t availabl e evidenc e strongl y suggest s a primarily Neogen e even t (Rundber g & Smalle y 1989; Jense n & Schmidt 1993 ; Jord t et al. 1995 ; Rohrman e t al. 1995 ; Rii s 1996 ; Martinse n et a l 1999). Early interpretation s suggeste d tha t uplif t o f southern Norwa y an d othe r margin s aroun d th e North Atlanti c wa s associate d wit h PaleoceneEocene break-up an d plume activity (e.g. Torsk e 1972; Co x 1993) . However , timin g o f doma l uplift a t c . 30Ma , i.e . 25-30M a afte r onse t of Eocen e volcanis m an d rifting , preclude s

significant synrif t uplift . Moreover , dynami c plume-generated uplif t i s transient , an d shoul d reverse int o subsidenc e afte r break-up . Perma nent uplif t ca n b e generate d b y magmati c underplating, but in the case of southern Norway there i s n o sig n of onshor e Cenozoi c intrusions. Eocene magmati c activit y too k plac e 300 400 km offshore . Riis & Fjeldskaa r (1992 ) propose d tha t mos t of the Norwegian uplift wa s caused by PliocenePleistocene isostati c readjustmen t t o glacia l erosion. However , thei r stud y showe d tha t additional tectoni c uplif t i s require d t o explai n the present-da y elevatio n o f morphologica l surfaces. The y explaine d thi s b y mantl e phas e transitions as a result of erosional unloading, but the dynamic s o f mantl e phas e change s driving this tectoni c uplif t componen t ar e a t presen t poorly understood . Non-tectoni c mechanism s that hav e been suggeste d (Eyle s 1996 ) focu s o n climatic deterioratio n an d sea-leve l changes . However, the kilometre-scale magnitude of uplif t is not readily explaine d b y eustasy. Others hav e propose d tha t Neogen e uplif t o f western Fennoscandi a i s a resul t o f therma l instability cause d b y larg e horizonta l tempera ture gradients between the Fennoscandian mainland and the Norwegian-Greenland Sea (Theilen & Meissne r 1979 ; Peulvas t 1985 ; Vagne s & Amundsen 1993) . Thi s woul d se t u p secondar y convection i n th e sub-lithospheri c mantle , causing th e ho t asthenospher e t o ris e an d sub sequently generat e tectoni c uplift. Thi s could be a viabl e model , an d wil l b e discusse d i n th e following sections . Finally, variou s workers have drawn attention to th e synchronou s timing o f margi n uplif t an d anomalous basin subsidence around Norway and the whol e Nort h Atlantic , suggestin g tha t coupled uplif t an d subsidenc e i s flexura l i n nature an d induce d b y intraplat e stres s fluctu ations (Cloetingh et al. 1990) . However, flexureinduced uplift i s not consistent with the observed correlation betwee n topograph y an d Bougue r gravity anomalie s (u p t o — 80 mgal) belo w th e uplifted region s (Fig . 1) . Althoug h th e gravit y is largel y isostaticall y compensate d (Ballin g 1980), ther e i s a mas s defici t belo w th e regions o f highes t elevatio n tha t canno t b e explained b y topograph y an d difference s i n crustal thickness .

Timing of denudation on the eastern Atlantic margin There are two principal types of constraint on the timing o f Norwegia n denudation : th e structural

TIMING AN D MECHANISMS O F NORTH ATLANTIC CENOZOIC UPLIF T

29

Fig. 1 . Bouguer gravity anomaly map for southern Norway, data from Sverige s Geologiske Undersokels e (1985) .

and stratigraphi c relationshi p o f Cenozoic strat a offshore (Jord t e t al. 1995 ) an d planatio n surfaces onshore , an d th e AF T analysis o f th e Norwegian basemen t (Rohrma n e t al . 1995) . AFT dat a fro m souther n Norwa y defin e a structural dom e wit h the youngest age s increas ing radiall y fro m c . 10 0 Ma a t se a leve l i n th e inner fjords , t o c . 17 0 Ma a t th e to p o f th e Jotunheimen peak s an d aroun d 20 0 Ma a t elevations les s tha n 500 m nea r th e shoreline s (Fig. 2) . Mea n trac k lengt h distribution s ar e more variable , bu t th e younge r age s generall y

correspond t o lo w mea n trac k length s (c. 11.6jjLm ) an d a lac k o f lon g (recent ) track s (Fig. 2) . Th e latte r indicat e fas t coolin g fro m temperatures a t th e lowe r en d o f th e partia l annealing zon e (c . 60-7 0 °C) t o surfac e tem peratures. Althoug h ou r younges t AF T sample s yield mixed ages, it is possible t o extract therma l history informatio n b y usin g state-of-the-ar t modelling techniques . Th e younges t AF T samples (AF T age c. 100 Ma, mean track lengths 11.6(xm) sugges t a predominantl y Neogen e onset of denudation (Rohrman et al. 1995) .

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M. ROHRMAN E T AL.

Northern Norwa y show s simila r AF T base ment age s althoug h the y ten d t o b e somewha t older (aroun d 16 0 Ma) wit h distribution s show ing longer track lengths (Hendriks & Andriesse n 2002). Meanwhile , Neogen e clasti c wedge s offshore fro m Lofote n indicat e mor e recen t erosion (Mokhtar i & Pegru m 1992) . A recen t AFT stud y o f Spitsberge n b y Blyth e & Kleinspehn (1998 ) place d significan t denudatio n around 3 6 Ma, whic h wa s mainl y attribute d b y those worker s t o rif t flan k uplift , followe d b y Pliocene-Holocene glacia l erosion . Th e dat a pattern yield s strongl y varyin g AF T age s (c. 27-5 6 Ma), sometime s withi n th e sam e locality. Thi s patter n suggest s tha t th e sample s experienced othe r processe s tha n simpl e uplif t and erosion, possibl y indicatin g migration o f hot fluids. Thi s interpretatio n i s supporte d b y th e resetting of zircon fission-track ages and vitrinite reflection data , i n th e proximit y o f unrese t samples a t simila r elevations . A possibl e explanation fo r these pattern s migh t b e volcanic and hydrotherma l activit y durin g Miocen e an d

Quaternary tim e (e.g . Vagne s & Amundse n 1993). Simila r fission-trac k pattern s hav e bee n observed i n th e Permia n Osl o Rift , wher e igneous activit y i s wel l documente d (e.g . Rohrman e l al . 1994) . However , a t se a leve l there seem s t o b e a sligh t tren d fro m wes t t o east alon g th e Longyearbye n fjor d (Centra l Spitsbergen), whic h suggest s a decrease i n AF T ages fro m aroun d 5 0 Ma nea r th e basemen t o n the wes t coas t t o 3 8 Ma i n th e Eocen e deltai c strata o f th e Centra l Basin . Thi s concur s with earlier studie s (N0ttvedt el al. 1992) , suggesting Neogene doma l uplif t base d o n geologica l observations, vitrinit e reflectanc e an d seismi c velocity data . Evidently , more AF T sample s o f east Spitsberge n ar e neede d t o verif y o r refut e this assumption. Much of the British Isles lacks the high rugged topography o f souther n Norway . Elevation s ar e more subdue d an d d o no t excee d 1 km i n larg e parts o f England , Wale s an d Ireland . Onl y th e Scottish highland s reach highe r elevations, up to 1343m. Publishe d AF T dat a fro m Scotlan d

Fig. 2 . Mea n elevatio n an d AFT isochrons o f souther n Norwa y draw n a t se a level , definin g a structura l dom e concordant wit h topography, fro m Rohrma n e t al. (1995). Typica l track length distributions for selected samples are show n o n the right .

31

TIMING AN D MECHANISMS O F NORTH ATLANTIC CENOZOI C UPLIF T

(Lewis e t al 1992 ) sugges t predominantl y Triassic-Jurassic ages . Onl y nea r th e Eocen e intrusions are younger ages found. However, data are too sparse to infer any details for this region. AFT dat a fro m norther n Englan d an d th e Iris h Sea are a sho w AF T age s betwee n 4 5 Ma an d > 400 Ma (Gree n e t al . 1993 , 1997 , 2001) , suggesting significant exhumation around 60 Ma. This uplif t i s possibl y centre d i n th e Eas t Iris h Sea an d decreasin g t o th e SE , concordan t wit h the Mesozoi c sedimentar y outcro p patter n o f southeastern England . Recently , Japse n (1997 ) suggested evidenc e fo r a Neogen e denudatio n phase base d o n AF T dat a an d compactio n studies fro m easter n Englan d an d th e wester n North Sea . H e estimate d tha t Paleogen e an d Neogene denudatio n wer e equa l i n magnitud e (around 1 km each) . However , th e spatia l pattern o f bot h Neogen e an d Paleogen e denudation fo r th e entir e Unite d Kingdo m requires furthe r study . Moreover, the area of highest denudation (East Irish Sea) is at present below sea level, in contrast to th e dome-shape d topograph y o f th e othe r regions (souther n Norway, northern Norway and Spitsbergen).

Quantification of Neogene uplif t and denudation i n southern Norway Regional change s i n surfac e elevatio n ar e th e combined resul t o f tectoni c uplift , erosio n an d the isostati c respons e t o erosion . T o us e th e present-day elevatio n o f a n uplifte d regio n t o constrain tectoni c processes , th e component s o f elevation chang e tha t resul t fro m erosio n an d isostatic reboun d must be quantified (Englan d & Molnar 1990 ; Gilchris t et al. 1994) . The tectonic uplift (W T) is related t o the present-day elevatio n

(Ho) an d th e amoun t of denudation (A£) by

where H { i s the initial elevatio n and / i s isostatic rebound:

where p c i s th e densit y o f th e erode d crusta l section, pa is sublithospheric mantle density, kE is the spatial wavenumber of erosional unloading, g is acceleratio n o f gravit y an d D i s flexura l rigidity (va n de r Bee k e t al. 1994 ; se e Table 1) . As D — > 0 , equatio n (2 ) simplifie s t o / = AEp c/pa, th e loca l isostati c solution . A s D — > o o (fo r a n infinitel y stron g lithosphere) , I — * 0. A regiona l analysi s o f tectoni c uplif t therefore require s a n estimat e o f H^ an d D , a s well as the ability to map out spatial variations in A£ (Abbot t e t al . 1997 ; Smal l & Anderso n 1998). Whereas th e AF T thermochronologica l dat a discussed abov e giv e u s a hig h tempora l resolution t o decipher the denudatio n histor y o f southern Norway , th e spatia l resolutio n o f ou r data i s rathe r coarse . W e therefor e us e th e elevation o f preserve d remnant s o f a platea u surface (th e Talaei c surface' ) t o spatiall y constrain th e amount s o f denudation , isostati c rebound, and tectonic uplif t i n southern Norway. The correlation o f plateau remnants and their use in reconstructin g landscap e developmen t i s a relatively hazardou s undertakin g becaus e o f th e general lack of temporal constraints (Brown et al. 1999; Summerfiel d 1999) . I n Norway , a s else where, ther e i s n o consensu s o n th e ag e o f th e surface remnants , whic h hav e bee n varyingl y interpreted as being of Jurassic to Paleogene age , nor o n thei r correlatio n (Gjessin g 1967 ; Torsk e

Table 1 . Parameter values employed Symbol PC Pa g

D E v f k a L

Description

Value

Crustal density Asthenospheric density Gravitational acceleratio n Flexural rigidit y o f the lithospher e Young's modulus Poisson rati o Ratio o f surface to base loadin g o f lithospher e Thermal conductivity Thermal diffusivit y Lithospheric thickness afte r thinning

2600 kg nT3 3250 kg m"3 9.8 m s ~2 8.9 X 1021 to 7. 2 X l 0 2 2 N m 10 n Nm~ 2 0.25 1 0.0006 cal (c m s 0.01 cm 2 s"1 80-110km

°cr '

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M. ROHRMAN ETAL.

Fig. 3 . Maps o f mean elevatio n o f souther n Norwa y (fro m ETOPO-5 global topograph y data) , elevation of the Palaeic surfac e (afte r Riis & Fjeldskaar 1992 ) an d amoun t of erosion, calculate d b y subtractin g mean elevation from th e elevatio n o f the Palaeic surface .

1972; Peulvas t 1985 ; Dor e 1992a ; Rii s & Fjeldskaar 1992 ; Riis 1996) . We follow the mos t recent correlatio n o f Riis & Fjeldskaa r (1992 ) an d Rii s (1996) , wh o suggested the Palaeic surfac e to be of Paleogen e age becaus e (1 ) ou r fission-trac k dat a sugges t

that samples from clos e to the Palaeic surfac e in the Hardangervidd a are a reache d surfac e tem peratures i n Paleogen e time s (Rohrma n e t al. 1995) and (2) the elevation of the Palaeic surface in th e Jotunheime n area i s consisten t wit h th e amount o f Neogen e denudatio n recorde d b y

TIMING AN D MECHANISMS O F NORTH ATLANTIC CENOZOI C UPLIF T

33

Fig. 4 . Correlatio n o f amoun t o f Neogen e denudation , calculate d b y subtractin g present-day topograph y fro m elevation of Palaeic surface (see Fig. 3) with apatite fission-track ages (from Rohrman etal. 1994,1995). Neogene denudation fo r fission-trac k sample s i s corrected fo r sampl e elevation wit h respect t o mea n elevatio n o f the 5 ' resolution topographic grid .

fission-track thermochronolog y o f sample s fro m close t o se a leve l (2. 0 ± 0. 5 km; va n de r Bee k 1995). Figure 3 show s th e present-da y elevatio n o f southern Norwa y a s well a s the elevatio n o f th e Palaeic surface. The latter was digitized fro m th e map of Riis & Fjeldskaar (1992) and interpolated using a continuous curvatur e algorithm (Smit h & Wessel 1990) . Th e amoun t o f denudatio n sinc e the formation of the surface (i.e. since the end of Paleogene time) can be calculated by subtracting the present-da y elevatio n H Q fro m th e elevatio n of th e Palaei c surface . The resul t show s a mean amount of denudation o f c. 400m, with maxim a of > 1000m i n th e inne r fjord s an d alon g th e northwestern coast . Th e patter n o f denudatio n that i s calculate d i n thi s manne r i s largel y consistent with the pattern deduced from fissiontrack thermochronology . Figur e 4 show s th e correlation betwee n th e amoun t o f denudatio n calculated using this approach an d the AFT ages from Rohrma n e t al. (1995) . Althoug h th e fission-track age s ar e mixe d ages , wit h minim a around 10 0 Ma, a clea r tren d emerge s wit h samples wit h th e younges t fission-trac k age s encountered in the regions of maximum Neogene denudation. Late Paleocen e an d earl y Eocen e marin e diatoms ar e encountere d o n planatio n surface s in Swede n an d Finlan d (Fenne r 1988) , whic h may b e correlate d wit h th e Palaei c surfac e i n

Norway. O n th e basi s o f thi s observation , Rii s (1996) suggested that th e surfac e wa s at or nea r sea level in early Cenozoic time. This assertion is consistent wit h th e offshor e sedimentatio n dat a (Jordt e t al . 1995) , whic h sugges t ver y lo w sediment inpu t fro m Fennoscandi a befor e lat e Oligocene time . W e therefor e assum e tha t th e Palaeic surface can be used a s a marker no t only of Neogen e denudation , but als o of uplift . If we suppose Hi — 0 in equation (1), then the present-day elevatio n o f th e Palaei c surfac e represents HQ + A E — MT + /• I n th e cas e o f local isostas y ( D — 0), th e amoun t o f isostati c rebound reduces to 7 = AEp c/pa (c. 0.8AE1 for the values in Table 1 ) and can be calculated directl y from th e patter n o f denudatio n (Fig . 3) . Th e tectonic uplif t U T the n equal s th e present-da y elevation o f th e Palaei c surfac e minu s th e isostatic uplif t / . Figur e 5 show s calculate d amounts of Neogene tectoni c uplif t an d isostati c rebound fo r th e loca l isostati c case . Tectoni c uplift i n this case has the sam e domal pattern as the mea n elevatio n an d th e elevatio n o f th e Palaeic surface , reachin g maximu m value s o f c. 1500 m fo r th e hig h mountai n area s (wher e Neogene denudatio n i s negligible) , wherea s isostatic reboun d reache s a maximu m o f c. 750 m i n th e inne r fjord s an d o n th e N W coast. However, because the isostatic response to denudation is generally regional instea d of local, the pattern and the amount of tectonic uplif t wil l

34

M. ROHRMAN ETAL.

Fig. 5 . Calculated tectoni c uplif t an d isostatic reboun d fo r a model o f local isostatic response t o erosion. I n this case isostatic reboun d equal s 0. 8 times th e amount of erosion; tectonic uplif t wa s calculated by subtractin g this amount fro m th e elevatio n o f the Palaei c surfac e (see text).

also be a function o f the flexural rigidity D of the southern Norwegia n lithosphere . Estimating flexural rigidity To make a realistic assessmen t o f the amoun t of tectonic uplift , th e flexura l rigidity , o r corre sponding equivalent elastic thickness T e, must be estimated. Flexura l rigidit y ma y b e readil y estimated b y a n analysi s o f th e coherenc e o f topography an d Bougue r gravit y anomalies . Gravity an d topograph y wil l b e coheren t a t long wavelength s bu t no t a t shorte r ones, th e wavelength a t whic h coherenc e break s dow n being dependen t o n D. The observe d coherenc e 7§ is define d a s (Forsyt h 1985)

where E 0(k} i s the average power of topography for a discret e wavenumber , Ei (k) i s th e averag e power of gravity for that wavenumber, C(k} is the

cross-spectral powe r o f gravit y an d topograph y and th e overba r indicate s averagin g ove r a wavenumber band . Fo r a mechanicall y aniso tropic lithospher e a simila r concept ca n b e use d (Simons e t al. 2000) , bu t i n thi s analysi s w e assume isotropy. The theoretical coherenc e fo r a plate loade d bot h at its surface and base i s given by (Forsyt h 1985)

where H T i s th e amplitud e o f surfac e deflection as a resul t o f surfac e loading , W T i s th e amplitude o f deflectio n o f th e bas e o f th e plat e (i.e. th e Moho ) a s a resul t o f surfac e loading , //B th e amplitud e of surfac e deflection resultin g from loadin g a t th e bas e o f th e plat e an d W B the amplitud e o f Moh o deflectio n resultin g from bas e loading . // T, // B, W T and W E depen d on th e wavenumbe r k an d flexura l rigidit y D

TIMING AND MECHANISMS O F NORTH ATLANTIC CENOZOI C UPLIFT

via

with / th e rati o o f surfac e t o bas e loading ; Ap = (p a - p c); ( f > = \ + (D/c 4/pcg) an d
35

& Cathle s 1991) . I t is , however , consisten t with regiona l T & estimate s fro m th e post glacial til t o f palaeoshoreline s (Fjeldskaa r 1997), whic h suggeste d T e ^ 20km, a s wel l as wit h a n independen t coherenc e stud y b y Poudjom Djoman i e t al (1999) , wh o foun d 8 < 7 e < 18k m fo r souther n Norway. We hav e calculate d th e isostati c respons e t o denudation adopting constant T e values of 10 and 20 km, using equation (2). The predicted isostatic rebound ha s a muc h smoothe r patter n tha n th e local isostati c case , reachin g a maximu m o f c. 700 m nea r th e N W coastlin e (Fig . 7) . Th e magnitude o f isostatic reboun d is , however , no t much smaller than for the local isostati c case, as the inferre d flexura l rigidit y i s no t ver y large . Resulting tectonic uplift pattern s are also shown in Fig . 7 ; th e model s incorporatin g flexura l rigidity sho w a simila r dome-shape d uplif t pattern a s th e loca l isostati c case . Uplif t i s centred o n th e region s o f highes t present-da y elevation, reachin g a maximu m o f c . 1500m . This surprisingl y hig h valu e i s relativel y insensitive t o th e adopte d flexura l rigidit y because, a s D increases , th e amoun t of isostati c rebound decreases but also spreads mor e toward the region s of highest present-day elevation and contributes t o th e uplif t o f thes e regions . Th e critical assumptio n i n thi s analysi s i s tha t th e Palaeic surfac e was a t se a level befor e Neogene uplift; i f i t wa s a t som e initia l elevatio n H[ th e inferred Neogen e uplift wil l be overestimated by the sam e amoun t (se e equatio n (1)) .

where £ i s Young' s modulu s an d ^ i s th e Poisson rati o (se e Tabl e 1) . The gravity data we used for this analysis are from th e Sverige s Geologisk e Undersokels e 1985; Fig . 1 ) and topography fro m th e ETOPO5 database . Gravit y an d topograph y dat a wer e projected ont o a 50 0 km X 720 km UT M grid . Data were clipped in the deep offshore area s and tapered towards the mean value along the sides of the grid , befor e bein g transforme d int o the wavenumber domai n usin g a fas t Fourie r transform algorith m (se e va n der Beek 1995) . Results of the coherence analysis are shown in Fig. 6 . Although a best-fittin g / an d T e ca n be estimated independentl y fro m th e dat a usin g a least-squares criterio n (Forsyt h 1985) , w e fee l that thi s ma y pu t to o muc h emphasi s o n a best-fitting numbe r fo r th e presen t dat a quality. Visual inspectio n indicate s a best-fi 21t T e T , „ between 1 0 an d 20k m ( D = 8. 9 X 10 t o Souther n Norway: mechanisms of uplif t 7.1 X 1022 Nm), assumin g / = 1 . This valu e is The relativel y lo w flexura l rigidit y o f th e at th e lo w en d o f estimate s o f flexura l rigidit y lithosphere (c. 10 22Nm) and high tectonic uplif t for Fennoscandi a fro m glacia l reboun d studies , strongly sugges t a n endogenou s caus e fo r th e which constrai n T e t o b e <50k m (Fjeldskaa r uplift. Thi s leave s u s wit h essentiall y tw o possible solutions : intraplat e compressio n an d dynamic mantl e u p welling. Intraplat e processe s require positiv e gravit y anomalie s belo w th e dome regions , i n contras t t o th e observations . Furthermore, intraplat e compression canno t generate th e amoun t o f uplif t required , a s i t i s an orde r o f magnitud e to o lo w (va n de r Bee k 1995). Mantle upwellin g ha s bee n propose d b y various worker s fo r th e Nort h Atlanti c margin s (e.g. Vagne s & Amundse n 1993 ; Stuevol d & Eldholm 1996 ; Rohrma n & van der Beek 1996) . The upwellin g mode l seem s plausibl e fo r Spitsbergen, wher e hig h hea t flo w 2 Fig. 6 . Result s o f coherenc e analysi s fo r souther n (c. 130mWm~ ) prevail s an d mantl e xenoliths are presen t i n Quaternar y alkal i basalts . Thi s Norway. Dots indicate observed coherence for average wavenumber bands, bars denote Ic r errors. Continuous model seem s less obvious fo r souther n Norway, lines are the various equivalent elastic thicknesses T c. where lo w surfac e hea t flo w (c . 40mW m )

36

M. ROHRMAN ETAL.

Fig. 7 . Calculated tectoni c uplif t an d isostatic reboun d fo r a model o f regional isostati c respons e t o denudation and differen t flexura l rigidities , correspondin g t o T e = 10k m ( D = 8. 9 X 1021 Nm; top ) an d T t — 20km (D = 7. 1 X 1022 Nm; bottom) . (See text for discussion.)

TIMING AND MECHANISMS OF NORTH ATLANTI C CENOZOI C UPLIFT

(Cermak 1979 ; Ballin g 1995 ) an d lac k o f volcanism ar e not readily consisten t wit h mantle up welling. However, a recent study by Goes et al (2000) proposed high sub-Moho temperatures in excess o f 1000° C base d o n P - an d S-wav e tomography belo w souther n Norwa y (Fig . 8) . These temperature s ar e simila r t o thos e i n regions suc h a s th e Massi f Centra l (Sobole v et al . 1996) , wher e a mantl e plum e seem s evident. I f thes e temperatur e estimate s ar e correct, th e discrepanc y betwee n surfac e hea t flow and elevate d temperature s a t dept h ca n b e explained by a time lag between equilibration of high lithospheric heat flow and surface heat flow (McGuire & Bohanno n 1989) . Assumin g con ductive heat loss, the change in surface heat flow Ag0 i s related t o a change i n temperature a t the base of the thinned lithosphere A7"i, the thickness of th e lithospher e afte r thinnin g L , an d th e diffusion time f , accordin g t o

where k i s therma l conductivit y an d a th e thermal diffusivit y (Tabl e 1) . For L between 11 0 and 80 km and AT^ from 400 °C to 500 °C, it takes around 6 0 Ma to observ e a c. 10mWm~ 2 heatflow change fo r souther n Norway . Upwelling materia l i s mos t probabl y associ ated wit h advection , therefor e w e ca n us e a 10mWm~ 2 heat-flo w chang e a s a conservativ e estimate. As no regional heat-flow anomalie s are observed i n th e souther n Norwegia n data , w e assume that any dynamic mantle upwelling must be younger than 60 Ma. Along with the evidence

37

of reduce d P an d S wave s belo w souther n Norway (Banniste r e t al 1991 ; se e also below) , this strongl y suggest s tha t th e souther n Norwegian structura l dom e wa s generate d b y activ e mantle upwelling or diapirism i n Neogene time. Seismic tomography of the North Atlantic One of the major advances i n our understanding of the Earth has been the advent and subsequent development o f seismi c tomography , a clas s o f imaging tha t now provides increasingl y detaile d 3D maps o f seismi c velocit y variations that can be related t o thermal an d chemical variation s i n the mantle . Althoug h i t ha s prove n relativel y easy t o imag e downwelling s (i.e . subductin g slabs), i t seem s muc h mor e difficul t t o imag e mantle upwellings (e.g. Grand et al. 1997) . This is becaus e upwelling s ar e likel y t o occu r i n aseismic region s and are not as well sampled by seismic data , especiall y i n th e shallo w mantle . Furthermore, us e o f firs t arrival s o f seismi c waves cause s a natura l bia s towar d fas t anomalies, becaus e annealin g o f wavefront s creates a tendenc y t o underestimat e slo w anomaly amplitudes . Another important issue is that resolution depends on data coverage, which is uneve n owin g t o th e spars e distributio n o f sources (i.e . earthquakes ) an d receiver s (i.e . seismological stations ) an d th e 3 D geometr y of the ra y path s alon g whic h th e seismi c wave s propagate. Mantl e structur e has remaine d unre solved beneat h larg e region s o f th e Nort h Atlantic because of absence o f recording station s and low levels of seismi c activity. With this in mind, we present P-wave velocity maps o f th e Nort h Atlanti c a t variou s depth s (Fig. 9), which provide a snapshot of the mantle

Fig. 8 . Temperature-depth plot of three geotherms belo w souther n Norway . The 40 mW m 2 geotherm (broke n line) is based on surface heat-flo w data . T P and T s ar e geotherms derived from inversion o f P-wave an d S-wav e seismic velocities , respectivel y (Goe s e t al 2000 ; Goe s pers . comm.) . Th e relativ e convergenc e o f T P an d T s should b e noted. VBS, volatile-bearin g peridotite solidus; AS , anhydrous peridotit e solidus; AD , adiabat.

38

M. ROHRMAN ETAL.

Fig. 9 . P-wav e seismi c tomograph y map s fo r th e Nort h Atlanti c a t variou s depths . Fo r th e uppe r mantl e low velocities (red ) correspond t o high temperatures, wherea s hig h velocities (blue) correspond t o low temperatures. EG, East Greenland ; 1C, Iceland; SN , southern Norway ; NN, northern Norway; IS, Irish Sea ; SC , Scotland .

as it is today. Th e image (Fig . 9 ) depicts P-wav e utio n tes t fo r th e velocit y ma p a t 150km , velocity map s throug h th e globa l mode l o f usin g a chequerboar d model . Resolutio n i s Karason & va n de r Hils t (2000 ) fo r th e Nort h rathe r poo r belo w Eas t Greenlan d (EG ) an d Atlantic region . Figur e 1 0 show s a resol - mos t o f th e Nort h Atlanti c Ocean . A bette r

TIMING AN D MECHANISMS O F NORTH ATLANTIC CENOZOIC UPLIF T

39

Fig. 10 . (a ) Latera l variation s o f P-wav e spee d a t 150k m dept h beneat h th e norther n Atlantic , (b ) Inpu t 'chequerboard' mode l fo r resolution test, (c) Result of 'chequerboard ' resolutio n test; as expected, ou r ability to image structur e i n th e uppe r mantl e beneat h th e oceani c region s i s very poo r owin g t o absenc e o f earthquak e sources an d seismologica l stations , bu t beneat h souther n Norwa y th e latera l resolutio n i s reasonable. Vertica l resolution i s poor because o f the smal l incidence angles of the seismi c wave s used for imaging in this region.

defined low-velocit y anomal y i s detected belo w Iceland (1C). This anomaly extends from near the surface to c. 400km, but a deeper structure could have been overlooke d owin g to insufficient dat a coverage (e.g . Kelle r e t al 2000) . Belo w southern Norwa y (SN ) a low-velocit y anomal y is observe d tha t extends fro m > 50 km dept h to > 250 km. The vertical resolutio n i s poor i n this part o f the model , bu t th e slo w anomal y agree s well wit h previous P- an d S-wav e tomographic

studies (Banniste r et al 1991 ; Zielhui s & Nolet 1994a, 1994b ; Marquerin g & Sniede r 1996 ; Bijwaard e t al 1998) . Although this is not likely to be well resolved, the model suggests that a thin high-velocity laye r overlie s a low-velocit y anomaly belo w souther n Norway , whic h i s consistent wit h ou r estimate s of T e and suggests a thinned and relatively weak lithosphere. Farther eastward we encounter the Baltic shield, evident as a thic k (>250km ) high-velocit y area . Th e

40

M. ROHRMAN ETAL.

change fro m a low-velocit y zon e i n th e uppe r mantle o f souther n Norwa y (SN ) t o th e hig h velocities o f th e Balti c shiel d i s dramati c an d suggests som e fundamenta l proces s tha t ha s s o far bee n neglecte d i n ou r curren t vie w o f plat e tectonics. W e remar k tha t simila r feature s hav e been observe d i n southeaster n Australi a (Zielhuis & va n de r Hils t 1996) . Here , a low velocity upper-mantl e anomal y is depicted belo w the mountain s (u p t o 2230m ) o f Easter n Australia, wherea s th e low-altitud e (<200m ) shield toward s th e west show s higher velocities . Regional tomographi c studie s (Bannister et al 1991) indicat e a low-velocit y anomal y belo w northern Norwa y (NN) , but i n ou r global mode l resolution i s too low to infe r an y details fo r thi s region. Simila r resolutio n problem s ar e observed belo w th e Britis h Isles , wher e fain t low-velocity anomalie s ar e presen t belo w th e Irish Se a (IS ) a t c . 200k m depth . Lo w velocities belo w th e Irish Se a were also observe d in previous models (Marquerin g & Snieder 1996 ; Bijwaard e t al 1998) . Scotlan d (SC ) also suffer s from th e aforementione d resolutio n problems , but generall y suggest s highe r velocities .

Discussion Our result s fro m integratin g geological , AFT , geomorphological, Bougue r gravit y an d seismi c tomography dat a strongl y sugges t tha t mantl e upwelling i s presen t belo w souther n Norway . Furthermore, i t seems eviden t that this upwelling was mos t activ e i n Neogen e tim e an d directl y responsible fo r generatin g mos t o f th e present day topography . Whethe r othe r area s experi enced simila r upwelling s i s stil l uncertain , mainly becaus e o f lac k o f sufficien t data . However, there are strong indications that similar mechanisms generate d th e domal topograph y o f Spitsbergen, norther n Norwa y an d th e Eas t Greenland coas t (Rohrma n & va n de r Bee k 1996). There are also some structura l difference s among th e domes . Wherea s Spitsberge n record s Neogene volcanism , souther n Norwa y i s com pletely devoi d o f an y Cenozoi c magmati c activity. Thi s i s probabl y simpl y constraine d b y the amoun t o f activ e lithospher e thinnin g b y the upwelling. Spitsbergen's lithospheri c thickness is estimated t o b e c . 50k m (Vagne s & Amundsen 1993), whereas th e sout h Norwegia n lithospher e is muc h thicker , probabl y c . 80-10 0 km. Th e Norwegian dom e thu s possibl y represent s a n early stag e of upwelling. Additionally, the P- and S-wave temperature s o f Goe s e t al . (2000 ) fo r Europe sugges t tha t divergenc e o f T P (P-wav e temperature) an d T s (S-wav e temperature ) a t depths < 100k m coul d indicat e th e presenc e o f

melt, a s show n fo r th e Massi f Centra l area . Southern Norwa y show s relativ e convergenc e between T P an d T s (Fig . 8) , possibly suggestin g absence o f melt an d therefor e n o volcanism. Apart fro m explainin g observe d uplif t pat terns, th e mantle upwellings also offe r a solution for offshor e anomalou s basi n subsidence , where asthenospheric materia l flow s fro m belo w th e basins (Nort h Sea , M0r e Basin ) towar d th e upwelling (souther n Norway) . Furthermore , i n this interpretatio n th e timin g o f upwellin g becomes les s o f a n issue , a s th e availabl e evidence suggest s tha t boundar y condition s ar e generated b y riftin g ove r a n anomalousl y ho t asthenosphere. Subsequen t uplift pulse s might be inferred b y change s i n upwelling flu x (Rohrma n & va n der Bee k 1996 ; Whit e & Lovell 1997) . Some concern s remain ; mantl e upwelling s seem t o follo w a distinc t sequenc e o f events , starting wit h surfac e uplift , followe d by basalti c volcanism an d finally therma l subsidence . Alternatively, ou r result s migh t sugges t tha t no t al l upwellings caus e surfac e volcanism , a s upwel ling temperatur e an d velocit y coul d b e to o low . Another questio n i s th e relationshi p betwee n plate movemen t an d upwelling . Traditiona l theory suggest s tha t upwelling s remai n station ary wit h respec t t o eac h other , bu t ou r result s suggest tha t this might not alway s be th e case . We than k Shel l U K Exploratio n & Productio n fo r sponsoring th e colo r figure s i n thi s articl e an d Sask i Goes fo r providin g th e temperatur e dat a belo w southern Norway

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RASMUSSEN, E . 1992 . Hydrocarbo n potentia l o f the Centra l Spitsberge n basin . In : VORREN , T.O. , BERGSAKER, E. , DAHL-STAMNES , 0.A. , HOLTER , E., JOHANSEN , B. , LIE , E . & LUND , T.B . (eds ) Arctic Geology an d Petroleum Potential. Nor wegian Petroleu m Societ y Specia l Publication , 2 , 333-361. PEULVAST, J.-P . 1985 . Post-orogeni c morphotectoni c evolution o f th e Scandinavia n Caledonide s durin g the Mesozoi c an d Cenozoic . In : GEE , D.G . & STURT, B.A . (eds ) Th e Caledonide Orogen — Scandinavia an d Related Areas. Wiley, Chichester , 979-995. POUDJOM DJOMANI, Y.H. , FAIRHEAD , J.D. & GRIFFIN ,

W.L. 1999 . Th e flexura l rigidit y o f Fennoscandia : reflection o f th e tectonotherma l ag e o f th e lithospheric mantle . Earth an d Planetary Science Letters, 174 , 139-154 . Rus, R 1996 . Quantificatio n o f Cenozoi c vertica l movements o f Scandinavi a b y correlatio n o f morphological surface s wit h offshore data . Global and Planetary Change, 12 , 331-358.

Rus, R & FJELDSKAAR, W. 1992. O n the magnitude of the Lat e Tertiar y an d Quaternar y erosio n an d it s significance fo r th e uplif t o f Scandinavi a an d th e Barents Sea . In : LARSEN , R.M. . BREKKE , H. , LARSEN, B.T . & TALLERAAS , E . (eds ) Structural and Tectonic Modelling and its Application to Petroleum Geology. Norwegia n Petroleu m Societ y Special Publication , 1, 163-185. ROHRMAN, M . & VA N DER BEEK, P . 1996 . Cenozoi c postrift doma l uplif t o f Nort h Atlantic margins: a n asthenospheric diapiris m model . Geology. 24, 901-904. ROHRMAN, M. , VA N DER BEEK , P . & ANDRIESSEN , P.A.M. 1994 . Syn-rif t therma l structur e an d post rift evolutio n of th e Osl o Rif t (S E Norway) : new constraints fro m fissio n trac k thermochronology . Earth and Planetary- Science Letters, 127, 39-54. ROHRMAN, M. , VA N DER BEEK , P.A. , ANDRIESSEN , P.A.M. & CLOETINGH , S . 1995 . Meso-Cenozoi c morphotectonic evolutio n o f Souther n Norway : Neogene doma l uplif t inferre d fro m apatit e fission trac k thermochronology . Tectonics, 14 , 704-718. RUNDBERG, Y. & SMALLEY, PC. 1989 . High-resolutio n dating o f Cenozoi c sediment s fro m th e norther n North Se a usin g 87 Sr/86Sr stratigraphy . AAPG Bulletin, 73, 298-308. SIMONS, F.J. , ZUBER , M.T . & KORENAGA , J . 2000 . Isostatic respons e o f th e Australia n lithosphere : estimation o f effectiv e elasti c thicknes s an d anisotropy usin g multitape r spectra l analysis . Journal o f Geophysical Research. 105 . 19163-19184. SMALL, E . & ANDERSON , R.S. 1998 . Pleistocen e relief production i n Laramid e mountai n ranges , wester n Unites States . Geology, 26, 123-126 . SMITH, W.H. R & WESSEL , P . 1990 . Griddin g with continuou s curvatur e spline s i n tension . Geophysics, 55 , 293-305. SOBOLEV, S.V. , ZEYEN , H. , STOLL , G. , WERLING , R . ALTHERR, R . & FUCHS , K . 1996 . Uppe r mantl e temperature fro m teleseismi c tomograph y o f French Massi f Centra l includin g effect s o f composition, minera l reactions , anharmonicity . anelasticity an d partia l melt . Earth an d PlanetaryScience Letters, 139, 147-163 . STUEVOLD, L.M . & ELDHOLM , O . 1996 . Cenozoi c uplift o f Fennoscandia inferred from a study of th e mid-Norwegian margin . Global an d PlanetaryChange, 12 , 359-386. SUMMERFIELD, M.A . 1999 . Geomorpholog y an d global tectonics : introduction . In: SUMMERFIELD , M.A. (ed. ) Geomorphology an d Global Tectonics. Wiley, Chichester, 3-11. " Sveriges Geologisk e Undersokelse , 1985 . Scandina vian Caledonide s gravit y anomal y map . In : GEE, D.G. & STURT , B.A . (eds ) Th e Caledonide Orogen—Scandinavia an d Related Areas. Wiley , Chichester. THEILEN, R & MEISSNER , R . 1979 . A compariso n o f crustal an d uppe r mantl e feature s i n Fennoscandia and th e Rhenis h Shield , tw o area s o f recent uplift . Tectonophysics, 61 , 227-242.

TIMING AND MECHANISMS O F NORTH ATLANTI C CENOZOI C UPLIFT TORSKE, T . 1972 . Tertiar y obliqu e uplif t o f wester n Fennoscandia; crusta l warpin g i n connectio n wit h rifting an d break-u p o f th e Laurasia n continent . Norges Geologiske Unders0kelse, 273 , 43-48 . VAGNES, E . & AMUNDSEN , H.E.F . 1993 . Lat e Cenozoic uplif t an d volcanis m o n Spitsbergen : caused b y mantl e convection ? Geology, 21 , 251-254. VAN DE R BEEK , P.A . 1995 . Tectonic evolution o f continental rifts: inferences from numerical modelling an d fissio n track thermochronology. Ph D thesis, Vrije Universitei t Amsterdam. VAN DE R BEEK, P.A . & ROHRMAN , M . 1997 . Passiv e margin uplift aroun d the North Atlantic region and its rol e i n Norther n Hemispher e lat e Cenozoi c glaciation: Comment . Geology, 25, 282 . VAN DE R BEEK , P.A. , CLOETINGH , S . & ANDRIES SEN, P.A.M . 1994 . Mechanism s o f extensiona l

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basin formatio n an d vertica l motion s a t rif t flanks: constraint s fro m tectoni c modellin g an d fission-track thermochronology . Earth an d Planetary Science Letters, 121 , 417-433 . WHITE, N. & LOVELL, B. 1997 . Measuring the pulse of a plume with the sedimentary record . Nature, 387 , 888-891. ZIELHUIS, A . & NOLET , G . 19940 . Dee p seismi c expression of an ancient plat e boundary in Europe. Science, 265 , 79-81 . ZIELHUIS, A. & NOLET, G . 19946 . Shear-wave velocity variations in the uppe r mantl e beneath centra l Europe. Geophysical Journal International, 111, 695-715. ZIELHUIS, A . & VA N DER HILST, R.D . 1996 . Mantl e structure beneat h th e easter n Australia n regio n from partitione d wavefor m inversion. Geophysical Journal International, 127 , 1-16 .

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Paleocene initiation of Cenozoic uplift i n Norway S. B. NIELSEN 1, G. E. PAULSEN 1, D. L. HANSEN 1, L. GEMMER 1, O . R. CLAUSEN 1, B. H. JACOBSEN 1, N. BALLING 1, M. HUUSE 2 & K. GALLAGHER 3 1

Department of Earth Sciences, Aarhus University, Finlandsgade 8, DK-8200 Aarhus N, Denmark (e-mail: sbn@ geo.aau.dk)

Department of Geology & Petroleum Geology, University of Aberdeen, King's College, Meston Building, Aberdeen AB24 3UE, UK ^Imperial College of Science, Technology & Medicine, South Kensington, London SW7 2AS, UK Abstract: Th e timin g o f Cenozoic surfac e uplif t i n N W Europ e relie s o n th e assumptio n that the sedimentary response in basins is synchronous with tectonic processe s i n the source areas. However , man y of the phenomena commonl y use d to infer recent uplif t may as well be a consequence of climate change and sea-level fall. The timing of surface uplift therefor e remains unconstraine d fro m th e sedimentar y recor d alone , an d i t become s necessar y t o consider th e constraint s impose d b y physicall y an d geologicall y plausibl e tectoni c mechanisms, whic h have a causal relation to an initiating agent. The gradual reversal of the regional stres s fiel d followin g th e break-u p produce d mino r perturbation s t o th e therma l subsidence o n the Norwegian Shel f an d in the North Sea. Pulses o f increased compressio n cannot b e th e caus e o f Cenozoi c lan d surfac e uplif t an d accelerate d Neogen e basi n subsidence.Virtually deformation-free regional vertica l movements could have been caused by change s i n th e densit y colum n o f th e lithospher e an d asthenospher e followin g th e emplacement of the Iceland plume. A transient uplift componen t was produced as the plume displaced dense r asthenospher e a t the base o f the lithosphere . Thi s componen t decaye d a s the plume material cooled . Permanen t uplif t a s a result of igneous underplating occurred in areas o f a thin lithosphere (som e Palaeozoi c an d Mesozoic basins ) or for lithosphere under extension a t the time of plume emplacement (th e ocean-continent boundary). In areas of a thicker lithospher e (Eas t Greenland, Scotlan d an d Norway) plume emplacement ma y have triggered a Rayleigh-Taylo r instability , causin g partia l lithospheri c delaminatio n an d associated transient surface uplift a t a decreasing rat e throughout Cenozoic time . A possible uplift histor y fo r the adjacen t land areas henc e reads: initia l transient surfac e uplift aroun d the break-u p time a t 5 3 Ma cause d b y plum e emplacement , an d permanent tectoni c uplif t caused b y lithospheri c delaminatio n an d associate d lithospheri c heating . Th e permanen t tectonic uplif t increase d throug h Cenozoic tim e a t a decreasing rate. Denudation acted o n this evolvin g topograph y an d reduce d th e averag e surfac e elevation , bu t significantl y increased th e elevation o f the summi t envelope. The marke d variation s in the sedimentar y response i n the basins were caused b y climatic variation s and the generally fallin g eustatic level. Thi s scenari o bridges th e ga p betwee n th e idea s o f Paleocene-Eocene uplif t versu s repeated Cenozoi c tectoni c activity : th e tectoni c uplif t histor y wa s initiate d b y th e emplacement o f th e Icelan d plume , bu t continue d throughou t Cenozoi c tim e a s a consequence of early plume emplacement, with climatic and eustatic control on denudation. The mechanism is consistent with topography, heat flow, crustal structure, and the Bouguer gravity o f Norway, and may be applicabl e als o to East Greenland .

Geomorphological evidenc e an d th e ag e an d e t al. 1990 , 1992) , throughou t Cenozoi c time , structure of sediments poin t to the occurrence o f Contemporaneously , adjacen t continenta l area s large-scale Cenozoi c vertica l movements i n NW an d thei r inne r shelve s suc h a s th e Britis h Isle s Europe (Rii s & Fjeldskaa r 1992 ; Rii s 1996 ; (Japse n & Chalmer s 2000) , norther n an d Stuevold & Eldhol m 1996) . Th e Norwegia n souther n Norwa y (Rii s 1996 ; Lidmar-Bergstro m continental shelf , th e Vikin g Grabe n an d th e e t al. 2000), southern Sweden (Lidmar-Bergstrom central Nort h Se a al l experience d subsidence , 1991) , as well as the Barents Sea and the margins and apparently at an accelerating rat e (Cloetingh o f th e Nort h Se a Basi n (Dor e 1992 ; Jense n & From: DORE , A.G., CARTWRIGHT, J.A., STOKER, M.S. , TURNER, J.P. & WHITE , N. 2002 . Exhumation o f th e North Atlantic Margin: Timing, Mechanisms an d Implications for Petroleum Exploration. Geologica l Society , London, Special Publications, 196, 45-65. 0305-8719/027$ 15.00 © The Geological Society o f London 2002.

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Schmidt 1992 ; Japse n 1998 ) experience d uplif t and denudation . A simila r patter n o f vertica l movements i s observed a t the western margi n of the Nort h Atlanti c (Larse n 1990 ; Japse n & Chalmers 2000 ; Johnson & Gallagher 2000) . Continued Cenozoi c basi n subsidenc e ca n generally be understood in terms of a continuation of Palaeozoic-Mesozoi c mainl y rift-initiate d basin subsidenc e (McKenzi e 1978) . However , much controvers y exist s abou t th e causa l relationships regardin g th e simultaneou s Ceno zoic uplif t an d denudatio n o f th e adjacen t lan d areas. Wa s ther e a n initia l Paleocene-Eocen e surface uplif t relate d t o th e magmaticall y dominated openin g o f th e Nort h Atlantic , followed b y a climatically controlle d denudatio n response (Rii s & Fjeldskaa r 1992) , o r hav e a number o f tectoni c pulses , particularl y i n Paleocene an d (mainly ) i n Neogen e time s (Rii s 1996; Stuevol d & Eldhol m 1996) , bee n activ e throughout Cenozoi c time ? The role of climate and eustasy The argument s fo r th e timin g o f Cenozoi c tectonics ar e predicate d o n th e assumptio n tha t the sedimentar y respons e i n basin s reflect s contemporaneous tectoni c processe s i n th e source areas . However , denudatio n depend s

strongly o n climat e (whic h o n a regiona l leve l controls precipitatio n an d vegetatio n cover) , o n the erosiona l bas e leve l define d b y th e eustati c sea level , an d th e connectivit y o f th e drainag e basin t o th e regiona l bas e level . The questio n is then t o wha t exten t climat e change s ma y hav e influenced th e denudatio n proces s i n th e Nort h Atlantic region . Among th e climat e variations , obviousl y th e Quaternary glaciation s hav e ha d a majo r denudational impac t i n N W Europ e an d i n Greenland. Othe r climat e variation s wer e mor e subtle; however, i t i s wel l documented tha t pos t Late Eocen e tim e wa s characterize d b y a significant globa l climat e coolin g an d a fallin g eustatic level of about 200-250 m (e.g. Haq el al 1987; Mille r et al 1998) . These change s can be correlated wit h a world-wid e chang e i n sedi mentation styl e fro m carbonate-dominate d environments t o siliciclastic-dominate d systems with hig h rate s o f sedimentation . Mor e so , th e mid-Miocene climat e an d eustati c change s signalled a furthe r world-wid e acceleratio n i n the siliciclastic flux (Donnelly 1982; Bartek el al. 1991). Th e widesprea d distributio n o f thes e contemporaneous phenomen a point s to a global scale cause, such as climate change and sea-level fall. Cenozoi c climat e changes i n th e Nort h Se a region have been addressed i n moderate detail by

Fig. 1 . Averaged Fennoscandia n Bougue r gravit y field. Redraw n afte r Ballin g (1984) .

PALEOCENE INITIATION O F CENOZOIC UPLIFT IN NORWAY

Buchardt (1978) , wh o derive d a 6 18 O recor d based o n macrofossils. His results are in general agreement wit h thos e derive d fro m calcareous nannoplankton studie s (Mille r e t al. 1998 ) an d show th e onse t o f Cenozoi c coolin g i n lat e Eocene-Oligocene times (se e Huus e 2002b) . Apart fro m blurrin g th e causa l relatio n between tectonic s i n th e sourc e are a an d sediment stratigraphy , th e dependenc e o f denudation o n climate ca n effec t a coupling betwee n climatic deterioration an d the uplif t o f mountain ranges. Molnar & England (1990) argued that the transition t o a coole r an d mor e erosiv e climat e enhances relie f i n mountai n range s an d raise s mountain crest s i n isostati c respons e t o th e removal o f load . Thi s i s i n contras t t o Whippl e et al . (1999) , wh o claime d tha t erosiona l processes i n genera l reduc e relie f an d tha t th e isostatic peak uplift i s negligible. However , if the Norwegian Caledonide s wer e indee d pene plained at the onset of the North Atlantic opening it is clear that significant relief has been produced in th e mea n tim e (Fig . 2 , below) . W e therefor e believe tha t the arguments of Molnar & England (1990) mus t als o b e considere d i n th e Nort h

47

Atlantic region . A furthe r discussio n ha s bee n given by Huuse (2002b) . Tectonics Irrespective o f th e denudationa l respons e i n basins, primary surfac e uplift mus t be caused by physically plausible tectonic mechanisms, which lend themselves to a quantitative description and hypothesis testing. During mid-lat e Cenozoi c time , mid-ocea n ridge compressiona l stres s produce d larg e offshore dome s o n th e Norwegia n margi n (e.g. th e Hellan d Hanse n an d Orme n Lang e Domes), an d ha s bee n propose d t o b e a n agent i n Neogen e surfac e uplif t o f Norwa y b y either (1 ) th e bucklin g o f a n elasti c plat e (e.g . Cloetingh e t a l 1992 ; Dor e 1992 ) o r (2 ) thickening o f th e continenta l lithospher e b y shortening. (1) Elasti c buckling would for the wavelength (200km) and amplitude (1 km) require forces far beyond the compressional strength of continental lithosphere. Bucklin g o f a n elasti c plat e i s therefore no t a possibl e mechanism . Lambec k

Fig. 2 . Topographic relief o f southern Norwa y wit h locations of topographic profiles .

48

S. B. NIELSEN ETAL.

(1983) argue d tha t a viscoelasti c plat e wit h erosion an d depositio n produce s long-wave length large-amplitud e deflection s that resembl e buckling fo r stresse s tha t are much les s tha n th e elastic bucklin g stresses . However , thi s mode l fails the gravity and structural tests in Norway, as ridges an d highland s woul d becom e associate d with Bougue r gravit y maxim a becaus e o f a n elevated Mono , wherea s basin s woul d becom e associated wit h Bouguer gravity minima becaus e of a depressed Moho . A s pointed ou t by Ballin g (1984), th e Bougue r gravit y o f th e Norwegia n highlands (Fig . 1 ) show s significan t minim a (deeper tha n — 80 mgal) wit h a n axi s followin g closely th e axi s o f the topography . (2) Th e uplif t o f Scotland , th e Norwegia n Caledonides an d Eas t Greenlan d i s virtuall y deformation fre e an d canno t b e cause d b y shortening o f th e continenta l lithosphere , a s th e following simpl e argument demonstrates (Brodi e & Whit e 1994) . Th e exces s crusta l thickness , x, necessary t o suppor t topograph y o f heigh t h i s given by x = /zp m/Ap, where p m is the density of the uppe r mantl e an d A p is th e densit y contras t between th e uppe r mantl e an d th e crusta l root . For pm = 3300k g m~ 3 and Ap = 400k g m" 3, x

becomes 8.2 5 km per km of topography h. For an initial crusta l thickness o f 3 5 km, th e horizonta l shortening neede d t o produc e 1 km o f topo graphy amount s t o 24% , an d th e fina l dept h t o Moho become s 44.2 5 km, which is similar to the present-day Moh o dept h unde r th e Norwegia n Caledonides. However , 24 % o f horizonta l shortening woul d hav e produce d significan t compression structures , which are no t observed . The reversal of the regional stress field that the Norwegian Shel f an d th e Nort h Se a mus t hav e experienced doe s hav e consequence s fo r th e subsidence pattern , althoug h no t t o th e exten t previously assumed . Belo w w e discus s thi s i n more detail . Vertical movement s o n a regiona l scal e occur a s a consequenc e o f change s i n th e density colum n o f th e lithospher e an d astheno sphere an d th e associate d isostati c adjustments , or a s so-calle d dynami c topograph y supporte d by vertica l stresse s a t th e bas e o f th e lithosphere, whic h ar e generate d b y flo w i n the mantle . Th e detectio n o f dynami c topo graphy, particularl y ove r th e oceans , remain s controversial (Lithgow-Bertellon i & Silve r 1998). I n thi s pape r w e limi t ou r scop e t o

Fig. 3 . (a) Present-day topography (60°N) , summi t envelop e (dashed line) , and initial surface befor e erosion ; (b ) initial an d present (dashed ) Moh o position; (c ) calculated Bougue r gravity .

PALEOCENE INITIATION OF CENOZOIC UPLIFT IN NORWAY

49

Fig. 4 . (a) Present-day topography (61°N), summit envelope (dashed line), and initial surface befor e erosion ; (b) initial an d present (dashed ) Moho position; (c ) calculated Bougue r gravity.

isostatically supporte d vertica l movements . Virtually deformation-free , isostaticall y sup ported an d positive vertica l movement s ma y b e caused b y (1 ) magmati c underplatin g a t crusta l levels (McKenzi e 1984) , (2 ) lithospheri c delamination (Bir d 1979 ; Housema n e t al. 1981), o r (3 ) emplacemen t o f a ho t plum e a t the bas e o f th e lithospher e (Skogsei d e t al . 2000). I n vie w o f th e magmatic-dominate d opening o f th e Nort h Atlanti c thes e litho spheric -asthenospheric mechanism s therefor e must b e considere d t o b e th e candidate s t o produce th e Cenozoi c long-wavelengt h lan d surface uplift . In th e followin g sectio n w e produc e a brie f catalogue o f mechanism s o f passiv e an d activ e surface uplif t an d presen t ne w result s fo r late r reference i n the discussion. Mechanisms of surface uplif t Passive mechanisms Isostatic adjustments caused by erosion of existing topography. Erosio n o f existin g topo graphy result s i n lowerin g o f th e averag e

topography but , a s propose d b y Molna r & England (1990) , Rii s & Fjeldskaa r (1992 ) an d Gilchrest e t al . (1994) , th e summi t envelop e height ma y increas e becaus e o f th e isostati c uplift following the removal o f mass between th e peaks o r interfluves . I n thi s sectio n thi s mechanism i s discusse d withi n th e framewor k of unloadin g o f a n elasti c lithospher e b y erosion s o a s t o asses s th e amoun t o f tectoni c uplift require d t o produc e th e topograph y o f southern Norway . Th e approac h i s simila r t o that of Riis & Fjeldskaar (1992) excep t that, for a give n amoun t o f erosion , w e determin e th e initial topograph y require d to reproduc e the present topography . Furthermore , w e calculat e the associate d Bougue r gravit y anomaly . Th e timing o f th e erosio n i s no t addresse d b y thi s procedure. Three profile s a t 60 , 6 1 an d 61.7° N (Fig . 2 ) were extracte d fro m th e 1 km X 1 km digita l topography databas e GLOB E (GLOB E Tas k Team et al 1999) . The summit envelope of each profile wa s constructe d b y passin g a smoot h curve through the summits contained in a belt of 40km widt h surroundin g eac h profile . Th e amount erode d i s th e differenc e betwee n th e

50

S. B. NIELSEN ETAL.

Fig. 5 . (a) Present-day topography (61.7°N) , summit envelop e (dashed line) , and initial surface before erosion ; (b) initial an d presen t (dashed ) Moh o position; (c ) calculated Bougue r gravity .

Fig. 6 . The correlation between Bougue r gravit y an d average topographi c heigh t for the three profiles : (a) befor e erosion; (b ) afte r erosion . Continuous line is the correlation derived by Balling (1984) .

summit envelop e an d th e presen t topography . Next th e pre-erosio n Moh o dept h an d topography i s determine d s o tha t afte r erosio n of th e specifie d amoun t an d isostati c com pensation, th e mode l topograph y agree s wit h the presen t topography . I t i s assume d tha t

isostatic compensatio n take s plac e a t th e crust-mantle interfac e wit h a densit y contras t of 400k g m~ 3 , an d tha t th e uppe r mantl e reacts b y flexur e t o loadin g wit h a n elasti c thickness o f 15km . Th e averag e crusta l density i s 2800k g irT 3.

PALEOCENE INITIATION OF CENOZOIC UPLIF T IN NORWAY

Figures 3-5 show the results of applying this procedure. I t appears tha t the process o f erosio n has raised th e summit s a maximum of c. 800 m. The amoun t i s directl y visibl e a s th e difference betwee n initia l an d presen t Moh o depths. Th e maximu m initia l topograph y required t o reproduc e th e presen t topograph y i s located close to the axis of maximum present-day topography an d i s i n th e rang e 1.3-1. 5 km. T o the west and east the initial topography gradually approaches zero . Th e initia l topograph y i s identical t o th e tectoni c uplif t tha t mus t hav e been produce d b y th e mechanis m responsibl e for th e Cenozoi c uplif t o f th e Norwegia n Caledonides, sav e fo r th e possibl e existenc e o f topography befor e th e actio n o f th e mechan ism. Fo r a prio r topograph y o f c . 300m , corresponding approximatel y t o th e topography that woul d emerg e afte r th e post-mid-Cretac eous sea-leve l fall , i f topograph y the n ha d been peneplaine d t o se a level , th e maximu m necessary amoun t o f tectoni c uplif t i s i n th e range 1-1. 2 km. I f positiv e topograph y i s allowed fo r durin g th e tim e o f maximu m se a level th e necessar y tectoni c uplif t become s correspondingly less . Figure 6 show s th e correlatio n betwee n topography an d th e Bougue r gravit y anomal y before (a) and after (b) erosion. The correlation is negative becaus e topograph y is compensated b y a crusta l root , a fac t tha t i s no t modifie d b y a relief-producing denudationa l process. The modelled correlatio n i s i n agreemen t wit h th e correlation determine d b y Ballin g (1984 ) fo r the wester n Fennoscandia n Shiel d (Norway) . The mode l Moh o depth s beneat h souther n Norway ar e qualitativel y i n agreemen t wit h th e depths derived b y Kink et al (1993) , but exhibit some deviations to the east in the Fennoscandian

51

Shield, probabl y becaus e o f breakdow n o f th e assumption of a laterally homogeneou s crust . hostatic and erosional response of basin margins to sea-level fall. I t i s generall y acknowl edged tha t th e eustati c leve l decrease d fro m mid-Cretaceous tim e t o th e presen t da y b y 200-250 m (Haq et al. 1987; Miller et al. 1998) , although smalle r amplitude s hav e bee n suggested (Komin z e t al . 1998) . Thi s eustati c fall expose d th e margin s o f sedimentar y basins, which may have been established under a former higher eustati c level , an d cause d erosiona l unloading an d passiv e isostati c uplif t (i.e . erosional rebound) . The subsidenc e diagra m o f Fig . 7 demon strates thi s mechanis m i n actio n o n a basi n margin. Th e initia l eustati c leve l a t 10 0 Ma i s 250 m, which is also the water depth at that time. Sedimentation ha s fille d th e basin to the brim at the onse t (6 0 Ma) o f a gradua l eustati c fal l o f 250m. A maximu m o f 750 m o f sedimen t i s accommodated b y th e initia l 250 m o f accom modation space because o f sediment loading and compaction o f the alread y existing sediment. As the eustatic fall exposes the sediment surface it is eroded. A t th e presen t da y a tota l o f 500 m o f sediment ha s bee n eroded , an d th e sedimen t a t the surfac e an d a t an y dept h belo w therefor e shows a n overburia l (sensu Japse n 1988 ) of 500m. Figure 8a- c show s th e result s o f combinin g the abov e principl e wit h therma l subsidence , lateral sedimen t transport, and a zone of tectonic inversion in a 1300k m profile. Th e profile coul d simulate an east-west North Sea profile at 56°N. It represent s th e post-mid-Cretaceou s fillin g from th e side s o f a n intra-cratoni c basi n

Fig. 7 . Subsidence diagram showing the effect o f falling se a level (dashed line) on the erosion of a basin margin.

52

S. B. NIELSEN ETAL.

dominated b y therma l subsidenc e i n th e centra l Th e subsidenc e diagra m (Fig . 8a) shows quiet parts, falling eustatic level, and localized tectoni c sedimentatio n an d subsidenc e unti l th e clino inversion (e.g. the Sorgenfrei-Tornquist Zone) at form s migrating from the basin margins reach the the easter n margin . basi n centre . Fo r late r referenc e i t i s note d tha t

Fig. 8 . Subsidenc e diagra m (a) , sedimen t structur e a t OM a (b) , an d overburia l (c ) fo r a 1300k m profil e representing a n intra-cratoni c basi n wit h thermal subsidenc e i n th e centre , fallin g eustati c level , an d sedimen t infilling fro m th e sides . Als o show n is a zone of structura l inversion.

PALEOCENE INITIATIO N OF CENOZOIC UPLIF T I N NORWAY

passive infillin g o f existin g accommodatio n space, quie t therma l subsidenc e an d sedimen t loading lea d t o a n apparentl y acceleratin g Neogene subsidence . Th e present-da y structur e of the sediments (Fig. 8b) shows truncation of the clinoforms toward s th e margin s a s wel l a s th e inversion structur e an d its marginal troughs . The present-day buria l anomal y alon g th e profil e (Fig. 8c ) varie s fro m O m i n th e basi n centre , where therma l subsidenc e an d sedimen t loadin g and compactio n outpace d th e reductio n i n accommodation spac e produce d b y th e fallin g sea level and sedimentation, to more than 1 km in the inversio n zone . Th e backgroun d valu e o f overburial on the basin margins is caused by the mechanism o f Fig. 7. The abov e principles , includin g tempora l variations i n sedimen t sourc e area s reflectin g the clockwise rotation of source areas throughout Cenozoic time, have been built into a large-scale post-mid-Cretaceous North Se a model. B y trialand-error runs , i n whic h th e initia l therma l anomaly an d th e sedimen t feedin g wa s varied , this model was brought into good agreement with the present-da y thicknes s distributio n of the Upper Cretaceou s an d Cenozoic unit s (Gemme r et al 2002 ; Nielse n 2002) . Here we show two of the principal predictions of this model, which can be compared wit h observations: chal k overburia l (Fig. 9 ) an d th e pre-Quaternar y geologica l ma p (Fig. 10).

53

The regional pattern of chalk overburial in the North Se a Basi n was quantifie d b y Japse n 1998 by comparing measured compressional velocitie s of chalk with a reference depth-velocity model . He foun d tha t th e overburia l increase s toward s the easter n an d wester n Nort h Se a margins i n a symmetrical patter n aroun d a centra l zon e aligned alon g th e Central Graben area , i n which chalk is at its maximum depth at present day. The overburial predictio n o f Fig . 6 reproduce s th e general patter n o f overburia l inferre d b y Japse n (1998). However , th e predicte d amplitude s ar e smaller tha n th e amplitude s o f Japse n (1998) , except i n th e inversion zones , wher e ver y larg e values of overburial can occur also in the model. The predicte d pre-Quaternar y ma p o f Fig. 10 shows marked similaritie s wit h the pre-Quaternary onshor e Denmar k (Sorgenfre i & Berthelse n 1954) an d with the compilation of Japsen (1998). The area l exten t o f th e Oligocen e outcro p o f onshore Denmark , however , i s to o larg e a s a consequence o f erosion t o sea level in the model of al l topograph y in th e basin . The curvatur e of the Cenozoi c laye r boundarie s adjacen t t o th e Sorgenfrei-Tornquist Zon e agree s wit h obser vations and is caused by the filling of the flexural foredeep an d the later erosional truncation. The conclusio n of this sectio n i s that much of the know n post-mid-Cretaceou s sedimentar y structure o f th e Nort h Se a Basi n ca n b e reproduced b y a mode l tha t include s passiv e

Fig. 9 . Predicte d overburia l o f chal k (o f lat e Cretaceou s an d Dania n age) . In th e centra l Nort h Se a thermal subsidence an d compaction of previous sediments keep pace with the reduction of accommodation spac e caused by falling se a level an d sedimentation. Fou r zone s of structural inversion have been included.

54

S. B. NIELSEN ETAL.

Fig. 10 . Predicted pre-Quaternar y geologica l map .

isostatic effect s i n respons e t o therma l cooling , sea-level fall , erosion , an d changin g sedimen t source areas . Th e questio n tha t remain s i s whether th e discrepanc y o f chal k buria l ampli tudes towards the margins is a consequence o f an unknown mechanis m o f activ e surfac e uplift , which is not part of this numerical model , o r can be explaine d b y th e referenc e chal k velocit y model (Huus e 2002a, 2002b) . Active mechanisms Changing in-plane stress. It is well known that fluctuating in-plan e stres s cause s fluctuation s o f relative wate r dept h o n passiv e margin s an d i n intra-cratonic basin s becaus e o f th e induce d vertical deflection s of the lithosphere (Cloeting h etal. 1985 , 1990 , 1992 ; Karner 1989) . However , it i s perhap s les s wel l establishe d whethe r a n increase i n compressio n cause s a depressio n o r an elevatio n o f th e basi n floor . Th e principa l reason for this controversy is that the thin elasti c plate model s applie d i n th e majorit y o f studie s require preloading in order to respond b y flexure to horizonta l stres s changes . Thi s preloadin g produces a n initia l deflection , the amplitud e o f which is enhanced i n compression o r reduced i n extension, a s th e produc t o f th e in-plan e forc e and initial plate curvature acts in the same way as a vertica l load . Th e resul t o f a thi n plat e compression o r extensio n experimen t i s henc e given a priori b y the modeller's preferenc e about initial loading .

An indication of what might happen when the tectonic stres s o n a continenta l margi n change s from extensio n to compression wa s presented by Braun & Beaumon t (1989) , wh o simulate d passive margi n formatio n b y a numerica l thermomechanical model . The y foun d tha t a sudden relie f o f th e extensiona l tectoni c stres s causes basin floor rise and rift flank subsidence a s a ne w balanc e betwee n in-plan e stres s an d buoyancy force s is established. Here w e pursu e thi s a littl e furthe r b y simulating th e formatio n o f a generi c passiv e margin b y lithospheri c extensio n i n Mesozoi c time, and reversal of the stress field at the time of opening o f th e Nort h Atlanti c (Paulse n e t al 2001). W e conside r th e inherite d strai n history, state of stress, and sediment and thermal loading of th e margi n a t the tim e of stres s reversal. Th e results also apply to rifts and intra-continental rift basins such as the Viking Graben and the central North Sea . The numerica l mode l i s a dynamica l large strain Lagrangia n finit e elemen t mode l wit h a n elasto-viscoplastic rheology . Plasticit y i s mod elled usin g non-associate d plasti c flo w wit h a Driicker-Prager yield criterion. Th e viscoelastic deformation i s modelle d b y a non-linea r Maxwell rheolog y wit h viscosit y dependin g o n temperature, materia l type , and deviatoric stress . In general, the type of deformation occurring at a point i n th e mode l depend s o n temperature , deviatoric stress, confining pressur e and material type. Figur e 1 1 show s th e initia l undeforme d

PALEOCENE INITIATION O F CENOZOIC UPLIF T IN NORWAY

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Fig. 11 . Finite elemen t mode l o f passive margin-rift formation . The model includes uppe r and lower crus t and mantle lithosphere. The boundary conditions are 0 °C at the surface and 1300 °C at a fixed depth. There is no flow of hea t acros s th e vertical boundaries . Loadin g b y sedimentatio n an d erosion occu r a t the surface . The botto m boundary condition is a hydrostatic pressure. A kinematic boundary condition causes extension or compression of the profile. Mode l parameter s follo w Paulse n et al (2001) .

Fig. 12 . Cross-section s o f margi n evolutio n wit h positions o f tracke d basemen t point s indicate d b y arrow s a t locations 0 , 50 , 9 0 an d 160km : OMa , initia l sedimen t structure ; 1 5 Ma, en d o f rifting ; 10 5 Ma, onse t o f compression; 15 8 Ma, final sediment structure . The deposition o f syncompressional sediment s has been retarde d between 2 0 and 120km .

56

S. B. NIELSEN ETAL.

finite elemen t gri d an d th e therma l an d mechanical boundar y conditions . Localizatio n of margi n formatio n i s achieve d b y assumin g a weaker (wet ) rheolog y i n th e regio n tha t i s t o become extended . Therma l weakenin g woul d yield simila r results , bu t th e exac t mod e o f weakening i s not the subjec t here . The margi n i s created by extensio n fo r 3 2 Ma by movin g th e righ t mode l boundar y a t a constant velocity t o the right (Fig. 12) . A smaller rate of extension woul d result in more synrif t an d less postrif t therma l subsidence , an d vic e versa . However, th e duratio n o f extensio n i s no t important t o th e principa l results . Afte r exten sion, therma l subsidenc e follow s fo r a period o f 85 Ma. Subsequently , th e stres s fiel d i s reverse d by moving th e right boundary t o the left . Thi s i s equivalent t o applyin g compressiv e stresse s associated wit h th e openin g o f th e Nort h Atlantic.

Figure 13 a shows th e forc e histor y applie d t o the profile . Th e forc e i s obtaine d a s th e depth averaged squar e roo t o f th e secon d invarian t of the deviatoric stres s a t the right model boundary multiplied b y th e lithospher e thickness . Thi s force measur e i s alway s positive , althoug h th e sign o f th e principa l horizonta l tectoni c stres s changes. Th e initia l puls e o f force produce s th e initial extension. Afte r cessatio n o f strainin g the force relaxes by viscous dissipation of deviatoric stresses i n th e ductil e part s o f th e lithosphere . The gentl e reversa l yield s a slow increase i n the force. A t 1 2 Ma a shor t duratio n pulse simulates the effect s o f reorganizatio n o f Nort h Atlanti c spreading. Th e tota l compressiv e forc e end s a t c. 1. 7 X 1012N m" 1 , whic h ca n b e compare d with a n estimate d magnitud e o f c . 3. 9 X 10 12 N m" 1 o f th e ridge-pus h forc e i n a n oceani c lithosphere o f 10 0 Ma ag e (Turcott e & Schuber t 1982).

Fig. 13 . (a) The forc e histor y applie d t o the profile; (b ) subsidenc e histor y without compression; (c ) subsidence history wit h compression; (d ) difference between (c ) and (b) .

PALEOCENE INITIATION O F CENOZOIC UPLIF T IN NORWAY

Figure 13 b an d c show s th e vertica l move ments o f fixed basement point s along th e profil e in th e case s withou t (Fig . 13b ) an d wit h stres s reversal (Fig . 13c) . The location of the basemen t points an d th e evolutio n o f cross-section s ar e shown in Fig. 1 2 at ag e OMa . The initia l margi n formatio n result s i n basi n subsidence a s tracke d b y basi n point s Okm , 300km an d 400km . Poin t 500k m show s th e formation o f th e rif t flan k durin g rifting . Poin t 725km register s th e far-fiel d vertica l movements. Following extensio n ther e i s therma l subsidence o n th e passiv e margin , enhance d b y sediment loading . Fro m compariso n o f Fig. 13 b and c i t become s apparen t tha t thi s genera l picture continue s throug h th e gradua l stres s reversal a t 5 3 Ma. Th e smal l subsidenc e pertur bations produce d b y th e stres s reversa l ar e enhanced i n Fig. 13d , which shows the difference between the subsidence curves of Fig. 13 b and c. The subsidenc e rate s o f point s 300k m an d 400km increas e slightl y i n th e initia l phas e o f stress reversa l becaus e o f th e smal l elasti c volume reduction associated wit h stress reversal and becaus e o f th e continuou s adjustment s to a new equilibriu m betwee n in-plan e force ,

57

buoyancy force s an d evolvin g stresses . Fo r continuing compressio n th e subsidenc e o f th e basin point s becomes retarded as compared with the cas e o f pur e therma l subsidence . Th e rif t flank poin t (500km ) subside s a t a n increasin g rate in the initial phase of compression, an d later becomes slightl y uplifted . The respons e t o th e compressiona l puls e a t 12 Ma depends o n the position alon g th e profile . Generally, ther e is a small elevation o f the basin floor, which , however , i n th e cas e o f poin t 300km i s followe d b y a mino r increas e i n subsidence. A larger compressiona l puls e would produce larger-amplitude deflections of the same shape. Our mode l applie s t o a passiv e margi n wit h laterally smoothl y changin g rheological proper ties. I n reality , th e strai n followin g stres s reversal wil l b e accommodate d b y th e reactivation o f majo r basemen t faults , resultin g in localize d shortenin g an d th e developmen t o f inversion structure s o r dome s suc h a s th e Helland Hanse n Dom e an d Orme n Lang e o n the Norwegia n continenta l margin . Margina l troughs wit h loca l subsidenc e ma y develo p i n connection wit h suc h inversio n zone s (Nielse n & Hanse n 2000) .

Fig. 14. Thermal and topographic response of lithospheric column to 10km intrusion: (a) temperature profiles; (b) topography; (c ) surfac e heat flux.

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Magmatic underplating. I t i s wel l know n tha t magmatic underplatin g provide s a mechanism o f immediate epeirogeni c uplif t (McKenzi e 1984 ; Brodie & Whit e 1994) . Figur e 1 4 show s th e thermal an d topographi c respons e o f a I D lithospheric colum n t o th e introductio n o f a magma a t th e bas e o f th e crust . Th e intrusio n causes immediat e surfac e uplif t followe d b y a minor therma l subsidenc e a s the intrusion cools . The uplift , /i , a s a resul t o f underplatin g o f thickness x * s given b y h = ( 1 — px/pa)x, wher e px is the density of the underplating materia l an d pa is the density o f the asthenosphere. Erosion of the surfac e i s no t considered . Th e hea t flow , initially a t 5 0 mW m~ 2 , increase s i n les s tha n 10Ma t o 6 2 mW m~ 2 . A t 5 5 Ma afte r th e intrusion the surface heat flow is essentially bac k to the initial value . The underplatin g mechanis m produce s uplif t in direct response t o the underplating history. It is followed b y mino r therma l subsidence . Depend ing o n th e densit y contras t betwee n th e under plating magm a an d th e mantle , c . 10k m o f underplating i s require d t o produc e a n initia l uplift o f 1 km. Lithospheric delamination. Delaminatio n o f the lithosphere , o r bas e lithospher e erosion ,

means los s i n some wa y o f a fraction o f the base of th e mantl e lithosphere . Thi s mechanis m ha s been applie d t o explainin g th e sudde n uplift s o f the Colorad o Platea u (Bir d 1979 ) an d th e Himalayas (Housema n et al 1981) . I n thes e cases delamination happen s a s a consequence o f thickening an d destabilizin g o f th e mantl e lithosphere i n collision . I n th e Nort h Atlanti c case delaminatio n mus t b e relate d someho w t o the interactio n o f th e Icelan d plum e wit h th e lithosphere. W e sugges t tha t a gravitationa l (Rayleigh-Taylor) instability (Conrad & Molnar 1997; Housema n & Molna r 1997 ) triggere d b y the arriva l o f th e low-viscosit y an d low-densit y Iceland plum e ma y hav e bee n responsibl e fo r convective remova l o f a fractio n o f th e mantl e lithosphere. Fo r later reference we present here a ID mode l o f th e therma l an d topographi c consequences o f this process . It i s assume d tha t th e lithospher e remain s a t the reduce d thicknes s afte r delamination . An y recovery o f the lithosphere thickness reduces th e topographic respons e o f the partial delamination mechanism. Fo r a ful l recover y o f lithospher e thickness th e topograph y return s t o it s initia l position, unles s perhap s th e delaminate d lithosphere wa s chemicall y buoyan t an d distinct from th e remainin g lithosphere . Th e presen t scenario henc e assume s tha t th e lithospher e i s

Fig. 15 . Therma l an d topographi c respons e o f lithospher e sufferin g 3 5 km o f delamination : (a ) temperatur e profiles (8 5 km plume); (b ) topograph y i n various cases; (c) surfac e hea t flux (85 km plume).

PALEOCENE INITIATION OF CENOZOIC UPLIFT IN NORWAY

delaminated afte r destabilizatio n b y plum e emplacement t o a ne w an d thinne r equilibriu m thickness, which may be true if the lithosphere is initially thicke r tha n th e presen t equilibriu m thickness. Figure 1 5 shows the therma l an d topographi c response t o delaminatio n ove r 5 Ma o f a I D lithospheric column . Air y isostas y i s assumed , and th e initia l topograph y i s zero . A t OM a th e plume arrive s a t th e bas e o f th e lithosphere , initially 16 5 km thick. Over a time spa n of 5 Ma the plum e grow s i n thicknes s t o 4 5 o r 85k m simultaneously with erosion of 35 km of the base of the lithosphere. Other growth times could have been chosen but do not significantly influenc e the principal results . Th e plum e ha s a n exces s temperature of 100°C , which decays after plum e emplacement wit h a tentativ e tim e constan t o f 40 Ma. Complet e modellin g o f th e deca y o f th e excess plume temperatur e involves modelling of whole-mantle convection , whic h i s beyon d th e scope o f thi s paper . Th e case s o f delaminatio n without plume thickening, but in the presence of the basal temperatur e anomaly, and the effec t o f the basa l temperatur e anomal y alon e ar e shown for comparison . The lithospheri c temperatur e profile s (Fig. 15a ) ar e fo r th e cas e o f 85k m o f plume . The tectoni c uplif t historie s (Fig . 15b ) sho w a

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phase o f rapi d initia l uplif t a s plum e materia l replaces dense r asthenospher e an d 3 5 km o f th e lower lithosphere , followe d b y increasin g tec tonic uplif t cause d b y heatin g o f th e remainin g lithosphere t o a ne w therma l equilibriu m situation. Th e magnitud e o f th e initia l uplif t depends strongl y on the plume thickness and the amount o f lithospheri c delamination . However , the plum e contributio n decay s ove r tim e a s th e plume materia l cools . Th e resultin g permanen t tectonic uplif t depend s onl y on the amoun t of lithospheric delamination. The uplift rat e following plum e emplacemen t an d delaminatio n depends o n the amoun t of delaminatio n an d th e initial plum e thickness . Fo r give n delaminatio n thickness th e rat e increase s wit h decreasin g plume thickness. The surface heat flux (Fig. 15c ) shows a minor slow increase. Unlike lithospheric stretching, whic h als o thin s th e lithosphere , delamination cause s n o rapi d squeezin g o f th e isotherms an d therefor e show s a muc h quiete r thermal response . Fo r late r referenc e i t i s note d that the initial, possibly rapid, disturbance of the lithosphere-asthenosphere system s cause s tec tonic surface uplift t o evolve with a time constant of c . 60 Ma. A Paleocene-Eocene delamination hence produce s increasin g cumulativ e tectoni c uplift a t a decreasin g rat e throughou t Cenozoic time.

Fig. 16 . Th e topographi c response of a lithospher e t o delamination . Curv e paramete r i s initia l lithospher e thickness.

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Figure 1 6 summarize s th e abilit y o f th e delamination mechanis m t o produc e permanen t tectonic uplift . Fo r simplicity , th e transien t components cause d b y plum e emplacemen t an d the exces s plum e temperatur e ar e lef t out ; they disappear anyway . Th e uplif t a t th e en d o f delamination represents , therefore , th e pur e initial delaminatio n uplif t fo r th e give n litho sphere thickness. Th e dependence o f initial uplif t on lithospher e thicknes s i s cause d b y th e dependence o f th e lithospheri c geother m o n lithospheric thickness . Fo r a thick lithosphere th e delaminated material has an average temperature that i s close r t o th e asthenospheri c temperatur e than i t i s i n th e cas e o f a thi n lithosphere . Replacing i t with asthenosphere therefore yields less excess buoyancy. The present-day values are the tectoni c surfac e uplif t afte r therma l relax ation fo r 5 0 Ma followin g delamination. At that time, a thi n lithospher e i s almos t i n therma l equilibrium, wherea s a thick lithosphere stil l has not adjuste d t o th e ne w therma l equilibriu m situation. Th e limitin g valu e o f surfac e uplif t after infinit e relaxation tim e depend s linearl y on the delamination thickness . Discussion Following Dor e et al (1999) , the North Atlantic continental break-u p a t 53 Ma was the culmina tion o f a c . 35 0 Ma perio d o f Permo-Triassic , (mainly late ) Jurassic , Early Cretaceous , 'mid' Cretaceous, and latest Cretaceous-Early Eocen e rifting in the area of the Caledonian orogen. After these event s the topograph y of N W Europ e wa s low an d separate d b y shallo w sea s overlyin g deep Palaeozoi c an d Mesozoi c sedimentar y basins. The fina l stag e o f lates t Cretaceous-Earl y Eocene riftin g le d t o break-u p an d becam e associated wit h intense magmatic activity as the Iceland plum e arrive d a t th e bas e o f th e Nort h Atlantic lithospher e a fe w millio n year s befor e break-up (Skogseid et al. 2000). The plume fron t may have extended mor e tha n 1500k m from th e plume centre , encirclin g th e Nort h Atlanti c Volcanic Province, including the British Tertiary Volcanic Province , th e Labrado r Se a volcani c provinces, an d th e Greenland , Scottis h an d Norwegian Caledonides (Skogsei d e t al. 2000) . In th e Britis h Isles , wher e age s o f volcani c rocks i n th e Britis h Tertiar y Volcani c Provinc e and larg e offshor e thicknesse s o f Paleocene an d Eocene sequence s bea r direc t witnes s t o th e timing o f th e uplift , ther e i s a goo d correlatio n between plum e emplacemen t an d clasti c sediment influx , whic h ha s bee n use d t o infe r surface uplif t variation s (Whit e & Lovel l

1997). However , regardin g th e surfac e uplif t of th e Norwegia n (and Scottish ) Caledonides , the direc t rol e o f th e plum e a s a n initiatin g agent a t th e tim e o f emplacemen t ha s no t bee n widely considered . O n th e Norwegia n con tinental shel f an d i n larg e part s o f th e Nort h Sea th e direc t associatio n betwee n sedimen t patterns an d tectoni c activit y i n th e sourc e areas yield s mino r Paleocene-Eocen e tectoni c activity wit h th e majo r phas e startin g i n Oligocene tim e (se e Huus e 2002b ) an d accelerating throug h Neogen e times . Thi s leaves a tim e ga p o f c . 3 0 Ma betwee n plume emplacemen t an d th e majo r phas e o f surface uplift , an d produce s majo r problem s explaining th e mos t recen t uplif t mechanism : why di d th e plume , whic h wa s mos t vigorou s and ho t durin g th e shor t tim e o f emplacement , after whic h i t starte d t o cool , tak e 3 0 Ma t o produce significan t surfac e uplift ? Sediment structure and fission tracks We believe that the direct association o f intensity of sedimentar y respons e wit h contemporaneous tectonic activit y i n th e sourc e area s i s proble matic. Molnar & England (1990) argued that the phenomena commonly used to infer recent uplif t may as well be a consequence of climate change. On th e basi s o f example s fro m th e Alps , Pyrenees and the Rocky Mountains of Colorado, Wyoming an d Utah , the y argue d tha t th e existence o f geomorphologicall y youn g land scapes wit h deeply incised valley s and a jagged relief, whic h usually i s taken as a n indication of recent tectoni c uplift , i s a consequenc e o f a change i n th e proces s o f erosion , fo r exampl e caused by climate change, and in the extreme by glaciers. Th e denudatio n proces s produce s apparent uplif t b y raisin g th e summi t envelope by passive isostatic compensation. The North Sea area di d feel the global climat e change from lat e Eocene int o Oligocene time s and did experience deep glacia l erosion an d a general sea-level fal l (Riis & Fjeldskaar 1992 ; Huus e 2002b). There fore, i n ou r opinion , th e timin g o f a n uplif t history based on sediment ages must be tempered with th e influenc e of climate an d erosional bas e level o n denudation rates. The influenc e o f climat e an d eustas y i s no t known in grea t detail , but the coolin g histor y associated wit h denudatio n ma y b e inferre d directly fro m apatit e fission tracks (AFTs). This technique therefore has been extensively applied in Norway (Rohrman et al. 1995 ) to infer recent cooling an d ha s contribute d t o th e concep t o f accelerating Neogen e uplif t o f th e souther n and northern Norwegia n domes. However, publishe d

PALEOCENE INITIATION O F CENOZOIC UPLIFT IN NORWAY

models fo r annealin g o f fissio n track s sho w a marked lac k of sensitivity to temperatures below 60 °C, althoug h i t i s clea r tha t som e annealin g does occur at lower temperatures ove r geologica l time scales . Whe n modellin g rea l data , thi s can often lea d t o th e spuriou s inferenc e o f recen t cooling fro m aroun d 60 °C (Gallagher & Brown 1999). Th e uncertaint y inheren t i n constraining denudation (an d an y associate d uplift ) become s most acut e i n recently glaciate d areas , o r where only 1 km o r les s o f materia l ma y hav e bee n removed. Thi s i s a proble m i n man y part s o f southern Scandinavia . Thi s recentl y recognize d issue ha s no t bee n addresse d i n AF T studie s of the area and means that the inferred temperature histories d o no t properl y represen t th e tru e cooling histor y produce d b y denudation . Th e recently revive d (U-Th)/H e technique , applie d to apatite , whic h ha s a lowe r temperatur e sensitivity tha n fissio n track s (e.g . Hous e e t al 1998; Wol f e t al . 1998) , ma y provid e th e ke y timing information in this region.

Mechanism of surface uplift In th e absenc e o f direc t constraint s fro m sediment age s an d coolin g histories , th e discussion o f th e timin g o f activ e Cenozoi c surfac e uplift aroun d th e Nort h Atlanti c an d th e simultaneous basi n subsidenc e become s mos t meaningful whe n the requirement of a physically plausible mechanis m i s considered . Compressional stres s is not a plausible agent . Our mode l exampl e show s tha t th e gradua l change from extensional to compressional stress , which th e peri-Nort h Atlanti c region mus t have experienced followin g the break-up, has had only minor influenc e o n th e subsidenc e o f th e basin s and uplif t o f th e lan d areas . Fo r physicall y reasonable stres s levels , stres s reversa l cause s a small enhancemen t o f th e rat e o f therma l subsidence, an d a sudde n compressiona l puls e caused by, for example, spreading reorganization causes a n elevatio n o f th e basi n floor . W e therefore think that the accelerating late Neogene tectonic subsidenc e derive d b y Cloetingh e t al . (1990, 1992 ) for the North Atlantic margins and the North Sea area is not related to compressional pulses from spreadin g reorganization. The spac e needed t o accommodate th e Neogene sediment s is muc h to o larg e t o b e explaine d b y compres sional forces. Rather, we suggest that the derived tectonic subsidence may be a consequence of the use o f to o shallo w palaeo-wate r depth s i n th e tectonic subsidenc e calculation . Palaeo-wate r depths ar e notoriousl y difficul t t o assess , ye t have significan t influenc e o n th e result s o f

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tectonic subsidenc e calculations . I n our opinion, the apparentl y acceleratin g lat e Neogen e sub sidence i s simpl y therefor e passiv e fillin g o f already existing accommodation spac e enhance d by loading-induced subsidence (see Fig. 8a) . We no w tur n ou r attentio n t o th e rol e o f th e Iceland plum e i n explainin g th e Cenozoi c surface uplift . W e adopt the plume emplacemen t scenario o f Skogsei d e t al . (2000) : a flattene d plume head with an excess temperature of 100 °C (density contrast of c. — 11 kg m~ 3 ) arrives a few million years before break-up time at the base of the Nort h Atlanti c lithosphere , whic h show s marked thicknes s variation s followin g th e Palaeozoic-Mesozoic rifting period . As describe d b y Skogsei d e t al. (2000) , area s of relativel y thi n lithospher e becam e associate d with th e thickes t plum e colum n an d therefor e experienced th e larges t initia l isostati c uplift . These ar e als o th e area s tha t ar e pron e t o decompression meltin g a s th e plum e materia l rises higher , wit h th e possibilit y o f magmati c activity i n th e for m o f floo d basalts , magmati c underplating o f th e crust , an d intrusions . Th e Norwegian continenta l shel f an d th e norther n North Sea , wit h a dens e dat a coverage , provid e suitable area s fo r quantificatio n o f th e sudde n isostatic uplif t inflicte d b y plume emplacement . In th e neighbourin g continenta l area s o f thicker lithospher e suc h a s wester n Fennoscan dia, Scotlan d an d Eas t Greenland , plum e emplacement cause d a transien t isostati c uplif t component a s ho t plum e materia l displace d denser asthenosphere . Furthermore , i f partia l delamination o f th e lithospher e too k place , permanent surfac e uplift woul d have resulted. It is therefor e importan t t o conside r th e physica l reason for delamination, and we propose that the Rayleigh-Taylor instabilit y i s a candidat e fo r explaining the process o f partial delamination . Rohrman & van der Bee k (1996 ) applie d this principle t o explainin g th e peri-Nort h Atlanti c domal uplif t pattern . Thei r varian t o f th e Ray leigh-Taylor instabilit y focuse s o n th e generation o f diapir s i n th e buoyan t lowe r layer, whic h ris e int o th e rigi d li d a t selecte d locations an d produc e doma l uplift . Thi s application utilize s localize d upwellin g and is analogous t o th e generatio n o f sal t diapir s o r thunderstorms. Th e timin g o f thei r scenari o requires tha t plum e emplacemen t beneat h th e Norwegian an d Scottis h Caledonide s too k plac e around 40-30Ma , whic h i s i n contras t t o th e early plum e emplacemen t scenari o o f Skogsei d et al. (2000). Our applicatio n o f th e Ra y leigh-Taylor instability t o th e lithosphere-asthenospher e boundary follow s th e original concept.

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Conrad & Molna r (1997 ) an d Housema n & Molnar (1997 ) hav e show n i n a linearize d analysis ho w disturbances a t th e lithospher e asthenosphere boundar y wit h a wavelengt h between 10 0 and 200k m ca n becom e unstable , resulting i n th e bas e o f th e lithospher e bein g swept into drips o r blobs o f descending material . This ultimatel y result s i n thinning of the mantl e lithosphere a s th e dro p detache s an d sink s through th e low-viscosit y asthenosphere , carry ing mantl e lithospher e away . Th e resul t i s localized downwelling , whic h feed s o n litho sphere materia l fro m th e sides . Ther e ar e n o localized asthenospheri c diapirs , bu t a rathe r broad asthenospher e upwellin g t o fil l th e spac e vacated b y th e los t lithosphere . Th e possibl e growth rat e o f the instabilit y ranges fro m 10" 14 s"1 t o higher , correspondin g t o evolutio n time s of 3 Ma or longer. Depending on the timing of the process i t is possible tha t decompression meltin g may take place in the upwelling asthenosphere. It may therefor e b e theoreticall y possibl e tha t magmatic underplatin g als o occurre d unde r th e adjacent lan d areas, explainin g part o f the uplift . Whether thi s has happened ma y be tested b y the teleseismic receive r functio n experiment s tha t are in progress i n souther n Norway. The above-quote d analysi s wa s base d o n continental lithospher e destabilize d b y com pressional thickening. However , w e suggest tha t the arriva l o f a hot , low-viscosit y an d buoyan t plume a t th e bas e o f th e Nort h Atlanti c lithosphere coul d hav e triggere d th e Rayleigh Taylor instabilit y i n location s o f relativel y dee p lithospheric roots or in areas of a steep gradien t of the lithosphere-asthenosphere boundary. Th e northern an d souther n Norwegia n domes , th e Scottish Caledonides an d East Greenland, i n this

model, woul d have undergone particularly severe delamination, perhap s becaus e th e lithospher e under these areas afte r the Palaeozoic-Mesozoic rifting perio d wa s lef t relativel y thick and with a steep gradien t toward s th e basi n areas . A t present, th e lithospher e i s know n t o b e thic k beneath th e Fennoscandia n Shiel d t o th e east . Therefore, a transitio n t o th e norma l thicknes s lithosphere on the Norwegian margin must occur somewhere unde r western Fennoscandia . Assuming tha t delaminatio n i s possible , w e now combine th e possible tectoni c uplif t histor y inflicted b y th e Icelan d plum e o n th e adjacen t land areas with the possibility of increasing rates of denudation induced by the changing Cenozoic climate an d the fallin g erosiona l bas e level . Figure 1 7 show s th e topographi c respons e t o 35 km of lithospheric delaminatio n an d 45 km of plume emplacemen t aroun d th e tim e o f Nort h Atlantic break-up. The initial topography is set to 200m, approximatel y correspondin g t o th e topography tha t woul d emerg e i f th e surfac e of the lithospher e wa s a t se a leve l durin g th e Cretaceous culminatio n of globa l se a level . It i s assumed tha t th e rat e o f denudatio n increase d through Cenozoic time , simulating the effect s o f climate cooling . Change s o f denudatio n rate i n this simplified mode l ar e located a t the EoceneOligocene transition , at mid-Miocen e time s and at the onse t o f glaciations . The heigh t o f th e averag e surfac e increase s rapidly durin g the phas e o f plume emplacemen t and delamination , an d continue s t o increas e a s the remainin g lithospher e i s heated. A s denudation start s t o remov e materia l an d th e rat e o f lithospheric heating decreases, the rate of surface uplift decrease s an d eventuall y starts t o declin e with th e progressio n o f denudation.

Fig. 17 . Topographi c respons e t o 35k m o f delaminatio n an d 45k m o f plum e emplacement , wit h varyin g denudation rat e included .

PALEOCENE INITIATIO N O F CENOZOIC UPLIF T IN NORWAY

The differenc e betwee n th e summi t heigh t envelope (Molna r & Englan d 1990 ) and th e average surfac e represents the average thicknes s of materia l remove d b y denudation . Th e rathe r limited reductio n i n average surfac e heigh t i s caused b y th e larg e isostati c compensatio n o f removed mas s an d th e continuou s tectoni c surface uplif t throug h Cenozoic tim e (Fig. 15b). In reality , erosio n dept h show s larg e lateral variations from almos t uneroded peaks to deeply incised valley s wit h bottom s fa r belo w th e average surfac e height . Th e summi t heigh t envelope therefor e show s the elevatio n a t which uneroded isolate d peak s could , i n principle , occur, if indeed the y exist. The summi t heigh t envelop e (Fig . 17 ) increases throughou t Cenozoic tim e wit h a rate that primarily reflects th e changes in the average denudation rate. In the case of Fig. 17 it reaches a maximum valu e o f c . 1700 m fo r a n averag e surface o f c . 700m . Othe r example s wit h different ratio s o f summi t envelope an d averag e surface heigh t ca n b e produce d t o fi t loca l features o f th e Norwegian , Scottish , o r Eas t Greenland topograph y b y varyin g th e initia l topography an d lithosphere thickness , delamina tion thickness , duratio n o f th e delaminatio n process, an d denudation rates. Gravity The relatively dee p an d narrow Bouguer gravity anomaly (Fig . 1) indicate s tha t topograph y i n western Scandinavi a i s compensate d mainl y a t Moho depth . I n souther n Norwa y i n particular , there i s a goo d correlatio n betwee n averag e surface topography and Moho depth. In northern Norway, topograph y compensatio n b y latera l density variation s o f th e crus t seem s t o pla y a more significan t role . A n outstandin g question, not addresse d i n detail i n this paper, i s how this gravity anomal y i s relate d t o th e possibl e mechanism of land surfac e uplift . The emergenc e o f th e clos e correlatio n between topograph y an d Bougue r gravit y (Balling 1984 ) is mos t easil y understoo d i f magmatic underplatin g a t basa l crusta l level s supports th e topography . However , fo r th e amount o f underplatin g neede d th e absenc e o f dykes an d volcani c rock s o f th e relevan t ag e a t the surfac e is curious. The ne t gravitationa l effec t o f partial delami nation of the lithosphere is a significant negativ e Bouguer anomaly . Superimposin g thi s anomal y on th e gravitationa l fiel d inherite d fro m th e Mesozoic riftin g phas e produce s a negativ e Bouguer fiel d correlate d wit h topography . However, becaus e o f th e dept h o f origin , th e

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anomaly become s lon g wavelength , unlik e th e relatively narro w anomal y observed . A preliminary forwar d calculation o f the gravity respons e shows tha t th e widt h o f th e gravit y anomal y associated wit h delaminatio n i s significantl y reduced whe n th e initiall y doma l topograph y is eroded t o increasin g dept h wit h increasin g distance fro m th e centr e o f th e uplift . Thi s denudation pattern brings the Moho closer to the surface b y isostati c compensatio n an d add s a positive gravity contribution, which narrows the originally wid e gravit y anomal y s o tha t i t becomes consisten t wit h observations . Th e model require s o f th e orde r o f 1.0-2. 0 km o f Cenozoic denudatio n i n th e coasta l area s o f Norway, i n concordanc e wit h the result s of Riis & Fjeldskaar (1992) an d Riis (1996) . Heat flow The therma l effect s associate d wit h magmati c underplating (Fig . 14c ) an d delaminatio n (Fig. 15c ) are insignifican t a t th e presen t day. We therefore should expect the surfac e heat flux in Norway and on the Norwegian passive margin to b e clos e t o th e value s o f a standar d passiv e margin. Thi s i s supporte d b y th e evidenc e presented b y Ballin g (1993) , wh o foun d hea t flux in Norway to be 50-60 mW m~2 , similar to the flu x i n th e res t o f Scandinavia . Th e well defined heat flux-oceani c ag e relationship foun d by Sundvor et al. (2000) only supports the picture of a generall y norma l passiv e margi n therma l regime. Conclusions In this paper w e have outlined a scenario for the Cenozoic surfac e uplif t histor y o f wester n Scandinavia, whic h i s directl y linked t o plum e emplacement an d is consistent with topography, heat flow , crusta l structure , an d th e Bougue r gravity o f western Scandinavia . Th e scenari o may als o be applicable t o East Greenland. The fundamenta l surface uplift mechanis m i s furnished b y partial lithospheric delamination by the Rayleigh-Taylo r instability , whic h wa s triggered b y th e arriva l o f th e Icelan d plum e a t the bas e o f th e Nort h Atlanti c lithosphere . Delamination preferentiall y occur s i n area s o f deep lithospheri c root s o r a t stee p slope s o f th e lithosphere-asthenosphere boundary . Th e dela mination mechanis m allow s fo r a dela y o f surface uplif t accordin g th e finit e tim e scal e fo r evolution o f th e instability . Transien t surfac e uplift continue d afte r cessatio n o f delaminatio n because o f lithospheri c heatin g t o a warme r equilibrium state.

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This scenari o call s fo r th e occurrenc e o f (1 ) significant climat e an d erosiona l base-leve l control o n denudation , fo r th e denudatio n response i n th e basin s t o becom e consisten t with th e observe d sedimen t pattern , an d (2 ) differential Cenozoi c denudation , wit h denuda tion depth s increasin g awa y fro m th e centra l highland areas towards the west coast of Norway, for th e Bougue r gravit y fiel d associated wit h lithospheric delaminatio n t o becom e consisten t with observations . Continuing 8 18O profiling in the PaleoceneOligocene interva l in the Danish sector, receive r function experiment s an d join t apatit e fission track an d (U-Th)/H e analysi s i n souther n Norway, an d numerica l modellin g o f th e geodynamic processe s an d Bougue r gravit y ar e expected t o contribut e t o resolvin g thes e remaining questions.

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A structural transec t betwee n the central Nort h Sea Dome and the South Swedish Dome: Middle Jurassic-Quaternary uplift-subsidence reversal an d exhumation acros s th e eastern Nort h Sea Basin OLE GRAVERSE N Geological Institute, University of Copenhagen, 0ster Voldgade 10, DK-1350 Copenhagen K, Denmark (e-mail: oleg@ geo.geol.ku.dk) Abstract: Th e Jurassic-Cenozoi c structura l evolutio n o f th e easter n Nort h Se a Basi n i s influenced b y the central North Se a Dome, the Danish Megablock , the Tornquist Zone and the South Swedish Dome. The central North Se a Dome is a composite dome comprising the Triple Junctio n Dome , th e Centra l Grabe n Dom e an d th e Frieslan d Dome . Th e Danis h Megablock, newl y recognize d here , i s a first-orde r tectoni c elemen t betwee n th e Centra l Graben and the Tornquist Zone . I n Jurassic-Cretaceous time it was tilted toward s th e east during uplif t o f the Central Graben Dome, whereas the movemen t wa s reversed during th e Cenozoic post-rif t subsidence . Contemporaneou s wit h th e westwar d tiltin g o f th e Danis h Megablock, th e Sout h Swedis h Dom e wa s uplifte d t o th e east . Th e uplift-subsidenc e reversal acros s the eastern North Se a Basin link s th e collapse of the Central Graben Dome and the til t reversa l of the Danis h Megabloc k wit h the uplif t of the Sout h Swedis h Dome . The uplif t followe d b y subsidenc e probabl y involved mas s flo w i n th e asthenospher e t o account fo r th e observe d balanc e betwee n post-rif t subsidenc e an d margina l uplift . Th e model explain s th e uplif t o f bot h th e Sout h Swedis h Dom e an d souther n Englan d a s th e result o f Cenozoic post-rif t subsidenc e o f the Mesozoic Centra l Grabe n Dome.

Uplift an d erosion ca n only be inferred from the rock record , a s materia l i s remove d durin g exhumation. A s a result , studie s o f sedimen t distribution in adjoining basins and geomorphological studie s o f uplifte d area s giv e a n indication o f th e balanc e betwee n uplif t an d subsidence (e.g . Rii s 1996) . Structura l analysi s may hel p t o identif y crusta l movemen t an d processes. Petrographi c an d geophysica l methods (vitrinite reflectance, analysis of density and soni c velocit y trends ) allo w estimate s o f maximum buria l t o b e mad e an d fission-trac k analysis point s t o th e subsidence-uplif t histor y (see Japse n & Chalmer s 2000 , an d reference s therein). A primary objective of this paper is to analyse the Cenozoi c uplif t an d erosio n o f th e easter n North Sea Basin and adjoining parts of the Baltic Shield i n souther n Sweden . Thi s require s analysis o f Mesozoi c riftin g i n th e centra l North Sea , whic h wa s th e precurso r t o th e Cenozoic post-rif t sag basin. The paper presents a structura l analysi s o f first-orde r tectoni c elements, namel y the centra l Nort h Sea Dom e (Ziegler 1990 ; Underbill & Partington 1993), the South Swedis h Dome (Lidmar-Bergstro m 1988 ,

1991, 1993) , an d the Danis h Megablock , whic h is establishe d her e a s a ne w tectoni c elemen t (Fig. 1) . Th e backbon e o f th e pape r i s a geological cross-sectio n tha t combine s a cross section base d o n a regiona l seismi c lin e (RTD 81-22) acros s th e centra l an d easter n Nort h Se a Basin wit h a cross-sectio n acros s souther n Sweden (Fig s 1 and 2) . Fro m th e cross-sectio n and regiona l map s i t i s possibl e t o demonstrat e basin evolutio n fro m initia l Jurassi c rifting , accompanied b y volcanicit y an d uplif t o f th e central Nort h Se a Dome , t o Cenozoi c basi n subsidence an d Neogen e uplif t o f th e Sout h Swedish Dome . Upper-crustal configuration o f the eastern North Sea Basin The presen t structur e o f th e easter n Nort h Se a Basin i s outline d b y th e depth-structur e o f th e uppermost pre-Zechstei n surfac e an d th e con figuration o f overlyin g sedimentar y basin s (Ziegler 1990 ; Vejbae k & Britz e 1994) . Majo r tectonic element s ar e th e Centra l Graben , th e Danish Megablock , th e Tornquis t Zone an d th e

From: DORE , A.G., CARTWRIGHT , J.A. , STOKER , M.S., TURNER , J.P . & WHITE , N . 2002 . Exhumation o f th e North Atlantic Margin: Timing, Mechanisms and Implications for Petroleum Exploration. Geologica l Society, London, Special Publications, 196, 67-83 . 0305-8719/027$ 15.00 © The Geologica l Society of London 2002 .

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Fig. 1 . Jurassic-Quaternar y tectoni c element s of th e easter n Nort h Se a Basi n an d adjoinin g borde r o f th e Fennoscandian High . Line indicates position of cross-section in Fig. 2. The Tripl e Junction Dome , the Centra l Graben Dom e and th e Friesland Dome are dome centres withi n th e composit e centra l Nort h Se a Dome. MFG, Moray Firth Graben.

South Swedis h Dom e (Balti c Shield ) rimmed by the Skagerrak- Kattegat Platform (Fig s 1 and 2). The crystallin e basemen t o f th e Sout h Swedis h Dome extend s westward s int o th e Tornquis t Zone an d th e Danis h Megablock , wher e i t subcrops the Mesozoic and Palaeozoic sediment s (EUGENO-S Workin g Grou p 1988 ; Michelse n & Nielse n 1993 ; Vejbae k & Britz e 1994) . Th e Danish Megabloc k i s thu s separate d fro m th e South Swedis h Dom e b y th e fault-bounde d Tornquist Zone . The post-Palaeozoi c sedimentar y cove r o f the Danish Megabloc k ca n b e divide d int o a Mesozoic serie s thinnin g toward s th e wes t an d an overlying Cenozoic interva l that thins towards the eas t (Fig s 2 an d 3) . Th e oppose d thinnin g directions relat e t o Jurassic-Cretaceou s domal uplift alon g th e Centra l Grabe n an d eastwar d tilting of the Danish Megablock. Thi s movemen t was reversed in Cenozoic time , when a westward tilting brough t th e uppermos t pre-Zechstei n surface bac k dow n t o subhorizonta l level . Th e uplift followe d b y subsidenc e demonstrate s th e impact o f th e centra l Nort h Se a Dom e (Ziegle r 1990) o n th e Danis h Megabloc k tha t occupie d the easter n par t o f th e composit e dome . Th e central Nort h Se a Dome, th e Danish Megabloc k and th e Sout h Swedis h Dom e ar e thu s majo r tectonic constituent s i n th e evolutio n o f th e central an d easter n Nort h Se a Basi n sinc e Jurassic time .

The central North Sea Dome The central North Sea Dome i s a Jurassic arch or composite palaeo-dom e situate d i n th e centra l North Sea. The subaerial part of the dome can be identified b y a regiona l earl y Middl e Jurassi c erosional unconformit y (th e 'Mid-Cimmerian ' unconformity) (Ziegle r 1990 ; Underbil l & Partington 1993) . The dome i s cut by the Central Graben, th e sout h Viking Graben an d the Mora y Firth Basin , meetin g i n a tripl e junction . Th e uplift o f th e Jurassi c dom e wa s initiate d during the Early-Middl e Jurassi c transitio n (lat e Toarcian-early Aalenia n time) , an d th e larges t area uplifte d abov e th e erosiona l bas e wa s apparently reache d durin g Aalenian-earl y Bajocian tim e (Ziegle r 1990 ; Underbil l & Partington 1993 ) persistin g unti l lat e Jurassi c time. The centre of the dome wa s situated above the grabe n tripl e junctio n an d th e Bajocian Bathonian volcani c centr e t o th e nort h (Fig . 4) . Whereas th e Tripl e Junctio n Dome wa s (partly) flooded durin g Lat e Jurassi c tim e (Underbil l & Partington 1993) , Palaeozoi c sediment s an d basement subcro p the base Cretaceous unconformity alon g th e margin s o f th e souther n Centra l Graben an d the Friesland Hig h (Fig. 4). Subcro p patterns o n th e pre-Cretaceou s geologica l ma p can b e use d t o identif y tw o Lat e Jurassi c dom e centres t o th e south : on e alon g th e Danis h Central Grabe n rif t axis , th e Centra l Grabe n

Fig. 2 . Geosection across the eastern North Sea Basin and the South Swedish Dome on the Baltic Shield. The offshore interva l is based on seismic lin e RTD-81-22 adopted from Vejbae k (1997); the onshor e Sout h Swedish Dome i s adopted fro m Lidmar-Bergstro m (1988) . Positio n o f sectio n lin e is indicated i n Fig. 1.

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Fig. 3 . Jurassic-Cenozoi c rotatio n o f th e Danis h Megabloc k i n th e easter n Nort h Se a Basin . Th e Mesozoi c movement wa s dominate d b y greate r subsidenc e i n th e eas t wit h uplif t an d earl y erosio n i n th e west . A t th e Mesozoic-Cenozoic transitio n th e movemen t wa s reversed ; th e bas e Mesozoi c surfac e subside d i n th e west , whereas uplif t an d erosio n dominate d i n the east .

Fig. 4 . Pre-Cretaceou s geologica l ma p o f th e Nort h Se a Basi n (Ziegle r 1990 ) with positio n o f th e Middl e Jurassic-Cretaceous dom e centre s o f th e centra l Nort h Se a Dome . (Se e text fo r furthe r explanation. ) A-B, Altmark-Brandenburg Basin; CG, Central Graben; FB, Farsund Basin; LS, Lower Saxon y Basin; MFC, Moray Firth Graben ; VG, Viking Graben .

UPLIFT-SUBSIDENCE REVERSAL, EASTER N NORT H SE A

Dome, an d one along the axial depression of the Friesland High , th e Friesland Dome (Fig . 4) . On th e basi s o f th e abov e observation s i t i s concluded tha t the central North Sea Dome was a composite dom e forme d b y th e Tripl e Junctio n Dome t o th e north , th e Centra l Grabe n Dome , and the Friesland Dome to the south (Fig. 4). The long axis of the oval-shaped dome s changes with the orientatio n o f th e Centra l Grabe n axis . Th e Triple Junctio n Dom e an d th e Frieslan d Dom e trend NW-SE, and are connected by the Central Graben Dome , whic h show s a n overal l north south tren d wit h th e Danis h Megabloc k o n it s eastern flank . Th e ris e o f th e composit e dom e was initiate d t o th e nort h i n th e tripl e junctio n area (earl y Middl e Jurassi c time) . Th e zon e o f uplift the n graduall y move d southwar d t o th e Central Graben Dome (earl y Late Jurassic time) and the Friesland Dome (late Late Jurassic time ) whereas th e Tripl e Junctio n Dom e t o th e nort h suffered deflation . The Danish Megablock The Danis h Megabloc k i s a crusta l bloc k extending fro m th e Centra l Grabe n i n th e wes t to the Tornquist Zone i n the east (Fig s 1 and 2). To the nort h i t is bounded b y the Farsund Basi n (Fig. 4) . Th e souther n boundary i s les s distinc t and it s position i s suggeste d toward s th e Lowe r Saxony an d Altmark-Brandenburg basins. During earl y Palaeozoi c time , th e are a tha t became th e Danis h Megabloc k forme d a n integral par t o f th e Baltic a Platfor m (EUGENO-S Workin g Grou p 1988 ; Berthelse n et al 1992) . Decouplin g o f th e Danis h Mega block fro m th e Baltic a Platform-Balti c Shiel d took place i n Triassic time . In Middle JurassicCretaceous time , the Danis h Megablock forme d the eastern block of the Central Graben Dome as part o f th e centra l Nort h Se a dom e comple x (Fig. 1) . The Jurassic-Cenozoi c evolutio n o f th e Danish Megabloc k document s th e evolutio n t o the eas t of the rif t axi s of the Centra l Grabe n Dome. Th e mai n trend s o f th e uplif t o r subsidence ca n b e evaluate d fro m regiona l erosional unconformitie s an d isopac h thick nesses. Analysi s o f th e Danis h Megabloc k ha s also included the pre-Cretaceous geologica l map (Fig. 4), as well as the Berriasian palaeotectoni c map an d isopac h map s o f Triassic , Lowe r Jurassic, an d Lowe r an d Uppe r Cretaceou s rocks (Ziegle r 1990) . The restored interval s (b y simple horizon flattening) illustrate the uplift and erosion of the Triassic deposit s along the Central Graben rift , an d sho w that the Jurassi c intervals only partl y covere d th e Centra l Grabe n Dom e

71

(Fig. 2) . Th e Centra l Grabe n Dom e wa s gradually floode d durin g Earl y Cretaceou s tim e and remained submerged during Late Cretaceou s time (Ziegle r 1990) . Th e westwar d thinnin g o f the Cretaceou s deposit s (Fig . 2 ) indicate s continued eastwar d tiltin g o f th e Danis h Mega block durin g Cretaceou s time . A schemati c model o f th e rotatio n o f th e Danis h Megabloc k is illustrated in Fig. 3 . In Cenozoic time the subsidence of the Danish Megablock shifte d t o the west along the Central Graben post-rif t therma l sa g basin (Fig . 2) . Th e onset o f th e Cenozoi c perio d thu s mark s a til t reversal o f th e Danis h Megabloc k toward s th e west, i n contras t t o th e eastwar d tiltin g i n Mesozoic tim e (Fig . 3) . Progressively olde r preQuaternary sediment s ar e encountere d toward s the N E unti l th e Precambria n crystallin e base ment o f th e Balti c Shiel d i s reache d (Jense n & Michelsen 1992 ; Japsen 1998). This suggests that the regional uplif t an d erosion alon g the easter n margin o f th e Nort h Se a Basi n ma y hav e corresponded in par t to the subsidenc e of the central North Sea . The South Swedish Dome The South Swedish Dom e i s a broad, dome-lik e structure identifie d b y th e topograph y o f th e present-day Precambria n surfac e i n souther n Sweden (Lidmar-Bergstro m 1988 , 1991 , 1993 , 1999). Th e highes t elevatio n o f c . 400m abov e sea leve l i s encountere d aroun d th e souther n Vattern Graben. Lidmar-Bergstrom (1988, 1991 , 1993) ha s distinguishe d three group s of palaeo surfaces: a sub-Cambria n peneplain , a pre Cretaceous surfac e o f suppose d Permo-Triassi c age that has 'sub-Mesozoi c hill y relief, an d the Paleogene Sout h Smalan d Peneplai n (Fig . 5a) . Segments of the sub-Cambrian peneplain and the South Smalan d Peneplai n radiat e ou t fro m th e centre o f the dom e an d ar e cu t ou t b y th e mor e steeply dippin g pre-Cretaceou s surfac e encoun tered alon g th e margin s (Fig . 5 a an d d) . A n exception to this general pattern is formed by an oval-shaped, block-faulted depression situated to the S E o f th e Vatter n Graben . Th e structur e is interpreted a s a cresta l collaps e structur e (Fig. 5a) . Lowe r Palaeozoi c cove r rock s ar e found i n a continuou s bel t t o th e eas t an d a s outliers i n th e centra l an d norther n par t o f th e dome. Th e steepe r inclinatio n o f th e pre Cretaceous surfac e relativ e to the sub-Cambrian peneplain indicate s tha t uplif t o f th e Sout h Swedish Dome postdates the development of the pre-Cretaceous surface . I t i s believe d tha t th e Late Cretaceous transgression may have resulted in a cover of chalk being deposited over southern

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Fig. 5 . The South Swedish Dome , (a) Morphological an d structural elements based on Lidmar-Bergstrom (1994). Position o f cross-sectio n show n i n (c ) i s indicated . SSP , South Smalan d Peneplain ; VG , Vatter n Graben . (b ) Structure of the pre-Cretaceous surfac e (reconstruction) . Elevation o f structure contours above present sea level, (c) East-wes t cross-sectio n o f th e Sout h Swedis h Dome . Reconstructio n o f th e block-faulte d sub-Cambria n peneplain cu t out by the pre-Cretaceous surfac e to the west. (Interpretation based on data from Lidmar-Bergstro m 1988). (d ) Interpretatio n o f th e pre-Cretaceou s surfac e befor e exhumatio n o f th e sub-Cambria n peneplain , (e ) Reconstruction o f the pre-Cretaceous surfac e and the sub-Cambrian peneplain before uplift of the South Swedish Dome.

UPLIFT-SUBSIDENCE REVERSAL, EASTERN NORTH SEA

Sweden befor e th e uplif t o f th e dom e (Norlin g 1994). The sub-Cretaceou s hill y relie f alon g th e dome margi n i s eroded dow n int o th e basemen t below th e sub-Cambria n peneplai n an d ha s a kaolinitic surfac e cause d b y dee p Mesozoi c weathering. B y contrast , th e sub-Cambria n peneplain i n th e centr e o f th e dom e doe s no t have thes e features . Thi s indicate s tha t th e basement i n th e centr e o f th e dom e mus t hav e been covere d b y Palaeozoi c sediment s wit h a thickness exceedin g th e maximu m weatherin g depth o f 20 0 m durin g Mesozoi c time (Lidmar Bergstrom 1993 , 1995) . Extrapolatio n o f th e more steepl y incline d pre-Cretaceou s surfac e across th e dome , i n combinatio n wit h th e weathering dept h an d th e thicknes s o f th e Lower Palaeozoi c cove r t o th e east , suggest s that th e pre-Cretaceou s surfac e i n th e centr e o f the dom e ma y hav e bee n a t leas t 100-200 m above th e sub-Cambria n peneplai n befor e th e post-Cretaceous erosio n (Fig . 5d). This valu e is in agreemen t wit h the altitud e o f the Palaeozoi c outliers i n th e wester n par t o f th e dome , whic h rise u p t o 200 m abov e th e sub-Cambria n peneplain. Th e uplif t o f th e pre-Cretaceou s surface a t th e centr e o f th e Sout h Swedis h Dome ma y thus have been o f the order o f 500 -

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600m abov e presen t se a leve l judge d fro m present-day elevatio n an d estimate d erosion . A reconstruction o f th e uppermos t pre-Cretaceou s surface o f th e Sout h Swedis h Dom e befor e erosion i s illustrated in both map view and crosssection (Fig . 5b and d). A restoration o f the preCretaceous surfac e befor e domin g i s shown i n Fig. 5e. The occurrenc e o f gravell y saprolite s i n th e Precambrian crystallin e basemen t i s confined to the cresta l collaps e structur e i n the easter n par t of th e Sout h Swedish Dom e (Lidmar-Bergstro m 1997; Lidmar-Bergstro m e t al 1997) . Th e development o f thes e saprolite s i s ascribe d t o weathering i n a coo l t o col d climat e i n Plio Pleistocene tim e (Lidmar-Bergstro m e t al . 1997, 1999) . Thi s illustrate s lat e uplif t an d fracturing o f th e dom e accompanie d b y renewed weathering , par t o f th e 720- 3 Ma denudation episod e o f Lidmar-Bergstro m (1996), resultin g i n exhumatio n o f th e sub Cambrian peneplai n an d th e sub-Cretaceou s hilly relief . Th e propose d timin g i s i n accordance wit h fission-trac k analysis , whic h indicates a lat e uplif t an d exhumatio n durin g the las t 2 5 Ma (Cederbo m e t a l 2000) . In the marine platform area in the Hano Bay to the sout h o f th e dome , th e thicknes s o f th e

Fig. 6. Neogene uplift o f the eastern North Sea Basin and the South Swedish Dome. Uplift in the Danish Basin is based on sonic velocities and vitrinite reflection (Jensen & Michelsen 1992; Japsen 1998; Graversen 1999). Uplift of the South Swedish Dome is based on uplift of the pre-Cretaceous surface modelled in this paper (Fig. 5b) added to c. 1 km uplift established at the eastern border of the Danish Basin and AFTA from Ivo (Cederbom 2002). (See text for discussion and references.)

Fig. 7 . Uplift-subsidenc e reversal across th e eastern Nort h Se a Basi n an d the Sout h Swedis h Dome . Th e Centra l Grabe n Dom e wa s uplifte d an d erode d durin g Jurassic time. Th e dom e wa s raise d abov e se a leve l t o a maximu m i n Lat e Jurassi c time an d wa s the n graduall y flooded durin g Early Cretaceou s time . Th e Cenozoi c post-rif t subsidence wa s centred abov e th e Central Grabe n wherea s th e Sout h Swedis h Dom e wa s uplifte d an d erode d t o th e east .

UPLIFT-SUBSIDENCE REVERSAL, EASTER N NORT H SE A

Cretaceous sediment s range s u p t o c . 650 m (Norling & Bergstro m 1987 ) (Fig . 6) . Thi s should b e regarde d a s a maximu m value , a s supposed Paleogen e (Paleocene? ) sediment s ar e encountered abov e (Kumpa s 1980 ; Norlin g & Bergstrom 1987 ; Svirido v e t al. 1995) . Also , erratics o f Paleocene-Eocene marine sediment s have bee n foun d i n easter n Scani a (Norlin g & Bergstrom 1987) , an d a marin e diatomit e o f Eocene ag e ha s bee n reporte d fro m Finlan d (Fenner 1988) , suggestin g tha t Paleocene Eocene sediment s wer e als o deposite d i n southern Sweden . A n estimat e o f th e thicknes s may b e obtaine d fro m th e Viborg- 1 an d Linde 1 well s i n th e Danis h Basi n t o th e west , wher e 200m an d 250m, respectively , o f Paleocen e Eocene sediment s hav e bee n drille d (Dinese n et a l 1977 ; Heilmann-Clause n 1995 ) (Fig . 6) . The Cretaceou s an d Paleogen e sediment s ad d up t o a total thicknes s o f u p t o 90 0 m tha t ma y have bee n deposite d abov e th e pre-Cretaceou s surface. A n evaluatio n o f th e uplif t shoul d als o take pre-compactio n thicknesse s an d wate r depth int o accoun t an d thi s ma y ad d anothe r 100-200 m. A Cenozoi c uplif t u p t o abou t 1 km ma y therefor e b e adde d t o th e pre Cretaceous surfac e o f th e Sout h Swedis h Dome. Thi s valu e shoul d b e regarde d a s a n absolute maximum , an d a realisti c figur e ma y be considerabl y less . Rii s (1996 ) estimate d Plio-Pleistocene erosio n i n southwester n Swe den o f th e orde r o f 500-1000m . Analysis o f reflectance data fro m th e Jurassic Fjerritslev Formatio n (Jense n & Michelsen 1992 ) and soni c velocitie s fro m th e Fjerritsle v For mation and the Cretaceous Chalk Group from the Danish Basi n an d th e Skagerrak-Kattega t Platform indicate s a n uplift o f the orde r o f 1 km or more (Japsen 1992, 1998; Jensen & Michelsen 1992; Michelse n & Nielse n 1993 ; Japse n & Bidstrup 1999) . Apatit e fission-trac k analysi s (AFTA) indicate s tha t Cretaceous-Paleogen e sediments amountin g to c. 650 m (Goteborg) and c. 1 km (Ivo) have been eroded fro m th e western and southern border o f the South Swedish Dome (Cederbom 2002) . On th e basi s o f th e structura l analysis , th e uplift i n th e centr e o f th e dom e shoul d b e expected t o excee d th e borde r uplif t b y 500-600 m. Although there are some differences between th e value s obtaine d fro m th e variou s methods, th e presen t morphologica l dom e i n southern Swede n rise s t o nearl y 400m an d thi s should b e regarde d a s a minimu m uplif t o f th e dome centr e abov e th e margina l areas . Th e uplift o f th e dom e centr e i n relation th e presen t sea level is estimated to be 500-600 m (Fig. 5b), and th e tota l uplift , whe n addin g th e missin g

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section alon g th e borde r o f th e dome , ma y ad d another 1000 m when the sonic an d fission-track analyses ar e take n int o account . Figur e 6 combines estimate s o f th e uplif t (missin g section) fro m soni c velocit y analysi s an d AFTA wit h th e reconstructe d structur e o f th e pre-Cretaceous surfac e o f th e Sout h Swedis h Dome (Fig . 5b) . Jurassic -Cenozoic uplift - subsidence reversal across the eastern North Sea Basin The structural evolution o f the eastern North Sea Basin, i.e . th e Danis h secto r an d souther n Sweden, i s dominate d b y th e Middl e Jurassic Cretaceous uplif t o f th e Centra l Grabe n Dom e followed by Cenozoic subsidence i n the west and uplift o f th e Sout h Swedis h Dom e i n th e east . This i s summarized in Fig. 7 . Mesozoic structure and evolution of the Central Graben Dome In Early-Middl e Jurassi c time , riftin g an d subsidence was concentrated along the Tornquist Zone in the east, whereas only minor subsidenc e and depositio n too k plac e i n the Danish Centra l Graben (Fig . 7) . Th e emergen t Centra l Grabe n Dome expande d i n Lat e Jurassi c tim e an d th e main depocentr e shifte d t o th e Centra l Grabe n whereas subsidenc e diminishe d i n the Tornquist Zone. Renewe d differentia l subsidenc e too k place i n the Tornquist Zone i n Early Cretaceou s time; th e Centra l Grabe n Dom e wa s graduall y flooded, an d subsidenc e i n th e Centra l Grabe n slowed down . Throug h Lat e Cretaceou s time , subsidence an d cresta l collaps e alon g wit h structural inversio n o f th e grabe n sediment s characterized th e Centra l Grabe n Dome . East ward tiltin g o f th e Danis h Megabloc k wa s stil l active an d structura l inversion wa s als o encountered t o th e eas t i n th e Tornquis t Zone . I n Jurassic-Cretaceous time, th e intensit y of basin subsidence shifte d severa l time s acros s th e Danish Megabloc k betwee n th e Centra l Grabe n and th e Tornquis t Zone . Thi s ma y indicat e a n active function o f the Danis h Megablock durin g tilting tha t wa s possibl y relate d t o a n episodi c rise o f th e Centra l Grabe n Dome . I n Mesozoi c time the Danish Megablock an d the Baltic Shield were decouple d alon g th e Tornquis t faul t zone . During th e Cenozoi c subsidence , however , th e Tornquist Zon e i n Skagerra k an d th e norther n part o f Kattega t wa s n o longe r active , an d th e Danish Megabloc k an d th e Balti c Shiel d ma y have acted together a s a single mega-unit.

76

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Fig. 8 . Pre-Quaternar y geolog y o f th e centra l an d souther n Nort h Se a Basi n an d margina l areas . Ther e i s a conspicuous symmetr y along the Cenozoic basi n axis situated above the Mesozoic Centra l and Viking grabens. Compiled fro m Ziegle r (1990 ) an d Japsen (1998) . S P FZ, Sol e Pit faul t zone ; TZ, Tornquis t Zone . Cenozoic structure and evolution of the North Sea Basin The Cenozoi c post-rif t sediment s for m a gentl e syncline wit h th e basi n axi s situate d abov e th e aborted Mesozoi c rift s (Fig . 8). Th e geologica l map o f th e Uppe r Cretaceous-Neogen e sequences exhibit s a symmetrica l patter n centred aroun d th e Neogen e interva l alon g the basi n axis . Pre-Mesozoi c basemen t crop s out alon g th e basi n margin s i n Norwa y an d Sweden, an d i n th e Britis h Isles . Th e basi n margins wer e uplifte d an d erode d wherea s th e basin centr e continue d t o subsid e (Japse n 1998).

In th e centra l Nort h Sea , Middle Miocen e t o Recent sediment s represen t abou t hal f o f th e Cenozoic sedimen t infil l (Nielse n e t al 1986 ; Michelsen e t al . 1998) , an d stron g tectoni c subsidence i s observe d i n Mid-Lat e Miocen e and Quaternar y time s (Koo i e t al . 1991 ; Clausen e t al . 1999) . Accumulatio n i n th e basin centr e (Fig . 9) i s base d o n lithologica l thicknesses i n well s (Nielse n & Japse n 1991) and isopac h map s (Nielse n et al. 1986 ) along the modelle d seismi c profil e (RTD-81-22 , Fig. 2) ; th e genera l tren d demonstrate s increasing accumulatio n an d sedimentatio n rates throug h Cenozoic tim e wit h a culmination in Pliocene-Quaternar y time . Th e Cenozoi c

UPLIFT-SUBSIDENCE REVERSAL, EASTER N NORT H SE A

77

Fig. 9 . Cenozoi c accumulation , erosio n an d averag e sedimentatio n rate s i n th e centra l Nort h Se a Basin , th e Skagerrak-Kattegat Platform, th e East Midlands Shel f an d southern Englan d (se e Fig. 8) . Hatched are a indicate s average Tertiar y sedimentatio n rate s i n the Danish an d Norwegian Centra l Graben . Erosio n i n southern Englan d assumes gradua l Paleogen e uplift. Compile d fro m Nielse n e t al (1986) , Nielsen & Japsen 1991 , Japsen (1997 ) and Japsen & Bidstrup (1999).

erosion (missin g overburden ) o f th e basi n flanks befor e Quaternar y tim e ha s bee n modelled b y Japse n (1997 , 1998 , 2000 ) an d Japsen & Bidstru p (1999) . Erosio n increase d towards th e basi n flanks , wher e i t exceed s 1 km, wherea s maximu m buria l o f th e Chal k Group stil l occur s i n th e centr e o f th e basin . I n the Sout h Swedis h Dom e an d onshor e area s i n southern England , th e tota l uplif t amount s t o 1.5-2km (Fig s 6 an d 9 ) (Japse n 1997) . The timin g o f erosio n ha s bee n modelle d b y Japsen (1997 ) an d Japsen & Bidstrup (1999). A t the S W borde r o f th e basin , maximu m buria l occurred o n th e Eas t Midland s Shel f i n mid Miocene tim e (Fig . 9 ) (Japse n 1997) . I n th e Danish area , th e erosio n wa s o f Miocen e Pliocene age , an d i n th e Skagerrak-Kattega t Platform o n th e N E borde r o f th e basin , th e erosion i s possibl y no t muc h earlie r tha n lat e Miocene tim e (Japsen & Bidstrup 1999) (Fig . 9). In the South Swedish Dome , th e late NeogeneQuaternary gravell y saprolite s ar e weatherin g

products o f th e crystallin e basemen t forme d during lat e uplif t an d exhumatio n i n a col d climate (Lidmar-Bergstro m e t al . 1997) . I n southern Englan d tiltin g an d uplif t i n Plio Pleistocene tim e i s recorde d (Watt s e t al . 2000; Westawa y e t a l 2002) . Lat e NeogeneQuaternary subsidenc e i n th e centra l Nort h Sea i s thu s contemporaneou s with , an d counterbalanced by , uplift o f the Sout h Swedis h Dome an d southern England (Fig. 9). Later uplif t an d erosio n ha s remove d direc t evidence fo r th e uplif t o f th e Balti c Shiel d an d evolution of the Sout h Swedish Dome. Mappin g of th e Cenozoi c sequence s i n th e easter n Nort h Sea Basi n indicate s tha t sediment s fro m th e Baltic Shield were routed into the basin from th e north and NE in Paleogene tim e (Danielsen et al. 1997; Michelse n e t a l 1998 ; Clause n e t a l 2000). I n Neogene time, the main transport route was fro m th e NE i n Miocene tim e an d fro m th e east i n Pliocen e time (S0rense n e t a l 1997 ; Clausen e t a l 1999) . Th e clockwis e rotatio n of

78

O. GRAVERSE N

the infil l direction s durin g Paleogene-Neogen e time i s believed t o indicat e tha t Cenozoi c uplif t of th e Balti c Shiel d wa s initiate d i n souther n Norway during Paleogene time and then changed to th e Sout h Swedis h Dom e i n Neogen e tim e (Riis 1996 ; Japse n 2000) . Lithosphere models of the Jurassic Cenozoic North Sea Basin The Mesozoic central Nort h Se a domes wer e rif t dominated, accompanie d b y volcani c activit y and underlai n b y a thin crus t (Ziegle r 1990) . In contrast, th e souther n Scandinavia n Cenozoi c domes are non-volcanic, developed alon g the SW rifted margi n o f th e uplifte d Precambrian Palaeozoic crystallin e basemen t wher e crusta l thickness reache s 50k m (Thyb o 1997 , 2000) . During Mesozoi c time , th e Centra l Grabe n Dome an d th e Danis h Megabloc k wer e decoupled fro m th e Balti c Shield-Platfor m along th e Tornquis t Zon e (Fig . 7). Afte r th e Late Cretaceous-Dania n inversio n differentia l subsidence-uplift i n the northern Tornquist Zone stopped, an d th e Cenozoi c subsidenc e o f th e Central Graben Dome was accompanied by uplif t of the South Swedish Dome. Thi s may indicate a deep-seated subcrusta l reorganizatio n an d a possible mutua l relationship between subsidenc e and uplift, a s both th e Central Grabe n Dom e an d the Sout h Swedis h Dom e ar e interprete d a s crustal structure s involvin g th e mantl e litho sphere and/o r th e asthenosphere . The Jurassic-Cretaceou s uplif t o f the Centra l Graben Dom e and the Cenozoic subsidenc e bot h amount t o c . 2k m alon g th e analyse d sectio n (Fig. 2) . Thi s valu e i s simila r t o th e propose d 1.5km uplif t o f th e Sout h Swedis h Dome . I t i s therefore suggeste d tha t ther e ma y b e a lin k between th e subsidenc e o f th e Centra l Grabe n Dome an d the uplift o f the South Swedish Dom e to the ENE . Models of synrift and post-rift basin evolution Existing model s o f synrif t an d post-rif t basi n evolution ar e base d o n th e McKenzi e (1978 ) concept o n sedimentar y basi n formation . Th e structure an d stratigraph y o f th e post-rif t basi n depends o n th e rheologica l propertie s o f th e lithosphere. Watt s e t al (1982 ) modelle d th e shape an d syntheti c stratigraph y o f post-rif t basins develope d o n eithe r a n elasti c o r a viscoelastic plat e overlyin g a wea k substratum . In the elastic model , th e post-rift basin develop s as a transgressiv e basin , wit h a so-calle d steershead geometr y i n cross-sectiona l view . By

contrast, i n th e cas e o f a viscoelasti c plate , th e post-rift basi n develop s a s a regressiv e basi n (Watts e t a l 1982) . Althoug h th e modelle d maximum subsidence is of the same order in both settings, the viscoelastic mode l result s in a postrift subsidenc e c . 20 % greate r tha n th e elasti c model. Th e effectiv e elasti c thicknes s o f th e lithosphere determines whether the sediments are locally supported , i.e . short-wavelength , Airy type isostati c compensatio n o f a viscoelasti c plate, o r supporte d b y longer-distance , latera l strength of the plate, i.e. long-wavelength, elastic compensation (Watts etal 1982 ; Barton & Wood 1984). Cloetingh e t a l (1985 ) an d Cloeting h (1986 ) introduced intraplat e stresse s a s a tectoni c mechanism i n continenta l basi n development . The mode l superimpose s fluctuatin g latera l stresses o n existin g basins situate d on a n elastic lithosphere a s a mechanism to explai n observe d vertical plat e movements . Th e intraplat e stres s model ma y b e viewe d a s a lat e short-ter m tectonic overprin t o n th e long-ter m McKenzi e (1978) concep t o f basi n evolutio n an d litho spheric flexur e modelle d b y Watt s e t a l (1982 ) (Kooi & Cloetingh 1989 ; Cloeting h e t al 1990 ; Cloetingh & Kooi 1992) . Previous models of Mesozoic-Cenozoic evolution of the North Sea Basin Previous models of synrift an d post-rift evolution of th e Nort h Se a are a wer e base d o n th e McKenzie (1978 ) concep t o f lithospher e exten sion applie d to the Mesozoic Centra l Graben rif t basin, supersede d b y th e Cenozoi c post-rif t thermal sa g basi n (Sclate r & Christi e 1980 ; Watts e t a l 1982 ; Barton & Woo d 1984 ; Cloetingh e t al 1985) . Bot h Watt s et al (1982 ) and, fo r example , Cloeting h e t a l (1990 ) hav e modelled th e post-rif t Nort h Se a Basi n t o hav e developed o n a n elasti c plat e exhibitin g a steershead geometr y fo r th e post-rif t sediments . Watts et al (1982 ) estimated the effective elasti c thickness a s th e dept h t o the 450° C isotherm ; i n the case of the North Sea Basin this is of the order of 30-35 km. On the basis o f gravity modelling, however, Barto n & Woo d (1984 ) arrive d a t th e conclusion that th e effectiv e elasti c thicknes s o f the lithospher e i s only o f th e orde r o f 5 km, an d that th e sedimentar y loa d wa s compensate d b y local, Airy-type isostasy, i.e. a viscoelastic plate. Intraplate stresse s hav e als o bee n propose d t o be involve d i n post-rif t basi n subsidenc e an d uplift o f th e margins . Cloeting h e t a l (1990 , 1992), Ziegle r (1990) , Koo i e t a l (1991) , Cloetingh & Koo i (1992 ) an d Dor e (1992 )

UPLIFT-SUBSIDENCE REVERSAL, EASTER N NORT H SEA

described th e gentle warping of the Base Tertiary surface during Cenozoic subsidenc e o f the North Sea Basin alon g wit h uplift o f southern Norway and th e Britis h Isle s a s initiate d b y sedimen t loading couple d wit h uplif t an d erosio n o f th e basin margins ; th e vertica l deflectio n wa s interpreted t o have bee n amplifie d b y intraplat e compression. The geometr y o f th e Cenozoi c post-rif t sediments ma y giv e th e impressio n tha t i t possess a steershea d geometr y (Fig . 2) . How ever, th e thic k interva l abov e th e Centra l Graben i s du e t o a combinatio n o f post-rif t compaction o f th e synrif t sediment s i n th e underlying graben , an d th e thic k Plio-Pleisto cene sectio n i n th e centr e o f th e Nort h Se a Basin (Japse n 1998) ; th e Tertiar y "steershorn " extending t o th e EN E (Fig . 2 ) i s mainl y related to erosio n durin g uplif t o f th e Sout h Swedis h Dome. Furthermore , i n th e cas e o f a steershea d geometry, th e syntheti c stratigraph y shoul d exhibit a n onlapping , transgressiv e basin , whereas observation s outlin e a n overal l regres sive basin-fil l progradin g fro m th e basi n margins (S0rense n e t al 1997 ; Michelse n et a l 1998 ; Clause n e t al 1999) . The Cenozoi c subsidence rat e i n th e centr e o f th e souther n North Se a Basi n increase d throughou t Neogene time (Fig . 9) (Barton & Wood 1984 ; Koo i et a l 1991). Thi s observatio n contrast s wit h mos t models o f post-rif t therma l subsidenc e tha t imply a decreasin g subsidenc e wit h age .

New model: Cenozoic subsidence linked with uplift of marginal domes The collaps e o f th e Centra l Grabe n Dom e caused reversa l o f th e til t directio n o f th e Danish Megablock . Wherea s th e Mesozoi c period witnesse d a n increasin g til t toward s th e ENE, th e til t wa s reverse d i n Cenozoi c time : the uppermos t pre-Zechstein-lowermos t Trias sic leve l alon g th e analyse d sectio n i s no w almost bac k t o subhorizonta l (Fig s 2 an d 3) . During Mesozoi c tilting , th e Danis h Mega block wa s disconnecte d wit h th e Balti c Shiel d along th e Tornquis t Zon e (Fig . 7) . I n Cenozoi c time, however , th e discontinue d faultin g alon g the Tornquis t Zon e indicate s tha t th e mega block an d th e shiel d acte d togethe r a s a singl e unit. Both pre-Quaternary geolog y an d the amount of erosio n outlin e a symmetrica l arrangemen t across the souther n North Sea area (Fig . 8) . The geometry illustrate s Cenozoi c subsidenc e alon g the basi n axi s an d uplif t o f th e margins .

79

Subsidence an d depositio n i n th e centr e (2km ) was counterbalanced b y uplift an d erosion in the South Swedis h Dom e an d souther n Englan d (1.5-2km) (Fig . 9) . Deposition alon g th e basin axis increase d ove r time , an d th e curv e o f deposition is mirrored by the curve of uplift. Th e Skagerrak-Kattegat Platform an d the East Mid lands Shel f alon g th e easter n an d wester n margins o f the Nort h Se a Basin both underwent uplift and/o r erosio n durin g Miocene-Quaternary time ; th e gradient s and magnitude s of uplif t correspond t o depositio n i n th e basi n centr e (Fig. 9) . Th e observatio n tha t depositio n wa s counterbalanced b y (pene)-contemporaneou s uplift and/o r erosio n i s interprete d t o indicat e that Cenozoi c post-rif t subsidenc e wa s linke d with marginal uplift . In a cas e wher e post-rif t therma l subsidenc e has induce d a flexura l bulg e i n margina l areas , uplift woul d equal only a small fraction (6-7% ) of the subsidence (Turcotte & Schubert 1982) . If this was the situation, uplift of the southern North Sea margins should only amount to c. 150m . In addition, Cenozoic post-rif t subsidenc e increased (Fig. 9) , wherea s therma l modellin g invoke s decreasing subsidenc e ove r tim e (McKenzi e 1978; Watt s e t a l 1982) . Thes e observation s indicate that thermal subsidence alone is not able to accoun t fo r th e observe d balanc e betwee n subsidence an d uplif t (Fig . 9) . I t i s suggested instead that the subsidence o f the Central Grabe n Dome, accompanie d b y th e backward tilt of the Danish Megablock , wa s associate d wit h a reorganized flo w patter n i n th e asthenospher e below the subsiding North Sea Basin (Fig. 10). In this case , mas s movements directed toward s the marginal dome s ma y accoun t fo r th e observe d uplift of the South Swedish Dome (Fig. lOa) . It is difficult t o evaluat e th e strengt h o f th e litho sphere during the collapse of the Central Graben Dome, a s vertical movement s of the lithosphere may be in interaction with, and supported by, the asthenosphere. The Sout h Swedis h Dom e an d th e Souther n Scandes Dom e i n Norwa y ar e situate d alon g individual segment s o f th e rifte d margi n o n the S W corne r o f th e emergen t Precambrian Palaeozoic crus t (Fig . 1) . Uplif t o f souther n Norway probabl y als o ha s a deep-seate d origin i n th e uppe r mantl e an d th e astheno sphere (Rohrma n & va n de r Bee k 1996) . Th e suggested interrelationshi p modelle d fo r th e South Swedis h Dom e an d th e Centra l Grabe n Dome may als o be propose d for the souther n Norway uplif t an d th e Tripl e Junctio n Dome . Initial uplif t i n souther n Norwa y too k plac e i n Paleogene tim e wherea s th e uplif t o f th e South Swedis h Dom e peake d i n Neogene -

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Quaternary tim e (Rii s 1996 ; Japse n 2000) . The differenc e i n timin g o f initia l Cenozoi c uplift ma y b e relate d t o th e migratio n o f earl y uplift an d subsidenc e o f th e Tripl e Junctio n Dome relativ e t o th e Centra l Grabe n Dom e t o the SE .

Discussion and conclusions The propose d mode l o f linke d subsidenc e an d uplift show s tha t a bette r understandin g o f Cenozoic uplif t i n th e Nort h Se a Basi n ma y b e obtained whe n a lin k i s mad e t o th e Mesozoi c evolution. Analysi s o f th e centra l Nort h Se a Dome has demonstrated th e composite characte r and differentia l uplif t an d subsidenc e o f th e dome. Riftin g an d uplif t wa s initiated i n Middle Jurassic tim e i n the nort h i n the Tripl e Junctio n Dome. Thi s dom e wa s floode d durin g Lat e Jurassic time as rifting an d uplift move d sout h to the Central Grabe n Dom e (i n early Late Jurassic time) an d then th e Friesland Dom e (i n late Lat e Jurassic time) . Th e Centra l Grabe n Dom e wa s flooded durin g Early Cretaceou s time . The large-scal e evolutio n o f th e Jurassic Cretaceous an d Cenozoi c sedimentar y basin s in the easter n Nort h Se a show s tha t th e Mesozoi c depocentres were situate d i n the Central Grabe n and alon g th e Tornquis t Zon e wherea s th e Cenozoic depocentre s wer e concentrate d i n th e central Nort h Se a abov e th e aborte d Mesozoi c rifts (Fig s 7 an d 8) . Th e change d subsidenc e pattern aroun d th e Mesozoic-Cenozoi c tran sition indicate s a n uplift-subsidenc e reversal . The change of the overall basin morphology wa s governed by the first-order tectonic elements, i.e. the Centra l Grabe n Dom e an d th e Danis h Megablock, th e Tornquis t Zone , an d th e Sout h Swedish Dome . Thes e structure s ar e severa l hundred kilometre s acros s an d involv e vertica l movements o f betwee n 1 an d 2km . Th e magnitude o f th e structure s an d th e movemen t suggest causa l mechanism s i n th e mantl e involving th e entir e overlyin g crust . Jurassi c doming wa s characterize d b y thinnin g o f th e crust an d th e developmen t o f rif t basin s wit h volcanic activit y (Ziegler 1990) , bu t by contras t the Cenozoi c dome s develope d o n a thic k Caledonian an d Precambria n crystallin e crust , and ther e ar e n o sign s o f volcani c activit y an d only limite d faultin g is observed (i.e . th e crestal collapse structur e o f the Sout h Swedis h Dome) . The Cenozoic domes all have a marginal position along th e wester n faulte d margi n o f th e Fennoscandian High . Th e domes appea r t o have been stationar y throughou t thei r existence , an d

this suggest s tha t th e crusta l structur e wa s th e locating factor . I f th e dome s ha d develope d above mantl e plume s ascendin g fro m th e dee p mantle, the y woul d hav e a mor e rando m distribution. The subdivision of the central North Sea Dome into individual dome centres connects the Danish Megablock wit h th e Centra l Grabe n Dom e (Fig. 1) . The til t reversa l o f th e Danis h Mega block a t the Mesozoic-Cenozoic transitio n thus suggests a mutua l linkag e betwee n th e sub sidence o f th e Centra l Grabe n Dom e an d th e uplift o f souther n Swede n (Fig . 7) . Th e accelerated subsidenc e of th e centra l Nort h Se a Basin durin g lat e Neogene-Quaternar y tim e (Fig. 9) (Nielsen etal 1986 ; Ziegler 1990 ) is thus contemporaneous wit h th e uplif t o f th e Sout h Swedish Dom e an d th e exhumatio n o f th e sub Cambrian peneplai n establishe d b y Lidmar Bergstrom (1995) . Fission-trac k thermochronol ogy supports a model of accelerated uplif t durin g Neogene tim e (Cederbom et al 2000 ; Cederbo m 2002). The Cenozoi c post-rif t subsidenc e i n th e central Nort h Se a Basi n doe s no t follo w th e McKenzie (1978 ) model . Instea d o f th e modelled decreasin g subsidenc e ove r time , th e observations outlin e increasin g subsidenc e (Fig. 9) . Th e Mesozoic-Cenozoi c evolutio n of the Nort h Se a are a involve d uplif t o r sub sidence of the orde r of 2k m or mor e of larg e crustal plates ; fo r example , th e Danis h Mega block coverin g th e easter n hal f o f th e souther n North Sea . Th e mode l propose d i n th e presen t paper attribute s uplif t o f th e Sout h Swedis h Dome t o subsidenc e o f th e easter n Nort h Se a Basin durin g collaps e o f th e Centra l Grabe n Dome. Th e lithospher e carryin g th e Danis h Megablock an d adjoinin g Balti c Shiel d wer e linked i n a commo n movemen t tha t involve d a reorganized flo w patter n i n th e asthenospher e (Fig. 10) . The uplif t followe d b y subsidenc e probabl y involved mas s flo w i n th e asthenospher e t o account fo r th e observe d balanc e betwee n post rift subsidenc e an d margina l uplift . I n thi s scenario th e impac t o f therma l subsidenc e may be suppresse d an d not easy t o model . The regressive character of the Cenozoic basin (Michelsen et al 1998 ) and limited occurrence of major fault s (Ziegle r 1990 ) suggest s tha t th e Cenozoic packag e was deposited o n a viscoelastic plate (Watts et al. 1982) . This conclusion is in accordance wit h th e modellin g o f Barto n & Wood (1984) , who suggeste d an effective elasti c thickness o f th e orde r o f 5km . Th e Cenozoi c basin show s increasin g subsidenc e rate s ove r time (Fig . 9) , an d hig h subsidenc e rate s i n lat e

UPLIFT-SUBSIDENCE REVERSAL, EASTER N NORT H SEA

Fig. 10 . Lithospher e mode l o f crusta l uplift-subsidenc e reversa l acros s th e souther n Nort h Se a Basin . CG , Central Graben ; N-DB , Norwegian-Danish Basin ; NSB , North Se a Basin ; S P FZ , Sol e Pi t faul t zone ; TZ , Tornquist Zone, (a) During the Cenozoic post-rift subsidence, faulting along the Tornquist and Sole Pit fault zones was discontinued. The subsidence was accompanied by reverse tilting of the flanks, whereas southern Sweden and southern Englan d wer e raise d alon g th e borders . I t i s suggeste d tha t Cenozoi c subsidenc e o f th e lithospher e resulted in a revised flow pattern in the asthenosphere directe d toward s southern Sweden and southern England, (b) Mesozoi c riftin g an d uplif t o f th e Centra l Grabe n Dom e wa s accompanie d b y crusta l thinnin g an d asthenosphere risin g belo w th e rifte d crust . Th e uplif t o f th e Centra l Grabe n Dom e wa s decouple d fro m th e adjoining area s alon g the Tornquist and Sole Pi t faul t zones . Neogene-Quaternary time ma y b e viewe d a s a natural par t o f th e observe d tren d a s a n alternative t o intraplate deformation . The bimoda l symmetr y o f th e Nort h Se a Basin alon g th e centra l Mesozoi c rift s an d Cenozoic basi n axi s implie s tha t th e propose d model o f th e Cenozoi c uplif t i n th e eas t alon g the borde r o f th e Fennoscandia n Hig h ma y b e mirrored i n th e wes t i n th e Cenozoi c evolutio n of Britai n (Fig s 8 an d 10) . Wherea s th e uplif t along th e easter n margi n o f th e Nort h Se a Basin ma y b e relativel y simple , th e uplif t i n Britain ma y interfer e wit h deformation s o f th e fault-bounded shel f basin s t o th e sout h an d west, a s wel l a s wit h th e effect s o f th e opening of the Atlantic . However , Neogene Quaternary uplif t o f th e Sout h Swedis h Dom e corresponds t o uplif t i n souther n Englan d (Fig. 9) , an d Paleogen e uplif t o f Scotlan d based o n geomorphologica l evidenc e (Hal l 1991) ma y counterbalanc e th e uplif t i n southern Norway .

I than k K. Lidmar-Bergstro m (Universit y o f Stock holm), who introduced m e to the South Swedish Dom e in Ma y 199 9 durin g a Nordi c researc h excursio n through souther n Sweden ; I als o than k T . Dore , R . Gatliff, E . Hakansso n an d a n anonymous reviewer for constructive review s an d languag e correction s o f a n early draf t o f the manuscript .

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Cenozoic inversio n an d uplift of southern Britai n DEREK J. BLUNDELL Department of Geology, Royal Holloway, University of London, Egham TW20 OEX, UK (e-mail: d. blundell @gl. rhul ac. uk) Abstract: Wherea s significant exhumatio n o f northern Britai n too k place during Paleocen e time, probabl y a s a consequenc e o f uplif t cause d b y a mantl e plume , Paleogen e basi n inversion an d uplif t i n souther n Britai n appear s t o b e a consequence o f Alpine tectonism. Recent publications demonstrat e that inversio n o f Mesozoic basins across souther n Britai n was accompanie d b y subsidenc e o f flankin g basin s i n areas that had previousl y remaine d stable. Structures observe d on seismic sections across the Weald Basi n in SE England revea l that inversion occurred locally by north-directed reverse movements on pre-existing normal faults that cu t down a t a low angle dee p into the basement. The overall effect o f inversion o f the Weal d Basin, however , i s a bulk deformatio n tha t produced a domal uplift, flanke d b y subsidence o f th e Londo n an d Hampshire-Diepp e basins . A 2 D finit e elemen t thermo mechanical model of continental lithospher e containing a region of reduced strengt h i n the crust simulates Jurassic-Early Cretaceous extension to form th e Weald Basin, followed by compression during th e Tertiary to produce it s inversion an d the flanking basins. Th e timin g of tectonic events across souther n Britain correlates with times when Alpine stresses were transmitted int o th e forelan d t o th e nort h sufficientl y wel l t o lin k them . Throug h mos t o f Tertiary time , th e landscap e o f souther n England wa s o f relatively lo w elevation an d low energy surfac e processes . However , lat e Neogen e uplift , generall y greate r i n th e west , appears t o hav e bee n par t o f a larger-scal e uplif t o f lan d area s wit h har d roc k a t surface , which ha s no obvious tectoni c explanation .

In simpl e terms , th e surfac e geolog y outcro p pattern o f th e Britis h Isle s show s th e younges t rocks i n th e S E an d th e oldes t i n th e nort h an d west. Th e leve l o f erosio n i s suc h tha t th e Phanerozoic successio n i s full y expose d a t surface, alon g wit h Precambria n unit s datin g back t o Earl y Proterozoi c time . Moreover , th e extent o f Tertiary sediment s expose d onshor e i s limited to southern England and that of Neogene sediments i s restricted t o eastern Eas t Anglia . In similar fashion , topographic elevatio n an d relief are generall y greate r i n the nort h an d west , and least in the sout h and east. Thi s rang e of surfac e geology an d topograph y contrast s wit h othe r regions o f N W Europ e wit h similar , 30k m crustal thickness , suc h a s th e Netherlands , northern Germany , Denmark and Poland. Brodie & White (1994 ) explained th e denudation of northern Britain as a consequence of 8 km of magmati c underplatin g relate d t o a mantl e plume beneat h N W Scotland , recognize d a t surface a s the Tertiary Volcanic Province. White & Lovel l (1997 ) correlated th e associate d uplif t and erosion from a succession of volcanic pulses with the depositio n o f large quantitie s of clastic sediment a s submarin e fan s i n neighbourin g basins (Nort h Sea , Wes t Shetlands ) durin g th e

period 62-54 Ma. Their estimate s of up to 3 km of exhumation, base d o n vitrinite reflectance and fission-track data, were independently confirme d by a careful stud y of denudation of the Iris h Se a Basin b y Rowle y & White (1998) . In this , the y estimated th e amount of denudation by calculat ing the stretching and thermal subsidence histor y of th e basi n t o determin e ho w muc h post-rif t sediment thicknes s i s missing . Th e mismatc h between th e present-day an d the predicted dept h to basemen t represent s th e amoun t o f denuda tion. They obtained value s for the East Irish Sea Basin rangin g betwee n 0. 4 an d 2.6km , wit h values u p t o 1.5k m an d 1.7k m fo r th e neighbouring Wes t Lancashir e an d Cheshir e basins, respectively . Doubt s remai n abou t th e extent of uplif t modellin g based o n fission-trac k and vitrinit e reflectanc e data , whic h depend s o n the value used for the geothermal gradien t at the time o f uplift , an d coul d b e halve d t o aroun d 1.5km. None th e less, there i s a strong cas e fo r explaining the exhumation of northern Britain as a result of an Early Tertiary plume, which moved away aroun d 5 4 Ma whe n Nort h Atlanti c sea floor spreadin g wa s initiated , leavin g th e are a above se a level t o the present day . However, th e influence o f th e plum e di d no t exten d acros s

From: DORE , A.G., CARTWRIGHT , J.A. , STOKER , M.S. , TURNER , J.P . & WHITE , N . 2002 . Exhumation o f th e North Atlantic Margin: Timing, Mechanisms an d Implications for Petroleum Exploration. Geological Society , London, Specia l Publications , 196 , 85-101 . 0305-8719/02/$15.00 © The Geological Societ y o f London 2002 .

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southern Britain , s o tha t othe r explanation s ar e needed t o accoun t fo r th e basi n inversio n an d uplift tha t i s observed . Thes e ar e generall y ascribed (e.g . Chadwick 1993 ) to th e deforma tional response t o lithospheric stresses generate d during the Alpine Orogeny that were transmitted laterally northwards. The purpose of this paper is to examin e th e natur e an d timin g o f basi n inversion across souther n Britain, and the Weald in particular , t o asses s whethe r the y ca n b e correlated wit h Alpin e event s an d wh y th e deformation i s apparentl y mor e pronounce d across souther n Britai n tha n elsewher e i n N W Europe. To understand the Cenozoic inversion of sedimentary basin s acros s souther n Britai n i t i s necessary firs t t o appreciat e thei r earlie r evol ution a s extensional basins.

1988) an d b y pur e shea r i n th e lowe r crust . Lateral ramp s an d strike-sli p escap e structure s created NE-SW-trendin g fault s t o th e wes t and NW-SE-trending fault s t o th e eas t (Fig . 2). I n the collaps e o f th e Varisca n Orogeny , sedimen tary basin s wer e initiate d during Permia n time , filled wit h terrestria l red-be d sequence s i n th e hanging wall s o f forme r thrusts , reactivate d a s low-angle norma l faults . Thes e wer e usuall y bounded b y steeper , short-cu t norma l faults , mainly downthrow n t o th e south . Extensiona l basins develope d throug h Triassi c tim e and , in the Wessex an d Central English Channel basins, were compartmentalize d b y NW-SE-trendin g faults inherite d from Varisca n structures. Further extension an d subsidenc e during Jurassic-Early Cretaceous tim e resulte d i n th e depositio n o f shallow marin e sequence s tha t includ e lime stones, sandstone s an d organic-ric h claystones . At th e sam e time , changin g stres s condition s Evolution of sedimentary basin s throug h resulted in the eastward migration of depocentre s the Mesozoic era within sub-basin s o f th e Wesse x an d Centra l The Celtic Sea, Bristol Channel, Central English English Channe l basins and th e developmen t of Channel, Wesse x and Weal d basin s (se e Fig. 1 new e n echelo n faul t set s an d transtensiona l for locations ) wer e buil t upo n a Varisca n movements o f th e earlie r norma l east-wes t structural framewor k (Fig . 2). Th e northwar d faults. Continuin g subsidenc e durin g Earl y movement o f th e Armorica n Bloc k toward s th e Cretaceous time resulted in the deposition locally Midland Crato n and London Platform resulted in of fluvial-lacustrin e deposits . The structura l and compression o f th e intervenin g crust . Thi s stratigraphic evolutio n of th e Wesse x Basi n ha s deformed dominantl y b y north-vergent , east - been describe d by a number of workers, notably west-trending thrust s an d fold s tha t develope d Chadwick (1986), Lake & Karner (1987), Butler (1998) an d Hawke s e l a l (1998) , an d ha s bee n above a mid-crusta l decollemen t (BIRP S & well summarized by Underhill & Stonely (1998), ECORS 1986 ; Chadwick 1986 ; Brooks e t al

Fig. 1 . Simplifie d geologica l ma p o f souther n Englan d showin g mai n fault s an d sedimentar y basins . SF , Sticklepath Fault; SH , Sout h Hewit t Fault; WF , Watchet-Cothelstone Fault ; PW, Purbeck-Wight Disturbance; CCH, Centra l Channel High. 3 , line o f section show n in Fig. 3; 4, SWT4 profile, bol d sectio n show n in Fig. 4; 6, line of BGS sectio n show n in Fig. 6. Inset: seismi c line s C78-02, C78-0 3 an d C78-04 ; B ? Brightlin g well.

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Fig. 2 . (a) Map showing the Variscan structural framework of southern Britain (after Lefor t & Max 1992) . VTF , Variscan Thrust Front , (b) Summar y of post-Variscan tectonic event s affecting souther n Britain.

who regarde d th e Permian-Lowe r Cretaceou s succession a s a singl e megasequence . It s uppe r bound i s marke d b y a regiona l unconformit y around Aptia n tim e (12 4 Ma), whic h truncate s successively olde r strat a toward s th e west , cutting dow n t o th e Permo-Tria s deposits . Th e Permian-Early Cretaceou s structura l and strati graphic evolutio n o f the Nort h an d Sout h Celti c Sea basin s ha s bee n describe d b y Petri e e t al.

(1989), tha t o f th e Bristo l Channe l Basi n b y Brooks e t a l (1988 ) an d tha t o f th e Centra l English Channel Basi n by Hamblin e t al (1992) . These basins all have similar histories. Through out thi s perio d th e neighbourin g stabl e blocks , such as the London Platform , Cornubia n Massi f and Armorica n Block , ha d remaine d a s positiv e regions wit h n o histor y o f subsidence . Durin g Late Cretaceous time , a broadl y westwar d

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progressing marin e transgressio n le d t o th e deposition o f sandston e an d clayston e followe d by th e Chal k Group , whic h extende d acros s the whole regio n an d wel l beyond , includin g th e former lan d areas . Thi s i s well illustrate d i n th e palaeogeographical atla s o f Ziegle r (1990) . Deposition o f Chalk i s known to have continued through t o Maastrichtia n tim e i n th e souther n North Se a an d th e Englis h Channe l (Rawso n 1992) bu t in most area s the youngest portion ha s been remove d b y erosion . A regional disconfor mity represent s a perio d o f uplif t an d erosio n between 7 4 an d 6 0 Ma, whe n th e earlies t recorded Tertiar y sediment s were deposited . Cenozoic inversion Extensive erosio n o f Chal k acros s souther n England le d t o th e developmen t o f a sub Paleogene surface , in some areas covered b y thin lateritic soi l deposit s (Gree n 1985) . Durin g Paleogene time , basi n inversio n i s eviden t i n various ways. In some areas it is concentrated on particular faults , wherea s elsewher e i t i s characterized a s a broad domal uplift (Chadwic k 1993), wit h periphera l basin s develope d acros s areas tha t had remained a s stable positive block s during Permian-Late Cretaceous times . Fault inversion The mos t spectacula r faul t inversio n i s th e Purbeck-Wight Disturbance , see n alon g th e Dorset coas t t o th e Isl e o f Wight . Underbil l & Paterson (1998) described th e structures in detail and demonstrated th e effect o f inversion upon the east-west-trending segmente d faul t system , linked b y rela y ramps , extendin g betwee n Abbotsbury in th e wes t and Whiteclif f Bay, Isle of Wight, in the eas t (Fig. 1) . Using information from seismi c sections , well s and surface geology , they showe d tha t th e faul t syste m acte d a s a normal growt h fault , downthrow n t o th e south , from Permia n t o Early Cretaceous time , but was inverted b y revers e movemen t durin g th e Tertiary. Althoug h th e faul t syste m remain s i n net extension for the Permian-Lower Cretaceous succession, invers e movemen t i s greates t wher e formerly extensio n was greatest. Deformation of incompetent hanging-wall strata against Chalk in the footwall, which acted as a semi-rigid buttress, implies tha t inversio n resulte d fro m northward directed stress . Movemen t o f hanging-wal l sediments relative to footwall Chalk and footwall deformation o f the Chalk can be used to estimat e the amoun t o f inversion , uplif t an d latera l shortening. Fo r example, fro m a seismic sectio n across th e Abbotsbury-Ridgewa y Faul t (Butle r

1998, Fig . 8 ) north of Weymouth, at least 350m Tertiary displacemen t ca n b e estimated . O n th e Isle o f Wight, uplif t o f a t least 1500 m has bee n recorded (Butle r 1998 , Fig . 13) . I n additio n t o localized uplift on inverted faults, there is a broad upwarp o f th e Wesse x Basi n o f betwee n 25 0 m (Chadwick 1993 ) an d 500m (La w 1998) . Some , at least, of the variation in structural style can be attributed t o th e presence , o r otherwise , o f sal t (Harvey & Stewart 1998) . Inversion i n th e Centra l Englis h Channe l Basin, to the south, is concentrated on the Central Channel Hig h (Fig . 1) , wher e La w (1998 ) ha s estimated ove r 1000 m o f uplif t fro m soni c velocity measurement s i n boreholes , supporte d by fission-trac k an d vitrinit e reflectanc e data . The structura l evolutio n o f th e Centra l Channe l High wa s investigate d b y Beele y & Norto n (1998) usin g tw o north-sout h high-resolutio n seismic section s acros s it , wit h wel l tie s fo r stratigraphic correlatio n an d dept h conversion . Sequential balance d sectio n restoratio n o f th e two true-scal e dept h section s resulte d i n estimates o f 1100-1500 m uplif t o f th e hang ing-wall sediment s ove r th e Centra l Channe l High relative to the footwal l to the sout h during Tertiary inversion . The fault geometr y require s a northerly di p o f 55 ° decreasin g t o 30 ° a t abou t 12 km depth. From this, and the deep structur e of the Wesse x Basi n identified byChadwic k (1986, 1993), Beele y an d Norto n propose d tha t th e inversion developed upon a linked set of crustalscale fault s tha t initiated as par t o f th e Varisca n thrust system, evolve d durin g basin extension as low-angle norma l fault s an d reactivate d agai n during Tertiar y tim e t o produc e th e fault controlled inversion . Thei r schemati c cross section i s reproduced i n Fig . 3 . The implication from thi s is a north-south shortenin g by c. 4km of som e 20 0 km o f crust. There ha s been considerabl e debat e abou t the timing o f inversion. The sub-Paleogen e regional unconformity, wit h differentia l erosio n o f th e Chalk, indicate s uplift acros s souther n England, but a s th e Uppe r Paleocene-Eocen e successio n within th e Hampshire-Diepp e Basi n i s itsel f involved i n th e faul t inversion , for exampl e o n the Isl e o f Wight , a t leas t par t o f th e inversio n must b e younger . Chadwic k (1993 ) argue d strongly i n suppor t o f a Miocen e ag e fo r th e main inversio n episode , 'correspondin g t o th e main Alpin e deformatio n event s a s continental collision occurre d betwee n Afric a an d Europe' . A successio n o f Tertiar y sediment s o f 650 m thickness i n th e Hampshire-Diepp e Basi n i s well exposed in Alum Bay at the western end and in Whitecliff Bay at the eastern end of the Isle of Wight. Restin g unconformabl y upo n a n erode d

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Fig. 3 . Schematic crusta l section by Beeley & Norton (1998) (locatio n show n in Fig. 1) , showing basin inversion on linke d crustal-scal e faults . CCH , Centra l Channe l High ; PW , Purbeck-Wight Disturbance ; WP , WardourPortdown structure , P, Vale of Pewse y Fault; VTF, Varisca n Thrust Front .

Chalk surfac e ar e re d mottle d soil s (Readin g Beds, c. 58 Ma) of Late Paleocene age . These ar e followed by an Eocene successio n comprising , in turn, th e marin e Londo n Clay , shallo w marin e and lagoona l sand s an d clay s wit h occasiona l coals o f th e Bracklesha m Group , an d marin e clays an d sand s o f the Barto n Group . Thes e are followed b y th e Solen t Grou p o f nea r shore an d freshwater silt s an d clays , an d a freshwate r limestone. Th e younges t o f thes e i s o f Earl y Oligocene age , c . 30 Ma. A s all these sediment s were deposited at or near sea level, sedimentatio n rates mus t hav e kep t pac e wit h subsidence , t o fill th e accommodatio n spac e withi n th e Hampshire-Dieppe Basin. A carefu l stud y b y Gal e e t al (1999 ) o f th e Eocene successio n expose d i n Whiteclif f Ba y established tha t clasts derive d fro m successivel y older horizon s occurre d i n successively younger strata, a s th e hanging-wal l sediment s o f th e Sandown Pericline , just a few hundred metres to the sout h (th e easter n segmen t o f th e inverte d fault syste m describe d b y Underhill & Paterso n (1998)), wer e bein g uplifte d an d eroded . Gal e et al (1999 ) foun d tha t uplif t occurre d i n tw o main phases . Uplif t o f 200-30 0 m occurre d between 47. 5 an d 4 4 Ma, wit h rates a t time s of 0.1 mma"1. A furthe r 200-300 m uplif t an d erosion occurre d durin g th e perio d 42-3 6 Ma, probably i n shor t pulse s o f les s than 1 Ma duration. Th e presenc e o f derive d fossil s an d clasts withi n th e uppermos t bed s i n th e succes sion suggest s tha t uplif t persiste d int o Earl y Oligocene time . Th e tota l uplif t durin g th e Eocene faul t inversio n o f som e 500 m i s a substantial part o f the displacemen t recorded o n faults withi n th e Wesse x Basin . However , th e Chalk an d overlyin g Reading Bed s an d Londo n Clay ar e vertica l i n bot h Alu m Ba y an d Whitecliff Bay , th e Bracklesha m Grou p ar e steeply dippin g to the north, but the dip reduces to a lo w angl e o f aroun d 5 ° through th e Barto n and Solen t Groups , som e 80 0 m fro m th e Chal k

outcrop. Clearly, this deformation, which has upended th e Paleogen e sequence , i s younge r than the 3 0 Ma sediments involved . Linked t o the east-west-trending fault system of th e Purbeck-Wigh t Disturbanc e vi a th e Lytton Cheyne y Faul t i s th e NW-SE-trendin g Watchet-Cothelstone Faul t (Fig . 1) . Reviewin g evidence o f faul t movements , Miliorizo s & Ruff ell (1998 ) conclude d tha t i t originate d a s a Variscan structure and has acted at various times subsequently as a strike-slip faul t i n both dextral and sinistral senses. East-west horsetail splays at either en d sugges t tha t th e mos t recen t move ments hav e bee n sinistral , bu t thei r timin g i s constrained only to post-Jurassic time. Sixty-fiv e kilometres t o th e west , th e NW-SE-trendin g Sticklepath Faul t (Fig . 1 ) appears t o hav e had a similar histor y o f strike-sli p movements . Seg mented alon g strike , relay s hav e acte d a s releasing bend s t o produc e pull-apar t basins , the Lundy, Petrockstow an d Bovey basins, whic h are fille d wit h terrestrial sediment s o f Eocenemid-Oligocene age . Thes e includ e fluvia l an d lacustrine gravels , sand s an d clay s wit h thi n lignite bands . Althoug h th e Bove y Basi n i s a t least 1000 m deep, an d the oldest strat a have not been sampled , th e natur e o f th e sediment s implies tha t th e tectonicall y drive n subsidenc e was matched by the rate of sedimentation so that accommodation spac e wa s filled . Thu s th e sediments dat e th e faul t movement . Hollowa y & Chadwic k (1986 ) argue d tha t thi s movemen t was sinistral, with 6 km horizontal displacement. However, compariso n wit h a scale d analogu e model (Doole y & McCla y 1997 ) suggest s tha t the sinistra l displacemen t wa s n o mor e tha n 2km. Th e age s o f sediment s i n th e smalle r Petrockstow an d Lund y basin s ar e simila r t o those i n th e Bove y Basi n an d impl y tha t movements o f th e Sticklepat h Faul t continue d to mid-Oligocene time , around 28 Ma. This may also hav e bee n th e cas e fo r th e Watchet Cothelstone Fault .

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Inversion by bulk deformation and uplift Although th e near-surfac e mechanis m o f basi n inversion ma y b e throug h invers e displacement s on individual faults, the bulk effect i s the uplift of a basin to a domal structure . This i s observed i n the North Celti c Se a Basin an d the Weald Basin . The SWA T 4 dee p seismi c sectio n (BIRP S & ECORS 1986), Fig. 4 illustrates the overall effec t of inversio n o f th e Nort h Celti c Se a Basin . Th e basin rest s o n th e hangin g wal l o f a low-angle , south-dipping, plana r faul t tha t ca n b e trace d from surfac e t o 15k m depth , wher e i t merge s with th e to p o f a highl y reflectiv e zon e o f subhorizontal reflectio n segment s tha t extend s down to the Moho. This fault can be correlated a t surface i n souther n Irelan d wit h th e Varisca n Thrust Front , bu t i t ha s clearl y bee n reactivate d in extension a s a low-angle normal fault t o allow space fo r th e formatio n o f th e Nort h Celti c Se a Basin a s a half-graben. Th e basin is filled with a thick Permo-Triassi c terrestria l red-be d succes sion an d a Jurassic-Cretaceou s marin e succes sion simila r t o tha t i n th e Wesse x Basi n (Petri e el al. 1989) . Late r compressiv e reactivatio n o f the majo r faul t ha s resulte d i n th e inversio n of faults withi n th e Nort h Celti c Se a Basin , producing localize d anticline s an d positiv e flower structures , togethe r wit h a broa d doma l uplift o f 1-2 km (Roberts 1989) . Robert s argue d that although som e inversio n may have occurre d

during Paleocen e time , th e majo r par t i s o f Oligocene age . Erosio n ha s resulte d i n th e outcrop of Jurassic strata at the sea bed. Flanking the inverte d Nort h Celti c Se a Basi n ar e sa g basins fille d wit h Paleogen e sediments , on e o f which i s show n i n Fig . 1 . It appear s fro m thei r association tha t thei r subsidenc e accompanie d the uplift o f the North Celtic Sea Basin. A similar configuration i s evident with the domal inversion of th e Weal d Basi n an d th e subsidenc e o f th e flanking London and Hampshire-Dieppe basins, which ar e no w examine d in more detail. Inversion of the Weal d Basin The Weal d Basi n initiate d a s a n easterl y prolongation o f th e Wesse x Basi n a t th e beginning o f Jurassi c tim e (20 8 Ma). I t devel oped as an extensional basin, subsiding by means of norma l growt h fault s o f mainl y east-wes t trend, th e mos t activ e of which were clos e t o its northern margi n agains t th e Londo n Platform , which ha d acte d a s th e undeforme d forelan d t o the Varisca n orogen . Thu s th e Weal d Basi n developed durin g Jurassi c tim e mainl y a s a n asymmetric basi n wit h stron g down-to-sout h normal faults alon g its northern margin, founded upon reactivated Varisca n thrusts, which acted as low-angle extensiona l detachments. An element of transtensiona l faul t movement s durin g thi s

Fig. 4 . Deep seismic profile SWAT 4 (BIRPS & ECORS 1986 ) (locatio n show n in Fig. 1) , interpreted t o show the Variscan Thrus t Fron t (VTF ) reactivate d a s a low-angl e norma l faul t t o provid e spac e fo r th e Mesozoi c Nort h Celtic Sea Basin i n its hanging wall , then inverted a s indicated b y structures within the basin. The broad uplif t o f the basin an d the flanking Tertiary basi n t o the sout h shoul d b e noted .

CENOZOIC INVERSION AN D UPLIFT OF SOUTHERN BRITAI N

period i s marked b y th e e n echelon geometr y of the normal faults in plan view and the presence of associated WNW-ESE-trendin g fault s (Fig . 1) . Within th e basin , lowe r Lia s marin e claystone s and limestone s res t directl y upo n a Devonian Carboniferous basement . Th e Jurassi c succes sion, althoug h simila r t o tha t i n th e Wesse x Basin, i s dominate d b y carbonat e rock s lai d down i n a shallo w marin e low-energy , wave dominated environment , indicativ e o f clea r se a and a limite d suppl y o f clasti c materia l fro m nearby land, particularly the London Platform to the north . Depositio n throug h Earl y Cretaceou s time o f marin e sandstones , mudstone s an d sabkha-type evaporite s testifie d t o continuin g subsidence, a t a rate greate r tha n i n th e Wesse x Basin, an d depositio n o f clasti c sediment s i n a nearshore environment . Sedimentation thu s kept pace wit h subsidence . A mid-Cretaceou s (124 Ma) unconformity marking th e terminatio n of thi s megasequenc e (compar e th e Wesse x Basin) represents a relatively small time gap and sedimentation resumed , i n continuit y wit h th e Wessex Basin , wit h marin e sandstone s an d claystones succeede d b y th e Chalk . Althoug h the younges t survivin g Chal k i s o f lat e Campanian ag e (8 3 Ma), i t i s likel y tha t deposition continue d throug h Maastrichtia n time t o aroun d 74 Ma, yielding a total thicknes s of Chal k o f abou t 500m . Durin g th e followin g 16 Ma, regiona l uplif t an d erosio n le d t o th e removal o f a substantia l portio n o f th e Chalk , ranging, throug h differentia l erosion , betwee n 100 an d 350m . A sub-Paleogen e surfac e developed, upo n whic h th e commencemen t o f subsidence o f th e Londo n Basi n t o th e nort h of the Weald, an d the Hampshire-Dieppe Basin to the south , i s recorde d i n th e depositio n o f terrestrial, intertida l an d shallo w marin e sedi ments of late Paleocene ag e (58 Ma). These wer e succeeded b y th e Londo n Clay , representin g marine conditions with no evidence o f shorelin e facies, whic h ma y hav e inundate d muc h o f southern Englan d betwee n 5 4 an d 5 0 Ma. Interbedded shallo w marin e sand s an d clay s were deposite d i n successio n throug h Eocen e time, th e younges t preserve d bein g o f lat e Eocene ag e (c. 30 Ma). The maximu m thickness of Tertiary deposits in the London Basin is 300 m and in the Hampshire Basin it is 650 m. Evidence that subsidenc e o f th e Londo n Basi n accompanied uplif t o f th e Weal d i s les s convincing tha n o n th e Isl e o f Wight. However , Gale et al. (1999 ) pointe d out tha t cher t clast s identified a s bein g derive d fro m th e Lowe r Cretaceous Hyth e Bed s (c . 14 0 Ma) foun d i n Lutetian ag e pebbl e bed s i n th e Londo n Basi n imply tha t th e Weal d wa s bein g uplifte d a t th e

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same tim e a s th e firs t phas e o f inversio n o f th e Isle of Wight (47.5-4 4 Ma). Inversion o f the Weald ca n be observed o n oil industry seismi c section s no w availabl e fo r research us e throug h th e U K Onshor e Geophy sical Library . Figur e 5 present s thre e north south true-scal e dept h converte d section s acros s the Weal d showin g th e principa l Jurassi c horizons interprete d fro m th e seismi c data : lin e locations ar e show n i n Fig . 1 . Faults hav e bee n migrated withi n th e vertica l plan e o f section . Line 78-0 2 wa s sho t wit h explosives , givin g greater dept h penetratio n tha n th e vibrosei s source use d fo r th e othe r lines . O n thi s line , reflections from low-angl e faults can be traced to 8km depth , wel l int o th e crystallin e basement . Similar south-dippin g fault s observe d o n th e adjacent line s cannot be traced a s far down. The low di p angle s o f thes e fault s sugges t tha t the y originated a s thrusts (probably Variscan) , which then acte d i n extensio n durin g Jurassi c tim e a s the Weal d Basi n subsided . The y wer e late r inverted during Tertiary time, as is evident on the fault-related uplif t structure s withi n th e Juras sic-Lower Cretaceou s succession . Th e sens e of movement o f inversio n i s t o th e north . Th e Brightling well , offse t 4k m eas t o f lin e 78-02 , provides stratigraphi c correlatio n wit h seismi c reflectors and , in addition , records a duplication by reverse faulting of the Lower Lias succession. Assuming th e faul t dip s a t 45 ° t o th e south , consistent wit h fault s observe d o n th e seismi c section, implies a 350m northward displacement as a result of Tertiary inversion, comparable with that on the Abbotsbury-Ridge way Fault. Further examples o f faul t inversio n i n th e Weal d wer e presented b y Butle r & Pulle n (1990) , wh o als o found NW-SE-trendin g fault s tha t ha d under gone significan t transpression . Although there is no direc t evidenc e o n th e timin g o f inversion , Butler an d Pulle n reporte d th e presenc e o f ferroan calcit e precipitate d o n th e crest s o f Tertiary uplif t structure s tha t hav e S r isotop e ratios correspondin g t o a Late Oligocene-Earl y Miocene ag e (c. 24 Ma). However, such evidence is highly suspect.

Inverse model of the Weald Although th e seismi c dat a indicat e tha t th e mechanism o f inversion i s based o n reactivatio n of low-angle , south-dippin g fault s tha t cu t dee p into th e crust , th e overal l effec t o f the inversio n of th e Weal d i s a broa d doma l uplif t o f som e 1500m, togethe r wit h th e downwarpin g o f th e flanking London and Hampshire-Dieppe basins, seen in Fig. 1 . Figure 6a is a simplified version of

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Fig. 5 . True-dept h interpretation s o f thre e north-sout h seismi c section s acros s th e Weal d (locatio n show n i n Fig. 1) . Formation tops : K , Kimmeridge; C, Corallian; O , Oxfordian; L, Lias; D , Devonian.

the cross-sectio n produce d b y th e Britis h Geological Surve y (BGS ) wit h thei r 1:25 0 00 0 series soli d geolog y maps , sheet s 50NO O an d 51NOO(British Geologica l Surve y 1988 , 1989) . The line of section i s located i n Fig. 1 . In Fig. 6a, individual faults shown on the BGS section have been remove d an d th e cross-sectio n ha s bee n smoothed accordingl y t o focu s o n th e effec t o f distributed bul k pur e shea r deformation , a s a basis fo r preparin g a n invers e model . Criticall y for th e modelling , th e deep , crustal-scal e fault s provide a zon e o f lo w strengt h i n th e crust , laterally confine d below th e Weal d Basin , with relatively stron g crus t beneat h th e flankin g basins. A 10: 1 vertical exaggeratio n emphasize s the doma l uplift . Isochron s hav e bee n adde d t o provide tim e marker s tha t monito r th e basi n evolution. Usin g a knowledg e o f Tertiar y stratigraphy an d landscape evolutio n base d o n a reconstructed cross-sectio n o f th e Weal d b y Jones (1999 , Fig . 3) , the cross-sectio n ha s bee n reconstructed bac k t o mid-Eocen e tim e (4 0 Ma, Fig. 6b) , a t whic h tim e th e flankin g basin s ha d subsided t o a maximu m extent . Becaus e strati graphic evidenc e indicate s shallo w seas , th e

reconstruction assume s th e wate r dept h i s negligible. No accoun t is take n of the lan d surface topograph y acros s th e centra l Weald , where faul t inversio n woul d hav e create d localized uplift . Reconstructio n bac k t o lat e Paleocene tim e (6 0 Ma, Fig . 6c ) whe n th e flanking basin s bega n t o subside , togethe r with the broa d doma l uplif t o f th e Weald , wa s accomplished b y strippin g off the younger strata and adjustin g th e geometr y o f th e underlyin g strata b y vertica l movement . N o accoun t wa s taken o f latera l compressio n fro m basi n inver sion. Th e uppe r surfac e o f th e cross-sectio n represents the sub-Paleogene surface. The cross section at latest Cretaceous time (68 Ma, Fig. 6d) represents the time when deposition of the Chalk had bee n complete d an d indicate s a unifor m thickness of Chalk across the section. The extent of uplif t and erosio n of Chal k betwee n 68 and 60 Ma (Fig. 6c) is based on the estimates of Jones (1999). Using the work of Butler & Pullen (1990) for th e Jurassic-Cretaceou s succession , th e reconstruction continue d to th e mid-Cretaceou s unconformity a t 12 4 Ma (Fig. 6e) and back to the initiation o f th e Weal d Basi n a t th e beginnin g

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Fig. 6 . Inverse model of Weald (location show n i n Fig. 1). Isochrons (Ma ) are show n in italics.

of Jurassi c tim e (20 8 Ma). I n thi s reconstruc tion, accoun t wa s take n o f th e effect s o f compaction an d decompaction o f the Jurassic Lower Cretaceou s successio n durin g subsi dence an d subsequen t uplift , bu t thi s wa s regarded a s unwarrante d fo r th e Uppe r

Cretaceous an d Tertiar y deposits . N o accoun t was take n o f an y deepe r crusta l thermo tectonic effects , a s th e purpos e o f th e invers e model wa s t o reconstruc t th e basi n geometr y in cross-section , whic h coul d ac t a s a templat e for forwar d modelling .

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Forward model The purpos e o f forwar d modellin g wa s t o lear n what cause d th e subsidenc e o f th e Londo n an d Hampshire-Dieppe basin s a t th e sam e tim e a s the Weald Basin was being inverted and uplifted . A 2 D thermo-mechanica l forwar d modellin g scheme develope d b y Nielsen & Hansen (2000) , based o n th e mode l o f Brau n & Beaumon t (1987), provide s a physica l basi s o n whic h t o simulate th e subsidenc e o f the Weal d Basi n an d its subsequen t inversion . Detail s o f th e math ematical expression s use d fo r th e mode l ca n b e found i n thes e publications . Figur e 7 show s schematically ho w the model works . The uppercrust, lower-crus t an d upper-mantl e lithospher e are define d by thei r physica l properties , suc h as density, Young' s modulus , Poisson' s ratio , compressional strength , cree p parameters , specific heat , hea t productio n an d therma l conductivity, appropriat e fo r quartz-rich , feld spar-rich an d olivine-ric h minera l rheologies , respectively, a s show n i n Tabl e 1 . Thei r behaviour i s dictated b y th e prevailing tempera ture, pressure an d strai n rate a t each poin t i n the model a s i t evolves , accordin g t o function s expressing th e response of the materia l t o stres s in a viscoelasti c o r plasti c manner . A constan t heat flux of 35 mW m~2 is applied t o the base of the mode l upwar d fro m th e underlyin g mantle , which result s i n a surfac e hea t flo w o f 60-70 mWm~2 , dependin g o n the thicknes s o f the crust. Buoyancy forces operate throughou t to maintain isostati c equilibrium . Th e mode l consists o f a finit e elemen t mes h tha t i s closel y

spaced i n th e uppe r crust , intermediat e i n th e lower crus t an d mor e ope n i n th e uppe r mantle. Time-variable condition s ar e impose d o n th e model, includin g hea t flo w upwar d fro m th e mantle, surfac e rate s o f erosio n an d sedimen t deposition, an d kinemati c boundar y conditions. These las t condition s impos e extensio n o r compression fro m on e end , whereas th e othe r end remain s pinned . Eustati c sea-leve l change s (Haq e t al. 1987 ) are als o incorporate d int o th e model. I n th e mode l fo r th e Weald , a centra l region o f 100k m widt h i s presen t i n whic h th e compressional strengt h o f th e uppe r crus t i s reduced t o 40% of the values on either side, given in Table 1 , and the creep parameter B of the lower crust i s reduce d t o 85% , asymmetrically northward, o f th e value s o n eithe r sid e (Fig . 7) . Whereas the reduction of compressional strength ensures plastic yieldin g at lower stres s levels, the reduction i n B lead s t o accelerate d cree p a s a result o f lowe r viscosities . The mode l wa s firs t subjected t o latera l extensio n a t 0. 1 mm a"1 between 21 0 an d 17 0 Ma, risin g t o 0. 5 mm a"1 between 16 5 and 15 5 Ma before reducing to zero by 14 5 Ma. Thi s simulate d th e Jurassic-Earl y Cretaceous extensio n require d t o creat e th e Weald Basin , amountin g t o 13km . Therma l relaxation followed for 8 5 Ma unti l compression began a t 6 0 Ma, risin g t o a rat e o f 0.2 5 mm a"1 between 5 0 an d 2 0 Ma, reducin g t o 0. 1 mm a * by 1 0 Ma, and continuing to the present, a total of 11.5km. Because th e model ignores faul t move ments, th e amount s o f extensio n an d com pression ar e overestimate d by a factor o f two s o that th e value s given above shoul d be halve d to

Fig. 7. Forward modelling schem e (Nielse n & Hansen 2000).

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Table 1 . Parameters for th e thermo-mechanical model o f th e Weald Parameter

Sediments

Upper crus t

Lower crus t

Mantle

Surface porosit y c/> 0 Density (k g m~ 3 ) Young's modulu s (Pa ) Poisson's rati o Creep paramete r n Creep paramete r B (MP a s 1/n) Creep activation energ y (k J mol"1) Compressional strengt h (MPa ) Thermal conductivit y (W m"1 K" 1) Specific hea t ( J kg" r KT 1) Heat productio n rat e (jx W m~ 3 )

0.60 2700 1011 0.25 3.10 208 135 17.5 2 1000 1.3

2700 1011 0.25 3.10 208 135 26.2 3 850 1.3

2900 1011 0.25 3.20 12.28 239 26.2 2.3 900 0.3

3300 1011 0.25 4.48 0.2628 498 52.4 4 1000 0.01

Sediment porosit y 0 varies wit h dept h z (m ) accordin g t o 4> = 0oexp(—z/2000) . Cree p activatio n energ y an d parameters n and B relate t o expressions fo r viscous deformation (Brau n & Beaumont 1987) .

be realistic . Th e progres s o f th e mode l throug h this deformationa l histor y wa s monitore d a t 1 Ma intervals . Sedimentatio n rate s o f u p t o 0.1 mm a~~ 1 ensured that sedimentation kept pace with subsidence . Accommodatio n spac e wa s largely fille d an d wate r depth s wer e small . Similarly, erosio n rate s wer e suc h tha t erosio n kept pac e wit h uplif t t o kee p surfac e elevatio n and relie f low . Th e maximu m erosio n rat e wa s O.lSmma^ 1 during the rapid fall o f se a level in Tertiary time . Figur e 8 present s th e forwar d model a t 124 , 68 , 60 , 4 0 an d OM a fo r direc t comparison wit h the inverse model, Fig . 6. Because the mode l i s an oversimplification of reality ther e i s a limi t t o th e exten t t o whic h i t should attemp t t o replicat e th e detail s o f th e inverse model. However, the main features of the model offe r ne w insight s int o th e histor y o f subsidence an d inversio n o f th e Weal d an d th e reasons wh y th e latte r wa s accompanie d b y th e formation of the London and Hampshire-Dieppe basins. The critica l factor i s the lo w strengt h of the crus t across th e Weald, sandwiche d between the high strengt h o f the Londo n Platfor m to th e north and the relatively high strength of the crust to th e south , i n th e easter n par t o f th e Englis h Channel an d norther n France . Thi s strengt h pattern wa s undoubtedl y inherite d fro m th e Variscan Orogeny and the area of weak crust was exploited durin g Mesozoi c tim e t o contro l th e locations of extensiona l basin s acros s souther n Britain. Th e model shows how the high strengt h of th e crus t i n th e area s flankin g th e Weal d requires th e developmen t o f compressiona l basins a s a necessar y accompanimen t t o th e updoming o f th e Weal d durin g inversion . Extension o f th e uppe r crust resulte d i n a 2 km rise o f th e Mon o beneat h th e Weal d Basin . Subsequent inversion thickened th e crust beneath

the basin so that the end result is a deepening of the Moh o t o 35k m dept h belo w th e zon e o f weakness. Th e Moh o beneat h th e flank s remained a t 34k m throughout . Thi s mode l contrasts wit h one proposed b y Cloeting h e t al (1990) in which lateral compressio n impose d on a rifte d basi n i n lithospher e tha t ha s vertically variable but laterally uniform rheolog y results in a deepening of the basin an d uplift o f the flanks, the opposit e o f what is observed i n the Weald . The mode l simulate s the subsidenc e histor y and geometr y o f th e Weal d Basi n durin g Jurassic-Early Cretaceou s time , whic h i s indi cated b y th e isochron s show n a t 12 4 Ma i n Fig. 8e . I t als o simulate s th e unconformit y a t 124 Ma a s a respons e t o th e cessatio n o f extension an d th e dominanc e o f therma l relaxation. Th e mode l the n simulate s erosion of the Chal k betwee n 6 8 an d 6 0 Ma, bu t fail s t o include the commencement of uplift of the Weald at this time. Between 60 and 40 Ma, uplift o f the Weald i s accompanie d b y subsidenc e o f th e London an d Hampshire-Diepp e basins . Between 4 0 Ma an d th e present , th e Weal d i s further uplifte d an d erode d dow n t o uppermos t Jurassic level , th e Weal d dom e become s mor e pronounced an d th e flankin g basin s ar e furthe r compressed an d partially eroded. Th e latter is in part due to the eustatic fall in sea level of around 100m sinc e Miocene tim e (Ha q et al. 1987) . Correlation of events in southern Britain with Alpine tectonics A very extensive literature chronicles the tectonic evolution o f th e Alp s a s a consequenc e o f th e convergence of Africa wit h Europe over the past 120 Ma (Dewe y e t a l 1989 ; Pfiffne r 1992) .

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Fig. 8 . Forward model of the Weald Basin, its inversion and the subsidence of the London and Hampshire-Dieppe basins produce d b y Hansen , based o n th e modellin g schem e show n in Fig . 7 . This ma y b e compare d wit h th e inverse model, Fig . 6 . Isochrons (Ma ) are show n in italics.

Ziegler e t al. (1995 ) examine d i n detai l th e effects o f the Alpine collisional history upon the foreland t o th e north . The y pointe d ou t tha t 'forces relate d t o collisiona l plat e interactio n appear t o b e responsibl e fo r th e mos t importan t intra-plate compressiona l deformations ' an d concluded tha t 'th e bul k o f Lat e Cretaceou s

and younge r intra-plate compressiona l defor mations observed i n western an d central Europe developed i n response t o stresse s tha t developed as a consequenc e o f collisiona l couplin g of th e Alpine an d Pyrenea n orogen s wit h thei r fore lands'. Th e transmissio n o f stresse s int o th e foreland thu s took place whenever they could not

CENOZOIC INVERSIO N AN D UPLIFT O F SOUTHERN BRITAI N

be accommodate d withi n th e oroge n o r th e foreland basin . Th e evidenc e fo r th e timin g o f such event s ca n b e foun d withi n th e forelan d basin, i n particula r th e Nort h Alpin e Forelan d Basin, wit h regar d t o north-directe d stresse s that coul d hav e affecte d souther n Britain . Figure 9 summarize s th e correlation , base d o n work b y For d e t al. (1999 ) an d For d (pers . comm.) o n th e stratigraphi c an d structura l evolution o f th e Nort h Alpin e Forelan d Basin . She ha s recognize d fou r stage s o f progressiv e deformation: (1) a Mid-Lat e Eocen e stag e (46-3 6 Ma), when the convergence o f the Apulian Plate with Europe, a t a rat e o f 15mma ~ , resulted i n th e migration of the flexura l basi n and the fron t o f a low-angle external orogenic wedge, allowing the northward propagation of stress into the foreland; (2) a n Early-Lat e Oligocen e stag e (33 23 Ma), whe n th e migratio n o f th e flexura l basin an d wedg e fron t slowe d significantly ; n o growth structure s developed an d shortenin g was accommodated withi n th e thicknes s o f th e orogenic wedge ; n o stres s woul d hav e propa gated int o the foreland; (3) a n Early-Mid-Miocene stage (16-11 Ma), when th e syste m remaine d i n muc h th e sam e state as in the previou s stage , wit h the wedg e

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front stationary , s o that littl e o r n o stres s would have propagated int o the foreland; (4) a Lat e Miocene-Pliocen e stag e (11 3 Ma): aroun d 1 1 Ma th e oute r orogeni c wedg e effectively collapse d an d compressiona l defor mation concentrated on the Jura fold belt , which detached o n high-leve l Triassi c evaporites , t o accommodate som e 30k m o f N W shortening ; beneath th e Jur a decollement , stres s withi n the lithosphere is likel y to have propagated int o the foreland. In additio n to Alpine tectonics, extensio n and subsidence withi n th e Europea n Rif t System , involving th e Lowe r Rhenis h Basin , th e Rhine , Rhone an d Limogne Graben , occurre d mainl y in Oligocene time through to Miocene tim e (Meie r & Eisbache r 1991) . A s th e grabe n ar e oriente d north-south, riftin g i s inferre d t o hav e resulted from east-wes t tension. As show n i n Fig . 9 , inversio n event s i n southern Britai n ca n b e correlate d i n tim e with events i n th e Alp s t o explain : (1 ) th e uplif t an d erosion o f th e Chal k betwee n 6 8 an d 6 0 Ma a s due to stres s generated b y convergenc e between Africa an d Europe ; (2 ) th e uplif t o f th e Weal d and subsidenc e o f th e Londo n Basi n an d th e uplift of the Sandown Pericline and subsidence of the Hampshire-Dieppe Basin in a succession of

Fig. 9. Correlation of Tertiary tectonic events between southern Britain and the Alpine Foreland. Isle of Wight

uplift phase s 1 and 2 identified by Gale e t al. (1999) .

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short pulse s o f deformation a s a consequence o f Alpine stresse s tha t could no t be accommodate d within the North Alpin e Forelan d Basin, stag e 1 ; (3) furthe r uplif t o f th e Weald , generatio n o f thrusts extendin g int o the Londo n Basi n cappe d by thrus t anticline s (e.g . Windsor), inversio n o n faults i n th e Wesse x Basi n an d tightenin g o f folds, foldin g o f Eocen e strat a (e.g . Isl e o f Wight), associate d wit h th e compressio n o f th e Jura fol d bel t an d northwar d propagatio n o f stress, stag e 4. The timin g o f strike-sli p faul t movement s o n the Sticklepath Fault (and probably the WatchetCothelstone Fault ) coul d correlat e wit h th e generation o f th e Europea n Rif t Syste m an d it s associated stres s field . Alternatively , bot h the y and the inversion of the Celtic Se a basins may be associated with changes i n spreading rate s in the North Atlantic . Nort h o f th e Varisca n Thrus t Front, th e S t George s Channe l Basi n i s no t inverted bu t is instead th e site of a Tertiary basi n flanking th e Nort h Celti c Se a Basin . Furthe r north, Robert s e t al (1999 ) identifie d majo r compression event s o f Eocen e t o Oligicen e ag e along th e Nort h Atlanti c margi n betwee n th e West Shetland s an d V0rin g tha t correspon d t o changes in spreadin g rate s an d plate boundarie s in th e Norwegian-Greenland Sea. Inversion event s i n th e souther n Nort h Sea, separated fro m souther n Britai n b y th e Londo n Platform, appea r t o hav e acte d independently . For example, Badley et al. (1989) interpreted the stratigraphic an d structural evidenc e depicte d o n seismic section s t o demonstrat e tha t a Lat e Cretaceous (8 8 Ma) compressiona l even t inverted th e NW-SE-trendin g Sout h Hewit t Fault (Fig . 1) , resultin g i n typica l harpoo n structures, an d a secon d compressiona l even t in Miocene time superimpose d a further inversion . Tertiary landscape evolutio n across southern Britain In hi s revie w o f Tertiar y landscap e evolution , Jones (1999 ) dre w attentio n to th e developmen t of a sub-Paleogen e surfac e (68-6 5 Ma) upo n which ther e develope d a residua l soi l cove r indicative o f tropica l conditions . Thi s i s pre served acros s th e Haldon Hills , near Exeter, an d plateau gravel s ar e preserve d i n Dorse t an d across Salisbur y Plai n (Gree n 1985) . Jone s agrees wit h Green tha t the preservatio n o f thes e soils an d th e lac k o f terrigenou s sedimen t offshore mean s that , throughou t Paleogen e time, souther n Englan d forme d a lan d are a o f low relie f an d lo w elevatio n (Green , pers . comm.), excep t i n localize d area s o f tectoni c

uplift. I n Neogen e time , thi s low-lyin g ground was floode d fro m th e eas t b y a marin e transgression a t 2. 6 Ma i n whic h th e Re d Cra g was deposited , succeede d b y th e Norwich Crag. These deposits have been subsequently deformed into basins and swells so that the Red Crag is now at — 50 m Ordnanc e datu m (OD) in Eas t Anglia but ove r H - 10 0 m O D a t th e wester n en d o f it s outcrop, 110k m t o th e west , nea r Bishop' s Stortford. Thi s i s jus t par t o f mor e extensiv e evidence o f Lat e Pliocene-Pleistocen e uplift , tilting an d deformatio n o f souther n England , which ha s place d th e lan d surfac e o f th e Chal k Downs o f S E Englan d a t 200 m OD , risin g t o surface elevation s o f 450 m O D i n th e wes t o f England. Uplif t o f Thame s terrace s (Madd y 1997) i s consistent with this general rise , which from thei r datin g indicate s a n uplif t rat e o f 0.07 mm a"1 durin g the past 2 Ma i n the middle Thames Valle y area . Walsh et al (1999 ) reviewed the evidence for a set of planation surfaces at elevations between 50 and 150 m O D developed acros s wester n Britain and Ireland , an d linke d the m wit h a numbe r o f early Neogen e sedimen t outlier s an d saprolit e bodies. Fro m this , they identified a sub-Neogene surface drape d abov e th e presen t topograph y o f western Britai n and Ireland that they regard a s a land surface with subdued topography of vertical relief measure d i n ten s o f metres . Thi s i s supported by a map o f Neogene an d Quaternary uplift o f norther n Englan d prepare d b y Frase r et al . (1990) . Th e pictur e thu s emerge s o f a British landscap e o f lo w elevatio n an d subdue d topography throug h Neogen e tim e unti l abou t 2.6 Ma whe n a broad-scal e uplif t occurred , culminating acros s th e wester n sid e o f England and Scotland . The exten t o f thi s uplif t ha s bee n quantified b y Clayto n & Shamoo n (1999) , fro m an analysis of the topography of Britain based on a 1 km squar e grid , a s a regional uplif t o f u p t o 300-400 m acros s Wales , th e Pennines , Lak e District an d Souther n Upland s of Scotland , an d up t o 500-600 m across th e Scottis h Highlands. In particular , the y calculate d tha t th e uplif t o f mountain peak s wa s amplifie d b y th e isostati c response t o denudationa l offloading, create d b y the dee p dissection o f topograph y relate d t o th e most resistan t rocks . Thei r calculatio n o f denudational isostati c uplif t account s fo r approximately hal f th e actua l mea n elevation , leaving the remaining uplift t o be explained. The late Neogene uplift an d deformatio n of southern England appear s t o hav e bee n par t o f a larger scale uplif t tha t affecte d al l th e lan d are a o f Britain. Clearly, there are exceptions, highlighted by Battiau-Quenc y (1999) , suc h a s th e 500-800 m contrast in relief between Anglesey,

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where elevation s ar e clos e t o se a level , an d th e mountains o f Snowdonia, whic h ar e separated by a major fault along the Menai Straits. Differential uplift ca n b e relate d t o boundin g fault s that ar e seismically active , indicativ e o f moder n differ ential movements . The mountain s o f Norwa y resulte d mainl y from lat e Plio-Pleistocen e uplif t an d ar e deepl y incised, wit h denudatio n accompanie d b y hig h sedimentation rate s i n offshor e basins , wher e deposits ar e u p t o 1500 m thic k (Rii s & Fjeldskaar, 1992) . Befor e 2. 5 Ma, th e lan d surface was one of low relief. Riis and Fjeldskaar calculated th e amoun t o f isostati c uplif t t o b e expected fro m denudationa l offloading , usin g flexural rigidit y wit h a n effectiv e elasti c thick ness of 22 km. They also allowed for the effect of phase change s a t dept h resultin g fro m pressur e changes owin g t o loadin g an d offloadin g a t surface. Thei r calculate d uplif t surfac e account s for les s tha n 70% o f the observe d elevation . In northern France, a general uplift is observed from th e Armorica n bloc k t o th e Rhin e Grabe n (Guillocheau et al 2000) . Differentia l uplif t i s responsible fo r large-scale relie f developmen t of the Armorica n massi f durin g Quaternar y tim e (Bonnet e t al 2000) . Lagarde e t al. (2000) hav e demonstrated ho w small-scal e deformatio n structures confir m tha t differentia l uplif t i s controlled b y faul t movements . The land areas of Britain, northern France an d Norway hav e rise n durin g th e pas t 2. 5 Ma t o create th e present-da y uplands . Fo r a lon g tim e before tha t th e landscap e wa s generall y o f lo w relief and low elevation. The uplands are all areas where har d rock s existe d a t surfac e befor e th e uplift 2. 5 Ma ago . Climat e chang e t o glacia l conditions increase d geomorphologica l energy , erosion an d dissectio n o f th e landscap e i n har d rock areas . Isostati c reboun d fro m denudationa l offloading enhance d topographi c rang e an d raised th e summi t level , bu t ca n accoun t fo r only 50-70 % o f th e observe d uplift . Uplif t i s differential an d faul t controlled . Presen t seismi city indicates that a number of the faults involved are currentl y activ e a s norma l faults . Th e caus e of th e uplif t i s unclear . I t i s difficul t t o explai n simply a s continenta l margi n uplif t originatin g from bod y force s a t th e continent-ocea n crus t boundary, a s thi s forme d wes t o f Britai n 3 0 Ma ago. No r ca n i t b e attribute d easil y t o plat e margin stresse s fro m th e Alp s o r Mid-Atlanti c Ridge, transmitte d laterally , whic h wer e muc h greater earlie r durin g Tertiar y tim e tha n i n th e past 2. 5 Ma. Differentia l erosio n betwee n har d and sof t rock , creatin g a differenc e i n surfac e elevation o f 400 m betwee n uplan d an d troug h from a n initial flat surface, produces a deviatoric

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stress fro m differentia l offloadin g o f lOMPa . This appear s t o b e sufficien t t o induc e faul t movements (Hardebec k & Hauksson 2001 ) and , hence, differentia l uplift . Thi s effec t i s on e o f strain relief, in addition t o isostatic rebound , and is aki n t o roc k burst s i n mines , bu t o n a large r scale. I t i s climaticall y induced , resultin g fro m the much increased rate s o f erosion fro m glacia l processes. Bu t it remains to be seen whether this is sufficien t t o accoun t fo r th e hithert o unex plained differenc e between th e observe d surfac e uplift an d th e amoun t calculate d a s du e t o isostatic rebound. The author is deeply indebte d to D. L. Hansen for using the forwar d modellin g progra m whic h h e an d S. Nielsen ha d develope d t o simulat e th e evolutio n of the Weald , Londo n an d Hampshire-Diepp e basins , and t o S . Nielse n fo r invitin g hi m t o a worksho p a t Aarhus i n Jun e 1999 , whic h provide d th e impetu s fo r much o f th e wor k i n thi s paper . H e ha s enjoye d an d benefited greatl y fro m stimulatin g discussion s wit h colleagues J . Rose, C . Green an d K. McClay, an d from helpful advic e fro m M . Ford, D. Roberts, S . Egan an d M. Butler . H e i s ver y gratefu l t o th e U K Onshor e Geophysical Librar y fo r th e provisio n o f seismi c sections across th e Weald, and to the Leverhulme Trus t for th e awar d o f a n Emeritus Fellowship , whic h mad e this research possible .

References BADLEY, M.E., PRICE , J.D . & BACKSHALL , L.C. 1989 . Inversion, reactivate d fault s and related structures : seismic example s fro m the souther n Nort h Sea . In: COOPER, M.A . & WILLIAMS , G.D . (eds ) Inversion Tectonics. Geologica l Society , London , Specia l Publications, 44 , 201-219. BATTIAU-QUENCY, Y . 1999 . Crusta l anisotrop y an d differential uplift : thei r role in long-term landfor m development. In : SMITH , B.J. , WHALLEY , W.B . & WARKE, PA . (eds ) Uplift, Erosion an d Stability: Perspectives on Long-term Landscape Development. Geologica l Society , London , Specia l Publications, 162 , 65-74 . BEELEY, H.S . & NORTON , M.G . 1998 . Th e structura l development o f th e Centra l Channe l High : constraints fro m sectio n restoration . In : UNDERBILL, J.R . (ed. ) Development, Evolution and Petroleum Geology of the Wessex Basin. Geological Society , London , Specia l Publications , 133, 283-298 . BIRPS & ECORS, 1986. Dee p seismi c reflectio n profiling between England , Franc e an d Ireland . Journal o f the Geological Society, London, 143 , 45-52 . BONNET, S. , GUILLOCHEAU , F. , BRUN , J.-P . & VA N DEN DRIESSCHE , J . 2000 . Large-scal e relie f development relate d t o Quaternar y tectoni c uplif t of a Proterozoic-Palaeozoic basement: the Armor ican Massif , N W France . Journal o f Geophysical Research, 105, 19273-19288 .

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CENOZOIC INVERSIO N AN D UPLIFT OF SOUTHERN BRITAI N UNDERBILL, J.R . (ed. ) Development, Evolution and Petroleum Geology of the Wessex Basin. Geological Society , London , Specia l Publications , 133, 187-197 . LEFORT, J.P . & MAX , M.D. 1992 . Structur e o f th e Variscan bel t beneat h th e Britis h an d Armorica n overstep sequences . Geology, 20 , 979-982. MADDY, D . 1997 . Uplift-driven valle y incisio n an d river terrac e formatio n i n souther n England . Journal o f Quaternary Science, 12 , 539-545. MEIER, L . & EISBACHER , G.H . 1991 . Crusta l kinematics an d dee p structur e o f th e Norther n Rhine Graben. Tectonics, 10, 621-630. MILIORIZOS, M . & RUFFELL, A . 1998 . Kinematics and geometry o f th e Watchet-Cothelstone-Hatc h Fault System: implications for the structural history of th e Wesse x Basi n an d adjacen t areas . In : UNDERHILL, J.R . (ed. ) Development, Evolution and Petroleum Geology of the Wessex Basin. Geological Society , London , Specia l Publications , 133,311-330. NIELSEN, S.B . & HANSEN , D.L . 2000 . Physica l explanation o f th e formatio n an d evolutio n o f inversion zone s and marginal basins. Geology, 28, 875-878. PETRIE, S.H., BROWN , J.R., GRANGER, P.J. & LOVELL , J.P.B. 1989 . Mesozoic histor y o f th e Celti c Se a Basins. In: TANKARD , A.J . & BALKWILL , H.R. (eds) Extensional Tectonics and Stratigraphy of th e North Atlantic Margins. Memoirs , America n Association of Petroleum Geologists, 46,433-444. PFIFFNER, A . 1992 . Alpine orogeny . In : BLUNDELL , D.J., FREEMAN , R . & MUELLER , S . (eds ) A Continent Revealed: the European Geotraverse. Cambridge Universit y Press , Cambridge , 180-190. RAWSON, P.P . 1992 . Th e Cretaceous . In : DUFF , P.McL.D. & SMITH , A.J. (eds) Geology of England and Wales. Geologica l Society , London , 355-388 . Rus, F. & FJELDSKAAR, W. 1992. O n the magnitude of the Lat e Tertiar y an d Quaternar y erosio n an d it s significance fo r th e uplif t o f Scandinavi a an d th e Barents Sea . In : LARSEN , R.M. , BREKKE , H. , LARSEN, B.T . & TALLERAAS , E . (eds) Structural and Tectonic Modelling and its Application to Petroleum Geology. Norwegian Petroleum Societ y (NPF) Specia l Publication , 1, 163-185.

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ROBERTS, D.G . 1989. Basin inversio n i n an d aroun d the Britis h Isles . In : COOPER , M.A . & WILLIAMS , G.D. (eds) Inversion Tectonics. Geological Society , London, Specia l Publications, 44 , 131-150. ROBERTS, D.G. , THOMPSON , M. , MITCHENER , B. , HOSSACK, J. , CARMICHAEL , S . & BJORNSETH , H.-M. 1999 . Palaeozoic t o Tertiar y rif t an d basi n dynamics: mid-Norwa y t o th e Ba y o f Biscay— a new contex t fo r hydrocarbo n prospectivit y i n th e deep wate r frontier . In : FLEET , A.J. & BOLDY , S.A.R. (eds ) Petroleum Geology o f Northwest Europe: Proceedings of the 5th Conference. Geological Society , London, 7-40 . ROWLEY, E . & WHITE , N. 1998 . Inverse modellin g o f extension an d denudatio n i n th e Iris h Se a an d surrounding areas . Earth an d Planetary Science Letters, 161 , 57-71 . UNDERBILL, J.R . & PATERSON , S . 1998 . Genesis o f tectonic inversio n structures : seismi c evidenc e fo r the developmen t o f ke y structure s alon g th e Purbeck-Isle o f Wigh t disturbance . Journal o f the Geological Society, London, 155 , 975-992. UNDERHILL, J.R. &STONELY , R. 1998 . Introduction to the development, evolution and petroleum geology of the Wessex Basin. In: UNDERHILL, J.R. (ed.) The Development, Evolution and Petroleum Geology of the Wessex Basin. Geologica l Society , London , Special Publications , 133, 1-18 . WALSH, P. , BOULTER , M . & MORAWIECKA , I . 1999. Chattian an d Miocen e element s i n th e moder n landscape o f wester n Britai n an d Ireland . In : SMITH, B.J., WHALLEY , W.B. & WARKE , PA. (eds) Uplift, Erosion and Stability: Perspectives on Long-term Landscape Development. Geologica l Society, London , Special Publications, 162,45-63 . WHITE, N. & LOVELL, B. 1997. Measuring the pulse of a plume with the sedimentary record. Nature, 387, 888-891. ZIEGLER, PA . 1990 . Geological Atlas o f Western and Central Europe. 2n d edition ; Shel l Internationale Petroleu m Maatschappij , Th e Hague. ZIEGLER, P.A. , CLOETINGH , S . & VA N WEES, J.-D. 1995. Dynamic s o f intra-plat e compressiona l deformation: th e Alpin e forelan d an d othe r examples. Tectonophysics, 252, 7-59.

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Landforms an d uplift in Scandinavia K. LIDMAR-BERGSTROM & J. O. NASLUND Department of Physical Geography and Quaternary Geology, Stockholm University, SE-106 91 Stockholm, Sweden (e-mail: karna@ natgeo.su.se) Abstract: Th e relatio n betwee n Scandinavia n landform s an d Cenozoi c uplif t event s i s examined by analysis of digital elevation data in a regional geological context as well as in a geomorphological proces s perspective . Re-expose d fla t sub-Cambria n an d sub-Mesozoi c hilly relief aids in deciphering uplift and erosional events. The highly dissecte d mountains of the Norther n Scande s (NS ) rise maximall y 1500 m abov e a slightl y tilte d lowes t leve l continuing i n the Muddus plain s eastwards at 300-550 m above sea level (a.s.L). This level is correlated with the lowest, slightly warpe d leve l of the Palaeic relief at 1000-1300 m a.s.L of th e Souther n Scande s (SS) , over whic h mountain s o f simila r heigh t rise . Thi s lowes t surface i s though t t o b e th e en d resul t o f Paleogen e erosio n t o th e genera l bas e level . Northern Scandinavi a wit h th e N S an d th e Muddu s plain s acte d a s a bloc k tha t wa s progressively tilte d t o th e SE , wherea s th e Souther n Scande s experience d continuou s doming, wit h a majo r uplif t even t o f abou t 1000 m i n Neogen e tim e causin g dee p valle y incision i n the uplifted plateau . Th e Sout h Swedis h Dome emerged from it s Palaeozoic and Mesozoic cover in Neogene time an d still retains well-preserve d re-expose d palaeosurfaces.

Uplift alon g continenta l margin s ha s latel y become a topi c o f commo n interes t fo r geologists, geochronologist s an d geomorpholo gists (Japse n & Chalmer s 2000 ; Summerfiel d 2000). Scandinavi a i s locate d clos e t o th e Atlantic margin . Its large-scal e relie f i s charac terized b y thre e domes , th e Norther n Scande s (NS) reachin g abou t 2000 m abov e se a leve l (a.s.L), th e Souther n Scande s (SS ) reaching 2500 m in south Norway, and the South Swedis h Dome (SSD ) reachin g 375 m i n sout h Swede n (Fig. 1) . Th e uplif t o f th e Scande s ha s bee n discussed sinc e th e beginnin g o f th e centur y (Reusch 1901 ; Ahlman n 1919 ) and Neogen e uplift o f th e S S ha s latel y bee n supporte d b y a fission-track stud y (Rohrma n et al 1995) . It has been suggeste d o n differen t ground s tha t th e domes have different uplif t historie s with a main uplift i n Paleogen e tim e i n th e nort h an d i n Neogene tim e i n th e sout h (Rii s 1996 ; LidmarBergstrom 1999) . I n thi s paper w e examine and compare th e topograph y o f th e dome s i n mor e detail an d correlate surface s between th e dome s to reveal areas of Neogene uplift and subsidence. Methods The relie f wa s examine d b y analysin g heigh t layer maps, slop e map s and topographic profiles . All map s an d profile s wer e constructe d fro m a digital elevatio n mode l (DEM ) o f Scandinavia .

For Sweden the elevation data have a true spatial resolution o f 50 0 m X 500 m (Nationa l Lan d Survey o f Sweden) , wherea s origina l elevatio n data fo r surroundin g area s ha d a resolutio n o f 1000m X 1000m (Staten s Kartver k i n Norway, and ETOPO5). Th e latter data were subsequently resampled to 500 m X 500 m. Both resolutions are suitable for the study of large-scale morphology. The major relief features are described with the aid o f a heigh t laye r map . Local topograph y i s analysed fro m slop e map s an d evaluate d i n a regional geologica l contex t a s wel l a s i n a geomorphological proces s perspectiv e following recent advances in knowledge on the effect of deep weathering in the shaping of relief (Thomas 1994). Three majo r palaeosurface s ar e identified an d correlated between the domes: the sub-Cambrian peneplain, a Mesozoi c surfac e and a Tertiar y surface. Th e profile s ar e located t o elucidate the suggested correlation s o f palaeosurface betwee n domes an d surroundin g terrain . Further , th e different degree s o f valle y dissectio n o f th e domes ar e examined. Analysis of maps Major shape of the Northern and Southern Scandes The Norther n Scande s for m an elongated dome , 1000 km long and 270 km or 165 km wide (at the

From: DORE , A.G., CARTWRIGHT, J.A, STOKER , M.S. , TURNER , J.P. & WHITE , N . 2002 . Exhumation o f th e North Atlantic Margin: Timing, Mechanisms an d Implications for Petroleum Exploration. Geologica l Society , London, Special Publications, 196, 103-116. 0305-8719/027$ 15.00 © The Geological Societ y of London 2002.

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Fig. 1. relief of Scandinavia. The following features, expressed in the landscape as a result of etching of geological strutures, should be noted: the Dellen meteorite impact (D), the Silijan ring meteorite impact (S), and

the plutoni c rocks o f th e Osl o rift . Th e location s o f Fig s 3- 6 an d 8 are indicate d by squares . MTFZ, M0re Tr0ndelag Faul t Zone; H , Hardangeroidda; O , Otta Valley; G, Gud brandsdalen valley.

LANDFORMS AN D UPLIFT IN SCANDINAVIA

600m o r 1000 m levels ) (Fig . 1) . The Souther n Scandes for m a mor e ova l an d somewha t ben t dome, 68 0 km lon g an d 400 o r 265 km wid e (a t the 600 m o r 1000 m levels) . Thu s th e S S ar e about 100k m wider and 300km shorte r than the NS. The N S ar e cu t b y valley s i n a NW-S E direction, leaving intact interfluves above 1000m a.s.l. wit h a maximu m widt h o f 20km . Th e

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maximum length of the interfluves is 40 km. This is in contrast to the SS, which have an area with 430km length from sout h to north above 1000 m a.s.l. Thi s elevated are a i s cut only by two major valleys, the Gudbrandsdalen valley and its majo r tributary valley , the Ott a valle y (Fig . 1) . Here, i t is possible to walk on the so-called Palaei c relief for ove r 300 km without descending into a valley below th e 1000 m level . I n th e east-wes t t o

Fig. 2 . Domes an d generalized palaeosurfaces of Scandinavia. The Palaeic relief i s located withi n the SS above about 1000 m a.s.l . (se e Fig . 1) . Locatio n fo r profile s 1 , 2 an d 3 i n Fig . 7 ar e indicated . Modifie d fro m Lidmar-Bergstrom (1999) .

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SE-NW directions , th e interfluve s ar e her e unbroken fo r u p t o 130km . Th e origi n o f th e Palaeic relief is discussed below . Thus th e dome s o f th e Norther n Scande s an d the Souther n Scande s sho w characteristicall y different shape s i n term s o f width , lengt h an d valley incision . I n addition , th e mai n valley s i n both area s are widened an d deepened t o varying degrees b y glacia l erosion . Scandinavian palaeosurfaces formed by etching and planation Four characteristi c landscap e types , forme d b y etching and planation, were studie d i n the maps : (1) sub-Cambrian peneplain; (2) undulating hilly relief (etc h surfaces) ; (3 ) plain s wit h residua l hills (Muddu s plains) ; (4 ) th e Palaei c relie f o f southern Norwa y (Fig . 2). Sub-Cambrian peneplain. Withi n th e Precam brian par t o f Fennoscandi a th e bedroc k wa s

denuded t o a n almos t leve l plai n a t th e en d o f Proterozoic tim e (Hogbo m 1910) . Vendian , Cambrian an d Ordovicia n strat a wer e succes sively deposite d directl y o n th e fla t basement . The surfac e wa s t o som e exten t overridde n b y Caledonian nappe s i n the west an d buried below thick cover s o f sedimentar y strat a i n th e eas t during lon g period s o f tim e (Koar k e t al. 1978 ; Zeck e t al . 1988 ; Lidmar-Bergstrom 1995; Cederbom e t al . 2000) . Thi s surfac e ha s bee n re-exposed an d ove r larg e area s i t ha s bee n totally obliterate d b y subsequen t denudation. In other areas it is still well preserved, and is called the sub-Cambria n peneplain . I t i s encountere d more o r les s intac t i n easter n an d south-centra l Sweden fro m se a leve l t o ove r 300 m a.s.l . (Lidmar-Bergstrom 1988 , 1996) and als o i n contact wit h Cambria n cove r rock s o n Hardangervidda, sout h Norway , a t abou t 1100-1350m a.s.l . (Schipul l 1974) . Th e subCambrian peneplain is met with along the eastern part o f th e NS , i n th e sout h at abou t 300m a.s.l and i n th e nort h a t ove r 1000 m a.s.l . (see

Fig. 3 . Slope an d height layer map of the SSD in combination with surrounding cover rocks . SSP . South Smaland Peneplain. A generalize d pictur e o f th e exten t o f th e sub-Cambria n peneplai n is show n b y th e blac k line . I t i s almost intac t in the S E up to 300m a.s.l . I n the norther n par t o f the ma p i t reaches fro m belo w Cambria n cove r rocks at 200m a.s.l. to summits furthe r sout h a t 350m a.s.l. Smal l area s i n the west have lo w relief an d Cambria n fissure fillings indicating long-lastin g Cambria n cover . Exhume d sub-Cretaceou s hilly relief extend s from belo w Cretaceous cove r rock s i n th e sout h an d west . Coordinate s ar e fro m th e nationa l grid o f Sweden .

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Fig. 4 . Slope ma p o f central Swede n an d southeastern Norway. Features t o be noted ar e the overall undulating , hilly relie f o f etc h character , th e Silja n Rin g (meteorit e impact) , th e volcani c rock s o f th e Osl o rift , an d th e exhumed sub-Cambrian peneplain with distinc t faults i n the southeaster n corner.

Ljungner 1950) . Th e presen t lan d surfac e i n Precambrian rock s i n Swede n coincide s wit h or is a s muc h a s 600 m belo w thi s surface , whic h can b e looke d upo n a s th e primar y peneplai n (Lidmar-Bergstrom 1995 , 1996) . Where th e subCambrian peneplai n i s wel l preserve d th e land scape show s a n extremel y fla t topograph y without residual hill s (Fig s 3 and 4). Undulating hilly relief. Alon g the eastern flan k of th e Souther n Scande s th e topograph y i s hilly and th e difference s i n bedrock compositio n an d structure ar e well expresse d i n the relie f (Fig s 1 and 4) , jus t a s withi n th e re-expose d sub Cretaceous relie f i n southernmos t Swede n (Fig. 3) . Th e latte r landscap e wa s forme d i n Late Mesozoi c tim e b y dee p weatherin g (etch ing) an d subsequen t strippin g o f th e weathering mantle (Lidmar-Bergstro m 1989) . Eas t o f th e Southern Scande s th e resul t o f denudatio n o f meteoric impact s such o as th e Lat e Palaeozoi c (Bottomley etal 1978;AberginWickman n 1988 ) Dellen structur e (Fig . 1 ) an d th e Cretaceou s (Deutsch et ai 1992 ) Silja n ring (Figs 1 and 4) is

clearly see n i n th e topography . Th e resistan t plutonic rock s i n th e Osl o fiel d giv e ris e t o massive hill s surrounde d b y lo w area s wit h Palaeozoic sedimentar y rocks (Figs 1 and 4). The relief a t th e souther n ti p o f Norwa y i s ofte n interpreted t o hav e a rathe r well-preserve d subCambrian surfac e o n interfluve s betwee n a fe w major joint-aligned valley s (e.g . Rii s 1996) . Th e peneplain i s no t intac t an d alon g th e coas t th e relief i s o f undulating character (Rudber g 1960 ) and classifie d a s a n etc h surfac e (Lidmar Bergstrom e t al. 2000) . Alon g th e wes t coas t of south Norwa y th e geologica l structure s ar e well expressed i n the relief, and this is mainly the case along th e entir e coas t o f wester n Norway . Mos t of th e easter n flan k o f th e Souther n Scande s i s located i n Swede n an d thi s par t contain s on e of Sweden's or e provinces . Thes e ore s ar e calle d soft ore s becaus e o f th e dee p weatherin g the y have experience d (Vivallo & Broma n 1993) . Other claye y weatherin g residue s ar e know n from thi s area , a s wel l a s alon g th e coas t o f Norway (Lidmar-Bergstro m e t al . 1999) . Dee p weathering an d subsequen t strippin g o f sapro lites ar e a majo r caus e fo r th e expressio n o f

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geological structures in the relief (Thoma s 1994) . The undulating hilly relief o n the eastern flank of the SS is thus interpreted t o be of etch character , maybe o f Mesozoic ag e (Reusc h 1903 ; Lidmar Bergstrom 1995 ; Cederbo m e t al 2000) . Th e structurally controlled relie f alon g the west coast of Norway is also thought to have an etch origin, but wit h furthe r incisio n alon g th e etche d structures afte r uplift . Th e Mesozoi c ag e o f th e etching eas t o f th e S S i s no t confirme d bu t possible (Cederbo m e t al 2000) . Along the coast of S W Swede n th e basemen t surfac e emerge s from belo w Mesozoi c strata , whic h dat e th e etching here . Grus saprolites an d landforms. Beside s th e remnants of kaolinitic saprolite s associate d wit h the re-exposed Mesozoi c undulatin g hilly relief, gravelly saprolite s (grus ) ar e o f commo n occurrence withi n man y part s o f Fennoscandi a (Lidmar-Bergstrom e t al 1999) . The y ar e interpreted t o hav e develope d mainl y i n Plio Pleistocene tim e bu t ma y dat e bac k t o Miocen e time. Withi n sout h Swede n the y ar e ofte n associated wit h a n etche d landscap e forme d a t the expens e o f th e re-expose d sub-Cambria n peneplain (Lidmar-Bergsto m e t al. 1997) .

Plains with residual hills. I n contras t t o th e dome flanks in the south, the relief east of the NS is characterized b y plains with residual hills, the so-called Muddu s plains , situate d mainl y a t 300-550 m a.s.l. (Fig. 5) (Wrak 1908) . Th e hills rise t o 300 m abov e th e plains . Th e plain s ar e interpreted t o hav e develope d fro m undulatin g hilly relie f (etc h topography ) b y pedimentatio n processes durin g mor e ari d period s o f Tertiar y time (Lidmar-Bergstro m 1995) . Th e plain s have acted a s base level s for the valleys that penetrate the mountain s i n th e wes t (Fig . 5). Individual plains ar e separate d b y lo w step s int o a number of separat e level s (Rudber g 1954 ; Lidmar Bergstrom 1996) . Simila r plain s (th e Sout h Smaland Peneplain , SSP ) wit h relativel y fe w and lo w residua l hills occu r at th e southwestern flank o f th e SS D abov e th e exhume d subCretaceous hill y relie f (Fig . 3). Th e plain s here cut of f th e re-expose d sub-Cretaceou s etche d relief an d ar e thu s o f Tertiar y ag e (Lidmar Bergstrom 1982) . Palaeic surface (or relief) of southern Norway. Traveller s i n souther n Norwa y ca n observe tha t mos t o f th e highe r groun d abov e about 1000 m a.s.l . is occupied by a high plateau

Fig. 5 . Slop e ma p o f region i n northern Swede n showin g th e Paleogene Muddu s plains in yellow (300-700 m a.s.l.) i n th e eas t an d thei r continuatio n a s valley s int o th e mountain s in th e west . Th e lin e mark s th e borde r between Precambria n basement an d Caledonian nappe s (se e Fig. 6).

LANDFORMS AND UPLIFT IN SCANDINAVI A

(Fig. 6). From the west deep valleys cut far inland into thi s plateau , name d th e Palaei c surfac e b y Reusch (1901) . The difference betwee n th e high plateau an d th e deepl y incise d valley s wa s interpreted t o reflec t a late uplif t (Reusc h 1901 ; Ahlmann 1919 ; Peulvas t 1978 , 1985) . Th e Palaeic surfac e i s separate d int o differen t level s thought to be induced by renewed valley incision as a resul t o f continue d warpin g durin g severa l Mesozoic-Paleogene uplif t event s (Lidmar Bergstrom e t al. 2000) . Th e Palaei c surfac e i s composed o f severa l surface s separate d b y distinct step s an d i s bette r referre d t o a s th e Palaeic relief. Th e lowes t leve l o f th e Palaei c relief, situated at about 1000-1200 m a.s.L, has a vast extent on the eastern side of the high dome in southern Norway . Thi s leve l ha s acte d a s bas e level for the river systems penetrating the higher ground (Fig s 1 an d 6 ) an d ca n b e followe d westwards alon g th e majo r rive r systems . Th e 1000m leve l occur s als o alon g the western sid e of sout h Norwa y (Lidmar-Bergstro m e t al . 2000). In detail, the Palaeic relief is characterized by slightl y undulatin g plains wit h residua l hill s with surroundin g pediment s (Fig . 6) . Dee p weathering an d pedimentatio n processe s ar e thought to have been importan t in the formation of thi s relief (Gjessin g 1967) .

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Zone of incised valleys Around the plateau wit h Palaeic relief i n the SS there is a zone with deeply incised valleys (Figs 1 and 6) . Th e valley s ar e deepe r o n th e wester n side, wher e the y exten d belo w se a level . A difference i n relief of up to 800 m between valley bottoms o f incise d valley s an d adjacen t lowes t shallow valle y leve l o f th e Palaei c relie f i s common. Th e valley s are glacially deepene d b y more tha n 1000 m i n th e Sognefjor d o n th e western, Atlanti c side , an d b y c . 250 m o n th e eastern side , e.g . i n th e Mjos a are a (Ahlman n 1919).

Comparison between the lowest level of the Palaeic relief and the Muddus plains The lowest level of the Palaeic relief constitutes a plateau strikingl y simila r t o th e Muddu s plain s but a t a highe r leve l (Fig s 5 an d 6) . The y bot h have acte d a s bas e level s fo r rive r system s penetrating westwards . I n detail , ther e ar e differences betwee n th e two , whic h ca n b e explained by differences i n bedrock composition. In th e Palaei c relie f depicte d i n Fig . 6 , th e

Fig. 6 . Slope and height laye r map showin g the lowes t leve l o f the Palaeic relief o f southern Norwa y i n yellow (900-1200m a.s.l.) . Th e lin e mark s th e borde r betwee n Precambria n basemen t an d Caledonia n nappe s (se e Fig. 5) .

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southern par t ha s develope d b y erosio n o f th e Caledonian nappe s an d now is mainly shape d i n Precambrian rocks . Th e associate d residua l hill s often includ e remnant s o f th e Caledonia n cover . Further t o th e N E (Fig . 6 ) the Palaei c relie f ha s formed i n highl y variabl e Caledonia n rocks , which hav e resulte d i n a somewha t differen t morphology wit h some resistant rocks givin g rise to hig h summits . I n contrast , th e correspondin g Muddus plain s in northern Swede n (Fig . 5 ) were formed o n Precambria n rocks , whic h ha d lon g since los t their Palaeozoic cover . In addition, the different appearanc e o f the relief i n Figs 5 and 6 is als o t o a minor exten t th e resul t o f th e us e of elevation dat a wit h differen t resolution s i n th e construction o f th e maps . However , mos t o f the differenc e i n appearanc e i s cause d b y th e difference i n geology . It i s likel y tha t th e Muddu s plain s o f northern Swede n an d th e lowes t leve l o f th e Palaeic relie f o f souther n Norwa y develope d at th e sam e leve l an d a t th e sam e time . Th e individual form s ar e th e en d resul t o f etching , stripping an d pedimentatio n i n war m climates . This unifor m lan d surfac e wa s subsequentl y deformed b y differentia l uplift . This , includin g the timin g o f events , i s furthe r discusse d below. Analysis of profile s The location s o f th e profile s (Fig. 2 ) ar e chose n to illustrat e th e correlatio n an d deformatio n o f the thre e palaeosurfaces , namely , th e sub Cambrian peneplain , th e Mesozoi c undulatin g hilly relief , an d the Tertiary plain s wit h residua l hills (correspondin g t o th e lowes t leve l o f th e Palaeic relief) . Th e vertica l positio n o f th e palaeosurfaces i s illustrate d i n th e profile s an d used fo r a discussio n o n uplift . Profile s 1 and 2 are use d t o discus s th e relationshi p betwee n th e SSD and the SS, as well as the deformation of the sub-Cambrian peneplain , th e inferre d Mesozoi c surfaces, an d the lowest Palaei c level . Profile 3 is used fo r discussing th e N S an d it s eastern flank . Profile 1 (Figs 2 and 7) The sub-Cambrian peneplai n extend s from below Cambrian cove r rock s i n th e SE . The peneplai n forms a low dome, constituting the SSD, which is interrupted b y the Vattern Graben , fille d wit h up to 1000 m o f Late Proterozoi c sedimentar y roc k (Axberg & Wadstei n 1980) , an d th e Hokensa s Horst. Palaeozoi c remnant s occu r o n th e slope s down t o Lak e Vaner n bu t d o no t occu r o n th e bottom o f th e lake . Furthe r t o th e N W th e peneplain i s dow n warped an d form s th e Vane r

Basin. Fro m Lak e Vaner n th e sub-Cambria n peneplain rises to the NW and can be followed in the summit s fo r som e distance . Profil e 1 the n follows a mor e westerl y directio n (Fig . 2) . Th e sub-Cambrian peneplain is met with again below Palaeozoic strat a i n th e Osl o rif t an d the n encountered a t abou t 1250-1300 m a.s.l . o n th e Hardangervidda. Here i t is identified wit h the aid of Palaeozoi c outliers . Th e peneplai n rise s towards the NW and thereafter it is downfaulted, where Caledonia n rock s mee t th e Precambria n basement along a steeply dipping front a t H in the profile (Fig . 7) . Northwest o f Lak e Vaner n th e sub-Cambria n peneplain i s replace d b y a hill y relie f (a n etc h surface), whic h i s tentativel y interprete d a s a Mesozoic surfac e (see above). This hilly relief is correlated wit h a n inferre d surfac e alon g th e highest summit s furthe r t o th e west . Th e weathering-resistant plutoni c rock s o f th e Osl o rift exten d above thi s surface. The present lowest level of the Palaeic relief at Hardangervidda partl y coincide s wit h bu t i s mainly slightl y belo w th e sub-Cambria n pene plain. I n th e westernmos t part s shallo w valley s are incise d i n th e lowes t level , her e a t abou t 1100m a.s.L , t o slightl y belo w 1000 m a.s.l . Structurally controlle d dee p valley s o f th e Hardangerfjord syste m penetrat e th e Palaei c relief i n th e NW . Dee p valleys , Tinnsj o an d Numedalen, ar e incise d alon g th e southeaster n flank. Profile 2 (Fig. 7) Profile 2 is identical to profile 1 in the SE but then continues straight towards the NW (Fig. 2). In the area o f hilly relief, NW of Lake Vanern , the sub Cambrian peneplai n ha s disappeare d an d i s me t with agai n belo w th e Palaeozoi c strat a i n th e Oslo rift. Thereafter, it directly disappears belo w the Caledonian rocks and has no further influenc e on the presen t topograph y alon g the profile. In this profile a correlation has also been mad e between th e undulatin g hill y relie f an d th e highest summit s o f th e SS . Th e inferre d Mesozoic surfac e ha s experience d domin g an d been uplifte d abou t 2000 m i n th e NW . Subsequently, i t has bee n successivel y dissected by major valleys, which has resulted in a stepped pattern o f th e presen t summi t surfaces (Lidmar Bergstrom e t al. 2000) , indicate d b y th e thre e summit level s in th e profile . A slightl y warpe d surfac e i s see n a t abou t 1000-1200m a.s.l. , dissecte d b y dee p valley s (Eikesdalen drainin g westwards ; Gudbrandsda len an d othe r valley s drainin g southeastwards). This i s th e lowes t leve l o f th e Palaei c relief ,

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Fig. 7 . Topographic profiles showin g Scandinavia n dome s and palaeosurfaces. Locatio n of profiles i s shown i n Fig. 2 . Profile 1 crosses th e souther n par t o f the SS , Vaner Basin an d SSD . B , Bergen ; H , Hardangerfjord ; Ha , Hardangervidda. Profil e 2 crosse s th e norther n par t o f SS , Vane r Basi n an d SSD . E , Eikesdalen ; G , Gudbrandsdalen. Profil e 3 crosses the NS and the Muddus plains in the east. O, Ofotfjorden; N , Norddalen; Cal , Caledonian.

comparable wit h th e lowes t leve l o f Hard angervidda in profile 1 . Above this lowest level, mountains rise to heights o f 1500-200 0 m a.s.l . The highest peaks of south Norway reach 150 0 m above thi s lowes t level , an d ar e indicate d b y a point showin g thei r vertica l positio n i n th e profile. Befor e the profile reache s th e Atlantic it crosses th e M0re-Tr0ndela g Faul t Zon e (MTFZ). Profile 3 (Fig. 7) The exhumed sub-Cambrian peneplain is seen at the easter n coas t an d ca n b e followe d i n som e

summits toward s th e west . Th e inferre d sub Cambrian peneplai n i s ben t down , belo w th e Caledonian rocks . The eastern parts of the NS are here formed in Precambrian basement rock with the Caledonian nappes followin g westwards. The highest mountains ar e forme d o f rock s relativel y resistan t t o deep weatherin g an d their summi t surfaces may date bac k t o Mesozoic tim e (Lidmar-Bergstro m 1996). Mesozoi c surface s hav e tentativel y been placed across the highest peaks along the profile. A lowes t Mesozoic(?) leve l i s tentatively shown as a surface inclined t o the east alon g the whole profile (se e Wra k 1908) .

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A fe w step s ar e incise d belo w th e sub Cambrian surfac e clos e to the coas t in the eas t and the n follo w westwar d th e mai n plain s wit h residual hill s (th e Muddus plains), which exten d to th e Scandes . Withi n th e mountain s thi s leve l can b e followe d a s th e bas e fo r valle y incision . Mountains ris e t o 1500 m abov e thi s level . Comparison and correlation of the Scandinavian domes from the profile data Profile 3 (Fig . 7 ) show s th e highl y dissecte d mountains o f the NS rising fro m a slightly tilted lowest level , continuin g i n th e Muddu s plain s eastwards. I n th e ma p analysi s (Fig . 5) , th e Muddus plain s wer e correlate d wit h th e lowes t level o f the Palaei c relie f i n the S S (Fig. 6) , an d they ar e show n a s the sam e surfac e i n profiles 1 and 2 (Fig . 7) . Th e correlatio n betwee n th e Muddus plains and the lowest level of the Palaei c relief i s strongl y supporte d b y th e fac t tha t th e vertical distanc e fro m thes e level s t o the highes t summits o f th e Scande s ar e 1500 m i n bot h regions (Fig . 7 , profile s 2 an d 3) . Thi s suggest s that th e tw o dome s wer e o f th e sam e heigh t before th e majo r uplift o f the SS . Discussion Uplift of the Scandes Riis (1996 ) suggeste d tha t th e uplif t alon g th e western margi n o f Fennoscandi a di d no t occu r simultaneously in the north and the south. On the basis o f offshor e geology an d identificatio n of a Paleogene surfac e it was suggested tha t the main uplift i n th e nort h occurre d i n Paleogen e time , whereas i t too k plac e i n Neogen e tim e i n th e south. Lidmar-Bergstrom (1999 ) agreed with this interpretation, a s a resul t o f analysi s o f th e relation between the mountains and the two types of relief tha t occur along the eastern flanks of the NS an d the SS . The Northern Scandes and their eastern flank. Th e shap e o f th e N S an d th e steppe d morphology o f th e Muddu s plain s indicate s tilting toward s th e east . Th e mai n relie f withi n the mountain s i s th e resul t o f valle y incisio n since th e initiatio n o f uplift , wherea s th e successively widene d oute r valley s forme d i n a tectonically relativel y stabl e environmen t clos e to a genera l bas e leve l i n th e eas t wit h th e Muddus plains as the end result. It is possible that the Precambria n surface , wher e no t i n contac t with th e sub-Cambria n peneplain , wa s expose d during Mesozoic tim e and that the Muddus plains

ultimately wer e forme d i n Paleogen e time , a s indicated by finds of redeposited marin e Eocen e diatomaceans (Cleve-Eule r 1941) . Continuin g apatite fission-trac k analysi s (AFTA) in the are a will shed mor e ligh t on this. So far, these studies suggest a til t o f th e are a alon g th e profil e beginning i n Cretaceou s tim e an d acceleratin g in Paleogen e tim e (Hendrik s & Andriesse n 2001). A ver y hig h relativ e relie f occur s wes t of th e Ofotfjorde n faul t line , and , i n detail , the relie f i s highl y irregular . I t i s likel y tha t this wa s originall y a n etc h surfac e o f Mesozoic (o r mayb e older ) age , a s Mesozoi c (or older? ) rock s ar e stil l preserve d i n a downfaulted basi n o n a kaolinitize d basemen t surface (Stur t e t al 1979) . Uplif t ha s cause d incison o f th e etche d structures . Analysi s o f fission-track dat a indicate s severa l phase s o f vertical movemen t i n thi s regio n (Hendrik s & Andriessen 2002) . The Southern Scandes. Th e surfac e form s o f the S S ar e cu t acros s bot h Caledonia n an d Precambrian basement . I n th e souther n par t o f the dome the present surface is relatively close to the sub-Cambria n peneplain , wherea s i n th e north th e sub-Cambria n peneplai n disappear s below th e Caledonia n nappes . Th e Mesozoi c surface inferre d in th e profile s indicate s that th e line throug h th e highes t summit s represent s a warped surfac e o f thi s age . I n th e Osl o rif t an d southeastwards i t become s mor e likel y tha t th e present relie f i s par t o f a re-expose d sub Cretaceous etc h surface . Apatit e fission-trac k modelling show s that this interpretation is viable (Cederbom e t al . 2000) . Profil e 2 crosse s th e MTFZ. Thi s i s a n ol d Caledonia n structur e that has subsequentl y bee n reactivate d (Dor e e t al . 1999). I t i s possibl e tha t vertica l movement s along th e MTF Z hav e affecte d th e Mesozoi c surface an d contributed to a slight asymmetry of the SS. The lowest level of the Palaeic relief has a position a t abou t 1000-1200 m a.s.l . compare d with th e tilte d plai n fro m whic h th e N S ris e a t 300-700 m a.s.l . I t i s a slightl y domed surface , which support s the ide a o f a continuation of th e Mesozoic doming . Th e Neogen e uplif t o f th e SS ca n thu s b e estimate d t o b e abou t 1000 m (Fig. 7 an d Lidmar-Bergstro m e t al . 2000) . Th e valleys alon g th e coas t ar e mainl y structurally controlled an d i t i s suggeste d tha t th e dee p incision followe d structure s etche d ou t durin g Mesozoic time . A major hinge line across central Scandinavia? A lin e fro m th e MTF Z toward s th e

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113

Fig. 8 . Slope map of centra l Scandinavia. The M0re-Tr0ndelag Fault Zone (MTFZ) an d the fault line s at the east coast of Sweden sugges t a hinge line in a SW-NE direction. This lin e marks the border betwee n th e undulating hilly relie f t o the sout h an d the Muddus plains t o the north.

WSW-ENE-trending faul t line s o f th e sub Cambrian peneplai n i n nort h Swede n approxi mately coincide s wit h th e border betwee n th e plains with residual hills (the Muddus plains) and the undulatin g hilly relief (Fig . 8) . In the region where thi s lin e crosse s th e mountai n chai n th e Precambrian basement has a low position (300 m a.s.l.)- Thi s i s i n contras t to bot h i n th e N S an d SS, wher e th e basemen t reache s abov e 1000 m a.s.l. Thi s lin e ma y approximatel y mar k th e border betwee n area s wit h differen t uplif t histories.

the cove r i s therefor e suppose d t o hav e bee n eroded befor e th e downwarp . A Neogen e ris e of th e SS D a t th e Oligocene-Miocen e bound ary wit h subsequen t erosio n o f it s cove r i s supported b y a fission-trac k stud y o f sout h Sweden (Cederbo m 2002 ) an d th e observatio n of a large amoun t o f erosio n o f Cretaceou s an d Paleogene strat a fro m nearb y Jyllan d an d th e Skagerrak-Kattegat Platfor m (Japse n & Bidstrup 1999) .

Summary of uplift an d relief development South Swedish Dome and Vdner Basin The SS D developed it s present shap e in Tertiary time, whe n th e basemen t successivel y wa s exposed afte r erosio n o f Palaeozoic cove r in th e north an d eas t an d Mesozoi c cove r i n th e sout h and wes t (Lidmar-Bergstro m 1991 , 1993) . W e suggest tha t i n connectio n wit h th e Neogen e uplift o f th e SS , th e Vane r Basi n wa s slightl y depressed an d th e SS D uplifted . N o Palaeozoi c rocks occu r o n th e botto m o f Lak e Yaner n an d

The Precambrian shiel d of Scandinavia has been covered by Palaeozoic rocks . In some parts they were erode d durin g Mesozoi c tim e an d etche d surfaces with undulating hilly relief were formed in th e basement . Mesozoi c denudatio n als o caused relie f developmen t acros s Caledonia n rocks an d th e Palaeozoi c cover . Jurassi c an d (probably mainly ) Cretaceou s strat a wer e the n deposited ove r the area to an unknown extent. In some case s the y wer e deposite d directl y on the etched basemen t an d protecte d it s Mesozoi c

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relief fo r a lon g time . Larg e part s o f th e Caledonian area s experience d continuou s relie f development withou t an y Mesozoi c temporar y cover. A t the end of the Paleogene period, plains with residua l hill s ha d forme d ove r th e easter n parts o f Scandinavi a wit h valle y system s penetrating mountainou s area s i n th e west . Th e mountains reache d abou t 1500 m abov e thi s lowest leve l o f th e Palaei c surface . I n souther n Norway th e sub-Cambria n peneplai n wa s successively re-expose d an d th e Precambria n basement her e experience d furthe r relie f development. The tectoni c uplif t differe d betwee n th e N S and th e SS . Th e N S mainl y acte d a s a block , which experience d simpl e tilt , wherea s th e S S were characterize d b y doming . Neogen e uplif t caused a sligh t furthe r til t i n th e nort h an d a major uplif t wit h continue d domin g i n th e S S with re-exposur e o f sub-Mesozoi c relie f alon g their flanks . Th e deepl y incise d valleys of the SS were mainl y forme d i n Neogen e tim e a s a consequence o f uplift . Subsequent , Lat e Cenozoic ic e shee t erosio n ha s furthe r widene d and deepene d thes e valleys. Neogene uplif t wit h a centr e sout h o f Lak e Vattern cause d th e formatio n o f th e SSD . Th e buried sub-Cambrian peneplai n was successively re-exposed. Th e re-exposure o f this surface at the top o f th e dom e cause d reactivatio n o f th e weathering system s alon g th e fractur e systems , which wer e successivel y expresse d i n th e topography a s joint aligne d valleys . Re-exposure of etched Mesozoi c relie f i n the S W caused total stripping o f th e remain s o f th e kaoliniti c saprolites b y th e SW-flowin g drainag e system , down t o c . 125 m a.s.l . Her e th e SS P gradually formed, governe d b y se a level a s the bas e fo r it s erosion. It s detailed form s probably developed i n semiarid climate s tha t promote d pedimentatio n (Lidmar-Bergstrom 1988) . Neogen e gru s sapro lites ar e stil l o f commo n occurrenc e i n th e eastern par t o f th e dom e abov e thi s leve l an d testify t o continue d etching. The lates t rise , probabl y i n Pliocen e time , caused th e re-exposur e o f th e sub-Cretaceou s etch surface s below 12 5 m a.s.l . in the sout h and west, an d successiv e re-exposur e o f th e extre mely flat sub-Cambrian rock surfaces in the north and east . Conclusions (1) Th e main uplift o f the Northern Scande s is older tha n th e uplif t o f the Souther n Scandes , a s the NS are considerably more dissected by deep valleys. A s a resul t o f th e Mesozoi c Paleogene uplift an d til t o f th e NS , th e Muddu s plain s

formed an d the y acte d a s bas e level s fo r th e development o f the deep valley s of the mountain range. I n Neogen e tim e th e are a experience d additional mino r uplif t an d tilting. (2) Th e mai n uplif t o f th e Souther n Scande s took place i n Neogene time , the uplift amounting to c . 1000m . Th e Neogen e uplif t wa s a continuation o f Mesozoic-Paleogene doming . (3) Th e Muddu s plains and their continuation in the valleys of the NS are tentatively correlated with th e lowes t leve l o f th e Palaei c relie f o f th e SS, because of strong similarities in morphology. (4) A hing e fro m th e MTF Z t o N E Swede n separates th e N S wit h th e Muddu s plain s fro m the S S an d th e undulatin g hill y relie f o n it s eastern flank. (5) I n Neogene tim e the South Swedish Dom e was elevated , wit h subsequen t developmen t o f the Sout h Smalan d Peneplain . A t th e sam e tim e the Va'ne r Basi n was downwarped . (6) Geneticall y interprete d landform s ar e important dataset s i n morphotectoni c analyses , complementary t o studie s o f th e sedimentar y records an d thermotectoni c evolutio n o f th e bedrock. The stud y was supporte d b y a grant fro m th e Swedis h Natural Scienc e Researc h Council . W e als o wan t t o thank P . Japsen, S . A. Cloeting h an d P . Andreissen fo r encouraging discussion s an d input . Elevatio n mode l data ove r Swede n courtes y o f Swedis h Nationa l Lan d Survey 2000 . Excerp t fro m GSD-elevatio n database , case no . L2000/646 . Elevatio n mode l dat a ove r Norway courtes y o f Staten s Kartverk . 3504 H0nefoss . Norway.

References AHLMANN, H.W . 1919 . Geomorphologica l studie s i n Norway. Geografiska Annaler. I. 1-20 . AXBERG, S . & WALDSTEIN . P . 1980. Distributio n of th e sedimentary bedroc k i n Lak e Vattern . souther n Sweden. Stockholm Contributions i n Geology. 3 4 (2). 15-25 . BOTTOMLEY, R.J. . YORK , D . & GRIEVE . R.A.F . 1978 . 40 Ar- 38 Ar ages of Scandinavian impact structures: I Mien and Siljan. Contributions to Mineralogy an d Petrology, 68 , 79-84. CEDERBOM, C. 2002. The thermotectonic development of southern Sweden durin g Mesozoic an d Cenozoi c time. In: DORE , A.G., CARTWRIGHT , J.A., STOKER , M.S., TURNE R & J.P., WHITE, N. (eds ) Exhumation of the North Atlantic Margin: Timing, Mechanisms and Implications for Petroleum Exploration. Geological Society . London . Specia l Publications . 196, 169-182 . CEDERBOM, C. , LARSSON , S.E., TULLBORG , E.-L . & STIBERG, J.-P . 2000 . Fissio n trac k thermochronol ogy applie d t o Phanerozoi c thermotectoni c events

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LIDMAR-BERGSTROM, K . 1999 . Uplif t historie s revealed b y landform s o f th e Scandinavia n domes. In : SMITH , B.J. , WHALLEY , W.B . & WARKE, PA . (eds ) Uplift, Erosion an d Stability: Perspectives on Long-term Landscape Development. Geologica l Society , London , Specia l Publications, 162 , 85-91 . LIDMAR-BERGSTROM, K. , OLLIER , C.D. & SULEBAK , J.R. 2000 . Landforms and uplift history of southern Norway. Global an d Planetary Change, 24 , 211-231. LIDMAR-BERGSTROM, K. , OLSSON , S . & OLVMO , M . 1997. Palaeosurface s an d relate d saprolite s i n southern Fennoscandia . WIDDOWSON , M . (ed. ) Palaeosurfaces: Recognition, Reconstruction and Palaeoenvironmental Interpretation. Geologica l Society, London , Specia l Publications . In : vo l 120,95-123. LIDMAR-BERGSTROM, K., OLSSON, S. & ROALDSET, E . 1999. Relie f feature s an d palaeoweatherin g remnants i n formerl y glaciate d Scandinavia n baemen t areas. In : THIRY , M . & SiMON-CoiNgoN , R . (eds) Palaeoweathering, Palaeosurfaces and Related Continental Deposits. Internationa l Associatio n o f Sedimentologists, Specia l Publications , 27 , 275-301. LJUNGNER, E. 1950 . Urbergsytan s for m vid fjall randen. Geologiska Foreningens i Stockholm Forhandlingar, 72, 269-300. PEULVAST, J.P . 1978. Le bourrele t Scandinav e e t le s Caledonides: u n essa i d e reconstitutio n de s modalites d e l a morphogenes e e n Norvege . Geographic Physique Quaterniaire, 32, 295-320. PEULVAST, J.P . 1985 . Postorogeni c morphotectoni c evolution o f the Scandinavia n Caledonide s durin g the Mesozoi c an d Cenozoic . In : GEE , D.G . & STURT, B.A . (eds ) Th e Caledonide Orogen — Scandinavia an d Related Areas. Wiley, Chichester, 979-995. REUSCH, H . 1901 . Nogle bidra g til l forstaaelse n a f hvorledes Norge s dal e o g fjeld e e r blevn e til . Norges Geologiske Unders0gelse, Aarbog (1900), 32, 124-263. REUSCH, H . 1903 . Glommen s bojnin g ve d Kongs vinger. Norges Geografiske Selskab, Aarbog, 14 , 96-102. Rus, F . 1996 . Quantificatio n o f Cenozoi c vertica l movements o f Scandinavi a b y correlatio n o f morphological surface s wit h offshore data . Global and Planetary Change, 12 , 331-357. ROHRMAN, M. , VA N DER BEEK , P. , ANDRIESSEN , P .

& CLOETINGH , S . 1995 . Meso-Cenozoi c morphotectonic evolutio n o f souther n Norway : Neogene doma l uplif t infere d fro m apatit e fission trac k thermochronology . Tectonics, 14 , 704-718. RUDBERG, S . 1954 . Vdsterbottens berggrundsmorfologi. Geographica , 25 . RUDBERG, S . 1960 . Geology an d geomorphology . In : S0MME, A . (ed. ) A Geography o f Norden. J . W . Cappelens, Oslo , 27-40 . SCHIPULL, K . 1974 . Geomorphologische Studien i n zentral Siidnorwegen mit Beitrdgen uber Regelungs- und Stuerungssysteme in der Geo-

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morphologic. Hamburge r Geographische r Studien , 31. STURT, B. , DALLAND , A . & MITCHELL , J . 1979 . The ag e o f th e Sub-Mid-Jurassi c tropica l weathering profil e o f And0ya , norther n Norway , and th e implication s fo r th e Lat e Palaeozoi c palaeogeography i n th e Nort h Atlanti c region . Geologische Rundschau, 68 , 523-542 . SUMMERFIELD, M . 2000 . Geomorphology an d Global Tectonics. Wiley , Chichester . THOMAS, M.F . 1994 . Geomorphology in th e Tropics. Wiley, Chichester . VIVALLO, W. & BROMAN , C. 1993 . Genesi s of the earthy ore s a t Garpenberg , sout h central Sweden .

Geologiska Foreningens i Stockholm forhandlingar, 115 , 209-214. WICKMANN, F.E . 1988 . Possibl e impac t structure s in Sweden. In : BODEN , A . & ERIKSSON , K.G . (eds ) Deep Drilling i n Crystalline Bedrock. Springer , Berlin, 1 , 299-327. WRIK, W . 1908 . Bidra g til l Skandinavien s reliefkro nologi. Ymer,28, 141-191 . ZECK, H.P. , ANDRIESSEN , P.A.M. , HANSEN , K. , JENSEN, P.K. & RASMUSSEN, B.L . 1988 . Palaeozoi c palaeo-cover o f th e souther n par t o f th e Fenno scandian shield—fissio n trac k constraints . Tectonophysics, 149 , 61-66.

Pattern an d timing of the post-Caledonian denudatio n o f norther n Scandinavia constrained b y apatite fission-trac k thermochronolog y BART W . H. HENDRIKS & PAUL A . M. ANDRIESSEN Department of Isotope Geochemistry, Faculty of Earth Sciences, Vrije Universiteit Amsterdam, De Boelelaan 1085, 1081 HVAmsterdam, The Netherlands (e-mail: henb@ geo.vu.nl) Abstract: Apatit e fission-trac k thermochronolog y ha s bee n use d t o stud y th e post Caledonian denudatio n histor y o f norther n Scandinavia . Post-orogeni c denudatio n progressively shifte d fro m th e interior of the continent towards the North Atlantic margin . The present-da y are a o f maximu m elevatio n i n th e Northern Scande s mountai n rang e ha s experienced continuou s denudatio n a t leas t sinc e Jurassi c time . I n Jurassic-Cretaceou s time, the area north an d east of this region experienced either no denudation a t all or some denudation followed b y a transient thermal event wit h a peak temperature in late Cretaceous time. Final denudation of the area to the east of the Northern Scandes probably starte d in late Cretaceous-Paleogene tim e an d possibl y accelerate d i n Neogen e time . Th e denudatio n history o f northern Scandinavi a ca n b e explaine d by scar p retreat o f a n uplifte d rif t flank . The patter n an d timin g o f denudation o f the Northern Scande s is different fro m tha t o f th e Southern Scandes , whic h experience d domal-style, late-stag e postrif t uplif t i n Neogen e time. Geomorphologica l observations , offshor e dat a fro m th e Atlanti c an d Barent s Se a margins, and scarce stratigraphical information fro m th e mainland are in general agreement with th e new thermochronological data .

Because o f the almost complet e absenc e o f postCaledonian sediment s o n th e Scandinavia n mainland, denudatio n ('uplif t o f rock s relativ e to th e surface ' accordin g t o th e definitio n o f Summerfield & Brow n (1998) ) o f Scandinavi a has bee n studie d mainl y b y analysi s o f geomorphology (Lidmar-Bergstro m 1993 , 1999) an d it s correlatio n wit h th e offshor e geology (Rii s 1996) . Low-temperatur e thermo chronology ha s bee n applie d successfull y i n southern Scandinavi a (Lehtovaar a 1976 ; Andriessen & Bo s 1986 ; Zec k e t al. 1988 ; Rohrman 1995 ; Hansen et al. 1996 ; Larson et a l 1999; Cederbo m e t a l 2000) , bu t unti l now n o comparable stud y has been undertake n i n north ern Scandinavia . The geomorpholog y o f Scandinavi a ca n b e characterized a s a plateau wit h two large dome s in th e wes t an d nort h an d a smalle r dom e i n southern Swede n (Fig . 1) . Th e sout h Swedis h dome reaches 37 7 m above sea level and is partly covered b y Cambrian an d Mesozoic cove r rocks . The souther n Norwegian dome, usually referred to a s th e Souther n Scandes , reache s 2469 m above se a leve l o n Jotunheimen . Th e mor e elongated norther n Scandinavia n dome , usuall y referred to as the Northern Scandes an d the study area here, has a maximum elevation of 2113 m on Kebnekaise.

Riis (1996 ) an d Lidmar-Bergstro m (1999 ) both conclude d tha t th e denudatio n histor y o f the Souther n Scande s i s differen t fro m tha t o f the Norther n Scande s mountai n range . Accord ing t o both studies , th e lates t uplif t phas e i n th e north wa s earlie r (i n lat e Cretaceou s t o Paleogene time ) tha n i n th e sout h (i n Neogen e time). Thi s i s in agreemen t wit h apatit e fission track dat a fro m Rohrma n (1995 ) an d fro m th e present study . Accordin g t o Rohrma n (1995) , the centra l Souther n Scande s experience d Triassic-Jurassic erosio n o f 2. 4 ± l.lk m an d a Neogen e denudatio n o f 2. 0 ± 0. 5 km, decreasing radiall y outwar d t o les s tha n 0.5km nea r th e coastline . West of the Northern Scandes mountain range, the Lofoten and Vesteralen island groups (Fig. 2) are part o f a hors t an d grabe n syste m tha t i s structurally very complex and poorly understood. According t o Riis (1996) , final uplif t o f this are a occurred in Neogen e time , wit h a stron g Plio Pleistocene component. Invers e modelling o f the small amoun t o f apatit e fission-trac k dat a w e have fro m Lofote n an d Vesterale n a t present , indicates considerabl e Neogen e denudatio n a s well. A detaile d stud y applyin g apatit e fission track analysi s and (U-Th)/He thermochronometry t o thi s region , aimin g a t unravellin g th e denudation histor y o f th e variou s structura l

From: DORE , A.G., CARTWRIGHT , J.A, STOKER , M.S. , TURNER , J.P . & WHITE , N. 2002 . Exhumation o f th e North Atlantic Margin: Timing, Mechanisms and Implications for Petroleum Exploration. Geological Society , London, Special Publications, 196, 117-137 . 0305-8719/027$ 15.00 © The Geological Societ y of London 2002.

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Fig. 1 . Map o f Scandinavi a an d surroundin g areas , showin g th e locatio n o f th e stud y are a an d indicatin g the Southern Scandes , Norther n Scande s and Sout h Swedish Dome .

blocks an d datin g faul t movements , i s i n progress.

Geological settin g During th e Caledonia n orogen y th e Fennoscan dian shiel d wa s overthrus t fro m th e (north)wes t (Roberts & Ge e 1985) . A t present , severa l windows i n th e Caledonia n cove r rock s expos e parts o f th e Fennoscandia n shiel d t o th e wes t of the Caledonia n fron t o n Lofoten , Vesteralen , Senja an d t o th e S W an d S E o f Narvi k (Geological Survey s o f Finland , Norwa y an d Sweden 1987) . In th e immediat e post-orogeni c period , i n Devonian time , fault-bounde d molass e basin s developed i n th e precurso r Norwegian-Green land rif t syste m an d continenta l molass e continued t o accumulate i n half-grabens border ing the remnant Caledonian mountains until midPermian tim e (Dor e & Gag e 1987) . Crusta l extension i n th e Norwegian-Greenlan d rif t system accelerate d durin g lat e Permia n an d Triassic time and resulted in low-relief doming of the flankin g areas , particularl y durin g lat e Triassic tim e (Ziegle r 1987) . Jurassi c regiona l uplift o n th e norther n margins of Laurenti a and

Baltica an d mid-Jurassi c reactivatio n o f th e margins an d othe r sourc e area s withi n th e Atlantic Rif t domai n she d thic k coars e clasti c deposits ont o surroundin g area s (Dor e 1991) . Major riftin g affecte d th e entir e Norther n Atlantic domai n a t th e Jurassic-Cretaceou s transition (Faleide et al. 1993). In late Cretaceous time, th e Barent s Se a domai n wa s tectonicall y decoupled alon g th e Senja-Hornsun d faul t system fro m th e Atlanti c Rif t domai n wher e downwarping continue d (Dore 1991) . Following break-up in the Norwegian-Greenland Sea at the Paleocene-Eocene transitio n (Srivastav a & Tapscott 1986) , th e western Barent s Se a margi n developed as a shear margin (Faleide et al. 1993) . According t o Rii s (1996) , lat e Cretaceous early Tertiar y uplif t o f norther n an d wester n Fennoscandia cause d dee p erosion , an d Paleo gene uplif t reache d a maximum value of almos t 1500m i n norther n Scandinavia . Neogen e tectonic uplif t cause d domin g i n souther n Norway and was of the order of 1 -1.5 km in the area of highest topography (Rohrman 1995) . Lofoten and Vesterale n experience d tw o phase s o f Tertiary uplif t (Stuevol d & Eldhol m 1996) , with th e Neogen e uplif t componen t bein g o f the orde r o f 1000 m (Rii s 1996) .

DENUDATION HISTORY O F NORTHERN SCANDINAVIA

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Fig. 2 . Map of the study area showing topographic elevation , locations of samples and transects, an d other features referred to in this paper. Locations of faults and basins after Faleide et al. (1993), Olesen et al. (1997) and Brekke (2000).

Plio-Pleistocene glacia l erosio n cause d unloading an d uplif t o f Fennoscandi a an d modified th e olde r fluvia l drainag e syste m t o a considerable extent, but did not totally obliterat e the landform s tha t ar e typica l o f uplifte d landmasses alon g passiv e continenta l margin s (Lidmar-Bergstrom et al. 2000). Glacia l erosio n also affecte d Svalbar d (Elyth e & Kleinspeh n 1998) an d th e wester n Barent s Sea , whic h wa s subaerial i n preglacia l times . Tectoni c uplif t played an important role in this region before and possibly als o durin g th e glaciation s (Dimaki s et al . 1998) . Postglacia l crusta l domin g i s estimated t o hav e reache d a maximu m uplif t value o f 850 m i n th e centr e o f Fennoscandi a (Gudmundsson 1999) .

Apatite fission-track thermochronology Annealing concept Fission track s are damag e zone s in the crysta l lattice that are formed by the spontaneous fission of 238U (Wagner 1968; Fleischer et al 1975) . The density o f spontaneou s fissio n track s i s pro portional t o th e elapse d tim e an d th e uraniu m content. Except for rapidly cooled rocks, fissiontrack age s ar e systematicall y younger compared with other radiogenic age determinations. This is a result of the instability o f fission tracks at high temperatures. Within a mineral-specific temperature range , calle d th e partia l annealin g zon e (PAZ), th e track s begi n t o annea l until they are

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completely erase d a t th e uppe r temperatur e boundary. Fo r apatite, th e mos t commonl y use d mineral, th e PA Z i s usuall y define d a s a temperature rang e betwee n c . 6 0 °C an d c. 130° C (Gleado w & Dudd y 1981 ; Naese r 1981; Wagner & Van den Haute 1992) , but even at roo m temperatur e low-rat e annealin g take s place (Donelic k e t al 1990) . Th e temperatur e range o f th e PA Z wil l b e differen t fo r apatite s with differen t chemica l composition s (e.g . Carlson e t al . 1999) . Fo r extrem e composition s the uppe r boundar y o f th e PAZ , meanin g tota l annealing, ma y b e a s low a s 9 0 °C or a s high as 200 °C (Ketcha m e t a l 1999) . Fissio n track s formed i n th e PAZ , an d track s tha t hav e bee n heated int o th e PA Z afte r formatio n a t lowe r temperatures, hav e reduce d lengths , compare d with thei r initia l lengt h / 0 o f betwee n 15. 8 an d 16.6 |xm. Track lengt h reduction withi n the PAZ results i n a mea n lengt h shorte r tha n / 0 an d a negatively skewe d distributio n toward s th e smaller tracks . Th e annealin g characteristic s o f fission tracks thus allow the reconstruction of the thermal history of a sample from th e track length distribution (Gleado w e t a l 1986a , 1986b) . Reheating o f a sampl e b y a transien t therma l event ca n b e inferred , bu t i t i s difficult , o n th e basis o f fission-trac k dat a alone , t o differentiate this fro m a purel y coolin g history . Thi s i s because th e fission-trac k syste m i s essentiall y one tha t records cooling , an d eve n i n a thermal history that has included some reheating, most of the record wil l come fro m th e cooling segment s (Gleadow & Brown 2000). Analytical procedure After minera l separation , apatit e grain s wer e mounted i n epoxy , polishe d an d etche d i n 7 % HNO3 a t 2 0 °C t o revea l spontaneou s fissio n tracks. Etc h tim e varie d betwee n 2 5 an d 40s , depending o n the etching rat e o f the apatite . Fo r all sample s th e externa l detecto r metho d wa s applied usin g micas tha t were heate d fo r 4 8 h at 600 °C to eras e an y existin g tracks . Al l mounts were irradiate d togethe r wit h dosimete r glas s CN5 in the LFR facility of the ECN at Petten, the Netherlands. Micas coverin g th e apatite s wer e subsequently etche d i n 48 % H F a t 2 0 °C fo r 12 min. Micas coverin g th e dosimeter glass wer e etched i n 48 % H F a t 2 0 °C fo r 3 0 min. Th e counting o f fissio n track s i n th e apatite s an d external detector s an d th e measuremen t o f confined fission-trac k length s wer e performe d at 100 0 X magnificatio n usin g 10 0 X dr y objective wit h a numerica l apertur e o f 0.90 . Only apatites with sharp polishing scratches were used fo r counting and measuring.

Age and track length statistics The fission-track ages were calculated using the £ age calibration method (Hurfor d & Green 1983) . A Rvalu e o f 35 5 ± 1 5 (la) wa s obtaine d fro m Mt. Dromedary , Fis h Canyo n an d Durang o apatite standard s usin g CN 5 dosimete r glass . All sample s displaye d P(x*) > 5% , indicating that th e dat a ar e consisten t wit h a singl e population o f age s i n eac h sampl e (Galbrait h 1981). Therefor e al l fission-trac k age s i n thi s paper ar e reporte d a s poole d ages , wit h a l a error. Initia l fission-track length was determined on confine d induce d track s o f Fis h Canyo n apatite, givin g a mea n valu e of 16. 3 (Jim wit h a standard erro r o f 0. 1 fjim an d a l a valu e o f 0.49 [Jim . Inverse modelling of fission-track data Inverse modellin g o f th e ag e an d lengt h measurement result s wa s carrie d ou t usin g th e AFTSoLVE 1.2.l a progra m (Ketcha m e t a l 2000). Compare d wit h AFTSoLVE , man y other fission-track modellin g program s (Corriga n 1991; Lut z & Oma r 1991 ; Gallaghe r 1995 ) tend t o generat e simple r historie s becaus e the y implicitly o r explicitl y favou r paths wit h fairl y few degree s o f freedo m (Ketcha m e t a l 2000) . AFTSoLVE yields a much wider array of possible modelling results . Th e annealin g mode l o f Crowley e t a l (1991 ) fo r F-apatite s an d th e annealing model of Laslett et al (1987) , which is based o n laborator y studie s o f Durang o apatite , have been used for inverse modelling. Compared with th e annealin g mode l o f Crowle y e t a l (1991), th e annealin g mode l o f Laslet t e t a l (1987) i s les s sensitiv e t o low-temperatur e annealing an d overemphasize s th e importanc e of recen t coolin g (va n de r Bee k 1995) . Th e annealing model of Laslett et al (1987 ) therefore often reveal s therma l historie s wit h a rapi d lat e cooling event , whic h i n man y case s i s a modelling artefact . However , th e par t o f th e thermal histor y belo w c . 6 0 °C i s onl y loosel y constrained b y an y annealin g mode l an d on e should alway s b e ver y critica l abou t an y interpretations based on this part of the modelled thermal history . A t hig h temperature s th e annealing mode l o f Crowle y e t a l (1991 ) i s much more retentive than the annealing model of Laslett e t a l (1987) . Extrapolatio n o f th e annealing mode l o f Crowle y e t a l (1991 ) t o geological tim e scale s predict s F-apatit e t o b e more resistant to annealing than the more Cl-rich Durango apatite , i n contras t t o geologica l observations (Gallagher et al 1998) . The Laslett et al (1987 ) annealing model is the most widely

DENUDATION HISTOR Y O F NORTHERN SCANDINAVI A

used annealin g model , bu t th e Crowle y e t al. (1991) mode l fo r F-apatite s ha s bee n use d i n many fission-trac k studie s i n Scandinavi a (Rohrman 1995 ; Larso n e t al . 1999 ; Cederbo m et al. 2000). To compare ou r results with studies using eithe r th e Laslet t e t al . (1987 ) o r th e Crowley e t al. (1991 ) annealin g model, al l dat a have been modelled wit h both annealing models. In addition, i t is possible t o test the sensitivity of the therma l historie s obtaine d fro m invers e modelling t o the annealin g model applied . Modelling strategies for AFTSoLVE have been outlined b y Ketcha m e t al . (2000) . Initia l trac k length was fixed at 16.3 jjurn . No other constraints were used than a present-day surface temperature of 0-10 °C. For samples S21 and P6, which were collected i n th e LKA B Kirun a an d Malmberge t mines a t 765 m an d 815 m belo w th e surface , respectively, a present-da y temperatur e o f 10-25°C wa s used . Th e oute r envelop e o f th e thermal historie s presente d i n thi s pape r repre sents therma l historie s tha t giv e a valu e large r than 0.05 fo r both the age goodness-of-fit an d the Kolmogorov-Smirnov tes t fo r th e lengt h distribution. Thi s typ e o f history cannot be rule d ou t by th e dat a (Ketcha m e t al . 2000) . Th e inne r envelope represent s therma l historie s wit h a value larger tha n 0.5 for both th e ag e goodness of-fit an d th e Kolmogorov-Smirno v tes t fo r th e length distribution . Thi s typ e o f histor y i s supported b y th e data . Also , th e best-fittin g thermal histor y i s depicte d fo r al l model s presented i n thi s paper . Durin g th e firs t invers e modelling run s fo r eac h sample , wit h th e Crowley e t al . (1991 ) annealin g mode l a s wel l as with the Laslett et al. (1987) annealing model, AFTSoLVE wa s restricte d t o produc e purel y cooling historie s alone . I f modellin g wit h thi s restriction yielde d onl y therma l historie s tha t cannot b e rule d ou t b y th e data , o r very fe w thermal historie s tha t ar e supporte d b y th e data , many differen t tim e interval s o f reheatin g wer e tested fo r eac h sample . Specia l attentio n wa s given t o mak e sur e tha t th e boundarie s o f th e time interva l did not forc e the mode l resul t in a certain direction , whic h coul d exclud e possibl e thermal histories .

Sampling Most sample s ar e concentrate d i n tw o transect s through th e Norther n Scande s mountai n range . Transect A-A' (Fig. 2) runs from Andenes on the northernmost ti p o f th e Vesterale n island s towards th e Gul f o f Bothnia . Thi s transec t includes tw o vertica l profiles . On e i s o n Kebnekaise, rangin g fro m 57 5 to 1530 m abov e sea level. The four samples in this vertical profil e

121

are all within 8 km in the horizontal direction. No post-Caledonian o r neotectonic faul t activit y has been reporte d fo r thi s area . Th e secon d vertica l profile i n transect A-A 7 is in the LKAB mine in Kiruna. Th e thre e lowermos t sample s i n thi s profile, fro m —344 m t o 196 m abov e se a level, are almost perfectly vertically aligned insid e the mine itself . Th e to p tw o samples , a t 50 2 an d 712m above sea level, are about 3 km to the NE of th e min e o n a nearb y hill , Luossavaara . Neotectonic activit y ha s bee n reporte d fo r th e Kiruna are a (Dehl s e t al . 2000) , bu t agai n n o other post-Caledonia n faul t activit y ha s bee n reported fo r this area to our knowledge. Another subsurface sampl e i n this transect wa s collecte d inside the LKAB min e in Malmberget, a t 815m below th e surface , a t 191 m belo w se a level . Transect B-B 7 (Fig . 2 ) run s betwee n Troms0 , Norway, an d Muonio , Finland . Thi s transec t includes a vertica l profil e o n Tromsdalstinden , SE o f Troms0 , fro m se a leve l t o 1238 m elevation. Thi s vertica l profil e include s fiv e samples, an d except for the sampl e a t se a level , which is about 8 km to the west of the summit of Tromsdalstinden, al l sample s ar e n o mor e tha n 3 km from eac h other in the horizontal direction. Two SW-NE-trending faults of unknown age lie within 5-1 0 km t o th e S E an d N W o f Tromsdalstinden (Zwaa n e t al . 1998) . Th e on e to th e N W o f Tromsdalstinde n exhibit s neotec tonic activit y (Dehl s e t al . 2000) . However , w e do no t kno w o f an y post-Caledonia n fault s tha t would cut an d might distur b the vertical profile . Block rotatio n betwee n th e tw o faults, however , may hav e cause d rotatio n o f th e vertica l profil e through time . In addition to the samples of the two transects, many surfac e sample s wer e collecte d acros s th e study area . Som e o f thes e sample s wer e take n inside fjords, whic h can be up to 2km deep. The fjords ar e glacially overdeepened valley s and the original fluvial incision may be much older than the Plio-Pleistocen e glaciation s (Lidmar-Berg strom e t al . 2000) . Becaus e o f th e perturbatio n effect o f eroding topography on isotherms i n th e crust (Stiiw e e t al . 1994) , th e result s fro m samples that wer e collecte d inside fjord s canno t be immediatel y interprete d i n th e sam e wa y a s those fro m 'normal ' surfac e samples . Several sample s wer e collecte d especiall y t o investigate th e effect s o f a pronounced negativ e gravity anomal y centre d aroun d Sulitjelma , i n the S W o f th e stud y are a (Olese n e t al . 1997) . In contrast , th e regio n encompassin g R0st , Vaer0y an d Lofote n i s characterized b y a strong positive gravit y anomaly . Sample s hav e bee n collected fro m thi s region , bu t result s ar e no t available yet .

Table 1 . Fission-track results Mean Number trac k of lengt h grains (|xm )

SE S D MTL MT L ((Jim) (jxm )

Number of lengths Roc

13.1 11.6 13.3 12.3 12.9 13.7 12.9 12.8 13.0 13.5 12.9 13.0 12.8 13.0 13.4 13.3 12.6

0.2 1. 0.2 2. 0.2 1. 0.2 2. 0.2 0.2 0.3 0.3 0.2 0.3 0.3 0.3 0.2 0.1 0.1 0.2

2 .3 .6

86 175 51 101 100 58 23 36 100 20 23 17 127 109 139 100 100

Gneiss Gneiss Granite Schist Granite Granite Schist Schist Gneiss Dolerite Volcanite Volcanite Iron or e Iron or e Iron or e Gneiss Iron or e Granodiorite Granite

46 51 53 40 47 49 54 27

13.7 13.7 13.3 13.2 13.4 12.8 13.2 14.0

0.2 0.2 0.1 0.1 0.1 0.1 0.1 0.1

.0 .1 .0 .2 .3 .3 .3 .1

28 40 101 100 183 204 204 150

Gneiss Gneiss Gneiss Gneiss Gabbro Schist Gneiss Granite

32 34 34 53 34 41

13.8 _ 12.9 14.3 13.0 13.0

0.1 0. 0. 1 1 O.I 1. 0.1 1. 0. 1 1.

8

18 107 200 100 100

Granitoid Volcanite Granite Gneiss Phyllite Gneiss

Latitude (N)

Pooled F T Longitude ag e ± S E P()C) (E) (Ma ) (% )

PS (MJ (106 tracks cm^ 2 )

Pi (N\) 6

(10 tracks cm~ 2 )

tracks cm )

10 360 5 380 12 1530 1070 815 575 450 712 502 196 -29 -344 418 -191 430 40

69.31 69.28 68.56 68.52 68.42 67.93 67.88 67.87 67.84 67.99 67.88 67.87 67.84 67.84 67.84 67.65 67.18 67.13 66.31

16.08 16.01 16.44 17.89 17.78 18.54 18.54 18.57 18.75 19.93 20.23 20.21 20.18 20.18 20.18 21.00 20.67 20.57 22.82

128 ± 1 4 142 ± 1 4 180 ± 2 1 134 ± 1 6 124 ± 1 2 220 ± 2 5 205 ± 2 4 157 ± 2 1 113 ± 1 4 243 ± 2 3 303 ± 3 6 268 ± 2 9 225 ± 2 2 251 ± 3 0 212 ± 2 4 322 ± 3 6 264 ± 2 5 318 ± 2 9 268 ± 2 3

23 83 28 76 9 30 36 91 30 32 26 94 80 100 45 100 17 8 79

0.368 (396) 0.767(553) 0.330(180) 0.393 (240) 0.802(615) 0.339(395) 0.133(349) 0.155(182) 0.263 (207) 1.937(1375) 0.773 (404) 0.346(508) 1.019(784) 0.270(359) 0.503 (428) 0.627 (536) 1.570(1449) 3.196(2220) 0.847(508)

0.491(528) 0.924(666) 0.308(168) 0.501 (306) 1.104(846) 0.271(316) 0.110(288) 0.167(196) 0.396(312) 1.352(960) 0.427 (223) 0.227(333) 0.763 (587) 0.176(234) 0.400 (340) 0.317(271) 1.045(965) 1.695(1177) 0.510(306)

0.957(12033) 0.957(12033) 0.957(12033) 0.957(12033) 0.957(12033) 0.994(12054) 0.951(11870) 0.951(11870) 0.951 (11870) 0.957(12033) 0.951(11870) 0.994(12054) 0.951(11870) 0.927(14206) 0.951(11870) 0.927(14206) 0.994(12054) 0.957(12033) 0.927(12904)

30 30 25 22 28 30 36 22 21 55 22 52 26 37 23 28 31 33 20

Transect B-B1 1238 Tl 753 T2 530 T3 262 T4 5 T9 15 Fl 465 F51 F48 300

69.61 69.61 69.60 69.61 69.63 69.26 68.84 68.35

19.15 19.11 19.11 19.07 18.95 19.92 21.17 22.90

230 ± 241 ± 189 ± 203 ± 216 ± 170 ± 247 ± 365 ±

26 26 1 9 20 22 1 5 22 34

84 12 81 75 96 60 53 24

0.216(442) 0.139(522) 0.492(783) 0.714(864) 0.273(655) 2.509(3235) 0.667(2408) 1.416(1956)

0.159(324) 0.097 (364) 0.460(732) 0.622(753) 0.213(511) 2.492(3213) 0.454(1639) 0.647(894)

0.951(11870) 0.951(11870) 0.994(12054) 0.994(12054) 0.951(11870) 0.951(11870) 0.951(11870) 0.95 1 ( 11 870)

Other F6 Fll F16 F28 F34 F37

69.79 70.03 70.7 1 69.47 70.86 70.01

20.94 23.07 24.59 25.85 29.11 29.17

214 ± 2 4 268 ± 3 6 205 ± 1 9 297 ± 2 6 269 ± 2 7 306 ± 2 8

50 98 61 15 33 84

0.478 (384) 0.35 1 (229) 0.573(1617) 1.111(3415) 1 . 1 7 1 (774) 1.044(2418)

0.393(316) 0.230(150) 0.472(1332) 0.628(1930) 0.766(506) 0.597(1382)

0.994 ( 1 2054) 0.994(12054) 0.95 1 ( 11 870) 0.951 (11870) 0.994(12054) 0.994(12054)

Sample name

Elevation (m a.s.l. )

Transect A N33 N34 N39 N2 N5 K9613 Kl K2 K3 S13 S23 S24 LI S21 L2 S16 P6 S18 GRM3

-A '

1 40 5 140 10 15

2

0.2 :

8 2 1 0 .6 .2 .3 .9 .7 .4 .7 .0 .8

.4 2 1 1

k type

DENUDATION HISTORY OF NORTHERN SCANDINAVI A

123

Results Apatite fission-track analysis has been performed on 4 3 samples . Despit e th e generall y lo w uranium concentratio n i n th e samples , reflecte d by low pi values (Table 1) , counting statistics fo r all sample s ar e a t a n acceptabl e level . Al l samples displaye d P(x 1) > 5% , indicatin g tha t the data are consistent with a single population of ages i n eac h sampl e (Galbrait h 1981) . Th e standard erro r o f th e poole d ag e wa s generall y lower tha n 10 % and neve r mor e tha n 13% , an d for al l sample s mor e tha n 2 0 grain s coul d b e counted. Th e lo w uraniu m concentratio n did , however, mak e i t difficul t t o obtai n enoug h length measurement s o f horizonta l confine d tracks for severa l samples . One hundre d length measurements are considered necessary to obtain the statisticall y reliabl e lengt h distributio n needed fo r confidenc e in th e invers e modellin g of th e fission-trac k dat a (Gris t & Ravenhurs t 1992). Si x sample s containe d onl y betwee n 5 0 and 10 0 confined track s (N33 , N39 , Kl , N13 , N15 and N49) but nevertheless inverse modelling was als o undertake n fo r thes e samples . Th e results ar e ver y simila r t o thos e fro m sample s nearby tha t containe d 10 0 o r mor e confine d tracks. Therefor e thes e model s ar e include d i n this paper , bu t shoul d b e considere d supportin g evidence for the nearby samples and conclusions should no t b e base d o n thes e sample s individu ally. Fo r sample s tha t containe d fewe r tha n 5 0 confined tracks , onl y the mea n trac k length can be use d becaus e invers e modellin g woul d result in larg e time-temperatur e field s rathe r tha n time-temperature paths . Sample s tha t d o no t have any track length information in Table 1 did not contain enough apatite to prepare a mount for the length measurements . The oldes t fission-trac k age s (>250Ma ) ar e from sample s mos t distant from th e Atlantic and southwestern Barent s Se a margins . Th e fission track ages in transect A-Ar range from 11 3 ± 1 4 to 32 2 ±36Ma (Fig . 3a) . I n transec t B-B ' (Fig. 3b ) th e fission-trac k age s rang e fro m 170 ± 1 5 t o 36 5 ± 3 4 Ma. Th e along-transec t variation o f th e fission-trac k age s i n transec t A-A/ an d transec t B-B 7 clearl y indicate s a decrease o f the fission-trac k age s from S E to NW. Closer t o bot h margin s th e fission-trac k age s decrease, bu t mor e s o towards th e Atlanti c margin than towards the SW Barents Sea margin. As a resul t o f complicate d therma l histories , reflected b y comple x confine d trac k lengt h distributions, ther e i s n o obviou s patter n fo r th e mean track length in the stud y area. The highest mean trac k length s wer e obtaine d fro m sample s F28 an d F4 8 (14. 3 an d 14. 0 |xm, respectively )

124

B. W. H. HENDRIK S & P. A. M . ANDRIESSE N

Fig. 3. (a) Fission-track ages for samples in transect A-A 7 , (b) Fission-track ages for samples in transect B—B' .

and these two samples ar e also among th e oldes t samples in the stud y area . Figure 4 plot s fission-trac k age s v . elevatio n for th e vertica l profile s o f Tromsdalstinde n (Fig. 4a) , Kebnekais e (Fig . 4b) an d Kirun a (Fig. 4c) . All fission-trac k age s i n Fig . 4 ar e displayed wit h 2c r errors . Ther e i s n o obviou s trend fo r the fission-trac k age s v . elevation i n the vertical profil e o f Tromsdalstinde n (Fig . 4a) . Near-invariant apatit e fission-trac k age s ove r elevation ranges of 1 -2 k m have been interprete d as th e resul t o f ver y hig h erosio n rates , bu t ca n also resul t fro m coolin g o f th e footwal l durin g normal faultin g (Gallagher e t al. 1998) . Becaus e the ag e o f mos t brittl e fault s i n th e are a o f Tromsdalstinden i s unknow n (Zwaa n e t al . 1998), i t i s difficul t t o determin e wha t mechan ism i s responsibl e fo r th e near-invarian t fission track age s o f th e Tromsdalstinde n vertica l profile. Th e fission-trac k ag e v . elevatio n plot s

for th e vertica l profile s o f Kebnekais e (Fig . 4b ) and Kirun a (Fig . 4c) sho w a n increas e o f th e apatite fission-trac k age s wit h highe r elevation . Unfortunately, man y o f th e sample s fro m th e vertical profile s of Tromsdalstinden, Kebnekais e and Kiruna yielded only small amounts of apatite that generally also were of poor quality and had a low concentratio n o f uraniu m (reflecte d by lo w values fo r p { i n Tabl e 1) . Therefor e th e trac k length informatio n fro m thes e sample s i s ver y limited, whic h severel y reduce s th e amoun t o f information tha t ca n b e extracte d fro m th e vertical profiles , suc h a s a n estimat e o f th e (palaeo)geothermal gradient . Interpretation of fission-trac k dat a and inverse modelling Thermal historie s obtaine d fro m invers e modelling ar e non-uniqu e an d th e uncertainties

Fig. 4 . Fission-track ages (2cr error) v. elevation for the vertical profile s o f (a) Tromsdalstinden , (b ) Kebnekais e and (c ) Kiruna .

DENUDATION HISTOR Y O F NORTHERN SCANDINAVI A

in the solution s for individual samples ar e large. However, th e combinatio n o f modelle d therma l histories o f samples fro m differen t location s an d different elevation s make s i t possibl e t o recon struct a regiona l denudatio n histor y wit h confidence. Becaus e th e annealin g model s use d for invers e modellin g o f apatit e fission-trac k data, i n thi s stud y th e Laslet t e t al (1987 ) annealing mode l an d th e Crowle y e t al . (1991 ) annealing model , d o no t accuratel y mimi c th e behaviour of fission tracks at temperatures below 60 °C, this part of the modelled therma l history is not considered i n the interpretation . Post-orogenic cooling Inverse modellin g o f sample s fro m transect s A-A' and B-B/ indicates post-orogeni c cooling progressively shiftin g toward s th e wes t (Fig s 5 and 6) . Thermal historie s fro m th e Laslett e t al. (1987) annealin g mode l (Fig s 5 b an d 6b ) ten d toward slightl y lowe r temperature s an d indicat e cooling t o commenc e somewha t earlie r tha n i n the therma l historie s fro m th e Crowle y e t al . (1991) annealing model (Fig s 5a and 6a). Taking into account , fro m bot h annealin g models , th e thermal histories that are supported by the data, it is clea r tha t the easternmos t sample s in bot h transects sho w coolin g fro m Devonia n tim e onwards. The westernmost samples di d not coo l to temperatures within the PAZ before Triassicearly Jurassi c time . Thi s Triassic-earl y Jurassic timing probabl y i s relate d t o domin g o f th e flanking area s o f the Norwegian-Greenland rif t system (Ziegle r 1987) . Th e hig h value s fo r th e mean track length of samples F28 and F48 (14. 3 and 14.0jjim , respectively ) reflec t tha t the y cooled rapidly through the PAZ in Carboniferous and Devonia n time , respectively , an d thereafte r have no t bee n reheate d int o th e PAZ . Thi s indicates rapi d post-orogeni c downwearin g o f the Caledonides , followe d b y a lon g perio d o f relative tectoni c stabilit y o f th e interio r o f th e continent. Jurassic-Cretaceous denudation of the Northern Scandes Samples fro m hig h elevation s i n th e Norther n Scandes mountai n range (K9613 , Kl ) recorde d continuous coolin g i n Jurassi c an d Cretaceou s times (Fig. 5). The Laslett et al. (1987) annealin g model (Fig . 5b ) agai n tend s toward s lowe r temperatures tha n th e Crowle y e t al . (1991 ) annealing mode l (Fig . 5a) , bu t bot h annealin g models indicat e a ver y simila r purel y coolin g trend. Although sample Kl i s from a lower level

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in the vertical profile on Kebnekaise than sample K9613, the thermal history solutions presented in Fig. 5 show that the temperature of Kl ha s been higher than that of sample K961 3 for most of its thermal history . Thi s does not make much sense , and shows that the modelled therma l histories of samples from which only such a small number of track length measurements could be obtained (58 confined track lengths for Kl), have to be treated with caution . Beside s sample s K l an d K9613 , samples N5 , N48 , N4 9 an d N6 0 als o recorde d cooling i n Cretaceous tim e (Fig s 5 and 7). They probably experience d Jurassi c coolin g too , bu t except for sampl e N60 , i n that period the y were still at temperatures too high to be constrained by apatite fission-trac k thermochronology . Whe n samples N48 , N49 and N60 were modelled wit h the Crowle y e t al . (1991 ) annealin g mode l an d only coolin g wa s allowed , fe w thermal historie s that ar e supporte d b y th e dat a wer e foun d fo r each o f thes e samples . Withou t thi s restriction , the mode l indicate d risin g temperature s fo r lat e Cretaceous-Paleogene tim e fo r thes e sample s (Fig. 7 a). With the Laslett et al. (1987) annealin g model, however , man y purel y coolin g historie s supported b y th e dat a wer e foun d (Fig . 7b) . Although the annealin g models do not agre e on the lat e Cretaceous-Paleogen e temperatur e history, bot h predic t coolin g durin g mos t o f Cretaceous time for samples N48, N49 and N60. For man y o f th e sample s t o th e eas t o f th e Northern Scandes , invers e modellin g run s wit h cooling only , usin g th e Crowle y e t al . (1991 ) annealing mode l a s wel l a s th e Laslet t e t al . (1987) annealin g model, generall y resulted only in therma l historie s tha t canno t b e rule d ou t b y the data . Ver y fe w therma l historie s tha t ar e supported by th e data were obtained in this way and fo r mos t o f thes e sample s non e a t all . Therefore reheatin g wa s allowe d durin g late r runs. However , i t i s importan t t o kee p i n min d that th e fission-trac k syste m i s essentiall y on e that record s cooling , an d tha t i t i s difficul t t o differentiate a thermal histor y tha t ha s included some reheatin g fro m a purel y coolin g histor y (Gleadow & Brown 2000) . Thermal historie s fro m th e Crowle y e t al . (1991) annealin g mode l indicat e tha t samples t o the eas t o f th e Norther n Scande s (GRM3 , S16 , P6, S21 and S13 in Fig. 5a ; F51 in Fig. 6a ; P2 in Fig. 7a) , ma y hav e experience d heatin g durin g Jurassic-Cretaceous time . Also , mos t o f thes e samples probably have experienced temperature s lower tha n 6 0 °C i n thi s period . Fo r sample s GRM3 an d SI3, th e Laslett e t al (1987 ) mode l also show s Jurassic-Cretaceou s reheating, wit h a lat e Cretaceou s pea k temperatur e somewha t higher tha n 6 0 °C (Fig . 5b) . Thi s mode l als o

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Fig. 5 . Transec t A-A 7. Result s o f invers e modellin g o f fission-trac k dat a wit h (a ) th e Crowle y e t al. (1991 ) annealing mode l an d (b ) th e Laslet t e t al . (1987 ) annealin g model , togethe r wit h resulting fit for trac k lengt h distributions. Dashed lin e in diagram of thermal histories defines envelope fo r thermal histories with acceptable fit; grey field shows envelope fo r thermal histories wit h good fit; continuous line inside grey field indicates bestfitting therma l history . Diagra m fo r trac k lengt h distributio n show s histogra m o f trac k lengt h measurement s (dotted line) and resulting fit from the best-fitting thermal history (continuous line). AGE, pooled fission-track age with 1 crerror; MTL, mean track length (in |xm); SD, standard deviation of mean track length (in jjim); N, number of track length measurements; F, fault. Frequency : relative frequency o f track lengths in 1 jxm bins. Locations of samples have been projected ont o transect line.

DENUDATION HISTORY OF NORTHERN SCANDINAVIA 125

Fig. 5. continued

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Fig. 6 . Transec t B-B'. Result s o f invers e modellin g of fission-trac k dat a with (a ) th e Crowle y e t al . (1991 ) annealing mode l and (b ) th e Laslet t et al . (1987 ) annealin g model , together wit h resultin g fi t fo r trac k lengt h distributions. (For other details, see caption of Fig. 5.)

indicates tha t all other sample s t o the east of the Northern Scande s wer e belo w 6 0 °C during th e last 15 0 Ma. Thus together, the annealing models indicate that , durin g Jurassic-Cretaceou s time , most o f the sample s t o th e eas t o f th e Norther n Scandes experience d temperature s o f aroun d 60 °C or less. Sample s GRM 3 and S1 3 probably were still inside the PAZ in late Cretaceous time . This mean s they may have experienced temperatures belo w 6 0 °C durin g Jurassic-Cretaceou s time an d a late Cretaceous therma l even t wit h a peak temperature somewhat higher than 60 °C, or that the y experience d a mor e o r les s stabl e temperature als o somewha t highe r tha n 6 0 °C during al l of Jurassic-Cretaceous time.

Also fo r sample s T3 , T4 , T 9 an d F l (Fig . 6 ) and fo r sample s F1 6 an d F3 7 (Fig . 7) , th e Crowley et al. (1991) and the Laslett etal. (1987) annealing model s indicat e that during Jurassic Cretaceous tim e the y wer e a t temperature s o f around 60 °C or less. For these samples, again, it was almos t impossible , wit h bot h annealin g models, t o obtai n purel y coolin g therma l histories tha t ar e supporte d b y th e data . Fo r samples F34 and F40, i n the northeastern corner of th e stud y area , i t wa s possibl e t o obtai n thermal historie s tha t ar e supporte d b y th e dat a with th e Crowle y e t al (1991 ) annealin g model when only cooling wa s allowed (Fig. 7a). But for these sample s i t wa s necessar y t o allo w fo r

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Fig. 6. continued

reheating t o obtai n therma l historie s tha t ar e supported b y th e dat a wit h th e Laslet t e t al. (1987) annealing model (Fig . 7b). However, both annealing model s indicat e tha t samples F3 4 and F40 experience d a temperatur e o f aroun d 6 0 °C or less , wit h mayb e som e coolin g o r reheating , during Jurassic-Cretaceou s time . Fo r sample s SE4 an d N13 , also , i t wa s difficul t t o obtai n thermal historie s tha t are supporte d by the dat a when onl y coolin g wa s allowed . Modelle d thermal historie s o f sample s SE 4 an d N1 3 (Fig. 7) , t o th e S W o f Troms0 , ar e simila r t o those o f sample s fro m th e vertica l profil e o n Tromsdalstinden (Fig . 6). The main difference i s that sample s SE 4 an d N1 3 generall y hav e a temperature somewha t highe r tha n tha t o f samples T3 , T 4 an d T9 . Fo r sample s SE 4 an d

N13 bot h annealin g model s indicat e tempera tures withi n th e PA Z i n lat e Cretaceous Paleogene time . Thi s mean s tha t sample s SE 4 and N1 3 experienced eithe r temperature s belo w 60 °C during Jurassic-Cretaceous time and a late Cretaceous-Paleogene therma l even t wit h a peak temperatur e insid e th e PAZ , o r a mor e o r less stabl e temperatur e als o withi n th e PA Z during al l of Jurassic-Cretaceous time. Samples fro m hig h elevation s in the Norther n Scandes mountai n rang e (K l an d K9613 ) an d samples from th e Atlantic margin (N5, N48, N49 and N60) , recorde d Jurassic-Cretaceou s cool ing. Thi s coul d b e th e resul t o f lowerin g o f th e geothermal gradient . Unfortunately , w e canno t obtain an y informatio n o n th e geotherma l gradient directl y fro m ou r fission-trac k data .

Fig. 7. Samples outside transects. Rresult ofinverse modelling of fission-track data with (a) the Crowley et al. (1991) annealing model and (b) the Laslett et al. (1987)

annealing model. togeter with resulting fit for track length distributions. (For other details, see caption of Fig. 5.)

Fig. 7. continued

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There is extensio n an d crusta l thinnin g offshor e (Brekke 2000) , bu t ther e i s n o indicatio n o f extension i n th e are a wher e thes e sample s wer e collected, an d therefor e on e woul d no t expec t a significant lowerin g o f th e geother m t o occur . This mean s i t i s unlikel y tha t lowerin g o f th e geotherm is the cause o f the observed cooling of samples Kl , K9613 , N5 , N48 , N49 an d N60. Generally, subsurfac e therma l effect s o f exten sion ar e largel y restricte d t o th e regio n under going extension , crusta l thinnin g and subsidenc e (Gallagher e t al. 1998) . Awa y fro m thi s region , the low-temperatur e therma l histor y o f rock s i s primarily controlle d b y denudation (Gallagher & Brown 1997 ; Summerfield & Brow n 1998) . I f cooling i s primaril y th e resul t o f denudation , i t follows tha t th e present-da y are a o f maximu m elevation an d par t o f th e Atlanti c margi n experienced Jurassic-Cretaceou s denudation . Most o f th e sample s fro m th e res t o f th e mainland wer e around o r below 6 0 °C during this period, and, at least in late Cretaceous time , som e probably experience d temperature s withi n th e PAZ. Thi s mean s tha t i n Jurassic-Cretaceou s time th e rest o f the mainlan d di d not experienc e denudation, or that following some denudation, it was affecte d b y a transien t therma l even t tha t reached it s pea k temperatur e i n lat e Cretaceou s time. A possible transien t thermal even t could be the deposition o f a sedimentary cover . This cover subsequently mus t hav e bee n removed , becaus e there ar e n o Jurassic-Cretaceou s sediment s o n the northern Scandinavia n mainland at present. A small amoun t o f Jurassic-Cretaceou s sedimen t is preserved, however , on And0ya, in the north of the Vesterale n island grou p (Dallan d 1980) . The amoun t o f denudatio n o f th e present-da y area o f maximu m elevatio n ca n b e calculate d from th e thermal histor y o f sampl e K9613 . Thi s sample staye d withi n th e PA Z fo r almos t th e entire Jurassic-Cretaceou s period , an d th e thermal historie s fro m th e Crowle y e t al . (1991) annealin g mode l an d th e Laslet t e t al . (1987) annealin g mode l ar e very simila r fo r this sample. Fo r samples N5 , N48, N4 9 an d N60 the differences betwee n th e therma l historie s fro m the Crowle y e t a l (1991 ) annealin g mode l an d the Laslet t e t al . (1987 ) annealin g mode l ar e significant. Fo r sampl e K l onl y 5 8 containe d track length s coul d b e measured . Therefor e th e amount o f denudatio n ha s no t bee n calculate d from thes e fiv e samples . Accordin g t o th e best fitting therma l histor y fro m th e Laslet t e t al . (1987) annealin g model , sampl e K961 3 experi enced 20-2 5 °C of Jurassic-Cretaceous coolin g (Fig. 5b) . The best-fittin g therma l histor y fro m the Crowle y e t al . (1991 ) annealin g mode l indicates c . 3 0 °C o f Jurassic-Cretaceou s

cooling (Fig . 5a). To calculat e th e amoun t o f denudation, the (palaeo)geotherma l gradien t ha s to b e known . Becaus e w e canno t calculat e th e (palaeo)geotherm fro m ou r vertica l profiles , w e used th e Mesozoi c geother m o f 2 8 ± 8 °C that Rohrman (1995 ) calculate d fro m fission-trac k data fro m a vertica l profil e o n Jotunheimen , southern Norway . Usin g thi s geother m t o calculate th e amoun t o f denudatio n o f th e present-day are a o f maximu m elevatio n i n ou r study area , w e arriv e a t a n estimat e o f 0.6-1.3 km wit h th e Laslet t e t al . (1987 ) annealing model , an d 0.8-1. 5 km wit h th e Crowley e t al . (1991 ) annealin g model . Takin g into accoun t the uncertaintie s from bot h annealing models, th e Jurassic-Cretaceous denudation of the present-day are a o f maximum elevation is estimated t o be 1 ±0.5 km. Tertiary denudation To obtai n detaile d informatio n abou t th e post Cretaceous therma l histor y o f ou r sample s fro m inverse modelling , th e time-ste p siz e fo r th e Cenozoic perio d wa s alway s kep t smal l compared wit h that for th e olde r par t o f th e therma l history. This wa s done by increasing the number of nodal points for this time interval. In this way, there ar e fewe r restrictions in term s o f wher e i n time a lat e coolin g even t ca n occu r (Ketcha m et al. 2000) . Inverse modellin g wit h th e Crowle y e t al . (1991) annealing model indicates that by the time of th e Cretaceous-Tertiar y transition , al l samples t o th e eas t o f th e Norther n Scande s were coolin g insid e th e PA Z o r wer e alread y below 60°C (GRM3 , SI8 , SI6, P6, S21 and S13 in Fig. 5a; F48 and F51 in Fig. 6a; P2 in Fig. 7a). The Laslet t e t al . (1987 ) mode l indicate s tha t only samples S13 and GRM3 were still inside the PAZ a t tha t time , an d tha t the y wer e coolin g (Fig. 5b) . Becaus e th e Laslet t e t al . (1987 ) annealing mode l i s renowne d fo r producin g anomalous lat e coolin g event s (va n der Bee k 1995), man y scenario s for the Cenozoi c therma l history o f eac h sampl e hav e bee n tested . I t wa s found tha t because, according t o this model, only samples S13 and GRM3 were still inside the PAZ at th e en d o f Cretaceou s time , thi s mattere d fo r the interpretatio n o f thes e tw o sample s only . When th e time-ste p siz e wa s decrease d onl y fo r Neogene tim e an d no t fo r Paleogen e time , th e Laslett et al. (1987) annealing model found many thermal historie s tha t ar e supporte d b y th e dat a that indicat e a Neogen e onse t fo r th e fina l cooling of samples S13 and GRM3. But when the time-step siz e was decrease d fo r all of Cenozoi c time, th e mode l n o longe r indicate d a Neogen e

DENUDATION HISTOR Y O F NORTHERN SCANDINAVI A

onset for the final cooling of the two samples, but tended more towards a Paleogene onset . Clearly , with th e Laslet t e t al (1987 ) annealin g model , scenarios with a Paleogene or a Neogene onset of the fina l coolin g phas e bot h produc e therma l histories that are supported by the data. However, one would certainly expec t that the Laslett et al (1987) annealin g mode l woul d indicat e a Neogene onse t o f th e fina l coolin g phas e fo r a sample that really experienced Neogen e cooling, even whe n the time-ste p siz e wa s decrease d fo r all o f Cenozoi c time . Thi s i s no t th e case , an d therefore i t i s considere d likel y tha t th e fina l cooling of samples S13 and GRM3 started before Neogene time . Thi s the n means tha t denudation of th e are a t o th e eas t o f th e Norther n Scande s probably resumed in late Cretaceous-Paleogene time. Thi s leave s ope n th e possibilit y o f a Neogene accelerate d cooling . For mos t o f th e sample s clos e t o the Atlantic margin (N 5 in Fig. 5; N48, N49, N13 and SE4 in Fig. 7 ) a Neogen e onse t fo r th e fina l coolin g phase i s predicte d b y bot h annealin g models , although th e Crowle y e t al . (1991 ) mode l als o allows for an earlier cooling of samples N13 and SE4. For sample N60 (Fig. 7a) the Crowley et al. (1991) mode l als o predict s a Neogen e onse t of cooling, wherea s th e Laslett e t al. (1987 ) mode l indicates a sub-PAZ temperature for this sampl e already i n Paleogene time . For sample s clos e t o the Barent s Se a margi n (Fl , T3 , T 4 an d T 9 i n Fig. 6 ; F16 , F34, F37 an d F4 0 i n Fig . 7 ) th e Laslett e t a l (1987 ) mode l indicate s tha t thes e samples wer e belo w 6 0 °C already i n Paleogen e time, o r tha t thei r fina l coolin g ou t o f th e PA Z started i n Neogen e time . Th e Crowle y e t al . (1991) annealing model predicts a Neogene onset of cooling for samples Fl (Fig . 6a ) and F34, F37 and F40 (Fig. 7a) as well. For samples T3, T4 and T9 (Fig. 6a) the Crowley et al. (1991 ) annealing model seem s t o indicat e a Paleogen e onse t o f cooling, bu t a Neogen e onse t o f coolin g i s certainly possibl e too , within th e limit s se t b y thermal historie s tha t are supported b y the data. Therefore fo r sample s o n the Atlanti c margin as well as on the Barents Sea margin, the Laslett et al. (1987) annealing model predicts a Neogene onset of the final cooling phase, or indicates that they wer e alread y belo w 6 0 °C. Th e Crowle y et al . (1991 ) annealin g mode l indicate s a Neogene phas e o f fina l coolin g fo r th e Atlanti c margin and for the part of the Barents Sea margin in th e northeaster n corne r o f th e stud y are a a s well. Although it seem s t o indicate a Paleogen e onset o f coolin g fo r th e par t o f th e mainlan d adjacent to wher e the Atlanti c and Barent s Sea margins com e together , i t certainl y doe s no t exclude the possibility of a Neogene onset of the

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final cooling phas e for this area. Assuming onc e again tha t coolin g i s th e resul t o f denudatio n (Summerfield & Brown 1998), it follows that the Atlantic an d th e Barent s Se a margin s probabl y experienced significan t Neogen e denudation . In th e par t o f th e mainlan d adjacen t t o wher e the Atlanti c an d Barent s Se a margin s meet , denudation ma y hav e starte d earlier . It should be noted that almost all samples from the margins are taken inside fjords, whic h may be incised a s deep a s 2 km. Fjords are probably not simply glacial features, but remnants of an older fluvial drainage system . Consecutiv e glaciation s preferentially exploite d th e pre-existing valley s (Lidmar-Bergstrom et al. 2000). It is not immediately clea r wha t th e coasta l sample s actuall y record: regiona l denudation , o r rapi d fluvia l incision an d glacia l erosio n tha t coul d locall y disturb the thermal structure of the crus t (Stiiwe etal 1994) . Lofoten and Vesteralen Samples from th e Lofoten and Vesteralen islands have thermal historie s ver y differen t fro m thos e of most other samples (Figs 5 and 7). Structurally this are a is ver y complex , wit h man y undate d faults an d probably many more brittle structure s than are currently known. These islands therefore are no w bein g studie d i n a separat e project , which aim s a t resolvin g th e timin g o f faultin g and denudation of the various structural blocks in this area with fission-track analysis and (U-Th)/ He thermochronometry . However , th e fe w available fission-trac k dat a (sample s N15 , N23, N33, N3 4 an d N39 ) sugges t tha t thi s are a experienced considerabl e Neogen e denudation , supporting the observations of Riis (1996) . Pattern of denudation Samples i n th e northeaster n corne r o f th e stud y area, close to the southern margin of the Barent s Sea, hav e fission-trac k age s i n th e rang e o f 200-300 Ma. These ages correspond t o those of samples fro m Lapland , eas t o f th e Norther n Scandes range, rather tha n to ages fro m sample s close t o th e Atlanti c margin , whic h ar e o f th e order of 90-150 Ma. The inverse models and the pattern of fission-trac k age s (Fig. 3) indicate that the greatest Mesozoi c and Cenozoic coolin g was experienced b y the part of the stud y area closes t to th e Atlanti c margin . I t i s eviden t tha t th e denudation o f th e Norther n Scande s mountai n range did not produce a domal pattern simila r to the one recognized i n the topography an d apatite fission-track dat a o f th e Souther n Scande s (Rohrman 1995) . Instead , th e Atlanti c margi n

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has experience d muc h mor e Mesozoi c an d Cenozoic denudatio n tha n th e present-da y are a of maximu m elevation , whic h in tur n was mor e affected tha n the interio r o f the continent .

erosion playe d in this region by removing earlie r Tertiary deposits , bu t make s i t ver y difficul t t o study the correlation wit h the Tertiary geologica l evolution onshore .

Correlation with offshore geology

Discussion

Because provenanc e i s unclea r fo r mos t o f th e sediments in basins surrounding the study area, it is difficul t t o stud y the relationshi p betwee n th e evolution o f th e onshor e par t o f th e stud y are a and th e adjacen t offshor e area . Th e comple x structural evolutio n o f th e offshor e area s als o obscures th e direc t connectio n betwee n on - and offshore. A furthe r complicatin g facto r i s th e limited time-resolutio n o f th e invers e model s from apatit e fission-trac k data , compare d wit h the detaile d stratigraphica l informatio n tha t i s available for mos t basins . However , despit e all this, ther e is a clea r lin k in spac e and tim e between th e off - an d onshor e geologica l evolution. Rapi d Triassic-Jurassi c denudatio n of samples in the Troms0 regio n (T3, T4, T9 and Fl) wa s contemporaneou s wit h depositio n o f Triassic-Jurassic sequences (Faleid e et al 1993 ; Reemst 1995 ) i n th e nearb y Troms 0 an d Hammerfest basin s (Fig . 2) . Th e Cretaceou s sequences i n these basins covered a t least part of the Finnmar k platfor m as wel l (Faleid e et al. 1993). Cretaceous deposits are now missing fro m the Finnmar k platfor m an d th e mainland , an d have bee n tilte d awa y fro m th e mainlan d i n th e basins furthe r nort h an d west . Th e inferre d Cenozoic denudatio n o f sample s fro m th e Troms0 regio n the n implie s a hing e zon e clos e to th e present-da y coastline . Th e Vestfjorden , Ribban, R0s t an d Harsta d basin s ar e locate d offshore th e Atlanti c margi n o f th e stud y are a (Fig. 2). All four basins contain thick Cretaceous deposits tha t ar e underlai n b y thinne r Jurassi c sediments (Faleid e e t a l 1993 ; Olese n e t a l 1997; Brekke 2000). Late Jurassic-early Cretaceous sediment s ar e als o preserve d o n And0ya , Vesteralen (Dalland 1980) . These sediment s thus were deposite d a t a tim e o f inferre d stron g denudation o f the present-da y are a o f maximum elevation an d th e Atlanti c margin , whic h indicates the y ma y hav e bee n derive d fro m thi s region o f Jurassic-Cretaceou s denudation . Except for the Harstad basin , the basins offshore the Atlanti c margi n o f th e stud y are a contai n relatively smal l amount s o f Tertiar y sediment s compared wit h the thic k Cretaceou s sequences . Thin Paleocen e sequence s ar e presen t i n th e Vestfjorden, Ribba n an d R0s t basins , whic h ar e unconformably overlai n b y Plio-Pleistocen e sediments (Brekk e 2000) . Thi s reflect s th e important rol e Neogen e denudatio n an d glacia l

Rohrman & va n de r Bee k (1996 ) explaine d th e domal styl e o f late-stag e postrif t uplif t o f th e Southern Scande s wit h an asthenospheri c diapi r model wherei n a ho t Icelandic ' asthenospher e layer meet s col d cratoni c lithosphere . Thi s mechanism woul d explai n mor e o r les s simul taneous domal-styl e uplif t o f th e Souther n Scandes, Norther n Scande s an d als o Svalbard . However, a comparison of the denudation history reconstructed on the basis of apatite fission-trac k data fo r th e Souther n Scande s (Rohrma n 1995 ) range wit h tha t o f th e Norther n Scande s (thi s study) make s clea r tha t ther e ar e importan t differences betwee n th e tw o i n bot h th e pattern and timin g of denudation . Inverse modellin g o f apatite fission-trac k dat a fro m th e Souther n Scandes suggest s two distinct phases of denudation (Rohrma n 1995) . Triassic-Jurassic denudation of c. 1.3-3.5 km, probably as a consequence of base-leve l lowerin g an d rif t flan k uplift , wa s followed by much slower denudation rates during Cretaceous-Paleogene time . Rapi d denudatio n caused b y tectoni c uplif t starte d fro m c . 3 0 Ma onward an d produce d th e doma l patter n recog nized i n the topography an d apatite fission-trac k ages. Therefore, there are similarities in timing of denudation fo r th e Souther n an d Norther n Scandes until late Cretaceous time, but certainly the patter n o f denudatio n an d probabl y als o th e timing o f denudatio n i n Cenozoi c tim e ar e different fo r the tw o ranges. Although denudatio n i s no t necessaril y con nected t o an y technicall y drive n surfac e uplif t (Summerfield & Brow n 1998) , i t i s temptin g to explain th e result s of th e fission-trac k dat a fro m the stud y area b y passive margin uplift followe d by scar p retreat . This type of model predict s the maximum amount of denudation near the margin and moderat e t o lo w amount s o f denudatio n i n the interior ; this characteristic produce s a strong gradient i n apatit e fission-trac k age , wit h th e oldest age s occurrin g i n th e interio r an d decreasing toward s th e coas t (Gallaghe r & Brown 1997 ; Gallaghe r e t a l 1998) . Rif t shoulders ar e thu s ofte n characterize d b y a marked large-scal e asymmetr y (Beaumont et a l 2000). Th e stud y area clearl y display s all thes e characteristics. Another important prediction of a scarp retrea t mode l i s tha t th e timin g o f maximum denudatio n decreases inlan d fro m th e coast towar d the final position of the escarpment

DENUDATION HISTORY OF NORTHERN SCANDINAVIA (Gallagher & Brow n 1999) . Nowher e o n th e Atlantic margin i n the study area is the samplin g density hig h enoug h t o convincingl y prov e tha t the margi n display s thi s characteristic . Sampl e N57 i s fro m c . 25k m furthe r fro m th e margin , and als o fro m 360 m highe r than sampl e N60 . Indeed, th e fission-trac k ag e o f sampl e N5 7 i s 30 Ma younge r tha n tha t o f sampl e N60 . Bu t o f course thes e ar e onl y tw o samples , an d thi s i s also th e only plac e o n the margin wher e w e can test thi s predictio n o f th e scar p retrea t model . Passive margin s ar e als o generall y associate d with thic k synrif t sediment s an d a postrif t unconformity (Brau n & Beaumon t 1989) . Although the Norwegian shel f is a very complex system o f basin s an d high s an d thi s complexit y makes i t difficul t t o recogniz e som e o f thes e characteristics readily , i t clearl y display s thes e features (Brekk e & Rii s 1987 ; Reems t 1995 ; Brekke 2000). I t appears therefor e that rift flan k uplift resultin g from passiv e margin development followed b y scar p retrea t ca n explai n th e reconstructed denudatio n history of the Northern Scandes range . This certainl y i s not the case fo r the peculia r denudatio n histor y o f th e Souther n Scandes (Rohrma n 1995) . The atypica l thermotectoni c developmen t o f the Souther n Scande s ha s implication s fo r th e accuracy o f th e estimat e o f th e amoun t o f Jurassic-Cretaceous denudatio n o f th e North ern Scandes . Th e Mesozoi c geother m o f 2 8 ± 8 ° C k m ~ 1 (Rohrma n 1995 ) fro m th e Southern Scande s regio n ha s bee n use d i n thi s calculation an d seems to be high compare d wit h the present-day geotherm of less than 20 °C km~l on th e Scandinavia n mainlan d (Ballin g 1990) . The use of too high a value for the geotherm wil l lead t o underestimatio n o f th e amoun t o f denudation. Th e estimat e give n in this paper fo r the amoun t o f Jurassic-Cretaceou s denudatio n of th e Norther n Scande s coul d therefor e b e to o low.

Summary and conclusions Post-Caledonian coolin g an d denudatio n i n northern Scandinavi a progressivel y shifte d fro m the interio r o f th e continen t toward s th e Nort h Atlantic margin . I n th e Norther n Scande s mountain range , th e present-da y are a o f maxi mum elevatio n ha s experience d continuou s cooling an d denudatio n a t leas t sinc e Jurassi c time. Th e combine d Jurassic-Cretaceou s denu dation of this region wa s 1 ± 0. 5 km. Except for most o f the North Atlanti c margin , th e norther n Scandinavian mainlan d eithe r wa s no t affecte d by denudatio n i n Jurassic-Cretaceou s time , o r experienced som e denudatio n followe d b y a

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transient therma l even t tha t reache d it s pea k temperature i n lat e Cretaceou s time . Thi s coul d indicate the occurrence o f a Jurassic-Cretaceous sedimentary cove r o n th e mainland . I t sub sequently mus t hav e bee n removed , becaus e there ar e n o Jurassic-Cretaceou s sediment s preserved o n th e norther n Scandinavia n main land. Final denudation of the eastern flank of the Northern Scandes had probably already started in late Cretaceous-Paleogen e tim e an d possibl y accelerated i n Neogen e time . Th e fission-trac k record o f th e Atlanti c an d Barent s Se a margin s indicates a Neogene onse t fo r the final phase o f cooling an d denudation , althoug h denudatio n may already have affected th e area where the two margins mee t befor e tha t time . Th e combine d Mesozoic-Cenozoic denudatio n wa s stronges t on th e Atlanti c margin . Th e interio r o f th e continent experience d th e leas t Mesozoic Cenozoic denudation . The patter n an d timing of denudation o f th e Norther n Scande s mountai n range i s differen t fro m tha t o f th e Souther n Scandes range , whic h experience d domal-styl e late-stage postrif t uplif t i n Neogen e time . Th e reconstructed denudatio n histor y o f norther n Scandinavia can be explained b y scarp retreat of an uplifte d rif t flank . This researc h i s funde d b y Nors k Hydro , Statoi l an d the Norwegian Petroleum Directorate. I t is part o f the NSG an d ISE S researc h school s an d ha s bee n performed i n the Department of Isotope Geochemistry of the Vrije Universiteit Amsterdam. We thank P. Green and P. Bishop for comments an d critical revie w o f the manuscript. We acknowledge the EC N a t Petten, Th e Netherlands, fo r irradiatin g the fission-trac k samples . We ar e gratefu l t o LKA B Kirun a an d LKA B Malmberget fo r providin g dee p sample s fro m thei r mines, an d t o O . Svenningse n for providin g us with samples fro m Kebnekaise . Thi s pape r i s Publicatio n 20010501 o f th e Netherland s Researc h Schoo l o f Sedimentary Geology .

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Along-slope variation i n the late Neogen e evolutio n of the mid-Norwegian margi n i n response to uplift and tectonis m 1

D. EVANS 1, S . McGIVERON2, Z. HARRISON 2, P. BRYN3 & K. BERG 3 British Geological Survey, Murchison House, West Mains Road, Edinburgh EH9 SLA, UK 2 Svitzer Ltd, Morton Peto Road, Great Yarmouth NR31 OLT, UK 3 Norsk Hydro ASA, PO Box 200, N-1321, Stabekk, Norway Abstract: A s par t o f th e Norwegia n Dee p Wate r Programme , a regional geologica l an d geophysical interpretation of the mid-Norwegian margin resulted in the establishment of a late Paleogene to Holocene stratigraphi c framewor k fo r the margin , and identification an d mapping o f a range o f possible geohazards, includin g slides . At the V0rin g margi n in th e north, ther e ha s bee n th e build-ou t o f a hug e progradin g wedg e o f sedimen t i n Plio Pleistocene times. The sediments of the wedge are assigned to the Naust Formation , whic h has been subdivide d int o eigh t unit s (A-H) , an d includes onl y a few palaeoslides. To the south o f the V0ring margin lies the Storegga Slide Complex and the North Se a Fan, an d the whole of this souther n regio n shows evidenc e of several major palaeoslides . The sediments are also referred t o the Naust Formation , which her e are subdivided into nin e units (O-W ) that hav e been partly correlate d wit h equivalen t Naus t unit s t o the north. The oldest Naus t unit in the south, Naus t Unit W, is largely made up of slide deposits that provide evidence for the earlies t identifie d large-scal e slop e instabilit y i n th e Storegg a Slid e Complex . Thi s instability wa s penecontemporaneou s wit h th e initiatio n o f th e progradin g wedg e t o th e north, an d both feature s ar e postulated to be the result o f uplift of the Norwegian mainland . In th e cas e o f th e Storegg a Slid e Complex , whic h lie s clos e t o th e mai n are a o f uplift , oversteepening o f th e margin , togethe r wit h seismicit y associate d wit h th e Ja n Maye n Fracture Zon e an d M0re-Tr0nderla g Faul t Complex , ma y hav e initiate d slidin g tha t ha s since occurre d intermittentl y u p t o Holocen e times , ove r a tim e interva l whe n ther e ha s additionally bee n muc h glacio-isostati c movement .

In th e south , th e shel f of f wester n Norwa y i s Britis h Geologica l Surve y acting a s consultants, narrow, but it widens considerable t o the north in A n understandin g o f margi n developmen t an d the Haltenbanke n an d beyon d th e Traenadjupe t slop e stability in the area is very important to the (Fig. 1). In the south the limit of the shelf north of hydrocarbo n industr y as deep-wate r exploratio n the Nort h Se a Fa n i s marke d b y th e scarp , o f continue s an d developmen t o f majo r ne w field s 290 km length (Bugge etal 1987) , of the Storegga begins . Slide, wherea s i n th e nort h a more gentl e slop e Th e summarize d result s o f Phas e I I o f th e leads down to the V0ring Basin, beyond which it projec t were published by McNeill e t al. (1998), rises slightl y at the V0ring Plateau. wh o described th e Cenozoic stratigraphic frame The mid-Norwegia n margi n becam e a n wor k and discussed geohazards relevant to the oil important are a fo r hydrocarbo n exploratio n industry . Bry n e t al . (1998 ) concentrate d o n following th e deep-wate r 15t h Norwegia n issue s o f slop e stability . Thi s assessmen t wa s round o f licensin g i n 1996 . T o gai n a bette r update d during Phase III of the project, for which understanding o f the shallo w geolog y an d slope man y additiona l dat a wer e available . Th e stability o f th e area , severa l companie s jointl y complet e databas e include d 1 0 000km o f ne w formed th e Seabe d Project , whic h wa s a an d reprocesse d seismi c an d high-resolutio n component o f th e Norwegia n Dee p Wate r seismic , mini-airgun , deep-to w boomer , 3 D Programme (Bry n e t al. 1998) . This projec t was seismi c dat a i n licenc e block s (Fig . 1) , towe d managed by Norsk Hydro, and one aspect of the ocea n botto m instrument (TOBI) sidesca n sona r project was geological and geophysical interpret- data , thre e Ocea n Drillin g Progra m (ODP ) o r ation, which was carried out in three phases. The Dee p Se a Drillin g Projec t (DSDP ) sites , thre e second an d thir d phase s wer e contracte d t o exploratio n well s an d fou r geotechnica l Svitzer Lt d betwee n 199 7 an d 1999 , wit h th e boreholes . From: DORE , A.G., CARTWRIGHT , J.A. , STOKER , M.S. , TURNER , J.P . & WHITE , N . 2002 . Exhumation o f th e North Atlantic Margin: Timing, Mechanisms and Implications for Petroleum Exploration. Geologica l Society , London, Special Publications, 196, 139-151 . 0305-8719/027$ 15.00 © The Geological Societ y of London 2002 .

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Fig. 1 . The bathymetry (i n metres) an d location o f the stud y area in relation t o the mid-Norwegian shel f an d the Storegga Slide . Als o shown ar e the 15t h Roun d licenc e block s (se e als o Fig. 6) , the location o f seismic sections , and th e landwar d limi t o f the slid e deposit s tha t for m par t o f Naus t Uni t W.

This wa s a regiona l interpretatio n i n whic h only th e outermos t portio n o f th e shel f wa s included i n th e are a o f study . The are a o f study covers ove r 100 000km2, and extends alon g the continental slop e fro m 62°15' N t o 68°N . Th e westward exten t o f th e are a i s approximatel y a t the V0rin g Escarpmen t i n th e nort h i n wate r depth o f abou t 1200-1300m , wherea s i n th e south i t extend s t o depth s o f 250 0 m withi n the Storegga Slide . It is important t o note that in this paper the term 'Storegg a Slide' refers only to the most recen t movemen t o n the slid e a s describe d by Bugge (1983 ) an d Bugge et al (1987 , 1988) , whereas th e ter m 'Storegg a Slid e Complex ' i s used to describe an area of longer-term instability centred around the Storegga Slide. It must also be emphasized tha t this paper is based on the results of th e Seabe d Projec t u p t o 1999 ; man y

commercial dat a hav e bee n collecte d sinc e that time, includin g 3D seismi c data , bu t th e newe r data ar e no t used here. Although th e projec t considere d th e whol e post-Paleocene succession , thi s paper wil l focu s on th e Plio-Pleistocen e history , wit h particula r emphasis o n th e lat e Pliocen e t o earl y Pleisto cene sedimentar y response. This includes , in the northern part of the area, the build-out of a major prograding wedge , whic h i s comparabl e wit h similar feature s observe d o n glaciate d margin s world-wide (e.g. Larter & Barker 1991 ; Clausen 1998; Kristofferso n e t al . 2000 ) an d ha s commonly bee n described a s a response t o uplif t of Scandinavi a (Pool e & Vorre n 1993 ; Hendriksen & Vorren 1996 ; Rii s 1996 ; Stuevold & Eldhol m 1996) . However , a marke d along slope contrast between the north and south of the

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study are a wil l b e described , an d som e possibl e reasons fo r these difference s discussed .

Stratigraphic framework The Stratigraphi c framework established fo r th e area i s base d o n seismi c interpretation , ag e information fro m th e projec t database , an d published information . Th e framewor k illus trated i n Fig . 2 i s adapte d fro m McNeil l e t al. (1998) following work in Phase III of the project. The formatio n names are those o f Dalland e t al. (1988), althoug h the age-rang e o f their Pliocen e Naust Formatio n ha s bee n extende d t o includ e the Pleistocene deposits . It ca n b e see n tha t th e Eocen e t o Oligocen e Brygge Formation extend s the length of the area . This formatio n was folded during late Eocene t o Oligocene compressio n an d inversio n (Dor e & Lundin 1996 ; Swiecick i et al. 1998) , an d Vagnes et al . (1998 ) hav e argue d tha t inversio n i s continuing t o th e present . However , younge r deposits d o not have this continuity. Kai units B and C ar e though t o n th e basi s o f similarit y of

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seismic characte r t o hav e bee n originall y continuous beneat h th e Storegg a Slide , bu t Kai Uni t A is a t present foun d onl y t o the nort h of Storegga . Th e shale-pron e Brygg e an d Ka i formations i n th e stud y are a wer e bot h deposited i n relativel y deep-wate r basins , probably largel y beyon d th e contemporar y shelf (Gradstei n & Backstro m 1996 ; Eidvi n et al . 1998) . I t ca n b e see n fro m Fig . 3 tha t these formation s ar e commonl y thi n o r absen t on th e shelf , althoug h the y typicall y hav e a combined thicknes s o f ove r 1000 m i n th e V0ring an d M0r e basins . Above the Kai Formation , ther e is a marke d change i n th e styl e o f sedimentation , fo r i n lat e Pliocene time s (Eidvi n e t al . 2000 ) th e larg e prograding wedge of the Lower Naust and Naust formations buil t ou t t o for m th e present-da y shelf. Thi s wedg e ha s resulte d i n th e advance ment o f the shel f by u p to 100k m locally i n th e north (Hendrikse n & Vorre n 1996) ; th e forme r extent in the south is unknown because of erosion at th e Storegg a Slid e Complex . Ther e i s als o a significantly differen t Stratigraphi c breakdow n

Fig. 2 . The seismostratigraphic framework establishe d for the area, illustrating th e time range of the Brygge, Kai , Naust and Lower Naust formations. This is a modification of the stratigraphy proposed b y McNeill e t al. (1998).

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Fig. 3 . Cross-section illustrating the general relationships of the Neogene stratigraphic unit s in the study area, with schematic extensio n acros s th e shel f t o th e S E based o n Rokoenge n e t al. (1995) . Th e locatio n o f th e portio n within th e stud y are a is close to tha t o f Fig. 4 , wit h th e schemati c extensio n toward s the Norwegia n coast .

to th e nort h an d sout h o f th e norther n flan k o f the Storegg a Slid e (Fig . 2) , wit h onl y a limite d seismic correlatio n establishe d betwee n th e tw o successions. North of the Storegga Slide Complex In the north, the basal componen t abov e the Kai Formation is the informal Lower Naust formation, a uni t tha t withi n the stud y are a form s th e thin ,

distal facie s o f th e earl y build-ou t o f th e prograding wedg e (Fig . 3). The formatio n is, however, locall y thicke r i n mounde d section s (Fig. 4), probably as a result of preferential finegrained sedimentatio n unde r th e influenc e o f contour currents . Th e Lowe r Naus t formatio n unconformably overlie s olde r deposits , an d i s itself unconformabl y overlai n b y th e down lapping unit s o f th e Naus t Formatio n o n th e slope, wherea s i n th e basi n it i s overlai n by th e

Fig. 4 . A seismic profile through the northern succession of the Naust Formation, indicating it s relationships wit h underlying formations . (Fo r location , se e Fig. 1. )

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distal facie s o f th e Naus t Formation. Th e Naust Formation ha s been subdivide d int o units A-H, and the discrete, westerly downlappin g package s of unit s B-H for m th e diachronou s bas e t o the formation. Thei r landwar d extent s are truncated at a glacia l unconformit y a t th e bas e o f Naus t Unit A , althoug h glacia l erosio n o f th e shel f probably pre-date s Naus t Uni t A (Haflidaso n et al. 1991) . Thi s glacia l erosio n surfac e wa s termed th e UR U (uppe r regiona l unconformity) by Hendriksen & Vorren (1996) . The Naus t Formatio n i n th e nort h i s u p t o 1500m thick , an d i t ha s bee n estimate d tha t its total volum e i n th e nort h i s 8 0 000km 3 (Evan s et al 2000) . A typical seismi c sectio n presented in Fig . 4 (se e als o McNeil l e t al 1998 ; Evan s et al. 2000) illustrate s the downlapping character of the units, and shows that each unit is made up of on e o r mor e package s an d i s separate d b y well-defined reflectors . Thes e package s ma y show bedding , bu t ar e generall y acousticall y structureless, althoug h adjacen t t o th e Storegg a Slide Complex they are locally of an acoustically well-bedded facies (Fig . 5) . From th e dat a availabl e t o th e study , i t i s apparent tha t there is evidence fo r only a limited number of major palaeoslide s withi n this sector ,

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and no evidence of such erosion i n the older par t of the succession. The Traenadjupet and NE Nyk slides ar e found in the far north of the study are a (Laberg et al 1999) , but these lie where the slope extends down to the Lofoten Basin to the north of the V0rin g margi n sector . A relativel y smal l palaeoslide at the northern flank of the Storegg a Slide has been describe d b y Evans et al. (1996) , and a part o f Palaeoslide-2 has been mappe d on the flan k o f th e Storegg a Slid e Comple x an d i n the far west of the study area. The onl y example on th e V0rin g slop e durin g thi s projec t i s th e Traenabanken Slid e t o th e nort h o f th e Hellan d Hansen licence area , a slide that took place afte r Naust Unit B times and commonly employed the top of Naust Unit C as a glide plane. The location of th e hea d wall o f this slid e i s evident i n Fig. 6 from th e remova l o f mid-Pleistocen e sediment s from the middle of the main northern depocentre. South of the Storegga Slide Complex The lates t movement s o f th e Storegg a Slid e a s described b y Bugge et al. 198 7 wer e Slides I , II and III, and these formed the present topograph y of the region. Slide I was described a s pre-dating the las t glaciation , wit h Slide s I I an d II I

Fig. 5 . A seismic sectio n acros s th e northern flank of the Storegga Slid e illustratin g th e largely acousticall y well bedded natur e of the Naust Formation immediatel y nort h of the sidewall. Palaeoslide s ca n be seen at depth belo w the Storegg a Slide . (Fo r location , se e Fig. 1. )

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Fig. 6 . Map showin g the thickness of middle Pleistocene sediment s (Naust units B and C) in the study area. The sediments have been wholly or partly removed fro m th e Storegga Slid e Complex region, and partially removed by the Traenabanken Slide . Also shown is the 2D seismic grid available to the study, and the 15t h Round deep-water licence blocks .

occurring abou t 7000-720 0 BP (Svendse n & Bondevik 1995 ; Bondevik ef a/. 1997) . However, the deposit s o f Slid e I ar e no t blankete d b y a well-defined sedimentar y cove r (Fig. 5) as would be expecte d i f tha t slid e pre-date d th e las t glaciation, an d it is now considered tha t all three

movements post-dat e th e las t glaciatio n an d probably al l occurre d a t aroun d th e sam e time . Haflidason e t al (2000 ) suggested that they took place between 6000 and 8000 BP. This suggestion is in accordance with an alternative interpretation originally propose d b y Bugge (1983) .

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To th e sout h o f th e norther n flan k o f th e Storegga Slide , Ka i Unit A is absent an d there is no direc t equivalen t o f th e Lowe r Naus t formation. A correlatio n ha s bee n establishe d between th e top of Naust Unit W in the souther n succession an d th e to p o f Naus t Uni t F i n th e north (Fig . 2) , a stratigraphi c leve l tha t fro m limited well evidence may approximate to the top of the Pliocen e sequence , but is though t mor e likely t o b e o f earl y Pleistocen e age . Figur e 5 shows the equivalence of the top of Naust Unit W (the oldest Naus t unit of the southern successio n in this region) with the top of Naust Unit F at the junction o f th e tw o succession s beneat h th e northern sid e wall of the Storegg a Slide . The characte r o f Naus t Uni t W i s i n marke d contrast to that of its coeval unit s (Naust units F, G and H) in the north. Th e base of the unit is in many place s severel y erosional ; th e exampl e i n Fig. 7 shows that one of the erosional scarp s at its base exceed s 250m s (c . 250m ) i n height . Although the scal e o f erosion i s small compare d with th e downcuttin g associate d wit h th e Holocene movement s o f th e Storegg a Slide , i t none the less represents erosion o n a large scale . Elsewhere th e Uni t W sediment s hav e th e

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slightly mounde d an d acousticall y opaqu e internal characte r tha t is commonl y associate d with slide deposits, although not all profiles show evidence o f a slide-related origin . The exten t of the slidin g i n Naust Unit W ha s been mappe d an d i s show n i n Fig . 1 , wherea s Fig. 7 shows the development of a headwall. The poorly defined trac e of the headwall lies seaward of th e present-da y Storegg a Slid e scarp , an d i s seen t o b e o f comparabl e length , indicatin g th e magnitude o f th e erosio n a t tha t time, althoug h the mass-movemen t ma y represen t man y separ ate events. It is considered that Naust Unit W was formed b y more than one process, but that majo r slide erosio n wa s a significan t component . Importantly, th e slidin g associate d wit h Naus t Unit W is the oldest so far clearly identified in the study area , an d ha s a clos e geographica l association wit h th e present-da y bathymetr y o f the lates t movement s o n th e Storegg a Slid e a s documented b y Bugg e e t al (1987 , 1988) . Bearing i n min d th e absenc e o f Ka i Uni t A an d the Lowe r Naus t formation , i t i s possibl e tha t earlier instabilit y may hav e caused th e irregula rities see n locall y beneath th e bas e o f the Naust Formation i n Fig. 7 , but this remains unclear.

Fig. 7 . Seismi c profil e fro m th e Storegg a Slid e Complex , illustratin g th e locall y erosiv e natur e o f the base of Naust Uni t W, as well a s deep erosion to the base of the sediment s deposite d during Holocen e movement o n the slide. (Fo r location , se e Fig. 1. )

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Fig. 8 . Th e relationshi p o f identifie d slide s an d palaeoslide s t o th e stratigraphi c subdivision s o f th e Naus t Formation. TJ , Tranadjupe t Slide ; NEN , N E Ny k Slide ; TB , Traenabanke n Slide ; TAM , Tampe n Slide ; FSE , Faeroe-Shetland Escarpment Slide ; 1 , 2, 3 and 4 are Palaeoslides-1, -2, -3 and -4. The Vigra and North Sea Fan Slide-1 were no t specificall y identified during thi s study , but wer e recognize d b y Evan s et al. (1996 ) an d King et al. (1996) within Naust units V-S.

The souther n successio n i n the North Se a Fan region t o th e sout h o f th e Storegg a Slid e i s o f comparable thicknes s t o the northern succession , but ther e ar e pronounce d difference s i n it s seismic character . Th e fa n becam e increasingl y important a s a depocentr e i n mid - t o lat e Pleistocene time a s vas t quantitie s o f sedimen t were carrie d t o th e shel f brea k alon g th e Norwegian Channe l (Sejru p e t al . 1996) . Th e units sho w a variet y o f acousti c characteristic s ranging fro m opaqu e t o wel l bedded , bu t a key characteristic o f the succession i s that it includes evidence fo r a numbe r o f majo r translationa l slides. Evans et al (1996 ) and King et al (1996 ) have described Nort h Se a Fan Slide-1, the Vigra Slide, th e M0r e Slid e an d th e Tampe n Slid e o n the fan. The presen t stud y ha s identifie d th e slidin g associated wit h Naus t Uni t W , th e Faero e Shetland Escarpmen t Slid e tha t lie s clos e t o th e eponymous escarpment , an d Palaeoslides-1 , -2 ,

-3 an d - 4 (Fig . 8) . I n th e presen t seismi c correlation, Palaeoslide- 2 is equivalent in age to the Traenabanken and M0re slides , and has been traced alon g th e norther n flan k o f th e Storegg a Slide a s wel l a s i n th e sout h (Fig . 6) . Isolate d thick remnants of mid-Pleistocene deposit s see n in Fig. 6 adjacent to the area of complete removal testify t o thei r probabl e forme r widesprea d presence befor e thei r remova l b y Palaeoslide- 2 and contemporaneous events. The precise exten t of these slides is unknown, largely becaus e significan t part s o f the m ma y have bee n remove d b y late r slid e erosion , bu t some wer e undoubtedl y ver y large , eve n compared wit h th e lates t Storegg a Slid e move ments, which are the most recent major erosional events in this M0re Basin region. All these slides are locate d aroun d th e Storegg a Slid e a s described byBugg e e t a l (1987) , an d i t i s clea r that the term Storegga Slide does not adequately describe th e lon g histor y o f movements . Th e

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term Storegg a Slid e Comple x i s therefor e proposed, wit h th e ter m Storegg a Slid e use d t o define onl y th e post-glacia l movement s an d consequent bathymetri c expressio n o f thos e events.

Discussion The abov e description s sho w tha t the histor y o f Plio-Pleistocene margi n developmen t ha s bee n significantly differen t eithe r sid e o f a line that is approximately equivalent to the northern flank of the Storegg a Slide . Thi s lin e i s coinciden t wit h the location of the Jan Mayen Lineament , whic h has controlle d th e tectoni c developmen t o f th e area sinc e Cretaceou s time s (Brekk e 2000) . I n particular, ther e was a marked differenc e durin g late Pliocene o r early Pleistocene time , when the prograding wedg e wa s initiate d i n th e nort h during th e depositio n o f th e Lowe r Naus t formation an d its mor e proxima l equivalent s o n the inner shelf, and Naust units F, G and H on the slope. However, a significant degree of progradation ma y als o hav e occurre d i n th e Storegg a region, an d ma y hav e bee n subsequentl y removed durin g Unit W or later erosion .

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The initiatio n o f th e wedg e represente d a fundamental chang e i n sedimentar y architectur e as the depocentre move d t o the inne r shel f fro m the earlie r deep-wate r basi n o f th e Ka i an d Brygge formations (Fig . 3) . Workers such as Riis (1996) an d Stuevol d & Eldhol m (1996 ) hav e related thi s chang e t o a perio d o f uplif t o f mainland Norway that led to increased erosio n of the mountainou s region s an d a n enhance d sediment suppl y to th e shel f an d th e generatio n of accommodation spac e seaward of a hinge line. Climatic deterioration , wit h mor e effectiv e fluvial erosion and the onset of upland glaciation, probably als o playe d it s par t i n th e increase d rate o f erosion, an d it is difficult t o separat e th e two force s (Lidmar-Bergstro m e t al . 2000) . Any mechanis m propose d fo r vertica l move ments i n th e mid-Norwegia n are a (e.g . Cloetingh e t al . 1990 ; Rii s & Feldskaa r 1992 ) needs t o b e consisten t wit h observation s indicating lat e Neogen e uplif t i n th e broade r North Atlanti c regio n (Stoke r 1995 , 2002 ; Andersen e t al. 2000; Chalmers 2000 ; Japse n & Chalmers 2000) . The V0rin g margi n woul d hav e bee n fe d b y sediments derive d fro m centra l Norway , which, according t o Rii s (1996 ) wa s a n are a o f lesse r

Fig. 9 . Ma p showin g th e genera l relationship s o f majo r slide s an d palaeoslides wit h structura l features , Plio Pleistocene uplif t (Rii s 1996 ) an d a zon e o f maximu m post-glacia l compressiv e stres s (Gudmundsso n 1999) . Surface slide s i n the nort h ar e fro m Laber g e t al. (2000) . MTFC, M0re-Tr0ndela g Faul t Comple x (Gabrielse n et al . 1999) , whic h i s thought t o be a particularly significan t zon e o f seismicity .

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uplift tha n southern an d northern Norwa y a t that time (Fig. 9) . Although evidenc e o f progradatio n in th e south i s lacking , i t i s clea r tha t a vas t quantity o f sedimen t wa s delivere d t o th e sea , and certainly le d to the pronounced progradatio n and seawar d advancemen t o f th e shel f brea k i n the north . At th e sam e tim e a s th e wedg e wa s bein g initiated, there was both sedimentatio n an d major sliding i n the M0r e Basi n t o the south , althoug h the Unit W sliding could have occurred a s late as Unit F times. Give n tha t the souther n part o f the area lies close t o the region o f maximum onshor e uplift i t i s likel y tha t ther e wa s a n abundan t sediment suppl y i n lat e Pliocen e t o earlies t Pleistocene times. Indeed , rapi d sedimen t suppl y may hav e bee n a facto r i n triggerin g th e slide s that characterize Naus t Unit W, as may have been the increas e i n slop e tha t i s likel y t o hav e bee n experienced a t the continental margi n a s a result of uplif t o f th e lan d an d subsidenc e offshore . Present evidenc e point s to late Pliocen e uplif t a s being a factor in the initiation of the slides at that time, a s wel l a s i n thei r continue d occurrenc e throughout Quaternar y time , bu t wha t wa s th e trigger fo r the slides ? Earthquakes ar e a commonl y quote d trigge r for slide s i n thi s regio n (e.g.Bugg e e t al. 1987 ; King e t a l 1996 ; Laber g e t al . 1999) . Figur e 9 shows that the Storegga Slid e Complex lie s along the Ja n Maye n Lineamen t (Blysta d e t al . 1995 ) and th e Ja n Maye n Fractur e Zone , a s ha s previously bee n note d b y Bugg e e t al . (1987) . Although thi s fracture zon e i s not a strong focu s of modern earthquake s (Dehl s et al. 2000), som e seismic event s hav e bee n recorde d alon g i t offshore, an d are particularly common wher e the lineament meet s th e Norwegia n coas t (Bungu m et al . 1991) . Brekk e (2000 ) consider s i t t o hav e acted a s a transfe r zon e durin g Cenozoi c time , when i t controlle d th e occurrenc e o f compres sional tectonics . Anothe r facto r i s tha t th e headwall (Fig . 9 ) i s broadl y coinciden t wit h th e still seismicall y activ e M0re-Tr0ndela g Faul t Complex (Gabrielse n e t al . 1999) . Ther e i s a clear relationshi p betwee n th e locatio n o f majo r structures or fracture zon e an d the position of the Storegga Slide Complex, and it is therefore likel y that ther e ha s bee n lat e Pliocen e o r earl y Pleistocene tectoni c movemen t i n th e vicinit y of th e slid e comple x tha t coul d hav e create d significant seismi c events . Althoug h th e Ja n Mayen Fractur e Zon e ma y n o longe r b e active , the lineamen t ha s continue d t o exer t a contro l over th e locatio n o f majo r slide s durin g Quaternary time . This indicatio n i s strengthene d b y a stud y o f the distribution of similar feature s on the V0rin g

margin an d farther north around Lofoten . Figur e 9 show s tha t ther e ar e n o crusta l lineament s o r fracture zone s cuttin g the V0ring margin , whic h lies adjacen t t o a zon e o f lesse r onshor e uplift . This i s th e regio n wit h littl e evidenc e o f palaeoslides; th e only significant mappe d featur e south of Traenadjupet is the Traenabanken Slide , which interestingl y lies o n th e projecte d lin e of the Gleipn e Fractur e Zone . To the north of the V0ring margin off Lofoten, the zon e o f majo r uplif t lie s clos e t o th e shel f break, an d thi s i s a margi n tha t display s ampl e evidence o f downslop e movemen t (Dowdeswel l & Kenyon 1997 ; Taylor et al. 2000). In particular there ar e two majo r slides , the Traenadjupet an d Andoya slide s (Laber g e t al . 1999 , 2000) , tha t respectively lie along the Bivrost Fracture Zone / Bivrost Lineamen t an d th e Senj a Factur e Zon e (Blystad e t al . 1995) . Th e ful l histor y o f thes e slides i s no t known , an d although , a s wit h th e Storegga Slid e Complex , thei r mos t recen t movements wer e durin g Holocen e time , olde r movements probably did occur but have yet to be fully documented . Nevertheless , the y d o con tribute t o a well-defined correlatio n betwee n th e location o f majo r crusta l structure s an d th e distribution o f larg e slide s alon g th e Norwegia n margin. In mid - t o lat e Pleistocen e times , glaciation s became stronge r and ice sheets extended onto the shelf wit h greate r frequenc y (William s e t al . 1988; Mangeru d e t al . 1996 ; Vale n e t al . 1996 ; Vorren & Laber g 1996) . Durin g thes e times , movements alon g th e Norwegia n margi n ar e likely t o hav e bee n increasingl y affecte d b y isostatic movement s relate d t o th e onse t an d removal o f th e ic e cove r fro m th e shelf . Thes e events increas e th e ris k o f seismicity , especially when superimpose d o n pre-existin g crusta l stresses relate d t o plat e tectonic s (Talbo t & Slunga 1988) . Although they noted that observed stresses i n Norwa y ar e consisten t wit h uplif t o f Fennoscandia, Fejersko v & Lindhol m (2000 ) considered tha t ridge-pus h associate d wit h sea floor spreading i s the primary caus e o f compres sional stress . Figur e 9 show s th e present-da y zone o f maximu m compressiv e stres s aroun d Norway a s a resul t o f post-glacia l uplif t a s calculated b y Gudmundsso n (1999) . Thi s zone , which i s likel y t o hav e bee n th e locu s o f earthquakes tha t coul d trigge r slides , cover s th e headwall o f al l thre e majo r Holocen e slides . I t has been estimate d that seismi c event s as stron g as M w 7. 9 ma y hav e occurre d i n Scandinavi a during the last ice retreat (Muir Wood 1988) , and given th e occurrenc e o f weak layer s i n the Plio Pleistocene sedimentar y column , slide s coul d well hav e bee n generate d b y thi s mechanism .

LATE NEOCENE EVOLUTION O F MID-NORWEGIAN MARGI N

Other mechanism s ma y als o hav e generate d slides, o r facilitated thei r initiation ; suc h factors include th e presenc e o f ga s o r ga s hydrate s (Bugge e t al 1987 ; Henrie t & Miener t 1998 ; Bouriak et al. 2000), although these factors may be o f mor e loca l importanc e tha n th e regiona l view taken in this paper.

Conclusions Late Neogen e uplif t o f Norwa y ha d a pro nounced influenc e o n margi n sedimentation , changing the pattern from one of slow depositio n in deep-water basins to more rapid sedimentatio n on th e inne r shel f a s a progradin g wedg e wa s initiated i n respons e t o increase d sedimen t supply an d th e generatio n o f accommodatio n space o n the shelf . Stratigraphic analysi s show s tha t ther e ha s been a significan t along-slop e differenc e i n th e latest Cenozoi c histor y o f margi n developmen t between th e V0rin g an d M0r e margins . Th e former i s characterize d b y th e depositio n o f a vast progradin g wedg e wit h littl e evidenc e o f major instability , wherea s th e latte r ha s a lon g history of sliding, which has largely removed any evidence o f wedg e development , an d ma y hav e received les s sedimen t late r in Quaternary time . The oldest identified sliding in the M0re Basin was penecontemporaneous wit h the earl y stage s of developmen t o f the progradin g wedg e in lat e Pliocene an d earliest Pleistocen e times, and both occurrences are considered t o be related to uplif t of th e mainland a t this time. Uplif t wa s greates t in the south adjacent t o the M0re margin , which has experience d a long history o f instability. As isostatic uplif t i s continuin g today, as suggeste d by Rii s (1996) , thi s ma y hav e implication s fo r present-day slop e stability . Three large slides on the Norwegian margin lie at the junctions of oceanic fracture zones with the continental crust , or alon g crusta l lineament s or major faul t zones . The y als o li e adjacen t t o th e zones of maximum onshore uplift. Thi s suggests strong structural control on the location o f slide s on this margin, although there is little evidence of modern seismicit y alon g the fractur e zones , and it ma y b e tha t th e M0re-Tr0ndela g Faul t Complex i s a particularly significan t structure . The extent of structural control is emphasized by the observation that the largest slide area with the longes t histor y o f movement , th e Storegg a Slide Complex , lie s a t th e conjunctio n o f th e largest oceani c fractur e zone , th e zon e o f maximum Plio-Pleistocen e uplift , a majo r faul t zone, an d th e zon e o f maximu m post-glacia l compressive stress .

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The author s woul d lik e t o than k th e oi l companie s involved in the Seabed Project (BP Norge, Esso Norge, Mobil, Norsk e Conoco , Nors k Hydro , Shel l an d Statoil) fo r the opportunity t o participate in the project and for their permission t o publish the data included in this paper . W e ar e als o gratefu l t o E . Gillespi e fo r producing th e diagrams . Th e pape r benefite d signifi cantly bot h fro m th e earl y comment s o f M.S . Stoke r and D. Long a s well a s those o f referees T . Eidvin an d A.G. Dore . The contributio n o f D.E. i s made wit h the permission o f th e Directo r o f th e Britis h Geologica l Survey (NERC) .

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Reconstructing the erosion history of glaciated passive margins: applications of in situ produce d cosmogenic nuclide techniques ARJEN P . STROEVEN1, DEREK FABEL 2, JON HARBOR 3, CLA S HATTESTRAND 1 & JOHAN KLEMAN 1 1

Department of Physical Geography and Quaternary Geology, Stockholm University, S-106 91 Stockholm, Sweden (e-mail: arjen@

geo.su.se)

2

School of Earth Sciences, University of Melbourne, Parkville, Vic. 3052, Australia

^Department of Earth and Atmospheric Sciences, Purdue University, West Lafayette, IN 47907-1397, USA Abstract: Offshor e sedimen t accumulation s provid e a n intriguin g recor d o f th e ne t sediment outpu t resultin g fro m geomorphologica l evolutio n o f th e circum-Atlanti c continental margi n sinc e the commencement o f Neogene glaciation . However, th e onshor e record of the timing, pattern and amount of bedrock erosion that produced these sediments is comparatively poorl y constraine d an d understood, although there ar e good genera l model s of glaciation history. The geomorphology of circum-Atlantic continental margin mountains, as assessed fro m remot e sensin g data and field observations, includes palimpsest landform s and landscapes that reflect a complex pattern of spatial and temporal variations in the impact of glacial , fluvia l an d periglacial processes. Perhaps mos t surprisin g is that, despite having been repeatedl y overridde n b y larg e ic e sheets , part s o f th e landscap e appea r t o be relict , with nonglacia l morphology . Thi s ha s importan t implication s bot h fo r glaciologica l conditions unde r ice sheets , an d for sedimen t sourc e area s an d erosion rates . Conventional dating an d analysi s hav e provide d a n excellent wa y t o begi n unravellin g th e timin g an d pattern o f erosion , landfor m development , an d possibl e landfor m preservatio n unde r ice. However, testin g hypothese s develope d fro m curren t models , an d addressin g critica l unresolved questions, requires additional approaches. The use of in situ cosmogenic nuclide production i n bedroc k i s a ne w approac h fo r investigatin g landscap e evolutio n i n mountainous areas . Wit h carefu l interpretatio n o f geomorphologica l settings , cosmogeni c nuclides can be used to determine apparent surface exposure age and landscape preservation, and constrain erosion depths and duration of burial by ice. Here we provide a framework for the interpretatio n o f cosmogeni c nuclid e concentrations i n bedrock surface s of landscape s affected b y glacial , fluvia l an d periglacia l processes , illustrate d wit h example s fro m th e northern Swedis h mountains . Thi s demonstrate s potentia l use s o f cosmogeni c nuclid e techniques, an d provide s a foundatio n for attempt s t o improv e geomorphologicall y base d reconstructions o f relict landscapes , t o reconstruct an d analys e the dynamic s of landscap e change in glacial times, and to define the consequences of different proces s regimes i n terms of erosio n patterns , sedimen t transport , an d th e suppl y o f sediment s tha t ar e deposite d offshore.

The evolutio n o f the circum-Atlantic continental 1996) . I n particular , curren t technique s o f margin sinc e th e commencemen t o f Neogen e estimatin g onshor e exhumatio n pattern s an d glaciation i s reflected in large offshor e sedimen t rate s d o no t resolv e margi n exhumatio n durin g accumulations (e.g . Solhei m e t al. 1996) . Th e lat e Cenozoi c tim e (Hendrik s & Andriesse n thickness o f offshor e sediments , importan t 2002) . quantities o f whic h wer e generate d b y glacia l Th e broad-scal e glaciatio n histor y o f th e processes, i s on e ke y ingredien t considere d i n circum-Atlanti c continenta l margi n i s wel l hydrocarbon exploration . However , th e onshor e understood i n genera l term s (e.g . Shackleto n record o f th e amount , timin g an d patter n o f e t al . 1984 ; Janse n & Sj0hol m 1991 ; Holeman n bedrock erosio n tha t wa s th e sourc e o f thes e & Henric h 1994 ; Mangeru d e t al . 1996 ; Janse n offshore sediment s i s no t wel l constrained , e t al . 2000) . I n Scandinavia , fo r example , ic e except at the mos t generalize d leve l (e.g . Rii s sheet s centre d wes t of the mountai n elevatio n From: DORE , A.G., CARTWRIGHT, J.A., STOKER , M.S. , TURNER, J.P . & WHITE , N . 2002 . Exhumation o f th e North Atlantic Margin: Timing, Mechanisms an d Implications for Petroleum Exploration. Geologica l Society, London, Special Publications, 196, 153-168. 0305-8719/027$ 15.00 © The Geological Society of London 2002.

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axis wer e th e dominan t for m o f glaciatio n form patter n that is both areally and altitudinally between 2. 0 an d 0. 7 Ma ag o (Ljungne r 1949 ; variable. The mos t prominen t valley s o f th e preglacia l Kleman 1992) . A n importan t threshol d wa s passed a t c . 0.9-0. 7 Ma (e.g . Porte r 1989 ; landscape wer e presumabl y exploite d b y earl y Raymo e t al 1989 ; Imbrie e t al 1992 , 1993 ; ice sheet s a s primary route s of ice drainage, an d Clark & Pollar d 1998) , afte r whic h ic e sheet s hence became primar y location s fo r subglacia l grew large r an d becam e centre d eas t o f th e bedrock erosio n an d valle y deepening . mountain elevatio n axi s (Klema n & Stroeve n Subsequent glaciation s would the n hav e prefer1997). However , wha t i s no t wel l establishe d i s entially exploite d thes e deepene d valleys , the timin g an d patter n o f erosio n an d landscap e because the y woul d hav e bee n mor e efficien t development associate d wit h thi s glacia l for ic e drainag e wit h eac h successiv e glaciatio n cycle (Sugde n & John 1976) . Th e implicatio n of chronology. It ha s lon g bee n recognize d tha t th e typica l this mode l i s tha t intervenin g highland s wer e geomorphology o f circum-Atlanti c continenta l covered b y relativel y stagnan t ice , ensuring tha t margin mountain s i s a patchwor k o f glaciate d subglacial froze n conditions , an d henc e land terrain and relict upland surfaces (Gjessing 1967 ; scape preservation , prevailed . Th e mos t visibl e alteration o f relic t uplan d surface s sinc e th e Sugden 1968 ; Sugden & Watt s 1977 ; Hall & commencement o f glaciatio n i n th e circum Sugden 1987 ; Ballantyn e 1994 ; Klema n & Stroeven 1997) . Fo r example , remnan t surface s Atlantic regio n ha s bee n th e deepenin g o f were recognize d i n th e Norwegia n an d Swedis h V-shaped valley s terminating in glacia l troughs. mountains earl y i n th e twentiet h centur y b y The deepenin g o f thes e V-shape d valley s Reusch (1901) , Wra k (1908 ) an d Ahlman n presumably happene d i n interglacia l times , a s (1919). Glacia l terrai n i s characterize d b y th e these river s adjuste d thei r profile s b y vertica l presence of U-shaped valleys , cirques, horn s and erosion toward s th e ne w base-leve l condition s arretes, valle y an d mountai n truncation , moun - produced b y troug h deepenin g durin g previous tain asymmetry , an d th e ubiquitou s presenc e of glaciation(s) (Ahlman n 1919 ; Rudberg 1992) . lakes (e.g . Sugden & Joh n 1976) . Relic t (o r The implication of this geomorphological history remnant) upland surfaces, on the other hand, are for interpretin g circum-Atlanti c continenta l characterized b y windin g V-shape d valleys , margin sediment s durin g Neogen e glaciatio n i s mountain symmetry , tors , weatherin g mantle s that offshor e sediment s wer e derive d primaril y and a n absenc e o f (water-filled ) roc k basin s from areall y an d altitudinall y restricte d sourc e (Figs 1 an d 2) . A patchwor k occurrenc e o f areas, rather than from equal erosion across much glacially scoure d an d relic t surface s indicate s of th e landscape . that lat e Cenozoi c subglacia l erosio n mus t have This proposed mode l of the geomorphologica l been areall y variabl e (Sugde n 1968 , 1974). This history i s base d o n fiel d an d remot e sensing patchwork patter n o f glacia l erosio n an d based mapping and interpretation of geomorpho preservation coul d eithe r resul t fro m restricte d logical feature s (e.g . Klema n & Stroeven 1997) . ice extents (glaciers limited to major valleys) or, This interpretation lead s to the following testabl e where we know that large ic e sheets covere d th e implications: (1 ) large-scal e relic t bedroc k mountains, fro m comple x pattern s o f ic e shee t morphology shoul d hav e surfac e exposur e age s much olde r tha n adjacen t glaciall y cu t surface s basal therma l regimes . On the basis of observations o f a patchwork of (potential ag e differences o f 10 5-106 years); (2) glacially scoure d an d relic t surfaces , typica l fo r bedrock erosio n o n glaciall y cu t surface s ha s glaciated passiv e margi n mountains , th e sub- been order s o f magnitud e highe r tha n o n relic t glacial therma l regime of average Quaternary ic e surfaces (rangin g fro m abou t 1 0 t o 10 3m); (3 ) sheets (Porte r 1989 ) was froze n o n th e upland s many bedroc k surface s hav e undergon e a and meltin g i n th e mai n valleys , wher e outle t complex histor y of multipl e burials (underneath glaciers an d ice-stream s forme d (Sugde n 1968, non-erosive ice ) and re-exposures . Thes e impli 1974). Relic t surface s ar e bes t preserve d a t cations, an d thu s th e large r mode l fo r th e intermediate elevations, lo w enough no t to hav e geomorphological history o f the circum-Atlanti c been covere d b y cirqu e glaciers , an d apparentl y mountains, can now be tested usin g an approach high enough not to have experienced melted-be d based o n measurin g multiple in situ cosmogenic conditions an d subglacia l erosio n durin g ic e nuclides in exposed bedrock . sheet overridin g event s (Klema n & Stroeve n This paper serve s to establish a framework for 1997). Hence , th e morphologica l erosiona l the interpretatio n o f i n situ cosmogeni c nuclid e impact o f glacier s an d ic e sheet s overridin g an d concentrations i n term s o f th e timing , pattern s expanding throug h thes e circum-Atlanti c con- and magnitud e of bedrock erosio n o f landscape s tinental margi n mountain s lef t a distinc t land - affected b y glacial , fluvia l an d periglacia l

IN SITU COSMOGENI C NUCLIDE S AN D PASSIV E MARGIN S

process systems . Formulating suc h a framework in advance minimizes the use of special pleading to interpre t forthcomin g dat a fro m glaciate d passive margi n mountains . The practica l appli cation o f thi s framewor k t o th e geologica l an d geomorphological 'traces ' o f th e last glacia l cycle establishes a set of attributes (typically, the timing, amoun t and patterns o f erosion) tha t can also serv e a s a basi s fo r interpretin g geo morphological evidenc e fro m prio r glacia l cycles. First , w e provid e a n overvie w o f cosmogenic nuclid e technique s an d theoretica l considerations. W e the n discus s th e geomor phological aspect s o f landscap e surfac e recon struction o n glaciate d passiv e margins . Finally , we illustrat e potentia l use s o f cosmogeni c nuclide technique s i n testin g erosio n historie s for glaciate d passiv e margins .

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Cosmogenic nuclid e productio n rate s var y with latitud e an d altitud e becaus e o f th e dissipation of cosmic radiatio n within the Earth's atmosphere an d th e dependenc e o f th e cosmic ray flu x o n the strengt h o f the Earth' s magneti c field (La i 1991 ; Robinso n e t a l 1995 ; Dunai 2000). Othe r variable s influencin g cosmogeni c nuclide productio n rate s ar e shielding o f incom ing cosmi c ray s b y surroundin g terrain , th e geometry of the sample surface, and erosion and/ or buria l o f th e sampl e sit e (La i 1987 , 1991). Provided th e geomorphologica l context o f th e sample is understood, appropriat e correction s fo r each of these variables ar e available (Nishiizumi etal 1989 ; Lai 1991; Dunne etal 1999 ; Fabel & Harbor 1999) . Surface exposure dating and erosion rates

With prolonge d exposure , cosmogeni c nuclide s accumulate withi n rock a s a function o f time and depth below th e surface . The time elapse d sinc e initial exposur e o f th e roc k surfac e ca n b e The cosmogenic radionuclides 10 Be and 26A1 and calculated fro m cosmogeni c nuclid e concen the cosmogenic stabl e nuclide 21Ne are produced trations i n th e rock , usin g know n rate s o f in rock s nea r th e groun d surfac e b y reaction s production. I f th e surfac e undergoe s erosion , with secondar y an d tertiary cosmic-ray neutrons depth profiles of nuclide concentrations, ratios of and muon s (Lai & Peters 1967) . Thes e nuclide s different nuclide s i n th e rock , an d nuclid e are commonl y use d i n studie s o f landscap e concentrations i n sediment s ca n al l provid e evolution (reviewe d b y Nishiizum i e t al. 1993; measures o f th e erosio n rat e (Cerlin g & Crai g Bierman 1994 ; Cerlin g & Crai g 1994 ; Fabe l & 1994; Grange r e t al . 1996) . Th e compariso n o f Harbor 1999) . Thi s i s because al l three isotope s concentrations of stable and radioactive isotope s are produced withi n a few metres o f the Earth' s can eve n facilitat e th e unravellin g o f comple x surface i n quartz, a ubiquitous mineral i n crustal histories o f exposure an d burial suc h as occurre d rocks and sediments, which has a simple 16 O and on glaciated passiv e margins. 28 Si target chemistry and a tight crystal structure Erosion o f a rock surfac e leads t o removal o f that minimizes diffusio n an d contamination . accumulated cosmogeni c nuclide s an d henc e a Applications of in situ produced cosmogenic 10 Be, 26 A1 and 21Ne

Fig. 1 . Landscap e typica l fo r glacia l erosion . Cirqu e glaciers hav e deepene d an d widene d bedroc k depressions i n the Part e Massif, Sare k Nationa l Park , northern Sweden , a mountai n massi f tha t wa s als o a prominent par t o f th e preglacia l landscape . However , the glacie r forefield , althoug h riddle d wit h lakes , ha s not bee n significantl y erode d belo w it s preglacia l elevation.

Fig. 2 . Landscap e shape d b y non-glacia l processes . Upland surfaces , Tarrekais e Massif, norther n Sweden , are though t t o retai n relic t morphology . Th e fluvia l valley, wit h interlockin g spurs , i s considere d t o b e younger in age and cut in response to (local) base-leve l lowering b y glacie r erosion . Thes e interlockin g spur s also show that glacial erosio n playe d no significant rol e in the evolutio n o f valley pattern .

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reduction i n nuclid e concentratio n (Fig . 3) . Because productio n decrease s roughl y exponen tially wit h dept h belo w th e surface , th e accumulated cosmogeni c radionuclid e concen tration in a mineral grai n record s the speed with which tha t grai n ha s bee n uncovered ; slowe r erosion rate s impl y longe r exposur e time s nea r the surface , an d thu s higher concentration s (Lai 1991). Therefore, t o calculate th e exposure tim e from a measure d cosmogeni c radionuclid e concentration i n a sampl e require s tha t th e erosion rate is known. If independent erosion rate evidence i s no t available , ther e i s n o uniqu e solution t o th e exposur e ag e calculatio n wit h a single radionuclide . However , a maximu m steady-state erosio n rat e ca n b e calculated . Th e cosmogenic radionuclid e concentratio n ca n b e used to calculate th e steady-state erosion rat e for the surfac e i f th e exposur e tim e ca n b e independently constrained , o r vic e vers a (se e Lai (1991 ) fo r equations). The measuremen t o f 10 Be an d 26 A1 concen trations in the sam e sampl e ca n provide a means of estimatin g bot h th e exposur e tim e an d th e steady-state erosio n rat e o f th e sampl e becaus e the ratio of the 26A1 and 10 Be concentrations i n an

eroding horizon changes sensitively with the rate of erosio n (La i & Arnold 1985) . Th e z °Al/luBe production rat e rati o i n quart z i s 6. 0 ±0.3 (Nishiizumi et al 1989) , regardles s o f th e absolute productio n rate . Becaus e 26 A1 decay s more rapidl y tha n 10 Be, th e 26 Al/10Be rati o decreases wit h increasin g exposur e time . Fo r a continuously expose d sampl e wit h n o erosion , the 26 Al/10Be rati o wil l follo w a smoothl y varying trajector y reachin g a n en d poin t where productio n an d radioactiv e deca y ar e balanced fo r bot h isotope s (Fig . 4 ; constan t exposure curve) . I f th e sam e sampl e i s subjec t to steady-stat e erosio n i t i s losin g mas s fro m the surfac e an d th e 26 Al/10Be rati o shoul d li e on th e steady-erosio n curv e a t a poin t determined b y th e erosio n rat e (Fig . 4) . Provided th e cosmic-ra y intensit y ha s remaine d constant th e 26 Al/10Be rati o wil l plo t betwee n these curve s (steady-stat e erosio n island ) fo r any simpl e exposure history under conditions of steady-state erosion . I t i s therefor e possibl e t o determine th e steady-stat e erosio n rat e an d exposure tim e o f a sampl e b y measurin g tw o cosmogenic radionuclide s wit h differen t half-lives i n the sam e sample .

Fig. 3. 10Be concentration v. depth for three differen t exposur e times and three different steady-stat e erosion rates. The first number i n the ratios is the exposure time (ka ) and the second i s the steady-state erosio n rate (cmka" 1). The curve s include 10 Be production by muon s and were calculate d accordin g to Granger & Smit h (2000) fo r a 10 Be productio n rat e o f 5. 1 atoms g"1 (SiO 2) yr" 1 an d a densit y o f 2.7gcirT 3. I f a glacia l erosio n even t removes 100c m o f bedrock (horizonta l dashe d line ) fro m th e surfac e afte r 10 , 10 0 and lOOOk a exposur e wit h zero steady-stat e erosio n (continuou s curves) , th e inherite d 10 Be concentration s in th e resultin g 'new ' surfac e are equivalen t t o apparen t 10 Be age s o f c . 2ka , c . 18k a an d c . 155ka , respectively.

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Fig. 4 . 26 Al/10Be ratio plotted agains t 10 Be concentration . 26 A1 (half-life 705 ka) decay i s relatively rapi d wit h respect to 10Be (half-life 150 0 ka), forcing the 26Al/10Be ratio of a sample to decrease exponentially over time (Lai & Arnold 1985 ; Klei n et al 1986) . If there is no erosion, 26 Al/10Be ratios will fall somewher e alon g the blue line, yielding an apparent exposure age. If there is steady-state erosion (measure d here in metres per million years), the ratio lies on the red steady-stat e erosion curv e a t a point determine d b y the steady-stat e erosion rate. The area between th e two curves is called th e steady-state erosio n islan d (La i 1991) . I f ratios plo t below tha t island, the n burial is inferred. Upon burial, ratios wil l fall i n the direction o f the dashed blac k arrow s an d burial tim e can be inferred whe n measure d ratio s are related to the 10 Be concentration in the sample (green lines) .

Erosion depths from cosmogenic radionuclide inheritance Passive continenta l margin s i n glaciate d regions, suc h as those i n Scandinavia , Scotland, Greenland, Canad a an d Antarctica , have appar ently repeatedly experienced hundred s of metres of erosio n underneat h outle t glacier s wherea s adjacent upland s remaine d unmodifie d (Sugde n 1968, 1974 ; Hall & Sugden 1987; Glasser & Hall 1997; Kleman & Stroeven 1997) . Hence, there is a remarkabl e spatia l variatio n i n subglacia l erosion tha t ma y b e reflecte d i n cosmogeni c radionuclide inheritance. Inheritance refers to the remnant cosmogeni c radionuclid e concentration from a prio r exposur e history . Thu s fa r ou r discussion ha s largel y deal t wit h cosmogeni c nuclide accumulation in the absence of, or under conditions o f steady-stat e erosion . Glacia l erosion i s a non-steady-stat e event . Whe n ic e overrides a rock surface it effectively shield s that surface fro m cosmi c ray s an d an y furthe r cosmogenic nuclid e accumulation . I f th e over riding even t i s erosive , cosmogeni c nuclide s accumulated befor e th e glaciatio n wil l b e removed. Dependin g o n th e dept h o f glacia l erosion, thi s cosmogeni c nuclid e remova l ma y

not b e complete. I n this case the surfac e retain s some o f th e cosmogeni c nuclide s and henc e a n inheritance fro m th e previou s exposur e even t (Fig. 3). The depth of erosion required to remove the entir e cosmogeni c nuclid e inventor y o f a previously exposed rock depends on the duration and erosio n rat e o f th e prio r exposur e histor y (Fig. 3) but is of the order of up to several metres. In case s wher e th e timin g o f glaciation s i s independently constraine d cosmogeni c nuclid e inheritance ca n b e used t o calculate the amount of roc k remove d b y th e glacia l even t (Brine r & Swanson 1998 ; Fabel & Harbor 1999) . Complex exposure and shielding histories Another challeng e i n datin g landscap e surface s on glaciate d passiv e continenta l margins is that the effects o f burial by ice in a complex exposure and burial history for long-lived surfaces need to be addressed . Man y bedroc k outcrop s o n glaciated passiv e margi n mountain s underwent a complex histor y of burials (typical duration of c. lOOka ) an d re-exposure s (typica l duratio n of c. 10 ka) (e.g. Kleman & Stroeven 1997) . We can approach unravelling complex burial histories by comparing multipl e cosmogeni c radionuclide s

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can be quantified. A comparison wit h PliocenePleistocene offshor e deposit s require s tha t th e amount, pattern and timing of this glacial erosio n is bette r understood . A s a firs t ste p w e nee d t o consider th e principle s fo r reconstructin g th e reference surface across areas that were dissected by glacial erosion . On th e larges t landscap e scal e tha t w e consider, tha t o f individua l mountai n blocks , glacial valley s an d th e mountai n foreland , reconstructing th e elevatio n an d shap e o f th e reference surfac e involve s considerations o f (1 ) the configuratio n o f th e preglacia l fluvia l drainage system , (2 ) fluvia l incisio n durin g interglacial periods, (3 ) glacial erosion of valleys by selective linea r erosion, (4 ) glacial erosion b y areal scouring , (5 ) localize d subglacia l modifi cation o f preglacia l uplan d surfaces , an d (6 ) lowering o f preglacia l uplan d surface s b y nonglacial processes . W e wil l firs t conside r th e significance o f eac h o f thes e erosiv e phase s fo r preglacial landscap e modification , an d the n consider th e applicatio n o f i n situ produce d cosmogenic nuclid e techniques to illuminate the magnitude, patter n an d timin g o f thes e erosio n phases.

with differen t productio n rate s an d half-live s (Bierman et al 1999 ; Fabe l & Harbor 1999) . The cosmogeni c nuclide concentratio n an d 26 Al/10Be rati o i n quart z a t th e groun d surfac e should plo t somewher e withi n th e steady-stat e erosion islan d (Fig. 4). If the surface is suddenly shielded fro m cosmi c radiatio n by ice, then 10 Be and 26 A1 production ceases . Radioactiv e deca y then lower s th e inherite d 26 Al/10Be rati o ove r time alon g th e dashe d blac k arrow s (Fig . 4 ) because 2e >Al decays faster than 10 Be (Lai 1991) . Burial tim e (gree n curve s i n Fig . 4 ) ca n b e inferred whe n measured ratio s ar e related t o the 10 Be concentratio n i n th e sampl e (Grange r & Muzikar 2001) . A s i n th e cas e o f obtainin g steady-state erosio n rates an d exposure duration , burial datin g requires steady-stat e conditions . A sample experiencin g a non-steady-stat e erosio n event afte r prolonge d exposur e a t lo w steady state erosion rates will also plot below the steadystate erosion island, but in this case burial cannot be inferred. This serve s t o illustrate the need fo r careful geomorphologica l interpretatio n whe n selecting sampl e sites . Although surfac e exposur e ages , erosio n rates , and buria l age s ca n b e obtaine d fo r well constrained situations , comple x surfac e exposur e histories ar e far more challenging . Interpretatio n of cosmogeni c radionuclid e concentration s i s much mor e complicate d i f a surfac e ha s experienced multipl e period s o f exposure , ero sion and burial, each of variable lengt h (Bierman etal. 1999). Stable nuclides such as N e provide an importan t additiona l too l t o assis t i n investigating comple x exposur e histories . Because 21 Ne i s stable , ther e i s n o reduction i n concentration durin g periods o f burial b y ice , i n contrast t o the radionuclides. Thus measurement s of 21 Ne provid e a n indicatio n o f tota l exposur e time, with loss owing to erosion, bu t not to burial. Measurement o f ! Ne permit s studie s o f exposure historie s beyon d th e limi t o f a fe w million year s imposed b y the l °Be and 26A1 halflives, an d compariso n wit h 10 Be an d 26 A1 concentrations shoul d provid e furthe r insigh t into buria l history . Thus , cosmogeni c 21 Ne ca n provide additiona l glaciologica l informatio n o f pertinence t o landscap e histor y an d landscap e surface reconstructio n (Niederman n et al. 1993) .

Configuration of the preglacial fluvial drainage system The preglacia l evolutio n o f th e fluvia l drainag e pattern o f th e passive-margi n mountain range in northern Scandinavi a ha s bee n controlle d pri marily b y uplift-induce d base-leve l change s i n the Balti c depressio n (Wra k 1908 ; Rudber g 1954; Lidmar-Bergstro m & Naslund 2002) . Th e eastern mountai n foreland i s characterize d b y a stepped 'plain s wit h residua l hills ' morpholog y generally a t 300-400 and 400-550 m above sea level (e.g . Frede n 1994 , pp . 50-51), referre d to as the Muddus Plains (Wrak 1908) . The Muddus Plains are the youngest preglacial fluvial surfaces of th e region , th e younges t o f whic h wa s th e last bas e leve l fo r th e evolutio n o f th e preglacial fluvia l drainag e syste m i n th e mountain rang e an d th e on e tha t mus t b e use d in an y reconstructio n o f upstrea m mountai n geomorphology.

Geomorphological principles for landscape surface reconstruction: examples from northern Sweden The presenc e o f relic t surfac e remnant s i n th e northern Swedish mountains provides a reference surface agains t whic h subsequen t glacia l erosio n

Fluvial incision during interglacial periods Fluvial incisio n o f th e preglacia l mountai n morphology continue d afte r th e Neogen e onse t of glaciation . Fo r example , Kleman & Stroeve n 1997, Figs 1 0 and 11 ) showed the distributio n of what ar e presumabl y Plio-Pleistocen e fluvia l valleys i n th e northwester n Swedis h mountains .

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These valley s ar e al l containe d withi n relic t th e size , dept h an d gradien t o f thes e fluvia l surface remnants , ar e lowere d relativ e t o thi s valleys reflec t th e amoun t o f tim e tha t wa s surface by < 300 m, and join prominent glaciall y available for cuttin g them . River s probabl y cut modified U-shape d valley s (Fig . 5) . Potentially , these valley s during interglacia l time s an d i n

Fig. 5 . (a) Stereogram o f the largest fluvial valley in the northern Swedish mountain range, Atnajakka, joining a glacial valley, the Teusajaure valley. The V-shaped valley has been lowered by 300 m below the valley bench at its mouth (panel b). The river, of 9 km length, has a concave valley profile and is graded to a level that is 110 m above the Teusajaur e valle y floo r (Klema n & Stroeve n 1997) . Abov e th e 'old ' valle y benc h (pane l c ) th e preglacia l surface extends towards the summits of the Karsatjakka mountain block. The Teusajaure valley, on the other hand, has been lowered by glacial erosion , initiatin g the accelerated downcuttin g of Atnajakka, leaving a typical valley floor (ice-moulde d bedroc k surface s an d elongate d lake s occupyin g overdeepenings ) an d valle y wal l glacia l morphology. Th e occurrenc e o f preglacial surfac e remnants helps i n reconstructing th e preglacial morphology . Cosmogenic nuclid e concentration s o f bedrock surface s i n eac h o f these thre e setting s (preglacial , glacia l an d fluvial) would potentiall y yiel d widel y differen t results . Thi s is because, on the basis of the geomorphologica l interpretation of Kleman & Stroeven (1997), (1) preglacial surfaces were preserved underneath cold-based ice and would tend to yield high concentrations of isotopes an d low surface erosion rates, (2 ) glacial surface s underwent intense modificatio n a t som e point durin g glaciation , resettin g th e cosmogeni c isotop e clock , an d (3 ) fluvia l surfaces have recorded maximu m surface lowering during interglacial period s and would tend to yield the lowest cosmogenic isotop e concentration s an d highes t surfac e erosio n rate s (Copyright : Nationa l Lan d Surve y o f Sweden 2002). (b ) Topographical ma p with profiles (Courtesy of the National Land Survey , 2002. Excerp t fro m GSD-elevation data , cas e numbe r L2000/646) . (c ) Topographica l profile s illustratin g potentia l cosmogeni c nuclide samplin g site s (dots , expecte d relativ e age s indicated ) an d possibl e preglacia l valle y reconstruction s (dashed). Fo r Profile s A-A ' an d B-B' , tw o alternative s hav e bee n given , reflectin g a n uncertaint y i n th e geomorphological interpretatio n o f the exten t of the preglacial surface .

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A. P. STROEVEN ETAL.

response t o base-leve l lowerin g a s a resul t o f glacial erosio n (Ahlman n 1919 ; Rudber g 1992) . Kleman & Stroeven (1997 ) hypothesized that the largest fluvia l valley , Atnajakk a (Fig . 5) , i s th e oldest valley , whic h implie s tha t th e U-shape d

valley whic h Atnajakk a joins , th e Teusajaur e valley, should be one of the oldest glacial valleys in th e mountai n range . Extendin g thi s logi c t o conclude tha t othe r glacia l valley s (som e o f which ar e much larger) must be younge r in age,

Fig. 6. (a ) Stereogra m o f the junction between a smal l glacia l valley , Vealevuomus , in the lowe r centre , an d a much larger glacia l valley , Rautas, in the upper centre (water-filled) . The floor of Vealevuomus is hanging 180 m above the floor of Rautas valley (panel b). Hanging valleys are typical features of glaciated landscapes (Sugden & John 1976 ) and ofte n a consequenc e o f selectiv e linea r erosion . Althoug h ice modifie d bot h thes e preglacia l valleys, it did s o more severel y i n Rautas valley, presumably because i t was deeper, wide r and better aligned fo r ice flow . Preglacia l surfac e remnant s occu r o n highe r surface s surroundin g thi s junctio n an d ca n ai d i n reconstructing th e preglacial valley topography (pane l c). Cosmogenic nuclid e concentrations o f bedrock surfaces in thes e valley s coul d potentiall y yiel d differen t results . This i s because , give n this differenc e i n topography , Rautas valle y ma y hav e bee n significantl y modifie d b y subglacia l erosion , whic h woul d rese t th e cosmogeni c isotope clock , wherea s basa l ic e i n Vealevuomu s ma y hav e remaine d belo w th e pressur e meltin g poin t an d inhibited erosion . Henc e cosmogeni c isotop e concentration s on multipl e isotopes coul d addres s difference s i n glacial erosio n rate s (Copyright : Nationa l Land Surve y o f Swede n 2002). (b ) Topographica l ma p wit h profile s (Courtesy o f the Nationa l Lan d Survey , 2002. Excerp t fro m GSD-elevatio n data , cas e numbe r L2000/646). (c ) Topographical profile s illustratin g potentia l cosmogeni c nuclid e samplin g site s (dots , expecte d relativ e age s indicated) an d possible preglacia l valle y reconstructions (dashed).

IN SITU COSMOGENI C NUCLIDES AND PASSIV E MARGINS

on th e basi s o f th e siz e o f th e fluvia l valley s draining int o them , represent s on e possibl e approach. However , thes e fluvia l valley s ma y be th e upper-reac h remnant s o f fluvia l valley s that wer e onc e muc h larger , bu t fo r whic h th e lower reaches have been truncated as subsequent glaciers widene d thei r valleys . A detaile d geomorphological stud y o f th e valle y benche s may revea l th e dept h t o which th e preglacia l valley wa s cu t befor e glaciall y induce d accelerated erosio n commenced . Glacial erosion by selective linear erosion Glacial modificatio n o f landscap e surface s occurs i n tw o distinctiv e patterns , b y selectiv e linear erosio n o f pre-existin g depression s o r areas les s resistan t t o erosion , an d b y area l scouring o f topographically les s variabl e terrai n (Sugden 1968 ; Sugde n & Joh n 1976 ; Harbo r 1995). Selective linear erosion occurs when there are extreme spatia l variation s in rates of erosio n under a n ic e sheet . Becaus e pressure-meltin g conditions a t th e bas e o f a n ic e shee t ar e necessary fo r bot h basa l slidin g an d extensiv e glacial erosion , an d thes e condition s ar e mos t common unde r thicke r ice , pre-existin g depressions an d the deepest o f preglacial valley s are lowere d preferentiall y b y glacia l erosion . This enhances relief an d has a positive-feedback effect o n subglacia l temperatur e an d pressure melting patterns (Oerlemans 1984 ; Maz o 1991) . Thus pre-existin g depression s an d valley s aligned wit h the ic e flow direction ar e preferentially eroded, an d other areas are subject to much less o r no erosion . Two feature s typica l o f passiv e margi n land scapes that can be explained in terms of selectiv e linear erosio n ar e shar p glacia l sur f ace-preglacial surfac e boundarie s (Fig . 5 ) an d hangin g valleys (Fig . 6) . The tota l relie f i n the norther n Swedish mountain s i s > 1 km bu t relie f withi n existing upland remnants is c. 600 m (Kleman & Stroeven 1997) . Hence , ther e wa s ampl e preglacial relie f t o establis h stron g subglacia l temperature gradient s favourin g preferentia l deepening o f pre-existin g fluvia l valley s o r preglacial depressions . I t is possible t o constrain the origina l dept h o f th e fluvia l valley s o r preglacial depression s base d o n (1 ) th e inferre d amount o f glacia l lowerin g fro m th e dept h o f 'glacial-age' fluvial valleys joining them, (2) the shape and gradient o f bordering relict slopes, and (3) fluvia l valle y gradient s betwee n rar e pre served 'upland ' preglacial-ag e valle y floors and the youngest lowland Muddus Plain. On the basis of these line s of reasoning i t is apparent tha t the larger glacial valley s have not been deepened b y

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more tha n c . 400 m (Fig . 5 b an d c) , althoug h some valley s appea r t o hav e experience d mor e erosion locall y (Klema n & Stroeven 1997) . Glacial erosion by areal scouring On a n ice-shee t scale , zone s o f area l scourin g occur wher e th e subglacia l therma l effec t o f topographical convergenc e an d divergenc e i s of insufficient magnitud e t o counterac t th e overal l subglacial meltin g regime . Thes e condition s occur wher e topograph y i s no t pronounce d compared wit h th e thicknes s o f th e ic e sheet , and where the heat of ice deformation suffices t o initiate basa l melting . Thes e condition s wer e apparently presen t wes t o f th e elevation axi s i n the norther n Swedis h mountains , probabl y because ic e accelerate d toward s th e fjorda l coast o f Norwa y (Klema n & Stroeve n 1997) . West o f thi s regio n o f area l scour , ic e flo w crossed a threshold i n basal topograph y an d ic e thickness, leadin g t o condition s o f selectiv e linear erosio n an d th e preservatio n o f high elevation perche d relic t surface s alon g th e Norwegian coas t (Wra k 1908 ; Dah l 1966 ; Peulvast 1985) . Reconstructing th e origina l elevatio n an d shape of landscapes that subsequently underwent areal scourin g i s especially challengin g becaus e there are few, if any, undisturbed surfaces against which t o compar e modifie d surfaces . On e approach is to view the areally scoured landscape as, at minimum, the result of stripping of material produced b y preglacia l chemica l weathering . Hence, th e minimu m amoun t o f erosio n wa s approximately equa l t o th e thicknes s o f th e weathering mantle s tha t wer e remove d (Glasse r & Hal l 1997) . Limite d observation s i n th e northern Swedis h mountain s indicat e a weath ering mantle thickness of < 10 m (e.g. Lundqvist 1985; Peulvas t 1985 ; Hirva s e t al 1988 ; Olse n et al. 1996 ; Re a et al 1996) . I n addition, glacial erosion ma y hav e extende d wel l belo w th e preglacial weatherin g base. Carefu l examinatio n of overridde n features , suc h a s degrade d cirqu e headwalls (Kleman & Stroeven 1997) , may yield an additiona l estimat e o f th e amoun t o f soli d bedrock removed . Finally , convergenc e mus t be sought betwee n thes e forme r estimate s o f weathering mantl e an d bedroc k erosio n an d a n interpolated surfac e betwee n th e westernmos t occurrences o f preglacia l remnant s i n Swede n and the easternmost outlier s in Norway. Localized subglacial modification of preglacial upland surfaces Detailed studie s o f surfac e morpholog y (e.g . Kleman 1992 ; Clarhall & Kleman 1999 ) indicate

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that subglacia l reorganizatio n o f som e area s within relic t uplan d surface s ha s occurred . Typically, thi s involve s sporadi c mino r erosion , transport an d depositio n o f th e weatherin g mantle (lee-sid e scarps , stoss-sid e moraines) , reshaping o f a weathering mantl e or til l int o till lineations (fluting) , erosio n o f bedroc k (striae) , and depositio n o f till . Withi n relic t area s evidence fo r subglacia l reworkin g i s generall y confined t o shallo w depression s (Klema n e t al. 1999). Boundarie s betwee n nonglacia l an d glacial surface s ar e sometime s extremel y shar p and hav e bee n interprete d a s longitudina l an d transverse subglacia l slidin g boundaries (Klema n & Borgstro m 1990 ; Klema n 1992) . I n the latte r case, erosio n occurre d o n th e le e sid e o f relic t hills (creatin g a scarp ) wherea s zone s o f deposition (o f th e erode d weatherin g mantle ) occurred o n the stos s side of relict hills (Kleman et a l 1999) . Currently , th e timin g o f landscap e modification ca n onl y be hypothesize d fro m th e association o f glacia l landform s (fo r example , the transvers e lee-sid e scarps ) wit h th e inferre d ice flow direction (and its age relative to other ice flow events).

processes o n slope s ar e ubiquitous , a s i s mountain-top detritu s o n uplan d flat s (e.g . Re a et a l 1996) . Th e presenc e o f preserve d periglacial phenomen a withi n uplan d relic t landscapes (e.g . Clarhal l & Klema n 1999 ) a s well a s activ e form s i n lowlan d location s indicates tha t landscape lowerin g b y periglacia l processes dominate s an d i s probabl y o f inter stadial age. Tors are also conspicuous features of periglacial landscapes. Generally, it is considered that tor s for m a s dee p chemica l weatherin g features (a n uneve n weathering surface) , whic h are subsequentl y uncovere d b y strippin g pro cesses (e.g. Linton 1955) . Suc h tors exhibit clear sheeting and rounded edges an d are found i n the northern Swedis h lowlands, where they survived despite ic e overridin g (e.g . Hattestran d & Stroeven 2002 ; Stroeve n et a l 2002) . However , tors i n th e norther n Swedis h mountain s exhibit an angula r frost-shattere d appearance , an d thi s type o f to r ha s previousl y bee n considere d t o have forme d o r t o hav e bee n uncovere d i n a periglacial environmen t (Palme r & Radle y 1961).

Lowering of preglacial upland surfaces by nonglacial processes

Application of in situ cosmogeni c nuclide techniques to passive margin erosion history

The lowerin g o f preglacia l uplan d surface s b y nonglacial processe s i s probabl y mainl y b y periglacial activit y a s indicate d b y (1 ) th e absence o f linea r feature s o f fluvia l origin , (2 ) the ubiquitou s presenc e o f gentl e convex concave slop e profiles , (3 ) th e presenc e o f sporadic permafrost , especiall y a t highe r elevations, whic h promote s periglacia l surfac e activity, (4) the ubiquitous presence o f active and relict periglacia l phenomen a acros s uplands , and (5) th e presenc e o f tor s o n interfluves . Th e absence o f fluvia l erosiv e feature s o n relic t upland surfaces , suc h a s dendritic rive r pattern s with interlockin g spurs , indicate s tha t fluvia l processes lef t n o significan t geomorphologica l imprint on the uplan d preglacial landscap e unti l accelerated fluvia l incisio n occurre d durin g interglacial periods . However , i n conjunctio n with th e sporadi c presenc e o f permafros t (Lundqvist 1962;Solli d e t a l 2000) , transpor t of solutes in (melt)water through the active layer probably promote d widesprea d bu t lo w rate s of long-term landscap e lowering . The mos t importan t processe s o f landscap e denudation i n thes e relic t areas , i t appears , ar e frost weatherin g an d periglacial slope processes, such a s solifluction . Fo r example , ston e stripe s and other sorted phenomena resulting from cree p

Measurements o f multipl e cosmogeni c nuclid e concentrations o f bedroc k surface s ca n poten tially be used to reconstruct th e timing and rates of landscap e chang e o f formerl y glaciate d passive continenta l margins . I n particular , the y can be used to (1) distinguish between surfaces of different age , (2 ) distinguis h betwee n surface s that experience d differen t erosio n histories , an d (3) addres s comple x exposur e an d shieldin g histories. O f these , calculation s o f long-ter m surface erosio n rate s ar e particularl y usefu l i n aiding reconstruction s o f th e origina l surfac e morphology acros s area s o f glacial erosion . The mos t fundamenta l postulat e tha t ca n b e tested usin g cosmogenic isotop e concentration s in bedroc k i s tha t 'preglacia l surfaces ' wer e preserved underneat h ic e sheet s an d thus can be used a s a referenc e horizo n agains t whic h t o measure th e magnitud e an d patter n o f glacia l erosion. I f these surface s wer e preserve d under neath cold-base d ice , the n th e cosmogeni c nuclide implication is that the difference betwee n the relativ e concentration s o f stabl e 21 Ne an d radioactive IO Be (i.e . th e 21 Ne-10Be contrast ) approaches th e maximu m possibl e give n th e glaciation histor y o f the passiv e margi n (Fig . 7 ) and that 26Al/10Be ratios are lower than expected for the other surfaces. If these surfaces are in fac t

IN SITU COSMOGENI C NUCLIDE S AN D PASSIV E MARGIN S

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Fig. 7 . Plot of the difference (%) between 21 Ne and 10 Be apparent exposure ages for a surface that has undergone a complex burial and exposure history imposed by the 8 18O record fro m DSD P Sit e 607 (top). The two curves are based o n the assumptio n tha t Mountain Ice Sheet s (MIS ) covered th e northern Swedis h mountain s when 8 18O was between 3.7 and 4.5%c, and Fennoscandian Ice Sheets (FIS) covered mountains and lowlands when 8 18O was >4.5%o. The grey shaded areas for both curves indicate 1 (Terrors if we assume that analytical precision fo r 10 Be and 21 Ne measurements ar e 5% and 20%, respectively .

not relic t a t all , an d wer e forme d eithe r i n postglacial time s o r b y significan t glacia l erosion, the n ther e shoul d b e n o relativ e difference betwee n thes e thre e isotopes ; tha t is , they should give the same apparent exposure age. There ar e thre e fundamentall y differen t bed rock surface s tha t coul d b e contraste d b y thei r cosmogenic nuclid e concentration s (Fig . 5) : relict surface s (unmodified , slightl y modified) , glacial surface s (selectiv e linea r erosion , area l scouring), an d fluvia l surface s (preglacia l age , glacial age). One would predict the following: (1) the unmodifie d relic t surface s woul d hav e th e highest concentration s o f isotopes , th e lowes t 26 Al/10Be ratios , th e highes t 2l Ne-10Be con trasts, an d th e lowes t surfac e erosio n rates ; (2 ) glacial surface s o f selectiv e linea r erosio n (U-shaped valleys ) woul d yiel d deglaciatio n ages o r olde r (dependin g o n th e amoun t o f glacial erosio n and , therefore , th e inheritanc e

signal), have 26Al/10Be ratios close to or equal to six, an d smal l 21 Ne-10Be contrasts ; (3 ) fluvial surfaces, if activel y forme d durin g interglacia l periods (Rudber g 1992 ; Klema n & Stroeve n 1997), includin g the presen t interglacial , woul d have th e lowes t cosmogeni c isotop e concen trations, 26Al/10Be ratios equal to six, 21Ne-10Be contrasts of zero, and the highest current surface erosion rates. If the stability of upland surfaces is supported b y th e isotopi c data , additiona l geomorphological interpretation s o f landscap e change ca n b e tested , an d preglacia l landscap e reconstructions based o n geomorphology ca n be strengthened. In th e cas e o f hangin g U-shape d valleys , cosmogenic nuclid e approache s provid e potential fo r differentiatin g between (1 ) glacia l valleys of different age , where the hanging valley is lef t 'hig h an d dry ' preserve d underneat h erosionally ineffectiv e ice , an d (2 ) valley s tha t

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are contemporaneou s bu t experience d differen t erosion rate s a s a resul t o f differin g subglacia l conditions (th e amoun t o f tim e availabl e fo r erosion and/or the rate o f erosion). I n the case of the Vealevuomu s an d Rauta s valley s (Fig . 6), if both valley s underwen t significan t erosio n during th e sam e glacia l even t (i.e . (2 ) above) , then bedroc k sample s fro m th e tw o location s should yiel d simila r exposur e ages , 26 Al/10Be ratios equa l t o six , and ^e-^B e contrast s of zero. However , i f basa l ic e i n Vealevuomu s remained belo w th e pressur e meltin g poin t an d inhibited erosio n (i.e . (1) above) , the n bedroc k samples i n thi s valle y shoul d hav e a n inheri tance signa l an d yiel d olde r apparen t exposur e ages, lowe r 26 Al/10Be ratios , an d large r 2 ^v[e-10Be contrast s tha n sample s fro m th e Rautas valley . Finally, cosmogeni c radionuclid e measure ments on bedrock samples ca n also contribute to resolving issue s o f glacia l modificatio n an d nonglacial lowerin g o f preglacia l uplan d sur faces. Fo r example , tor s ca n b e use d t o infe r minimum rates of preglacial landscap e lowering . Cosmogenic nuclid e concentration s o f multipl e isotopes from tor summit flats should indicate the timing o f to r exposur e an d long-ter m surfac e erosion rates . Thes e long-ter m surfac e erosio n rates on tors provide minimum long-term erosio n rates fo r th e surroundin g summi t flats . Thi s i s because th e erosio n rate s o f tors, assumin g their formation or uncovering occurred i n a periglacial environment, hav e t o be lowe r tha n those o f the surrounding summi t flats , otherwis e th e tor s would no t exist . A cosmogeni c radionuclid e study o f upland terrai n i n the Rock y Mountains, USA, yielde d a long-ter m surfac e lowerin g rat e of 10 m Ma" 1 fo r summi t flat s (Smal l e t al. 1997). Thi s valu e appear s reasonabl e fo r north ern Swedish condition s whe n compared wit h the height o f individual tors abov e thei r surrounding (generally < 10m ) and the amount of erosion by fluvial and glacial processes on adjacent surfaces (an orde r o f magnitude higher) . A framewor k fo r interpretin g cosmogeni c isotope concentration s i n term s o f patterns , rates an d timin g o f landscap e change , a s i s presented here , an d develope d befor e interpret ation o f actua l data , give s maximu m credibilit y to dat a interpretatio n an d help s i n effort s t o resolve difference s i n regional dataset s t o arriv e at a commo n interpretation . Th e practica l application o f thi s framewor k t o th e geologica l and geomorphologica l trace s o f th e las t glacia l cycle, wil l hel p establis h a se t o f attribute s (typically, erosio n rate s an d patterns ) tha t serv e as a workin g mode l fo r interpretation s furthe r back i n time .

Concluding remarks : uneve n sedimen t production o n glaciated passive margin s The glacia l chronologie s o f severa l passiv e margin mountain s ar e wel l understoo d a t a general leve l (e.g . Mangeru d e t al . (1996) , Kleman e l al . (1997 ) an d Klema n & Stroeve n (1997) fo r Scandinavia) . A patchwor k o f glaciated terrai n an d preglacia l relic t uplan d surfaces i n thes e mountain s reflects (1 ) the tota l erosive impac t of the lat e Cenozoic glacier s an d ice sheet s tha t covere d th e mountains , an d (2 ) that thi s subglacia l erosio n mus t hav e bee n areally restricted . What is not well established is the timin g an d patter n o f erosio n an d landscap e development associate d wit h thi s glacia l chron ology, a n issu e o f critica l importanc e whe n attempting a comparison o f th e onshor e erosio n histories an d offshor e sedimen t accumulation s (e.g. Glasse r & Hall 1997) . Conventional datin g an d analysi s hav e provided a n excellen t wa y t o begi n unravellin g the timing an d patter n o f erosion , landfor m devel opment, an d possibl e landfor m preservatio n under ice . However, attempt s t o utiliz e geomor phology an d remot e sensin g technique s hav e often bee n frustrate d b y th e limitation s o f conventional datin g techniques . W e describ e a new approac h fo r investigatin g landscape evol ution in mountainous areas, the in situ production of cosmogeni c nuclide s i n bedroc k surface s o f landscapes affecte d b y glacial, fluvial (preglacial or interglacial ) an d periglacia l (interstadial ) process system s (Cerling 1990 ; Nishiizumi et al. 1993; Bierma n 1994 ; Cerlin g & Crai g 1994 ; Small & Anderson 1995 ; Fabel & Harbor 1999) . Cosmogenic nuclide s produced i n rocks nea r the ground surface by reactions with cosmic rays can be use d t o determin e apparen t surfac e exposur e age an d landscap e preservation , an d constrai n erosion depth s an d duratio n of burial b y ice. This pape r present s a coherent framewor k for interpreting measurement s o f i n situ produce d cosmogenic nuclide s i n bedroc k i n term s o f timing, rates and patterns of landscape change on glaciated passiv e margins . Th e applicatio n o f cosmogenic nuclid e techniques to the reconstruction o f preglacia l surface s provide s ne w infor mation t o complemen t tha t obtaine d usin g traditional geomorphologica l approache s (Andersen & Nesj e 1992 ; Nesj e & Whillan s 1994; Rii s 1996 ; Glasse r & Hal l 1997) . Th e cosmogenic nuclid e techniqu e ha s previousl y been use d t o investigat e aspect s o f landscap e change i n glaciate d region s suc h a s th e ag e o f specific glacia l event s (e.g . Phillips e t al . 1990 , 2000; Broo k e t al. 1995 ; Ston e e t al 1996 ; Ivy Ochs e t al . 1997 ; Bierman e t al . 1999 ; Jackson

7W SITU COSMOGENI C NUCLIDE S AN D PASSIV E MARGINS et al 1999 ; Schafe r e t al 1999) , erosio n rate s over glacia l surface s (Nishiizum i e t a l 1989 ; Brook e t a l 1995 ; Brine r & Swanso n 1998) , landscape denudatio n rate s (Smal l e t a l 1997 ; Summerfield e t a l 1999) , an d th e elevatio n o f former ice sheet surface profiles (see Brook et al 1996; Ston e e t a l 1998 ; Acker t e t a l 1999 ; Kaplan e t a l 2001) . However , compare d wit h these othe r studies , a strengt h o f th e structur e presented her e i s tha t i t i s develope d before , rather tha n as a response to , dat a interpretation . We no w hav e a framewor k withi n whic h t o interpret nuclid e concentration s i n term s o f apparent exposur e ag e an d surfac e erosio n rates, an d whic h provide s a n overvie w o f sampling strategie s t o determin e th e timing , rate and pattern of landscape change on glaciated passive margins. This is particularly important as complex burial-shielding and exposure histories have probabl y affecte d al l sample s wher e substantial preservation length has been inferred from geomorphologica l evidence . Given the number of glacial cycles (c. 40) that have affected th e landscape in the last 2.7 Ma, we cannot expect cosmogeni c nuclid e techniques to resolve individua l events beyond the last glacia l cycle. However , throug h cosmogeni c datin g o f key landscap e element s w e ca n establis h th e pattern an d selectivit y o f erosio n b y th e las t ic e sheet, thereb y establishin g a mode l fo r th e erosional functionalit y o f ic e sheet s i n rugge d terrain. Onc e suc h a model i s erected , i t ca n b e used to provide insight into process pattern s and interrelationships acros s th e glacia l an d inter glacial cycle s o f th e Quaternar y period . Thi s provides a soli d foundatio n fo r attempt s t o improve geomorphologicall y base d reconstruc tions o f preglacia l surfaces , reconstruc t an d analyse th e dynamic s o f landscap e chang e i n glacial time , an d defin e th e consequence s o f different proces s regime s i n term s o f erosio n patterns, sedimen t transpor t an d th e suppl y o f sediments tha t ar e deposite d offshore . Th e thickness o f offshor e sediment s i s on e ke y ingredient considere d i n hydrocarbo n explora tion. However , despit e th e fac t tha t importan t quantities of sediment were generated by glacia l processes, curren t technique s o f estimatin g onshore exhumatio n pattern s an d rate s d o no t resolve margin exhumation during late Cenozoi c time (e.g.Lidmar-Bergstrom & Naslund 2001). It is thes e offshor e sediments , ofte n use d i n generalized recontruction s o f passiv e margi n development (e.g . Riis 1996 ) that , in th e future , will provide the link between offshore patterns of passive margi n sedimentatio n an d onshor e reconstructions o f passiv e margi n evolution . Furthermore, studie s o f th e erosiona l histor y o f

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upland landscape s wil l als o ai d i n resolvin g climate debate s wher e the sourc e are a o f erode d sediments i s o f ke y importance , suc h a s th e mechanisms leadin g to , an d th e implication s o f Heinrich event s fo r circum-Atlanti c climat e change an d landform development .

This paper has benefited fro m a detailed formal review by P . Bishop , an d fro m comment s b y K . Lidmar Bergstrom (informa l review) and A.G. Dore. Funding for par t of the research reporte d her e was provided by the Swedish Natural Science Research Counci l (NFR), Grants G-AA/G U 12034-30 0 an d G-AA/G U 12034 301, an d b y th e U S Nationa l Scienc e Foundation , Grant OP P 9818162. This manuscrip t was complete d while J.H. was supported by the New Zealand-United States Educationa l Foundatio n a s a Fulbrigh t Senio r Scholar.

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The thermotectonic developmen t o f southern Sweden durin g Mesozoic an d Cenozoic tim e CHARLOTTE CEDERBOM The CRUST Consortium, Department of Geology and Geophysics, Edinburgh University, Grant Institute, King's Buildings, Edinburgh EH9 3JW, UK (e-mail: cederbom@ glg.ed.ac.uk) Abstract: Lat e Carboniferous-Earl y Mesozoi c exhumatio n o f souther n Swede n ha s previously bee n trace d usin g apatit e fission-trac k thermochronology . I n addition , th e morphotectonic development of the region has been studied usin g geomorphology . The aim of thi s stud y i s t o attai n furthe r knowledg e o f th e Mesozoi c an d Cenozoi c thermotectoni c development o f souther n Swede n b y integratin g result s fro m thes e methods. Well-dated re-exposed palaeosurface s an d sedimentar y record s in the surroundin g area s were used a s constraints i n the modelling of apatite fission-track data from th e Precambrian basement. Th e obtaine d modelle d therma l historie s sugges t tha t souther n Swede n ca n b e divided into three main tectoni c areas associated with differen t cooling histories. In Triassic and Jurassic time, lo w to moderate exhumation i n the central part wa s accompanied by more rapid exhumatio n i n the S E and NW. Additionally, individua l block movements may hav e occurred i n th e NW . I t ha s als o bee n possibl e t o estimat e th e heatin g effec t o f renewe d Cretaceous-Paleogene burial t o 20-35 °C on the west an d SE coasts. Final Cenozoic unroofing o f the basement is indicated b y the modelled thermal histories. Areas aroun d th e souther n ti p o f Lake Vatter n togethe r wit h th e S E coast experience d th e most pronounce d exhumatio n compare d wit h th e surroundin g parts .

Southern Swede n i s characterize d b y a Precambrian basement , whic h emerge s a s a dome-shaped structur e fro m belo w lowe r Palaeozoic cove r rock s i n th e nort h an d eas t and Mesozoi c cove r rock s i n th e sout h an d west. Large-scal e Phanerozoi c event s i n th e area ca n therefor e b e identifie d onl y by therma l and relie f studie s of the basement . Palaeozoi c heating o f centra l an d souther n Swede n a s a result o f th e developmen t o f a Caledonia n foreland basi n ha s bee n establishe d (Larson et al 1999 ; Cederbo m e t al. 2000 ; Cederbo m 2001). I n addition , large-scal e Palaeozoi c t o Early Mesozoi c exhumation , perhap s accompanied b y tectonism , ha s bee n demon strated base d o n apatit e fission-trac k (AFT ) data fro m souther n Swede n (Cederbo m 2001) . From anothe r direction , Mesozoi c an d Tertiar y morphotectonic event s have been inferre d base d on studie s o f th e relie f an d it s relatio n t o remnants o f the cove r rock s (Lidmar-Bergstro m 1994, 1996) . I n thi s study , result s fro m AF T thermochronology ar e integrate d wit h geo morphological dat a t o attai n furthe r insigh t into th e Mesozoic an d Cenozoic thermotectoni c development o f souther n Sweden .

Palaeosurfaces The palaeosurface s o f souther n Sweden (Fig . 1 ) have been analyse d and roughly dated by mean s of their relative position, remnants of Palaeozoi c and Mesozoi c cove r rocks , an d saprolit e occurrences (e.g . Lidmar-Bergstro m 1982 , 1988, 1995) . I n addition , a mode l fo r th e long term morphotectoni c evolutio n o f souther n Sweden ha s bee n suggeste d (Lidmar-Bergstro m 1994, 1996) . Th e mai n type s o f palaeosurfaces, defined an d interprete d b y Lidmar-Bergstrom , are presente d in Fig . 1 and summarize d belo w (see Lidmar-Bergstro m 1996 , Fig . 2) . In th e nort h an d east , Cambria n strat a res t unconformably o n Precambrian basemen t rocks . The relie f o n this basement surfac e is extremely flat. I t i s possibl e t o trac e thi s exhume d Sub Cambrian Peneplai n (SCP ) fro m belo w lowe r Palaeozoic deposit s in the Lake Vattern and Lake Vanern are a an d alon g th e eas t coast , t o th e summits i n th e centra l par t o f souther n Swede n (Fig. 1 ; Lidmar-Bergstrom 1988) . A hill y etc h surfac e emerge s fro m belo w Upper Cretaceous cove r rocks in the SE and SW (Fig. 1) . The hill y relief i s associated wit h thick

From: DORE , A.G. , CARTWRIGHT , J.A. , STOKER , M.S., TURNER , J.P . & WHITE , N . 2002 . Exhumation o f th e North Atlantic Margin: Timing, Mechanisms an d Implications for Petroleum Exploration. Geologica l Society , London, Special Publications, 196, 169-182 . 0305-8719/027$ 15.00 © The Geological Societ y of London 2002 .

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Fig. 1. Simplified map of palaeosurfaces an d Phanerozoic sediments in southern Sweden (modified fro m LidmarBergstrom 1996) . Sample localities for apatite samples modelled in this study (•) ar e shown together with sample localities fo r additiona l apatit e sample s no t use d i n thi s stud y (O) (Larso n e t al. 1999 ; Cederbom el al. 2000 ; Cederbom 2001) . A, B and C indicat e the three mai n tectonic area s discusse d in the tex t (borders marked with bold dashed lines) . The elevations fo r the two highest points in the area are given in metres above sea level (a.s.L). CDF, Caledonia n Deformatio n Front ; FBZ, Fennoscandian Border Zone; SKP , Skagerrak-Kattega t Platform.

kaolinitic saprolit e remnants , consisten t wit h deep weatherin g durin g war m an d humi d conditions (Lidmar-Bergstro m 1995) . Thi s typ e of relief continue s alon g the entire west coast and is interpreted a s a re-expose d sub-Mesozoi c palaeosurface (Lidmar-Bergstro m 1994) .

A disintegrated part of the SCP is documented at c . 200 m elevatio n i n th e centra l par t o f southern Swede n (Fig . 1) . Farthe r south , a n almost horizonta l plain with low relief i s define d as th e Sout h Smalan d Peneplai n (SSP) . Thi s surface truncate s the sub-Mesozoi c etc h surfac e

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Fig. 2. Extensive lineaments (a), open linear structures (b) and sharp and straight linea r structures , shar p line s (c), i n souther n Swede n (modifie d fro m Tire n & Beckholmen 1992) .

(Lidmar-Bergstrom 1982 ) at abou t 100-125 m above se a leve l (a.s.l. ) an d i s therefor e inter preted a s a younge r surface . Frequent remnant s of gravelly saprolites, consistent with weathering during cold and humid conditions, are associated with thes e surface s (Lidmar-Bergstro m el ai 1997). Th e ag e o f th e gravell y saprolite s i s difficult t o constrain, but they are certainly olde r than Late Weichselian time , and their position in areas protecte d fro m glacia l erosio n give s support fo r a t leas t a Plio-Pleistocen e age . Possibly, the y ha d forme d alread y i n Miocen e time (Lidmar-Bergstro m e t al 1997) . Th e SS P and th e disintegrate d par t o f th e SC P ar e bot h interpreted a s Tertiary surface s that were formed by stepwis e exhumatio n an d exposur e o f th e basement durin g war m an d dr y condition s

(Lidmar-Bergstrom 1991) . Furthermore , th e vast exten t o f th e SS P indicate s a considerabl e time fo r it s formation. Faults and lineaments The regiona l occurrenc e an d exten t o f Phaner ozoic fault s i n souther n Swede n i s no t wel l documented. Ahli n (1987 ) presente d a stud y of Phanerozoic fault s i n th e northwester n part o f southern Sweden , base d o n fiel d wor k an d th e reconstruction o f th e SCP . Frequent NE-SW trending Phanerozoi c fault s wit h a n offse t o f <100m hav e bee n observed , an d i n Lak e Vanern, a n offse t o f 'som e 10 0 metres' i s recorded (Ahli n 1987) . I n addition , large-scal e north-south-trending fractur e zone s appea r

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along the eastern sid e of Lake Vatter n (Persson & Wikman 1986) . Th e lak e i s propose d t o b e a graben structur e (e.g . Lin d 1972 ; Persso n e t al 1985), an d i t contain s a c . 1000 m thicknes s o f Late Proterozoi c sediment s (e.g . Vida l 1984) . However, i n most area s of southern Swede n only lineaments an d fractur e zones , whic h appea r either o n aeromagneti c map s and/o r i n the field, have been mapped . I t is unknown whether som e of the m constitut e fault s o r not . The y ar e presented i n a fe w 1:5000 0 bedroc k map s (covering c . 20 % o f th e investigatio n area ) an d six 1:25000 0 temporar y genera l bedroc k map s published b y th e Swedis h Geologica l Survey , most o f which ar e not digitized . Phanerozoic fault s and fractures have also been interpreted fro m th e Relativ e Relie f Ma p o f Sweden (Lantmateriverke t 1986) , contou r map s and vertica l topographi c profile s (Elvhag e & Lidmar-Bergstrom 1987; Lidmar-Bergstro m 1991, 1996) . Relativ e relief maps are constructed by computerize d hil l shading , i.e . obliqu e illumination o f a terrai n mode l tha t i s base d o n topographic map s (Elvhage & Lidmar-Bergstrom 1987). Fro m th e Relativ e Relie f Map , th e topographically well-expresse d fractur e lines , rather tha n th e complet e fractur e pattern , wer e identified (Elvhag e & Lidmar-Bergstro m 1987 , p. 347) . I n comparison, o n the basi s o f the sam e Relative Relie f Map , Tire n & Beckholme n (1992) presente d a n analysi s o f th e tectoni c rock bloc k patter n i n souther n Swede n togethe r with a detaile d analysi s o f th e variou s fractur e sets. T o date , thi s i s th e onl y publishe d compilation o f linea r structure s for th e whol e of southern Sweden . In Fig . 2 , thre e type s o f structure s ar e illustrated (extensiv e lineaments , ope n linea r structures, shar p an d straigh t linea r structures) , which were presented an d interpreted by Tiren & Beckholmen (1992) . Th e uncertaint y of individual structure s i s high , bu t som e genera l trend s can be discerned fro m th e lineament pattern. The plot o f extensiv e lineament s (Fig . 2a ) illustrates that a bel t o f c . 50k m width , includin g Lak e Vattern, run s i n a N20 E direction , dividin g southern Swede n int o tw o halve s (Tire n & Beckholmen 1992) . Southward s th e bel t widen s and deflect s westwards , ou t int o th e sea . T o th e south, th e orientation o f lineaments is influenced by th e NW-SE-oriented Fennoscandia n Borde r Zone (Fig . 1) . Open structure s (Fig . 2b) , whic h represent flexures , th e terminatio n o f th e SC P and wide , trough-lik e valleys , ar e scattere d throughout souther n Sweden . I n contrast , shar p and straigh t linea r structure s (Fig . 2c ) ar e recognized mainl y i n th e northwester n par t o f southern Sweden. The surface nature of the sharp

lineaments i s no t known , bu t the y ar e probabl y the younges t structure s detecte d fro m th e Relative Relie f Ma p (Tire n & Beckholme n 1992). Modelling of AFT data The dating method Fission track s ar e create d b y th e spontaneou s fission o f 238 U an d ar e forme d continousl y through time . The y becom e instantaneousl y annealed a t hig h temperatures , an d th e numbe r of track s (i.e . th e trac k density ) i s therefor e proportional t o the time that has passed sinc e the grain coole d belo w a certai n temperatur e an d track retention started. For apatite, this temperature varie s b y severa l ten s o f degree s aroun d 100°C depending on the duration of heating and the minera l compositio n (Gleado w & Dudd y 1981; Naese r 1981 ; Carlson e t al 1999) . Furthermore, i t i s possibl e t o investigat e the thermal histor y i n apatit e b y studyin g th e distribution o f trac k lengths . Relativel y slowly cooled basemen t sample s ar e characterize d b y wide, ofte n negativel y skewe d trac k lengt h distributions wit h shorte r mea n trac k length s and highe r standar d deviation s compare d with , for example , volcani c sample s tha t hav e coole d very rapidly. Mixed track length distributions are formed whe n track s fro m earlie r event s hav e survived partia l reheatin g and a ne w population of tracks from th e las t period o f cooling is added to the older generation of tracks (Gleado w et a l 1986a, 1986b) . Forward modelling of AFT data Forward modellin g of AF T dat a i s the testin g of potential therma l historie s b y comparin g mod elled an d observe d AF T results . Th e annealin g model fo r apatit e presente d b y Ketcha m e t a l (1999) an d th e AFTSolv e modellin g progra m constructed b y Ketcha m et al (2000 ) wer e use d in thi s study . I n contras t t o earlie r publishe d apatite annealing models (i.e. Laslett et al 1987 ; Carlson 1990 ; Crowle y e t a l 1991 ; Laslet t & Galbraith 1996) , the Ketcham et al. (1999) model attempts to account for the variation in annealing behaviour tha t exist s betwee n differen t apatite s as a resul t o f compositiona l variations . Whe n comparing th e Ketcha m e t a l (1999 ) annealing model wit h the precursor models, it turned out to be mor e sensitiv e t o bot h low-temperatur e an d high-temperature annealin g tha n it s precursors . Generally, significantl y lowe r Lat e Palaeozoi c Early Jurassi c an d Paleogen e temperature s were obtained whe n usin g th e Ketcha m e t a l (1999 )

Table 1. Modelling results

Sample number Area A B14 S9819a S9819b P9910 S9620 S9621 P9909a P9909b S9622 S9623 9901 AreaB S9624a S9624b P9906a P9906b P9906c S9626 9802 9807 9808 Area C P9904 9905 9903a 9903b SA9629a SA9629b 9902 9805

Modelled MTL (ixm)

No. of grains

Observed FTA (Ma)

Modelled FTA (Ma)

GOF

No. of lengths

20 27 27 20 23 20 20 20 10 20 20

313 ± 25 208 ± 10 208 ± 10 161 ± 9 313 ± 17 220 ± 14 149 ± 8 149 ± 8 211 ± 16 188 ± 10 175 ± 12

322 209 214 160 313 226 150 149 218 188 176

0.72 0.92 0.55 0.94 1.0 0.68 0.86 0.98 0.65 0.99 0.90

50 100 100 100 100 110 59 59 100 100 90

13.6 12.7 12.7 12.7 13.1 12.7 13.1 13.1 12.4 12.8 13.6

± 0.2 ± 0.2 ± 0.2 ± 0.2 ± 0.2 ± 0.2 ± 0.2 ± 0.2 ± 0.2 ± 0.2 ± 0.1

13.7 12.7 12.6 12.7 13.2 12.6 13.1 13.1 12.5 12.8 13.6

± 0.2 ± 0.2 ± 0.2 ± 0.2 ± 0.2 ± 0.2 ± 0.2 ± 0.2 ± 0.2 ± 0.2 ± 0.2

0.58 0.31 0.30 0.26 0.50 0.51 0.71 0.74 0.64 1.0 0.39

11 11 20 20 20 20 21 20 20

261 ± 261 ± 213 ± 213 ± 213 ± 282 ± 231 ± 240 ± 231 ±

17 17 14 14 14 15 12 12 11

264 262 215 212 216 283 230 240 232

0.86 0.97 0.87 0.91 0.81 0.93 0.90 1.0 0.96

92 92 100 100 100 100 97 100 100

13.0 13.0 12.4 12.4 12.4 13.3 13.4 13.1 13.6

± 0.2 ± 0.2 ± 0.2 ± 0.2 ± 0.2 ± 0.2 ± 0.1 ± 0.2 ± 0.1

13.0 13.1 12.6 12.5 12.5 13.4 13.4 13.1 13.5

± 0.2 ± 0.2 ± 0.3 ± 0.2 ± 0.2 ± 0.2 ± 0.2 ± 0.2 ± 0.2

0.89 0.97 0.11 0.70 0.81 0.45 0.38 0.51 0.27

20 14 20 20 21 21 20 23

196 ± 12 188 ± 15 165 ± 10 165 ± 10 172 ± 9 172 ± 9 160 ± 9 189 ± 9

198 188 165 169 174 171 161 188

0.84 0.96 0.99 0.63 0.86 0.93 0.90 0.90

80 22 52 52 166 166 100 100

13.0 12.8 12.8 12.8 13.5 13.5 12.3 12.9

± 0.2 ± 0.2 ± 0.2 ± 0.2 ± 0.1 ± 0.1 ± 0.2 ± 0.1

13.0 12.9 12.8 13.1 13.5 13.4 12.4 12.9

± 0.2 ± 0.2 ± 0.2 ± 0.2 ± 0.2 ± 0.2 ± 0.2 ± 0.2

0.81 0.76 0.77 0.29 0.60 0.51 0.62 0.44

Observed MTL

K-S test

No. of grains, number of grains dated for each sample ; FTA, apatite fission-trac k age ; GOF, goodness o f fit; no. of lengths, number of track length measurements in eac h track lengt h distribution; MTL, mea n track length ; K- S test , Kolmogorov-Smirno v test . Th e datin g results were originall y publishe d b y Cederbo m e t al (2000 ) (fo r SA9629) and by Cederbom (2001) .

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C. CEDERBOM

Fig. 3. Modelled thermal histories for apatite samples fro m souther n Swede n (fo r sample localities , se e Fig. 1). Only th e 'discernibl e part' o f the modelle d Phanerozoi c history i s illustrate d fo r eac h sample .

annealing model compared with when the Laslett et al (1987 ) o r th e Crowle y e t al (1991 ) annealing mode l wa s adopted . A detaile d comparison betwee n th e annealin g model s wa s presented b y Ketcha m e t al. (1999) . The datase t use d i n thi s stud y doe s no t contain an y kineti c parameter s (i.e . trac k length orientation s o r etc h pi t widths) , an d th e

initial trac k lengt h wa s therefor e se t t o a fixe d value o f 16. 2 (jum. I t i s importan t t o hav e i n mind tha t th e therma l historie s presente d i n this stud y ar e no t unique . I n forwar d modelling, a s wel l a s i n invers e modelling , somewhat differen t therma l historie s tha t hav e not bee n trie d ma y als o fi t th e observe d AF T results.

THERMOTECTONIC DEVELOPMEN T O F SOUTHERN SWEDEN

Fig. 3. Continued

175

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C. CEDERBOM

In general, 2 0 grains should be dated to obtain a statisticall y reliabl e AF T ag e and , similarly, 100 track length s should be measured. Modelle d thermal histories for samples having fewer single grain ag e determination s and/o r trac k lengt h measurements (se e Table 1 ) must be interprete d with caution. Two statistical methods are used to gauge how well th e dat a an d th e mode l result s fit . Th e Kolmogorov-Smirnov test (K-S test) is used to compare th e measure d an d th e predicte d trac k length distributions and the GOF is the 'goodness of fit' between the age data and the age predicted by th e model . Fo r bot h statistics , a probability value >0. 5 implie s tha t th e modelle d therma l history i s 'supported ' b y th e dat a an d a probability value between 0. 5 and 0.05 indicates that th e modelle d therma l histor y i s 'no t rule d out' b y th e dat a (Ketcha m e t al 2000) . There fore, a value of 0.05 is regarded a s the lower limit for acceptance, although a value > 0.5 was aimed for when modelling. A detailed description of the statistical method s ha s bee n give n b y Ketcha m et al (2000) . Geological and geomorphological modelling constraints In modelling , severa l constraint s wer e set . Because al l sample s wer e collecte d a t o r clos e to the SCP , they ar e known to have experience d near-surface temperature s durin g Earl y Cambrian time . B y th e en d o f th e Caledonia n orogeny, souther n an d centra l Swede n wer e covered b y thick foreland basin deposits (Larso n et a l 1999 ; Cederbom e t a l 2000 ; Cederbo m 2001), resultin g i n tota l annealin g o f al l AF T samples i n souther n Sweden . Thi s heatin g wa s not se t a s a constraint whe n modelling, bu t wa s necessary t o obtai n a fi t betwee n observe d an d calculated AF T data . Additionally , heating t o at least 9 0 °C was required . Documented Cretaceou s remnants , lyin g directly o n basemen t o n th e wes t an d S E coas t of souther n Sweden , i n combination wit h a subMesozoic relief , revea l tha t thes e area s experi enced surface-leve l temperature s a t c . 14 0 Ma. The Cretaceou s palaeosurfac e temperatur e wa s estimated t o hav e bee n c . 3 0 °C. Surfac e leve l conditions continue d unti l Lat e Cretaceou s sedimentation starte d a t c . 8 5 Ma. I n severa l cases, the final cooling event was assigned an age of a t leas t 1 5 Ma, becaus e thi s i s th e expecte d minimum tim e fo r th e Tertiar y etc h surface s t o have develope d (Lidmar-Bergstro m 1991) . Samples fro m area s wher e th e SC P is preserve d probably experience d fina l exhumatio n i n

Neogene time . Th e recen t (OMa ) surface-leve l temperature was set to c. 2 0 °C. Hig h surface level temperature s (3 0 °C an d 2 0 °C, respect ively) were deliberately set to attai n a minimum value o f th e modelle d Cretaceous-Paleogen e reheating. The AFT dataset Modelled therma l historie s fo r 2 1 apatit e samples wer e achieve d usin g AF T dat a fro m Cederbom e t a l (2000 ) an d Cederbo m (2001) . The samplin g localitie s ar e presente d i n Fig . 1. All samples were collected fro m th e outcropping Precambrian basement (granitoids and gneisses), and the y al l experience d tota l annealin g during Late Palaeozoi c and/o r Earl y Mesozoi c tim e (Cederbom e t a l 2000 ; Cederbo m 2001) . Th e AFT ages fo r the 2 1 apatite samples used in this study al l passe d th e ^ tes t (Cederbo m e t a l 2000; Cederbom 2001), indicating that the grains in each sampl e belong t o a single population. There ar e mino r difference s i n AF T result s between th e sample s tha t ma y deriv e fro m kinetic variations among the apatites. The apatite composition ha s no t bee n examined , s o compo sitional variation s canno t b e excluded . Never theless, al l sample s passe d th e ^ tes t an d no major differenc e i n etch pit widt h wa s observe d when analysin g the grains. There i s no correlation between elevation and AFT age in southern Sweden. Instead, three areas characterized b y differen t trend s i n th e AF T results ca n b e discerne d (Cederbo m 2001) . Th e southeastern par t o f souther n Swede n (are a C ) is characterize d b y AF T age s younge r tha n 200 Ma. A SW-NE-trendin g bel t (are a B ) includes sample s wit h age s betwee n 23 1 an d 282 Ma, wherea s are a A i n th e northwester n part o f souther n Swede n represent s sample s with age s rangin g fro m 14 9 t o 31 5 Ma (se e Fig. 2 ; Cederbo m 2001) . Non e o f th e sample s have mixe d trac k lengt h distributions . T o summarize, area s A , B an d C ar e define d based o n difference s i n th e AF T result s that ar e too larg e t o b e cause d solel y b y compositiona l variations. Modelling results The modelling result s for the 2 1 apatite samples are shown in Fig. 3 and Table 1 . The Phanerozoi c thermal histor y wa s modelle d fo r al l samples. However, only the discernible part of the thermal history i s presente d i n Fig . 3. Whe n modellin g samples P9904 and 9905, for example, heating to at least 90 °C for the 200-300 Ma time interval is required for both samples (see Fig. 3). However,

THERMOTECTONIC DEVELOPMENT OF SOUTHERN SWEDEN heating abov e 9 0 °C is no t accepte d fo r sampl e P9904 during this time interval, wherea s heatin g above 9 0 °C make s n o differenc e i n th e calculated AF T results o f sample 9905 . There ar e mino r difference s betwee n th e modelled therma l historie s tha t ma y deriv e from kineti c variation s amon g th e apatites . However, severa l significan t observation s ca n be made . First, th e modellin g result s ca n b e separate d into thre e group s supportin g th e existenc e o f areas A, B and C in Fig. 1 (see Cederbom 2001) . The samples from are a C have relatively uniform cooling histories . The y al l cooled belo w c . 90 °C during Earl y Jurassi c tim e an d the y ar e characterized b y coolin g rate s o f c . 15° C pe r 10 Ma. Are a B is characterized b y sample s wit h older AF T age s tha n th e sample s i n are a C. Likewise , modelle d therma l historie s for the samples i n are a B indicate muc h earlie r coolin g and muc h lowe r coolin g rates . Coolin g belo w 90 °C had occurre d b y Earl y Triassic time i n the SW, and a s earl y a s Late Carboniferou s time i n the NE . Wit h tw o exception s th e coolin g rate s decrease fro m c . 5° C pe r 1 0 Ma i n th e S W t o c. 3 °C per 1 0 Ma in the NE. Samples S962 6 and 9808 a t th e margin s o f are a B sho w slightl y higher cooling rates, 7 °C per 1 0 Ma and 8 °C per 10 Ma, respectively . Are a A include s sample s showing a remarkabl e heterogeneity . Th e poin t of tim e whe n th e sample s coole d belo w 9 0 °C ranges fro m Lat e Carboniferous t o Late Jurassi c time. Additionally , th e coolin g rate s var y between c . 4 0 °C pe r 1 0 Ma an d c . 3° C pe r 10 Ma. Second, i t i s possibl e tha t al l sample s wer e reheated durin g Late Cretaceou s time . However, such reheatin g ca n b e establishe d onl y fo r samples collecte d fro m th e sub-Mesozoi c palaeosurface, whic h ar e know n t o hav e experienced surfac e condition s befor e and/o r during Cretaceou s time . Th e modellin g result s for sampl e 990 2 an d 980 5 indicat e c . 35 ° reheating i n S E Sweden , wherea s heatin g of th e order of c. 20 °C is indicated for the samples fro m the west coast (i.e . sample B14 , 9901 and 9808 ) (Figs 1 an d 3) . Modellin g o f th e remainin g samples show s tha t Cretaceou s reheatin g o f southern Swede n a s a whol e i s possible , bu t cannot b e take n a s certain . Alternativ e therma l histories are presented fo r a few samples (P9909 , S9624, P9906 , 9903 , SA9629 ) i n Fig . 3 an d Table 1 . Third, Oligocene-Miocen e fina l coolin g i s possible i n al l thre e area s accordin g t o th e modelling. Fo r occasional samples , Lat e Eocen e final cooling is also allowed, althoug h it is never required.

177

Discussion The forwar d modellin g o f AF T result s fro m southern Swede n confirm s th e conclusion s o f Cederbom (2001) , i.e . tha t souther n Swede n experienced differentiate d coolin g afte r th e Caledonian orogen y an d th e relate d forelan d basin formation. Furthermore, a pattern illustrating loca l variation s in both onse t o f cooling an d cooling rat e appear s an d is discussed below . The Palaeozoi c an d Mesozoi c coolin g wa s probably a resul t o f erosio n o f cove r rock s (unroofing), bu t change s i n th e geotherma l gradient ma y hav e als o ha d a n influence . I t i s likely that the thermal conductivity and heat flow of th e crystallin e basemen t an d sedimentar y cover have differed, bot h regionally an d through time during the Palaeozoic an d Mesozoic eras . It can be noted that Triassic and Jurassic sediments deposited i n shallow-wate r an d fresh-wate r environments ar e recorde d withi n th e Fennos candian Borde r Zon e i n southernmos t Swede n (Norling & Bergstro m 1987 ; Guy-Ohlso n & Norling 1988) , supportin g th e ide a o f a n Earl y Mesozoic cove r i n souther n Sweden . Th e observed Cretaceou s reheatin g wa s probabl y caused b y sedimen t deposition , an d th e fina l cooling even t i s explaine d b y Cenozoi c fina l exhumation. The onse t of unroofin g is not necessaril y the same a s th e poin t i n tim e whe n th e sample s cooled belo w 9 0 °C. I f a highe r temperatur e regime originall y prevailed , exhumatio n ma y have starte d muc h earlie r tha n whe n fissio n tracks started to accumulate. According to earlier published conodon t alteratio n an d organi c maturation studies of Lower Palaeozoic remnants (Bergstrom 1980 ; Buchard t e t al 1997) , however, a temperatur e regim e abov e 9 0 °C i s not likely fo r the area betwee n Lak e Vaner n and Lake Vattern . Triassic and Jurassic exhumation After th e Caledonia n orogen y an d the formation of a forelan d basin , unroofin g started . Th e basement, wit h it s heav y pil e o f sedimentar y strata, ma y no t hav e acte d a s a singl e entit y i n southern Sweden. The first recorded unroofin g i s indicated i n areas A and B. Within area A , Lat e Carboniferous, Earl y Jurassi c an d Lat e Jurassi c first recorded exhumatio n timing s ar e obtained . In area B the recorded exhumatio n started in Late Carboniferous t o Earl y Triassi c time , wit h a younging tren d toward s th e SW . Early Jurassi c first recorded exhumatio n is obtained i n are a C . A compilation of modelled coolin g rates for three points in time is presented in Fig. 4 together with

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C. CEDERBO M

Fig. 4 . A speculativ e reconstruction o f the Mesozoic an d Cenozoic tectoni c development i n southern Sweden. A heterogeneity i n Triassic-Jurassic exhumatio n betwee n th e thre e mai n tectoni c area s o f souther n Sweden , and between individua l samples i n the northwestern par t o f southern Swede n ca n be discerned. Oligocene-Miocen e final exhumation was mos t pronounce d aroun d th e souther n tip o f Lak e Vatter n an d o n th e S E coast.

THERMOTECTONIC DEVELOPMENT OF SOUTHERN SWEDEN the probabl e exten t o f Uppe r Siluria n t o Uppe r Jurassic sediment s (i.e . forelan d basi n relate d sediments). A s mentione d above , mino r differ ences betwee n th e modelle d therma l historie s may b e du e t o compositiona l variation s amon g the apatites . Nevertheless , Fig . 4 illustrates tha t there is a heterogeneity i n Mesozoic exhumation among the thre e area s of souther n Sweden . In the beginning of Triassic time (Fig. 4a), it is clear tha t mos t o f area s A an d B wer e stil l covered by thick, foreland basin related deposits . Slow to moderate exhumatio n rates are recorde d in area s A an d B , wit h a n increasin g tren d towards th e SW . A t th e sam e time , th e Skagerrak-Kattegat Platfor m (SKP ) (Fig . 1 ) basement wa s exposed; seismi c offshor e record s (e.g. Vejbae k 1997 ) revea l that the basement wa s exposed befor e Triassi c sedimentatio n started . The AF T dat a fo r sampl e B1 4 als o suppor t surface temperatur e conditions . I n contrast , th e AFT results for sample 9808 reveal temperatures of 9 0 °C, indicatin g thic k coverin g o f th e basement i n th e S W i n Earl y Triassi c time . T o solve thi s contradiction , additiona l studie s o f both th e sedimentar y recor d an d tectonics clos e to shor e o n th e SK P ar e needed , togethe r wit h denser AF T sampling. It is probable, but not established, that area C and centra l Swede n wer e covere d b y sediment s in Earl y Triassi c time . I n Earl y Jurassi c tim e (Fig. 4b) , however, i t is evident that are a C was covered b y thic k piles o f sediment , a s tempera tures > 90 °C are recorded b y the AFT data and exhumation is recorded i n the area. Cooling as a result of exhumation is also detected i n all thre e areas for Late Jurassi c time (Fig . 4c) . It i s interestin g t o not e tha t sample s fro m th e NE-SW-trending area B are consistent with slow to moderat e coolin g rate s throughou t Triassi c and Jurassi c time . Meanwhile , area s B an d C behaved i n a differen t manner , an d larg e variations ar e recorde d withi n are a A . I n are a A, bloc k movement s ma y hav e cause d uneve n erosion an d redeposition o n downfaulted blocks. It i s possibl e tha t are a C acte d a s a temporar y store for reworked sediment s from th e NW . Th e unroofing seem s t o hav e accelerate d i n earlies t Jurassic time , whe n are a C experience d rapi d exhumation, followed by block uplift s i n area A. Before o r durin g Earl y Cretaceou s time , a NW-SE-trending axi s o f elevation developed , resulting i n exposure o f the west an d S E coasts . The Palaeozoi c cove r ha d bee n remove d i n th e Bastad an d Kristiansta d area s o n the wes t coas t and th e S E coas t whe n Earl y Cretaceou s sedimentation starte d (Lidmar-Bergstro m 1982). I n Fig . 4d , area s wit h expose d basemen t are illustrate d together wit h th e probabl e extent

179

of Phanerozoic , pre-Cretaceou s sediment s b y that time ; erosio n o f the Uppe r Siluria n t o Lat e Jurassic sediments may have caused exposure of lower Palaeozoi c sediments . However , mos t parts o f th e sout h Swedis h basemen t surfac e were stil l protecte d fro m weatherin g (Lidmar Bergstrom 1991) . The Cretaceous and Paleogene cover Late Cretaceous an d Paleogene deposit s covere d the souther n par t o f th e Fennoscandia n Shield , but th e exten t of these sediment s t o the nort h is not established. The modelling results reveal that the west and SE coasts of southern Swede n wer e buried, and they illustrate that the central part of southern Swede n ma y als o hav e bee n covere d (Fig. 4e) . A n estimat e o f th e thicknes s o f thes e deposits o n the west an d SE coast ca n be made . The modelled therma l histories indicat e that the Cretaceous-Paleogene cover was thicker o n the SE coast (sample s 9902 , 9805) tha n on the west coast (9808 , 9901) , a s c . 3 5 °C an d c . 2 0 °C temperature increase s ar e recorde d fo r th e S E and west coasts, respectively. A n estimate of the thickness o f thes e sediment s involve s a n unconstrained estimat e o f th e geotherma l gra dient. I f the geotherma l gradien t durin g Cretac eous tim e wa s 3 0 °C km" 1, a sedimentar y thickness o f c. 650m in the west an d > 1000 m in the S E is indicated. In comparison, soni c data and basi n model s fo r dat a fro m Danis h well s support th e idea tha t the Cretaceous-Paleogene cover o n th e SK P exceede d 1000 m thicknes s before Neogen e uplif t an d erosio n (Japse n & Bidstrup 1999) . Cenozoic exhumation Where th e SC P is well preserved i t cannot have been re-exposed until Late Tertiary time , but it is uncertain whe n th e basemen t wa s re-expose d within th e highes t part s t o th e sout h an d S E of Lake Vattern. According to the modelled thermal histories, th e basement , wit h o r withou t a n additional Cretaceou s cove r o n to p o f th e remnant Palaeozoi c cover , di d no t reappea r a t the surfac e unti l a t leas t Oligocen e o r Miocen e time. Areas with a well-preserved SCP were still covered b y sediment s whe n th e SS P wa s developed (Fig . 4f) . I n Fig . 4f , th e relativ e amount o f fina l exhumatio n indicate d b y th e modelling result s i s als o shown . Thi s i s a ver y speculative picture , a s severa l alternativ e mod elled historie s matc h th e AF T data . Larg e exhumation i s indicate d fo r th e souther n ti p o f Lake Vattern , wher e th e highes t altitude s ar e found today , and fo r the coast i n the SE .

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Cenozoic uplif t an d domin g o f souther n Sweden ha s bee n supporte d b y Lidmar-Berg strom (1996 , 1999 ) an d Japse n & Bidstru p (1999). Cenozoi c uplif t o f souther n Swede n ha s also bee n discusse d i n a broade r perspective , together wit h Cenozoic uplif t an d doming o f the southern Scande s i n souther n Norwa y (e.g . Rohrman e t al 1995 ; Rii s 1996 ; Lidmar Bergstrom 1999) . Th e modelle d AF T dat a indicate tha t large-scal e Cenozoi c exhumatio n of souther n Swede n ha s occurred , whic h i s consistent wit h Cenozoi c uplift . However , th e speculative exhumatio n patter n illustrate d i n Fig. 4 f doe s no t suppor t symmetrica l domin g of the basement i n southern Sweden, as not only the central part o f souther n Sweden , but als o th e SE coast, experience d a larg e amoun t o f lat e exhumation. Large-scale Phanerozoic tectonism The mai n difference s i n th e AF T dat a an d th e modelled therma l historie s fo r sample s fro m areas A , B an d C ar e to o larg e t o b e explaine d solely by compositional variation s in the apatites. Neither i s long-term partia l annealin g a probabl e explanation fo r the differences , accordin g t o the observed trac k lengt h distributions . I f th e base ment in souther n Swede n coole d as an entit y during Lat e Palaeozoi c t o Lat e Jurassi c time , large-scale Cretaceou s an d Cenozoi c tectoni c movements mus t hav e occurred . Thi s seem s unlikely considering th e present-day topography . Alternatively, souther n Swede n experience d differentiated exhumatio n accompanie d b y tectonism durin g Lat e Palaeozoi c t o Lat e Jurassi c time. I n addition , larg e difference s i n th e AF T results an d th e modelle d therma l historie s recorded fo r sample s withi n are a A indicat e individual block movements i n the northwester n part o f souther n Sweden . The proposed extent of the three main tectoni c areas A , B an d C i s base d o n a coars e gri d o f apatite sample s (se e Fig . 1) , an d additiona l sampling is required before the existence of these areas ca n b e verified . A comprehensiv e investi gation o f th e occurrenc e an d exten t o f Phaner ozoic fault s i n souther n Sweden ha s no t bee n published. However , the study area where severa l Phanerozoic fault s wer e observe d b y Ahli n (1987) i s situate d withi n are a A . I n addition , two fracture zone s ar e mapped alon g th e easter n side o f Lake Vattern (Persso n & Wikman 1986) . One o f the m i s situate d betwee n th e sampl e points S9623 and S9624. Block movements hav e been observe d alon g th e easter n sid e o f the lake (Persson & Wikman , 1986) , bu t individua l blocks hav e not been identified.

It is also interestin g t o compare th e suggeste d extent of the three areas, which is based solel y on AFT data , wit h th e genera l trend s i n th e lineament pattern presented by Tiren & Beckholmen (1992) (Fig . 2). The shar p lines , interpreted as the younges t features (Fig. 2c) ar e foun d i n a restricted are a south of Lake Vanern, corresponding mor e o r les s t o are a A . Furthermore , th e NW-SE-trending bel t o f extensiv e lineament s that divide s souther n Swede n int o tw o halve s (Fig. 2a) may have a correspondence i n the areal extent o f are a B . The existenc e o f thre e mai n tectoni c unit s corresponding to areas A, B and C and individual block movement s withi n a t leas t northwester n Sweden i s supporte d b y th e studie s o f faults , fracture zone s an d linea r structures published to date. However , large-scal e Phanerozoi c tectoni c block movement s o f the orde r o f > 100 m hav e previously been suggeste d only for Lake Vanern (Ahlin 1987) .

Conclusions On their own, published AFT data have revealed a genera l pictur e o f the thermotectonic develop ment i n souther n Swede n fo r Palaeozoi c an d Early Mesozoi c time . However , b y forwar d modelling th e AF T dat a i n combinatio n wit h studies o f palaeosurface s an d relief , furthe r information o n exhumatio n rates , sedimentar y thicknesses an d th e Mesozoi c t o Cenozoi c thermotectonic histor y has been derived . Southern Swede n ca n b e divide d int o thre e main tectonic areas, which were characterized by different onse t o f recorde d unroofin g an d different exhumatio n rate s durin g Lat e Palaeo zoic t o Lat e Jurassi c time . Individua l bloc k movements i n on e o f thes e areas , i.e . th e northwestern par t o f souther n Sweden , ar e suggested fo r this time interval. There i s a contradictio n betwee n th e thermo tectonic developmen t i n S W Swede n an d o n the SKP furthe r west . Thic k deposit s covere d th e basement i n SW Sweden during earliest Triassic time, wherea s th e basemen t o n th e SK P wa s exposed. T o solv e thi s contradiction , additiona l studies of offshore sedimentar y records and nearshore tectonics , an d dense r AF T samplin g ar e needed. Southern Swede n experience d reburia l during Late Cretaceou s an d Paleogen e time . A temperature ris e o f c . 3 5 °C an d c . 2 0 °C ha s bee n detected fo r th e wes t coas t an d th e S E coast , respectively. Th e temperatur e differenc e indi cates tha t th e sedimentar y cove r wa s thicke r o n the west coast tha n o n the S E coast.

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fission track s i n fluorapatite . Geochimica e t Cosmochimica Acta, 55, 1449-1465. ELVHAGE, C . & LIDMAR-BERGSTROM , K . 1987 . Some working hypothese s o n th e geomorpholog y o f Sweden i n th e ligh t o f a ne w relie f map . Geografiska Annaler, 69A , 343-358. GLEADOW, A.J.W . & DUDDY , I.R . 1981 . A natura l long-term trac k annealin g experimen t fo r apatite . Nuclear Tracks, 5 , 169-174. GLEADOW, A.J.W. , DUDDY , I.R. , GREEN , P.F . & HEGARTY, K.A . 19860 . Fissio n trac k length s i n the apatite annealin g zone and the interpretation of mixed ages . Earth an d Planetary Science Letters, 78, 245-254. GLEADOW, A.J.W. , DUDDY , I.R. , GREEN , P.F . & LOVERING, J.F . 19866 . Confine d fissio n trac k lengths i n apatite : a diagnosti c too l fo r therma l history analysis . Contributions t o Mineralogy an d Petrology, 94, 405-415. GUY-OHLSON, D . & NORLING , E . 1988 . Upper Jurassic Litho- and Bio stratigraphy ofNW Scania, Sweden. Sverige s Undersokning Serie Ca, 72. JAPSEN, P. & BIDSTRUP , T. 1999. Quantification o f late Cenozoic erosio n i n Denmark based o n soni c dat a and basi n modelling . Bulletin o f th e Geological K. Lidmar-Bergstro m i s gratefull y acknowledge d fo r Society o f Denmark, 46 , 79-99. fruitful criticism . Th e manuscrip t wa s improve d b y KETCHAM, R.A. , DONELICK , R.A . & CARLSON , W.D. comments fro m K . Gallagher, A.G . Dore, H . Sinclai r 1999. Variabilit y o f apatit e fission-trac k annealin g and a n anonymou s reviewer . Thi s projec t ha s bee n kinetics: III . Extrapolatio n t o geologica l tim e financially supporte d b y th e Nuclea r Wast e an d scales. American Mineralogist, 84 , 1235-1255. Management Co . (SKB ) an d th e Roya l Swedis h KETCHAM, R.A. , DONELICK , R.A. & DONELICK , M.B. Academy o f Sciences . 2000. AFTSolve : a progra m fo r multi-kineti c modeling o f apatit e fission-trac k data . Geological Material Research, 2 (1), 1-32. References LARSON, S.A. , TULLBORG , E.-L. , CEDERBOM , C.E . & STIBERG, J.-P . 1999. Sveconorwegian an d Caledo AHLIN, S . 1987 . Phanerozoic fault s i n th e Vastergot nian forelan d basins i n th e Balti c Shiel d reveale d land basin area , S W Sweden. Geologiska Foreninby fission-track thermochronology . Terra Nova, 11, gen i Stockholms Forhandlingar, 109 , 221-227. 210-215. BERGSTROM, S . 1980 . Conodonts a s paleotemperatur e tools i n Ordovicia n rock s o f th e Caledonide s an d LASLETT, G.M . & GALBRAITH , R.F . 1996. Statistical modelling o f thermal annealing o f fission tracks i n adjacent area s i n Scandinavi a an d th e Britis h apatite. Geochimica e t Cosmochimica Acta, 60 , Isles. Geologiska Foreningen i Stockholms 5117-5131. Forhandlingar, 102 , 377-392 . BUCHARDT, B. , NIELSEN , A.T . & SCHOVSBO , N.H. LASLETT, G.M. , GREEN , P.P. , DUDDY , I.R . & GLEADOW, A.J.W . 1987 . Thermal annealin g o f 1997. Alu n Skifere n i Skandinavien . Geologisk fission tracks i n apatit e 2 . A quantitativ e analysis. Tidsskrift, 3 , 1-30. Chemical Geology, 65, 1 — 13. CARLSON, W.D . 1990. Mechanisms an d kinetics o f LIDMAR-BERGSTROM, K . 1982 . P re-Quaternary apatite fission-trac k annealing . American Geomorphological Evolution in Southern Mineralogist, 75 , 1120-1139. Fennoscandia. Sverige s Undersoknin g Seri e C , CARLSON, W.D. , DONELICK, R.A . & KETCHAM , R.A. 785. 1999. Variabilit y of apatit e fission-trac k annealin g LIDMAR-BERGSTROM, K . 1988 . Denudation surface s kinetics: I. Experimental results . American Minerof a shiel d are a i n sout h Sweden . Geografiska alogist, 84 , 1213-1223 . Annaler, 70 A (4), 337-350. CEDERBOM, C.E . 2001. Phanerozoic, pre-Cretaceou s thermotectonic events in southern Sweden reveale d LIDMAR-BERGSTROM, K . 1991 . Phanerozoic tectonic s in southern Sweden. Zeitschiftfiir Geomorphologie, by fissio n trac k thermochronology . Earth an d Neue Folge Supplement, 82 , 1-16. Planetary Science Letters, 188 , 199-209 . CEDERBOM, C.E., LARSON , S.-A. , TULLBORG, E.-L. & LIDMAR-BERGSTROM, K . 1994 . Morphology o f th e bedrock surface . In : FREDEN , C . (ed. ) Geology. STIBERG, J.- P 2000. Fissio n trac k thermochronol National Atlas o f Sweden. SN A Publishing , ogy applie d t o Phanerozoi c thermotectoni c event s Stockholm, 44-54. in centra l an d souther n Sweden . Tectonophysics, LIDMAR-BERGSTROM, K . 1995 . Relief an d saprolite s 316, 153-167 . through time on the Baltic Shield. Geomorphology, CROWLEY, K.D. , CAMERON, M . & SCHAEFER , R.L . 1991. Experimental studie s of annealing of etche d 12,45-61.

Final exhumatio n o f souther n Swede n durin g Cenozoic tim e i s supporte d b y th e modellin g results. This is consistent with previous studies of Cenozoic uplif t o f souther n Sweden (Rii s 1996 ; Lidmar-Bergstrom 1999 ; Japse n & Bidstru p 1999). However, the pattern of final exhumation obtained fro m th e modellin g o f AF T dat a doe s not suppor t a symmetrica l domin g o f th e basement i n southern Sweden . Finally, large-scal e Phanerozoi c tectoni c movements i n souther n Swede n canno t b e rejected base d on faul t an d fracture zon e studies published t o date . Additionally , ther e ar e similarities betwee n majo r feature s observe d i n the lineament pattern of southern Sweden and the suggested areal extent of the three main tectonic areas. Nevertheless , a mor e detaile d an d bette r constrained patter n o f th e tectoni c development in souther n Swede n demand s dense r apatit e sampling an d a regional study o f the occurrence of Phanerozoic faults .

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LIDMAR-BERGSTROM, K. 1996 . Long ter m morpho tectonic evolutio n in Sweden. Geomorphology, 16 , 33-59. LIDMAR-BERGSTROM, K . 1999 . Uplif t historie s revealed b y landform s o f th e Scandinavia n domes. In : SMITH , B.J. , WHALLEY , W.B . & WARKE, P.A . (eds) Uplift, Erosion an d Stability: Perspectives on Long-term Landscape Development. Geologica l Society , London , Specia l Publi cations, 162 , 85-91. LIDMAR-BERGSTROM, K. , OLSSON , S . & OLVMO , M . 1997. Palaeosurface s an d associate d saprolite s i n southern Sweden . In : WIDDOWSON , M . (ed. ) Palaeosurfaces: Recognition, Reconstruction and Palaeoenvironmental Interpretation. Geologica l Society, London, Special Publications, 120,95-124. LIND, G. 1972 . The gravity and geology o f the Vattern area, souther n Sweden . Geologiska Foreningen i Stockholms Forhandlingar, 94, 245-257. LANTMATERIVERKET. 1986 . Sveriges Relief. Lantmateriverket, Gavle. NAESER, C.W. 1981. The fading of fission tracks in the geologic environment—data from deep drill holes . Nuclear Tracks, 5, 248-250. NORLING, E . & BERGSTROM , J . 1987 . Mesozoic an d Cenozoic tectoni c evolutio n o f Scania , souther n Sweden. Tectonophysics, 137 , 7-19 .

PERSSON, L . & WIKMAN , H . 1986 . Provisoriska Oversiktliga Berggrundskartan Jonkoping, Map Sheet and Description. Swedis h Geological Survey Af39. PERSSON, L. , BRUUN , A . & VIDAL , G . 1985 . Berggrundskartan Hjo SO, Map Sheet and Description. Swedis h Geologica l Surve y A f 134. Rus, F . 1996 . Quantificatio n o f Cenozoi c vertica l movements o f Scandinavi a b y correlatio n o f morphological surface s with offshore data. Global and Planetary Change, 12 , 331-357. ROHRMAN, M. , VA N DER BEEK , PA. , ANDRIESSEN ,

P.A.M. & CLOETHING , S . 1995 . Meso-Cenozoi c morphotectonic evolutio n o f souther n Norway : Neogene doma l uplif t inferre d fro m fissio n trac k thermochronology. Tectonics, 14, 704-718. TIREN, S.A . & BECKHOLMEN , M . 1992 . Rock bloc k map analysi s o f souther n Sweden . Geologiska Foreningen i Stockholms Forhandlingar, 114 , 253-269. VEJB^EK, O.V . 1997. Dybe strukturer i sedimentcere bassiner. Geologisk Tidsskrift , 4 . VIDAL, G. 1984 . Lake Vattern. Geologiska Foreningen i Stockholms Forhandlingar, 106, 397 .

Neogene uplif t an d erosion of southern Scandinavia induced by the rise of the South Swedish Dome PETER JAPSEN 1, TORBEN BIDSTRUP 1 & KARNA LIDMAR-BERGSTROM 2 1 Geological Survey o f Denmark an d Greenland (GEUS), 0ster Voldgade 10 , DK-135 K0benhavn K, Denmark (e-mail: pj@ geus.dk) 2 Department of Physical Geography and Quaternary Geology, Stockholm University, SE-10691 Stockholm, Sweden Abstract: Basi n modelling an d compaction studie s based o n sonic data from the Mesozoi c succession i n 68 Danish well s wer e use d t o estimate th e amoun t of section missin g du e to late Cenozoi c erosion . Th e missin g sectio n increase s graduall y toward s th e coast s o f Norway and Sweden from zero in the North Sea to c. 500 m in most of the Danish Basin, but over a narro w zon e i t reache s c . 1000 m o n th e Skagerrak-Kattega t Platfor m i n northernmost Denmark . The increasing amount of erosion matche s the increase in the hiatus at th e bas e o f th e Quaternary , wher e Neogen e an d olde r strata ar e truncated , an d th e Mesozoic successio n i s thus found to have been mor e deepl y burie d by c. 500 PaleoceneMiocene sediment s i n large part s o f the area . Thes e observation s sugges t tha t th e onse t of erosion occurre d durin g th e Neogene , an d tha t th e Skagerrak-Kattega t Platfor m wa s affected b y tectonic movement s prio r to glacial erosion . I n southern Swede n just east of the Kattegat, th e expose d basemen t o f th e Sout h Swedis h Dom e attain s altitude s o f almos t 400 m. Th e formatio n o f th e Dom e starte d i n th e Lat e Palaeozoic , bu t geomorphologica l investigations hav e le d t o th e conclusio n tha t a ris e o f th e Dom e occurre d durin g th e Cenozoic. W e fin d tha t th e patter n o f lat e Cenozoi c erosio n i n Denmar k agree s wit h a Neogene uplif t o f the Sout h Swedis h Dom e an d o f the Souther n Scande s i n Norway. Thi s suggestion i s consisten t wit h majo r shift s i n sedimen t transpor t direction s durin g th e lat e Cenozoic observe d i n the eastern Nort h Sea, an d with formation of a new erosion surfac e as well as re-exposure o f sub-Cambrian an d sub-Cretaceous surface s in southern Sweden . Th e Neogene uplif t an d erosio n o f souther n Scandinavi a appear s t o hav e bee n initiate d i n tw o phases, a n early phase of ?Miocene ag e and a better-constrained late r phase that began in the Pliocene. Neogen e uplif t o f the Sout h Swedis h Dom e wit h adjoining areas i n Denmark fit s into a pattern o f late Cenozoi c vertica l movement s aroun d th e North Atlantic .

Recognition of the Neogene uplift an d erosion of 2k m i n severa l wells (Jensen & Schmid t 1992, Denmark an d Swede n is difficul t becaus e of it s 1993 ; Japsen 1993 , 1998 ; Michelsen & Nielsen regional extent , an d becaus e th e effect s ar e 1993) . This paper reports the results of a study of overprinted by the erosion during the subsequent th e lat e Cenozoi c erosio n o f cove r rock s i n Quaternary glaciation s (Fig . 1) . Consequently , Danis h well s locate d outsid e th e lat e Cenozoi c only few relevant observations were presented in depocentr e in the centra l North Sea an d outside the literatur e befor e th e 1990s . Studie s o f th e th e Bornhol m are a i n th e souther n Baltic Se a Miocene Vejl e Fjor d Formatio n le d Larse n & (Fig . 2) (Japsen & Bidstrup 1999) . Estimate s of Dinesen (1959 ) t o conclud e tha t considerabl e erosio n have been based on basin modelling and parts o f Fennoscandia , includin g no t onl y soni c dat a fro m severa l stratigraphi c units , basement bu t als o sedimentar y formations , resultin g i n maximu m value s o f c . 100 0 m, were erode d i n Neogen e time . Spjeldnae s whic h are considerably lower than those reported (1975) foun d tha t uplif t o f th e Fennoscandia n i n the above earlier studies . Shield in late Oligocene-Miocene time resulted Al l studie s find erosion t o increas e fro m th e in a significan t chang e i n th e sedimentar y easter n North Sea towards the coasts of Norway environment an d i n th e drif t o f th e coastlin e an d Sweden . This increasing amount of erosio n towards the SW . matche s the increase in the hiatu s a t the base of The effec t o f lat e Cenozoi c erosio n i n th e Quaternar y succession, where Neogene and Denmark ha s bee n quantifie d b y a numbe r of olde r strat a ar e truncated . Thes e observations workers, wh o estimate d erosion to be fro m 1 to sugges t that the onset of erosion occurred during From: DORE , A.G., CARTWRIGHT, J.A. , STOKER, M.S., TURNER, J. R & WHITE , N. 2002 . Exhumation of th e North Atlantic Margin: Timing, Mechanisms and Implications for Petroleum Exploration. Geologica l Society, London, Special Publications, 196, 183-207 . 0305-8719/02/$15.00 © The Geological Society of London 2002.

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Fig. 1. Pre-Quatemary geology of sourthern Scandinavia and landforms of the bedrock across the South swedish

Dome. Th e hiatu s at the base o f the Quaternary (c . 2.4 Ma, Zagwijn 1989 ) and the change i n sediment transport direction fro m Oligocene t o Pliocene tim e agrees wit h Neogene uplif t o f the South Swedish Dome. Compar e th e increasing ag e o f th e Quaternar y subcro p toward s th e expose d basemen t i n Norwa y an d Swede n wit h th e corresponding deepening o f the estimated erosio n (Fig. 13) . The mountains of the Southern Scandes constitut e the main part of southern Norway. Modified after Freden (1994), Vejbae k & Britze (1994), Lidmar-Bergstrom (1996), Japsen (1998 ) an d Clausen e t al. (1999).

Neogene time , an d indicat e tha t uplif t an d erosion hav e affecte d no t onl y Norwa y an d Denmark, bu t als o souther n Swede n (Japse n 1993). The geological recor d o f south Scandina via i s thus of grea t importanc e t o understandin g the Neogen e developmen t o f th e whol e o f Scandinavia an d th e Atlanti c margin s a s such . In thi s are a i t i s easie r tha n i n mos t place s t o

compare dat a fro m th e Cenozoi c sedimentar y cover (partl y onshore ) wit h observation s fro m exposed basemen t wher e pre-glacia l landform s are well preserved. Distance s are small, and data as well a s geoscientific studie s are abundant. The present study concludes that the Mesozoic succession has been c. 500 m more deeply buried than toda y i n mos t o f th e area , wher e th e

Fig. 2 . Locatio n maps , (a ) Plac e name s an d profile s ABC D (Fig . 15 ) and EF (Fig . 14) . (b) Locatio n o f th e 6 8 Danish and three Norwegian wells used in the study, (c) Structural elements. Basement highs indicated with grey. See als o wel l locatio n ma p o f Nielsen & Japsen (1991) .

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succession generall y i s overlai n b y Paleogen e strata, an d tha t a sectio n o f mainl y Paleocene Miocene ag e must have been remove d fro m thi s area durin g lat e Cenozoi c time . W e combine th e model o f late Cenozoic erosio n i n Denmark wit h observations fro m th e expose d basemen t i n Sweden, jus t eas t o f th e Kattegat . Here , geomorphological investigation s hav e le d t o th e conclusion that a rise of the South Swedish Dom e (the upland s o f Smalan d wit h altitude s u p t o 400 m) occurred durin g Cenozoi c tim e (Lidmar Bergstrom 1995 , 1996) . W e therefor e fin d tha t Neogene uplif t o f souther n Scandinavia , centre d around th e Sout h Swedis h Dome , coul d explai n both th e geophysica l an d th e geomorphologica l observations.

Erosion estimated fro m soni c dat a Derivation of normal velocity-depth trends A norma l velocity-dept h tren d (velocit y base line), VN(Z) , describe s i n a functiona l for m ho w the soni c velocit y o f a relativel y homogeneou s sedimentary formatio n saturate d wit h brin e increases wit h dept h whe n porosit y i s reduce d during norma l compaction . Th e pressur e o f th e formation i s hydrostatic durin g norma l compac tion, an d th e formatio n i s a t maximu m buria l depth; i.e. the thickness of the overburden has not been reduce d b y erosio n (e.g . Bula t & Stoke r

1987; Hilli s 1995 , Japse n 1998 , 2000) . Simpl e boundary condition s fo r suc h trend s ar e that th e normal velocit y at the surface equals the velocity of th e sedimen t whe n i t wa s firs t deposited , an d that velocit y a t infinit e dept h approache s th e matrix velocity of the rock whereas the velocitydepth gradien t approache s zero . The derivatio n o f a norma l velocity-dept h trend involve s thre e step s o f generalization : (1 ) identification o f a relativel y homogeneou s lithological unit ; (2 ) selectio n o f dat a point s representing norma l compaction ; (3 ) assignment of a functional expressio n t o the velocity-dept h trend. Velocity baselines ma y thu s be difficul t t o establish, and different trend s have been assigned to identica l unit s b y differen t worker s (compar e Bulat & Stoke r (1987 ) an d Japsen (2000)) . (1) 'Relativel y homogeneous ' refer s t o thos e properties tha t are important for the macroscopi c acoustic behaviou r o f th e unit . However , w e d o not always know if data from a well represent the typical developmen t o f th e uni t o r whic h mineralogical difference s may b e o f importanc e for its acoustic behaviour, e.g. the clay content in sandstones o r chalks. (2) 'Norma l compaction ' ma y b e a difficul t condition t o prove , becaus e w e d o no t alway s know if formation pressure is hydrostatic or if the formation ha s been burie d deepe r befor e erosio n (Fig. 3) . Lat e Cenozoi c erosio n alon g th e margins o f th e Nort h Se a Basi n an d over -

Fig. 3 . Buria l anomaly , dZ B(ra), relativ e t o a norma l velocity-dept h trend , V N. Uplif t an d erosio n reduc e th e overburden thicknes s and result i n overcompaction expressed as anomalously hig h velocitie s relative t o present day depth (negative dZB). However, post-exhumational burial, BE, will mask th e magnitude o f the missing section , Azmiss (Eq . 2) . Undercompactio n a s a resul t o f rapi d buria l an d lo w permeabilit y cause s overpressure , AP comp (MPa), an d lo w velocitie s relative t o dept h (positiv e dZ B). Modifie d afte r Japse n (1998) .

NEOCENE UPLIFT AND EROSION O F SOUTHERN SCANDINAVI A

Fig. 4 . Outlin e o f th e derivatio n o f th e norma l velocity-depth tren d fo r th e Nort h Se a Chal k (V£ h, Eq. 3 i n Appendix) . Th e shallo w par t o f th e tren d i s constrained b y soni c dat a fro m pelagi c carbonat e deposits of Recent age (a). The deeper part of the trend is defined by the upper boun d for interval velocit y dat a from well s wher e th e Chalk i s at maximum buria l an d at hydrostatic pressure (b). (a ) Sonic logs from pelagic carbonate deposit s o f Eocene to Pleistocene ag e drilled in hol e 807 , Ocea n Drillin g Progra m (ODP ) Le g 13 0 (Shipboard Scientifi c Part y 1991) , an d th e Chal k Group i n th e Danis h Stenlille- 6 (locatio n show n i n Fig. 2 ) an d th e Karl- 1 well s (centra l Nort h Sea ; onl y the uppe r thir d o f th e lo g i s shown) , (b ) Interva l velocity v . mid-poin t dept h fo r th e Chal k Grou p fo r wells wher e thic k Quaternar y an d Neogene sediment s are present in areas with limited or no overpressure an d for well s wit h maximu m velocit y fo r z > 2000 m (5 5 out of 845 wells in Chalk velocity database). In (a), th e

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pressuring a s a resul t o f rapid , lat e Cenozoi c burial i n th e basi n centr e hav e resulte d i n a systematic variatio n o f buria l anomalie s (se e below). Thes e anomalie s ar e withi n ± 1 km for the Uppe r Cretaceous-Dania n Chal k an d lower Cenozoic sediment s whereas the upper Cenozoic sediments ar e clos e t o norma l compactio n (th e anomalies fo r th e Cenozoi c sediment s ar e calculated relativ e t o a baselin e fo r marin e shale; se e Eq. 4 in the Appendix) (Japse n 1998 , 1999). Thu s a regression lin e fitte d t o velocity depth data may not represent a physical model of the subsurfac e if suc h anomalie s ar e no t take n into account. (3) 'Assignmen t of a functional expression ' to an observed velocity-depth trend is not straightforward becaus e o f th e limite d rang e o f an y dataset. A n observe d tren d may , however , b e extrapolated t o range from th e surfac e to infinit e depth provide d tha t th e extrapolate d tren d complies wit h th e abov e boundar y conditions . Assignment o f a n arbitrar y mathematica l vel ocity-depth relation fo r a formation may lead t o identification o f a n erroneou s baseline , fo r example, th e frequentl y applie d formul a fo r shale trends , t t = l/V = aexp(—biz), tha t pre dicts bot h velocit y an d velocit y gradien t t o increase toward s infinit y wit h depth (t t is transit time (sm" 1), an d a (sm" 1 ) an d b (m ) ar e parameters). Erosion may thus be underestimated if suc h a trend is applied t o singl e data points at great depth. Formulation of velocit y baseline s is thu s not an arbitrar y choic e o f mathematica l function s and regressio n parameters , bu t shoul d b e considered a s settin g u p a physical mode l fo r a given lithology . First , baseline s shoul d b e established fo r formation s tha t ar e relativel y homogeneous wit h regar d t o macroscopi c acoustic properties , e.g . chal k o r marin e shale dominated b y smectite-illite . Second , baseline s should reflec t norma l compaction , an d buria l anomalies relativ e t o th e tren d shoul d b e i n agreement wit h othe r estimate s o f erosio n an d overpressure. Consequently , a baselin e fo r a

depth-shift shoul d b e note d betwee n th e thre e soni c logs tha t al l represen t pelagi c carbonate s o f ver y uniform composition . Th e shif t i s suggeste d t o b e caused b y overcompactio n a s a resul t o f remova l o f overburden alon g th e margi n o f th e Nort h Se a Basi n during lat e Cenozoi c tim e (Stenlille-6 ) an d b y undercompaction a s a resul t o f rapi d buria l i n th e central Nort h Se a durin g lat e Cenozoi c tim e (Karl-1 ; Chalk formatio n overpressur e i s 15MP a i n a nearb y well). (Compar e Fig . 3. ) Modified afte r Japsen (1998 , 2000).

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Fig. 5 . Interva l velocit y v . mid-poin t dept h fo r th e Chalk Grou p i n th e well s studied , an d th e revise d normal velocity-depth trend for the Chalk, V^h (Eq . 3 in Appendix ) (Japsen 2000). Mos t dat a points reveal high velocities relative to the normal trend, and this is suggested generally to be due to overburden reduction. Estimates o f erosio n base d o n chal k soni c dat a correlate well with estimates based on basin modelling (Fig. 12a) . The dashe d lin e indicate s th e origina l baseline of Japsen (1998).

given litholog y shoul d b e constraine d a t th e surface b y th e velocit y o f recen t deposit s o f th e sediment, a t shallo w depth s b y th e lowe r boun d for velocity-dept h dat a fo r whic h th e effec t o f overcompaction a s a resul t o f erosio n i s minimum, an d at greater depths along th e uppe r bound fo r dat a fo r whic h th e effec t o f under compaction a s a resul t o f overpressurin g i s minimum. Third , th e mathematica l formulatio n of baseline s shoul d b e constraine d b y simpl e boundary condition s a t the surface and at infinit e depth. The shap e o f baseline s fo r unifor m sedi mentary formation s reflect s th e fac t tha t th e compaction processe s depen d o n th e minera logical compositio n o f th e formation s (th e derivation o f norma l velocity-dept h trend s fo r three unifor m formation s i s outline d i n th e Appendix; se e Fig. 4). (1) Th e chal k baselin e reveal s a moderat e velocity increas e fo r depth s les s tha n 1 km, whereas the velocity gradien t increase s a t greater depths unti l i t i s graduall y reduce d a t depth s below 1.5k m (Fig . 5; Eq . 3 i n th e Appendix) . This variatio n i s i n agreemen t wit h th e preservation o f chal k porositie s o f c . 40 % t o depths of 1 km during normal compactio n befor e the onse t o f calcit e cementatio n an d th e consequent increas e o f velocit y (Borr e & Fabricius 1998 ; Japsen 1998) .

Fig. 6 . Interva l velocit y v . mid-poin t dept h fo r th e Lower Jurassi c F- I Membe r in th e well s studied , an d the norma l velocity—dept h tren d fo r Lowe r Jurassi c shale, V^ h, suggeste d t o b e characteristi c for marine shale dominate d b y smectite-illit e (Eq . 4 i n Appendix). Al l dat a point s revea l hig h velocitie s relative to the normal trend, and this is suggested to be due t o overburde n reductio n an d t o a hig h conten t o f sand or kaolin in shale deposited close to the exposed basement o f th e Scandinavia n Shiel d durin g earlies t Jurassic time. Estimates of erosion based on sonic data for th e Lower Jurassic shal e are overestimated relative to estimate s base d o n chal k dat a i n N E Denmar k (Fig. 12b).

(2) Th e moderat e velocit y gradien t o f th e baseline fo r marin e shal e dominate d b y smectite-illite ma y also b e relate d t o miner alogical compositio n (Fig . 6; Eq . 4) . Smectite illite particle s ar e separate d b y wate r molecule s (Bailey 1980) , an d thi s interlaye r wate r i s adsorbed t o the particles eve n during deep buria l (van Olphe n 1966) . Japse n (1999 , 2000 ) argue d that th e wate r adsorbe d o n th e smectite-illit e particles coul d lea d t o wea k mechanica l grai n contacts, an d thu s t o th e lo w soni c velocit y observed fo r th e marin e shal e a t depth . (3) Th e baselin e fo r th e continenta l Bunte r Shale (Lowe r Triassic ) i s foun d t o b e simila r t o the normal trend of the Bunter Sandstone (Fig. 7 : Eq. 5 in the Appendix). Japsen (2000 ) suggeste d that thi s similarit y coul d b e relate d t o th e hig h kaolin conten t o f th e Bunte r Shale . Kaoli n ha s little adsorbed wate r and it builds up thick flake s that ar e u p t o a thousan d time s large r tha n smectite-illite particle s tha t ar e separate d b y water molecule s (Baile y 1980 ; Lindgreen, pers .

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189

estimating erosion , and that the trend fo r marin e shale give n b y Eq . 4 i s i n agreemen t wit h th e suggestions o f Scherbau m (1982 ) an d wit h tha t of Hanse n (1996 ) fo r th e uppe r c . 2k m o f th e shale trend . Suggeste d baseline s fo r th e Bunte r Shale diffe r significantly , bu t th e tren d given b y Eq. 5 agrees with the results of Marie (1975 ) an d Bulat & Stoke r (1987 ) fo r th e velocit y interva l from whic h mos t dat a ar e availabl e (se e discussion b y Japsen (2000)) . Burial anomaly and missing section Velocity-depth studie s hav e prove n usefu l because the y ar e base d o n easil y accessibl e data wit h a wide area l coverage, an d thus allo w for settin g up simple constraints on both physical and geologica l parameters . Over - an d under compaction relativ e t o a baseline may be studie d by computin g buria l anomalies , dZ B (m) , fo r a formation a s th e differenc e betwee n presen t depths an d depth s correspondin g t o norma l compaction fo r th e measure d velocit y ( z an d Z N (m), respectively , V i s soni c velocit y (ms" 1 ); Fig. 3 ) (Japsen 1998 ) (Eq . 1) : Fig. 7 . Interva l velocit y v . mid-poin t dept h fo r th e Bunter Sandston e an d Bunte r Shal e i n th e well s studied, an d th e Bunte r Shal e trend , V^ sh, whic h i s suggested t o b e characteristi c fo r lithologie s domi nated b y quart z and/o r kaoli n (Eq . 5 i n Appendix) . Considerable lithologica l variation s ar e likel y withi n these formations, bu t the plot show s a number o f dat a points clos e t o th e norma l tren d an d other s tha t plo t above th e tren d generall y a s a resul t o f overburde n reduction. Estimate s o f erosion base d o n Bunter Shal e and Sandstone dat a correlate wel l with estimates base d on Chalk data (Fig. 12c) . The clear distinction betwee n the trend for the continental Bunter Shale an d the trend for marin e Lowe r Jurassi c shal e i n Fig . 6 shoul d b e noted.

comm.). A shal e dominate d b y well-packe d flakes of kaolin coul d thus acquire rock physica l properties simila r t o thos e o f a consolidate d sandstone. A n alternativ e explanatio n could , however, b e relate d t o dominanc e o f quart z i n both the Bunter Sandstone an d the Bunter Shale. The shar p increas e o f velocit y aroun d 2 km fo r these formations could thus be related t o onset of quartz cementation , whic h ha s bee n reporte d t o develop a t depths belo w 2. 5 km in , for example , the Nort h Se a (Bj0rlykk e & Egeber g 1993) . Further studie s ar e require d t o clarif y thes e issues. Finally, i t shoul d b e note d tha t th e chal k baseline give n by Eq. 3 is in agreement wit h that of Hilli s (1995 ) fo r th e depth s relevan t fo r

A baselin e ma y b e give n a s a linea r trend , V — VQ + k-z, wher e V Q i s velocit y a t th e surface an d & (ms"1 m"1, or s"" 1) is the velocity gradient (se e Eq . 3) . Th e buria l anomal y thu s becomes (Japse n 1993 , 1998) :

where A z i s laye r thickness , A 7 (s ) two-wa y traveltime thickness an d z t i s depth t o the top of the layer . A baselin e ma y als o b e formulate d a s a constrained, exponentia l transi t time-dept h trend, t t = (tt 0 - rr 00)e~z/fc2 + #«> , wher e t t = l/V ( s m"1) is transit time, # 0 and ttoo are transit time a t th e surfac e an d a t infinit e depth , respectively an d b 2 (m ) i s a n exponentia l constant (se e Eq . 4) , I n thi s cas e th e buria l anomaly may be approximate d b y the following expression i f laye r thicknes s an d velocit y gradient ar e moderate (Japse n 1999) :

Low velocitie s relativ e t o dept h giv e positiv e burial anomalies , whic h ma y indicat e under compaction a s a resul t o f overpressur e (se e Japsen 1998) . Hig h velocitie s relativ e t o dept h give negativ e buria l anomalies , whic h ma y b e caused b y a reductio n i n overburde n thicknes s

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P. JAPSEN ETAL.

('apparent uplift' , Bula t & Stoke r 1987 ; 'ne t uplift an d erosion' , Rii s & Jense n 1992 ; 'apparent exhumation' , Hilli s 1995) . I t must , however, b e note d tha t an y post-exhumationa l burial, B E (m) , wil l mas k th e magnitud e o f th e missing overburde n section , Az miss (m) , an d w e obtain where th e minu s indicate s tha t erosio n reduce s depth (Hilli s 1995 ; Japse n 1998) . Whether a buria l anomal y i s a measur e o f erosion o r i s cause d b y othe r factor s (e.g . lithological changes ) i s subjec t t o a n integrate d evaluation o f th e are a i n question . Apar t fro m being i n agreemen t wit h othe r estimate s o f erosion, th e buria l anomalie s shoul d als o correspond geographicall y t o th e exten t o f a section missin g fro m th e stratigraphi c record . Erosion ma y b e underestimate d i f th e compac tion proces s is reversibl e whe n the loa d of the overburden i s reduced . However , w e fin d tha t previous estimates of erosion based on sonic data from th e stud y are a ar e exaggerate d b y u p t o 1000m, mainly becaus e o f the above-mentione d problems wit h identifyin g vali d baseline s fo r homogeneous formation s (se e sectio n 'Compari son wit h othe r studies') . Overestimatio n o f erosion fro m soni c dat a thu s seem s t o b e a typical proble m i n suc h studies , an d under estimation becaus e o f th e reversibilit y o f th e compaction proces s seem s t o b e onl y o f mino r importance. Summing up , a larg e area l an d stratigraphi c data coverag e i s o f crucia l importanc e fo r evaluating th e validit y o f estimate s o f erosio n from soni c data . Moreover , bot h velocit y base lines an d estimate s o f erosio n shoul d b e i n agreement wit h stratigraph y i n th e area , for mation overpressur e i n adjacen t basin s an d estimates o f erosio n base d o n independen t methods. I t i s thu s a n importan t tes t o f th e validity o f th e derive d baseline s tha t th e know n overpressure i n th e centra l Nort h Se a ma y b e predicted fro m buria l anomalie s relativ e t o th e trends fo r chal k an d fo r marin e shal e (Japse n 1998, 1999) . Sonic data Velocity-depth data fro m 6 0 Danish well s form part o f th e databas e fo r thi s stud y (Fig . 2) . Th e data were presented by Nielsen & Japsen (1991), apart fro m th e Ida- 1 an d Jelling- 1 wells . Fifty two o f th e well s have interva l velocitie s fo r th e Chalk Group , 3 1 fo r th e F- I Membe r o f th e Lower Jurassi c Fjerritsle v Formation (Michelse n

1989), an d 2 2 fo r th e Bunte r Sandston e o r th e Bunter Shal e (Bertelse n 1980) ; 4 2 well s hav e data fro m th e Chal k a s wel l a s fro m th e pre Chalk interva l (Fig s 5-7) . Dat a fro m thre e Norwegian well s wit h Chal k velocit y dat a ar e included t o support the contouring. Chalk buria l anomalie s wer e calculate d relative t o th e revise d norma l velocity-dept h trend fo r th e Chal k developed b y Japse n (1998 , 2000) (Eq . 3) . Anomalie s fo r th e pre-Chal k formations wer e calculate d relativ e t o baselines suggested by Japsen (2000). Burial anomalies for the Lowe r Jurassi c F- I Membe r ar e calculate d relative t o th e shal e tren d give n b y Eq . 4 , an d those fo r th e Bunte r Sandston e an d th e Bunte r Shale relative to the Bunter Shale trend are given by Eq . 5 . A singl e burial anomaly as a result of late Cenozoic erosion was estimated on the basis of th e availabl e soni c dat a fo r eac h well , an d corrected fo r the Quaternary reburial to obtain an estimate o f th e missin g sectio n (Eq . 2 ) (se e Japsen & Bistrup (1999) for details).

Erosion estimated from basin modelling Model description and input data The basin development of the study area has been modelled fro m 3 5 wells wit h a commercial, I D forward modellin g progra m (Yiikle r 1978 ; Iliff e & Dawson 1996) . The program starts simulation of th e geologica l developmen t fro m th e bas e o f the sedimentar y sectio n an d perform s a calcu lation of parameters such as formation thickness, pressure, temperatur e and vitrinite reflectance as a functio n o f time. From th e geologica l inpu t a t the wel l locatio n (thicknes s an d age s o f sediments, estimate d magnitud e an d timin g o f erosion an d estimate d heat-flo w history ) th e program calculate s vitrinit e reflectance , tem perature an d pressur e a s a function o f time. The calculated value s o f thes e parameter s fo r th e present-day situatio n ar e the n compare d wit h data that can be divided int o two groups: (1) data that constrai n th e therma l history : present-da y temperature (BHT , botto m hol e temperatures ) and thermal maturity indicators (mainly vitrinite reflectance values in the study area); (2) data that constrain th e compactio n o f th e sediments : pressure an d porosit y (n o overpressur e i s encountered withi n the stud y area). If th e measure d an d calculate d values do no t match, the input parameters ar e changed and the program is run again, until a satisfactory match is obtained. A genera l matc h t o man y dat a point s and a laterall y consisten t heat-flo w an d erosio n model wa s preferre d becaus e o f th e regiona l

NEOCENE UPLIFT AND EROSION O F SOUTHERN SCANDINAVI A

character o f th e study , th e varyin g dat a qualit y and the sometimes conflicting values . Chronostratigraphic event definition. Th e pro gram quantifie s al l importan t processe s a s a function o f time , an d th e basi n developmen t i s thus defined in terms of chronostratigraphic unit s valid for the entire area (mode l layer s or events). The numbe r o f event s i n th e geologica l model , including period s o f deposition , non-depositio n and erosion, must be chosen in such a way that all major change s ca n b e describe d an d relate d t o existing stratigraph y whil e avoidin g excessiv e calculation times . A total of 45 events was chosen to describe th e geological developmen t withi n th e stud y are a from Cambrian time until the present (a time step of 250 ka and a depth step of 25 m). The duration of the events is shorter in the Cenozoic interval to describe th e rapi d change s i n thi s period . Th e lithologies use d i n th e modellin g wer e kep t constant fo r th e sam e even t i f n o geologica l information dictate d otherwise . Th e mai n litho types for the events are: Cenozoic event s 30-45 (excluding Dania n time ; 60- 0 Ma): san d an d shale wit h occasiona l coals, mor e shal y toward s the bas e o f th e succession ; Lat e Cretaceous Danian events 26-29 (96-60Ma): Chalk; Early Cretaceous event s 24-25 (129-96 Ma): mixture of silt, marl and shale; Late Jurassic events 20-23 (152-129 Ma): shal e and siltstone; Mid-Jurassic events 17-1 9 (178-152Ma) : mixe d sandstone , siltstone and shale; Earl y Jurassi c event s 13-1 6 (210-178 Ma): shale,siltand sandy shale ;Triassic events 10-1 2 (250-21 0 Ma): sandstone , sand y shale t o shal e wit h carbonate ; locall y salt ; Lat e Permian even t 9 (256-250 Ma): clean sal t if the layer i s thick , otherwis e a mixtur e o f shale , anhydrites an d salt ; Cambrian-Earl y Permia n events 1-8 (570-256 Ma): shale and sandstone. Sparse temperatur e an d vitrinit e reflectanc e data ar e availabl e belo w uppermos t Triassi c level, and even fewe r below lowermos t Permia n level. Th e model fo r each wel l was extrapolate d to basement b y the use of seismic dat a where no well dat a wer e available . Palaeo-surf ace-temperatures. Th e palaeo-sur face-temperature is the averag e temperatur e at the sediment-water interface during a particular period. Estimate s o f palaeo-surf ace-temperature were modifie d fro m Buchardt (1978 ) (Fig. 8) . Heat-flow model. Th e heat-flo w histor y wa s constrained by two datasets only: (1) the presentday temperatur e i n th e wells , whic h combine d with the given lithologies (thermal conductivities)

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defines th e present-da y hea t flow ; (2 ) value s of vitrinite reflectance, whic h define the heat flow at the tim e o f maximu m buria l fo r a give n formation. A simpl e heat-flo w mode l tha t honour s thes e values i s estimated fo r eac h wel l base d o n a n assumption o f a constan t hea t flo w o f 1 Hea t Flow Uni t (HFU) fro m Cambria n time (even t 1) until the beginning o f Oligocene time (even t 34 ) (1 HFU = 42mWm~ 2 ). This valu e is typical for the constrained par t o f the heat-flo w history , but the exact choice i s of minor importanc e becaus e only the heat flow at maximum burial affects th e vitrinite reflectanc e an d mos t o f th e stud y are a has bee n affecte d b y erosion . Th e heat-flo w model fo r th e Oligocene-Recen t tim e interva l was modifie d t o matc h present-da y temperatur e and vitrinite reflectance data, but also to result in smooth heat-flo w variation s i n tim e an d spac e (Fig. 8) . The hea t flo w use d i n th e mode l i s th e hea t flow at the base of the sedimentary succession. In contrast t o th e backgroun d hea t flo w fro m th e upper mantle , th e hea t flo w a t thi s leve l i s affected b y transien t effect s fro m sedimentatio n and erosio n (Vi k & Hermanru d 1993 ) an d b y local variation s in the geometry an d lithology of the sedimentary units (e.g. sal t domes). Transient

Fig. 8 . Plots o f palaeo-surface-temperatures an d heatflow histor y use d i n th e modellin g o f Hyllebjerg- 1 well. A hea t flo w o f 1 HFU ha s bee n use d unti l Oligocene time , afte r whic h heat flow was allowed t o change smoothl y t o matc h calibratio n data . Th e palaeo-surface-temperatures ar e th e sam e fo r al l wells i n th e stud y (Buchardt , 1978) . (Se e th e corresponding calibratio n plot i n Fig. 9. )

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P. JAPSEN ETAL.

NEOGENE UPLIF T AND EROSION O F SOUTHERN SCANDINAVI A

heat-flow variatio n fro m 0. 9 t o 1. 1 HFU i n th e sedimentary sectio n wa s modelle d b y assumin g deposition ( 1 km pe r 5 Ma) followe d b y erosio n (also 1 km pe r 5 Ma) o n to p o f a sedimentar y succession o f 5 km an d assumin g constan t hea t flow o f 1HF U a t th e base o f a 50k m thic k basement sectio n (usin g th e PetroMo d 2 D software, versio n 6.1) . Erosion model. Vitrinit e reflectanc e i s mor e sensitive t o temperatur e tha n t o time , an d therefore depends mainly on maximum tempera ture an d no t o n th e timin g o f maximu m temperature. In this study , therefore, assessmen t of the timing of erosion was based on the existing stratigraphy and on general geologica l consider ations becaus e n o fission-trac k o r (U-Th)/H e data were available . A model o f erosion startin g in lat e Miocen e tim e an d continuin g unti l lat e Quaternary time was used in the basin modelling throughout th e are a (se e th e discussio n i n th e section 'Timin g o f maximu m buria l an d subsequent erosion') . Model calibration and results Basin modellin g wa s performe d fo r 3 5 well s where the heat-flow an d erosion mode l coul d be calibrated agains t temperature an d vitrinit e data (Figs 2 an d 9) . Th e estimate d heat-flo w mode l varies smoothl y wit h tim e fo r individua l well s (e.g. Fig . 8) , an d i n spac e fo r differen t tim e intervals (Fig . 10) . T o matc h th e stee p vitrinit e reflectance gradient s observe d i n thi s are a (Fig. 9) , lo w heat-flo w value s hav e bee n use d for central and northern Jylland during maximum burial (partl y a s a resul t o f transien t effect s caused b y sedimentatio n befor e maximu m burial). T o matc h present-da y temperatures , high heat-flo w value s hav e bee n use d fo r th e most recen t developmen t i n norther n Jyllan d (partly because of transient effects reflectin g late erosion). Calibration o f th e mode l assume d th e depo sition o f overburde n sediment s an d thei r subsequent remova l i n al l well s studied , apar t from th e L- l an d S- l well s (Fig . 11) . Th e thickness o f th e missin g sectio n ca n onl y b e estimated withi n a rang e o f possibl e solutions , and this range is constrained by th e data quality (e.g. vitrinite data), by th e latera l consistenc y of

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the mode l o f erosio n an d hea t flow, and b y th e general geological understanding of the area. The uncertainties o n th e estimate s o f erosio n ar e o f the orde r o f 100-200m . Quantification o f late Cenozoic erosion in Denmark Estimates of missing section based on different data Estimates of the missing sectio n base d o n Chal k sonic data and on basin modelling are similar; the correlation coefficien t i s 0.8 1 fo r 2 4 well s i n common an d th e mea n differenc e betwee n th e estimates i s 30m , standar d deviatio n 130 m (Fig. 12a) . The estimate s o f th e missin g sectio n based o n Chal k an d Triassi c soni c dat a ar e als o rather similar , bu t th e scatte r i s greater ; th e correlation coefficien t i s 0.72 fo r th e 1 9 wells in common an d th e mea n differenc e betwee n th e estimates i s 10m , standar d deviatio n 210 m (Fig. 12c) . Th e estimate s of the missin g sectio n based o n Chalk soni c dat a ar e generally smalle r than thos e base d o n Lowe r Jurassi c soni c data ; the correlation coefficien t i s 0.71 for the 27 wells in common an d the mean difference betwee n the estimates i s 150m , standar d deviatio n 270 m (Fig. 12b) . This difference i s suggested to be due to lithologica l variation s withi n th e Lowe r Jurassic sequenc e a s discussed below. Only i n N E Denmar k doe s th e magnitud e o f the buria l anomal y base d o n soni c dat a fo r th e pre-Chalk section generall y excee d the estimate s from Chal k sonic data and from basin modelling; e.g. difference s o f 500-1000 m betwee n anomalies base d o n soni c dat a fo r th e Chal k and th e Lowe r Jurassi c units . A possibl e interpretation o f thi s differenc e coul d b e tha t the Lower Jurassic units in that area experience d maximum buria l befor e th e depositio n o f th e Chalk (se e Japse n 2000) . I t i s not , however , possible t o identif y a majo r hiatu s i n th e stratigraphic recor d correspondin g t o a dee p erosional event tha t could explai n th e difference between th e Chal k an d th e pre-Chal k buria l anomalies. Maximu m buria l o f th e Mesozoi c succession i s thu s suggeste d t o hav e occurre d during Cenozoic tim e in all wells studied. Lateral lithological variation within the Lower Jurassic sequenc e i s a likel y caus e fo r th e hig h

Fig. 9 . Plot s o f calculate d an d measure d value s o f (a ) vitrinit e reflectanc e an d (b ) temperatur e fo r th e well s B0rglum-l, F-l, Hyllebjerg-1, 0rslev-l, S-l an d T0nder-2. The slow increase of vitrinite reflectanc e wit h dept h for, fo r example, th e B0rglum- l wel l shoul d be noted . Thi s is suggeste d t o be partly du e to a transient therma l effect relate d to deposition followed b y erosion during Cenozoi c time.

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Fig. 10. Maps of heat flow at three times during the late Cenozoic perio d resultin g fro m th e basi n modelling . The hea t flo w i s assume d t o b e constan t unti l Oligocene tim e ( 1 HFU), bu t throughou t Oligocene Recent tim e th e heat flow is allowed t o vary smoothly to matc h calibratio n data , (a ) Presen t day ; (b ) 2 Ma before present ; (c ) 1 0 Ma before present .

Fig. 11 . Maps o f erosion a t three interval s during late Cenozoic tim e resulting from th e basin modelling. The existing stratigraph y constrains the tempora l development of the erosion, (a ) Quaternary time ; (b) Pliocen e time; (c ) lat e Miocen e time . A mode l o f erosio n starting in late Miocene time has been used in the basin modelling throughou t the area .

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estimates o f erosio n base d o n soni c dat a fro m these sediment s compare d wit h estimate s fro m both basin modelling an d from Chal k soni c data. During earlies t Jurassi c tim e a shallo w marin e environment prevaile d clos e t o th e expose d basement o f th e Scandinavia n Shield , an d correspondingly a hig h conten t o f san d an d kaolin ha s bee n reporte d fo r the Lowe r Jurassic shale i n thi s are a (Pederse n 1985 ; Lindgree n 1991). The relatively high velocity of the Lowe r Jurassic sequenc e in NE Denmark coul d thus be due t o th e hig h conten t o f both san d an d kaoli n (see the section 'Derivatio n of normal velocit y depth trends ' an d Appendix) . Furthermore , th e general agreemen t betwee n th e erosiona l esti mates based on Chalk sonic data and those based on basin modelling i n northern Jylland, as in the rest of Denmark, suggests that estimates based on Chalk dat a ar e to be preferred t o those based o n Lower Jurassic data. The hig h velocitie s o f th e F- I Membe r hav e previously bee n suggeste d t o reflec t maximu m burial before Lat e Cretaceous-Paleogen e inversion withi n th e Sorgenfrei-Tornquis t Zon e an d the subsequent removal of 1 km of Chalk (Fig. 1 ) (Japsen 1993 ; Michelse n & Nielse n 1993) . Removal o f a thic k Chal k sectio n durin g th e inversion is , however , unlikel y becaus e seismi c reflectors withi n th e Chal k sectio n onla p a n anticlinal structur e alon g th e Sorgenfrei-Torn quist Zon e i n th e wester n part s o f th e Kattega t (Liboriussen e t al, 1987) . Syndepositiona l growth o f th e inversio n structure , wit h pea k movements durin g mid-Cretaceou s times , implies tha t onl y a thi n Chal k sectio n wa s deposited i n the inversion zone . Magnitude of late Cenozoic erosion The sectio n remove d b y lat e Cenozoi c erosio n has bee n estimate d fo r 6 8 Danis h well s o n th e

Fig. 12 . Correlatio n betwee n estimate s o f missin g section, (a ) Estimate s base d o n Chal k soni c data ,

AZ^SS, v . estimates fro m basi n modelling , AZ^ SS. (b ) Estimates based on Chalk sonic data, AZj^ss, v. estimates based o n soni c dat a fo r Lowe r Jurassi c F- I Member , AZJ^. (c ) Estimate s base d o n Chal k soni c data , AZ^SS, v . estimate s base d o n soni c dat a fo r Lowe r Triassic Bunte r Sandston e an d Bunte r Shale , AZ^ riss. The goo d correlatio n betwee n estimate s base d o n Chalk velocitie s an d o n basi n modellin g shoul d b e noted. Estimate s base d o n dat a fo r Lowe r Jurassi c shale ar e overestimate d relativ e t o estimate s fro m Chalk dat a i n N E Denmar k becaus e o f lithologica l variations i n th e shale . Th e line s illustratin g th e 1: 1 relationship betwee n th e estimate s ar e shown . Wel l names ar e give n fo r well s wit h result s o f basi n modelling show n i n Fig . 9 .

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Fig. 13 . Map of the section missing as a result of late Cenozoic erosion based on estimates from basi n modelling and sonic data from 6 8 Danish and three Norwegian wells. A succession of about 500 m of post-Chalk sediments is missing from larg e parts of the area. Towards the NE, c. 1000 m are missing where the Chalk is absent or deeply eroded o n an d alon g th e Skagerrak-Kattega t Platform.

basis o f a combinatio n o f result s fro m basi n modelling i n 3 5 well s wit h compactio n studie s based on velocity-depth data in 60 wells (Figs 13 and 14 ) (Japsen & Bidstrup 1999) . (1) Th e estimate d missin g sectio n reache s 1000m i n a numbe r o f well s i n N E Denmark , on o r clos e t o th e Skagerrak-Kattega t Plat form, wher e th e Chal k i s deepl y truncate d o r absent (Fig . 1) . Fo r th e Frederikshavn- 1 well , both method s indicat e a missin g sectio n o f c. 1000m , wherea s th e maximu m base d o n basin modellin g alon e i s 1200 m i n th e Hans- 1 well, an d th e maximu m fro m soni c dat a alon e is 1300 m i n th e Saeby- 1 well . Th e latte r valu e is, however , base d o n dat a fro m th e Lowe r Jurassic sequence , an d appear s t o b e over estimated whe n compare d wit h result s fro m neighbouring wells . (2) Th e missin g sectio n i s foun d t o b e jus t over 500 m i n man y Danis h well s locate d i n a broad ban d fro m N W t o S E wher e th e Uppe r Cretaceous-Danian Chal k i s preserved . (3) Missin g section s betwee n 25 0 an d 500 m are mappe d i n th e wester n an d souther n par t o f the stud y area. Erosio n estimate s o f onl y 100-250m ar e foun d i n sout h westernmost Jylland. Estimate s o f 250 m ar e clos e t o th e accuracy o f th e methods , an d th e Mesozoi c succession i s thu s probabl y clos e t o norma l compaction i n thi s area . Th e tw o easternmos t

Danish well s with Mesozoic sediment s found to be at maximum burial are the L-l an d S-l wells . The absenc e o f upper Pliocen e sediment s i n th e L-l indicate s non-depositio n rathe r tha n a n episode o f erosion (Laurse n 1992) . The magnitud e o f erosio n increase s clearl y across th e N E sid e o f th e Sorgenfrei-Tornquis t Zone, wher e estimate s o f erosio n ar e a s hig h as on the Skagerrak-Kattegat Platform to the north, c. 1000m . I n th e souther n par t o f th e inversion zone, estimate s o f erosio n ar e simila r to values found sout h of the zone, c. 500 m. This differenc e indicates stronge r tectoni c movement s o n th e Skagerrak-Kattegat Platfor m tha n in the are a t o the south. Along the S Wedge o f the Skagerrak-Kattegat Platform, th e missin g sectio n o f 1000 m i s suggested t o be c. 50 0 m of Cenozoic sediment s as i n the Danis h Basin to th e sout h plus a Chalk section o f c . 500m . Farthe r nort h o n th e Platform, th e missin g Cenozoi c sectio n i s suggested t o b e c . 250m , an d th e missin g Chalk sectio n c . 750m . O n th e basi s o f thes e assumptions a profil e alon g Jyllan d ha s bee n reconstructed t o th e situatio n befor e Neogen e uplift an d erosio n (Fig . 14) . Th e reconstructe d profile reveal s th e know n Chal k depocentr e (thickness <2km ) sout h o f th e Sorgenfrei Tornquist Zon e a s wel l a s on e N E o f th e Zon e (reconstructed thicknes s < 1 km) correspondin g

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to a predictio n base d o n thermo-mechanica l modelling (Gemme r 2002) . We have estimated th e thickness of sediment s deposited withi n th e stud y are a durin g th e Cenozoic epoch s fro m th e know n stratigraphy , the estimatio n o f th e tota l missin g sectio n an d from th e abov e assumption s abou t remove d chalk cover : n o Pliocene sediment s ar e found t o have been deposite d withi n the stud y area apar t from th e S W par t wher e thicknesse s ar e estimated t o hav e reache d 300m ; a Miocen e depocentre < 110 0 m thick is found t o have been located wes t o f Jyllan d wit h thicknesse s gradually decreasin g t o 100 m toward s th e N E and east and 300 m towards the SE; an Oligocene depocentre < 600 m thick i s found t o have bee n located N W o f Jylland ; i n th e res t o f th e are a thicknesses rang e fro m 5 0 t o 100m ; Eocen e deposits ar e foun d t o hav e covere d th e entir e area, excep t fo r th e S W par t o f th e Nort h Sea , with thicknesse s rangin g fro m 5 0 to 150m ; Paleocene (excludin g th e Danian ) deposit s ar e found t o hav e covere d th e entir e are a wit h thicknesses o f 50-100m, locally 150m .

Timing of maximum burial and subsequent erosion Along th e margin s o f th e Nort h Se a Basin , th e Chalk wa s a t maximu m buria l befor e Neogen e erosion (Japsen 1998). This conclusion applies to the majorit y o f wells i n the presen t stud y wher e Chalk i s overlai n b y Paleogen e sediments , an d where soni c dat a fo r th e Chal k indicat e a previous greater dept h of burial. Thus maximum burial o f th e Chal k mus t hav e occurre d durin g Cenozoic tim e afte r earl y Cenozoic burial. (1) Th e timin g o f erosio n ca n b e furthe r detailed i n souther n an d centra l Jyllan d an d offshore west of Jylland. Here erosion must postdate th e depositio n o f offshor e t o shorefac e sediments o f th e uppe r Miocen e Gra m For mation; fo r example, durin g Plio-Pleistocen e time (Rasmusse n 1961 ; Rasmussen , pers . comm.). Thi s erosiona l even t thu s matche s th e basin-wide hiatu s a t th e bas e o f th e Plio Pleistocene deposit s tha t ar e younge r tha n 2.4 Ma (the late Cenozoic successio n is complete only i n a narrow zone i n th e centra l North Sea ) (e.g. Zagwijn 1989 ; se e Japsen 1998) . Fig. 14 . (a ) Profil e o f th e post-Triassi c successio n i n Jylland, and reconstructions at (b) 2 Ma and (c) 1 0 Ma before presen t (befor e uplif t and erosion increasing to 1 km towards the NE). The (now partl y eroded) Chalk depocentre o n th e Skagerrak-Kattega t Platfor m

should b e noted . DB , Danis h Basin ; RFH , Ring k0bing-Fyn High ; SKP , Skagerrak-Kattega t Plat form; STZ , Sorgenfrei-Tornquis t Zone . Locatio n shown i n Fig. 2 . (Compare Japsen (1993). )

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(2) Th e timin g o f erosio n i s les s constraine d on an d alon g th e Skagerrak - Kattegat Platform , where th e Chal k i s deepl y erode d o r absen t (Fig. 1) . Th e pronounce d movement s o f th e Platform documente d her e may, however, lead to the suggestio n tha t erosio n o n the Platfor m wa s initiated earlie r tha n i n th e res t o f th e area ; fo r example, b y uplif t durin g mid-Miocen e time . This suggestio n i s further supported b y th e shif t in sedimen t transpor t directio n durin g mid Miocene tim e observe d i n th e northeaster n North Se a Basi n a s discusse d belo w (Fig . 1 ) (Clausen et al 1999) . The sedimentar y deposit s i n th e stud y are a were thu s erode d durin g lat e Cenozoi c tim e subsequent t o thei r maximu m burial , whic h generally mus t have taken place durin g Neogene time. The erosion i s likely to have been initiated in tw o phases , a n earl y phas e o f ?Miocen e ag e and a better constrained late r phase that began in the Pliocene . Th e lat e even t ha s affecte d a vast area tha t reaches far into the North Sea , wherea s the extent of the early phase must be restricted t o areas close to Scandinavi a where n o Miocene i s present, e.g . th e Skagerrak-Kattega t Platfor m and souther n Sweden . The lat e Cenozoi c erosio n wa s followe d b y Quaternary reburial , whic h i n onshore Denmar k occurred late r tha n c . 0. 3 Ma, th e ag e o f oldes t Quaternary sediment s (Knudse n 1995) . Th e ag e of the oldest Quaternary deposits is progressively older farthe r int o th e Nort h Sea , an d i n th e westernmost par t o f the Danis h sector, sedimen tation wa s almos t continuou s durin g lat e Cenozoic tim e (se e Konrad i 1995 ; Japse n 1998) .

Comparison with other studies Three studie s fro m th e earl y 1990 s foun d th e missing sectio n t o be substantiall y greater alon g the Sorgenfrei-Tornquis t Zon e tha n suggeste d here (Jense n & Schmid t 1992 ; Japse n 1993 ; Michelsen & Nielse n 1993) . Th e studie s b y Japsen (1993 ) an d b y Michelse n & Nielse n (1993) wer e base d o n dat a fro m th e Lowe r Jurassic F-I Member, which , as suggeste d i n th e previous section , ma y lea d t o overestimate d erosion toward s th e N E becaus e o f lithologica l variations withi n th e unit . Furthermore , th e difficulty i n definin g th e absolut e leve l o f baselines fo r differen t lithologie s contribute s t o the uncertainty of the estimated burial anomalies. Jensen & Schmidt (1992, 1993 ) overestimate d erosion b y a n average o f c. 450 m relative to th e results presente d her e fo r th e 1 4 Danis h well s in common ; th e maximu m overestimat e wa s 1000m. Th e buria l anomalie s o f Jense n &

Schmidt (1992 , 1993 ) wer e estimate d fro m vitrinite reflectance, density an d soni c data . Japsen (1993 ) overestimate d erosio n b y a n average o f c . 250 m relativ e t o th e result s presented her e fo r th e 3 1 well s i n common ; th e maximum overestimat e wa s 1000m . Th e esti mates of erosion wer e based o n the same data for the F-I Member as used here, and were calculated relative t o th e linea r velocity-dept h tren d fo r Lower Jurassi c shal e i n th e N W Germa n Basi n determined b y Scherbau m (1982) . Thi s tren d deviates les s tha n 100 m fro m th e shal e tren d applied her e fo r th e relevan t velocit y interva l (Eq. 4) . Michelsen & Nielse n (1993 ) overestimate d erosion b y c . 650 m relativ e t o th e result s presented her e fo r th e seve n well s i n common ; the maximu m overestimat e wa s 1300m . Th e burial anomalie s wer e calculate d fo r th e F- I Member relativ e t o a simpl e exponentia l transi t time-depth tren d establishe d b y thos e workers . This tren d result s in overestimate s o f erosio n o f up t o 300 m fo r th e relevan t velocit y interva l relative to the trend applied here (Eq . 4). Japsen (1998 ) overestimate d erosio n b y a n average o f c . 200 m relative to th e presen t study for th e 5 1 well s i n common ; th e maximu m overestimate wa s 500m . I n th e presen t study , Chalk burial anomalies are calculated relative to the revised Chalk baseline (Eq. 3), which, for the main velocit y interval , i s shifte d c . 200 m towards mor e shallo w depth s relativ e t o th e Chalk tren d applied by Japsen (1998) . Huuse et al. (2001) found erosion to be several hundred metres less whe n estimated from strata l geometries alon g a north-south seismi c sectio n west o f Jyllan d tha n whe n estimate d fro m th e maximum buria l studie s o f Jense n & Schmid t (1993) an d Japse n (1998) . Thi s conclusio n regarding the S-l wel l to the south of the section is in agreement with the present study, where the drilled sectio n i n thi s wel l i s foun d t o b e a t maximum buria l today . Thi s interpretation , which i s supporte d b y vitrinit e data , i s als o a consequence of the revision of the Chalk baseline as discusse d i n th e Appendi x (Fig . 4) . W e find , however, tha t a missin g sectio n o f c . 500 m o f Miocene sediment s i s compatibl e wit h th e stratigraphy aroun d th e norther n F- l well , where n o Miocene sediment s are present today. Neogene uplift of southern Scandinavia and of the South Swedish Dome Lidmar-Bergstrom (1999) discussed the geomorphological evidence concerning the uplift history of th e thre e surfac e domes i n Scandinavia . The

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Northern an d Souther n Scande s ar e th e mos t prominent o f th e domes . Th e Souther n Scande s constitute the main part of southern Norway and have a maximum elevation o f 2540 m above se a level (a.s.l.) . Ther e i s a genera l consensu s tha t major uplif t o f the Southern Scandes occurred i n Neogene time (Peulvast 1985; Jensen & Schmidt 1992; Rohrma n e t al 1995 ; Riis 1996) . Thi s timing of uplift i s in accordance with a Neogene onset o f erosio n i n norther n Denmark . Th e Southern Scande s wer e als o affecte d b y a Paleogene uplif t phas e (Clause n et al. 2000 ; Lidmar-Bergstrom e t al. 2000). Denudation surfaces of the South Swedish Dome The Sout h Swedish Dome, whic h is the smallest of th e Scandinavia n domes , culminate s N E o f Kattegat, jus t sout h o f lak e Yattern , an d ha s a maximum elevation of 380 m a.s.l . (Fig s 1 and 15) (Lidmar-Bergstro m 1988 , 1996) . O n it s northern an d easter n flanks , th e sub-Cambria n peneplain ca n b e trace d u p t o it s summit s (see

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Fig. 16) . This surfac e i s extremely fla t an d wa s formed durin g a period o f major denudation over all o f th e Balti c Shiel d i n lat e Proterozoi c tim e (e.g. Hogbo m 1910 ; Hogbo m & Ahlstro m 1924). I t i s stil l wel l preserve d i n part s o f eastern an d south-centra l Swede n (Rudber g 1954; Lidmar-Bergstro m 1996) . I n contrast , the sub-Cambrian peneplai n ha s bee n destroye d o n the souther n an d wester n flank s o f th e Sout h Swedish Dome, wher e th e Precambria n base ment was re-exposed during the warm and humid Mesozoic climat e (Lidmar-Bergstro m 1989 , 1995). Dee p weatherin g too k plac e an d thic k kaolinitic saprolite s wer e formed , and , afte r partial erosio n o f th e saprolites , a n undulating, hilly relief developed. This sub-Cretaceou s hill y relief ca n b e see n u p t o 125 m a.s.l . toda y a s a result o f preservatio n unde r a long-lastin g Cretaceous cove r (Lidmar-Bergstro m 1982 , 1996). Wher e th e sub-Cambria n peneplai n i s well preserved , i t mus t hav e bee n protecte d during Mesozoi c tim e b y a Palaeozoi c cover . Consequently, th e lates t uplif t o f th e Sout h Swedish Dom e wit h re-exposur e o f th e subCambrian peneplai n di d no t occu r unti l som e

Fig. 15 . Profile o f cover rocks an d basement topograph y an d across souther n Swede n fro m the Kattega t t o th e Baltic. A first phase o f Neogene uplif t an d erosio n le d t o the formatio n o f the Sout h Smalan d Peneplai n an d a second phas e t o the re-exposure o f the sub-Cretaceou s hill y relief (i.e . during mid-Miocene an d Pliocene time) . Location show n i n Fig. 2. SCP, Sub-Cambrian Peneplain . Modified fro m Kornfal t & Larsson (1987) , Britze & Japsen (1991) , Japse n & Langtoft e (1991a , 1991b) , Lykke-Anderse n (1991) , Lidmar-Bergstro m (1995 ) and Vejbaek (1997) .

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Fig. 16. Block diagrams illustrating the development of bedrock relief in southern Sweden between the Kattegat and the Baltic (see Figs 2 and 15) . Neogene uplift an d erosion is assumed to have been initiated i n two phase s (?mid-Miocene, Pliocene) . The fina l diagra m illustrates ho w th e Sout h Swedis h Dom e is composed of surfac e facets fro m widel y differen t periods . Surface feature s ar e exaggerate d an d Quaternar y deposit s are no t show n in th e fina l diagram . SCP, Sub-Cambria n Peneplain; SCr, Sub-Cretaceou s hilly relief ; SSP , Sout h Smalan d Peneplain. Modified fro m Nielse n & Japsen (1991), Frede n (1994) , Buchard t e t al. (1997) and Vejbae k (1997).

time durin g th e Cenozoi c perio d (Lidmar Bergstrom 1991) . The Sout h Smalan d Peneplai n i s a n almos t horizontal an d ver y fla t erosio n surfac e wit h only few, low , residua l hills , whic h extend s SW o f th e cresta l par t o f th e Sout h Swedis h Dome (Fig . 1 ) (Lidmar-Bergstro m 1988 , 1996). Th e peneplai n cut s of f th e incline d sub-Cretaceous hill y relief , an d therefor e developed durin g Cenozoi c tim e b y erosio n

by river s flowin g toward s th e sout h an d wes t down t o a bas e leve l tha t correspond s t o th e present leve l o f abou t 125 m a.s.l . (Figs 1 5 and 16) (Lidmar-Bergstro m 1982) . Th e formatio n of th e Sout h Smalan d Peneplai n destroye d th e sub-Cretaceous hill y relief . Th e peneplai n mus t have continue d acros s Cretaceou s rock s i n th e west wher e th e Kattega t i s foun d today . Towards th e NE , th e Sout h Smalan d Peneplai n was cu t int o th e sub-Cambria n peneplain ,

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Fig 16. - contiunued

leaving south - an d west-facin g erosiona l scarp s and hig h residua l hill s (150m ) i n th e basemen t surface. In th e cresta l par t o f th e Sout h Swedis h Dome, severa l occurrence s o f weatherin g products suggeste d a s being o f Plio-Pleistocen e age hav e bee n describe d (Lidmar-Bergstro m et al 1997) . Thes e gravell y saprolite s diffe r

markedly fro m th e kaoliniti c saprolite s associ ated wit h th e pre-Cenozoi c weathering . Th e stripping o f thes e weatherin g mantle s appear s to b e responsibl e fo r muc h o f th e hill y relie f incised i n th e uplifte d sub-Cambria n peneplain . Thus th e erosio n o f th e las t remnant s o f Palaeozoic rock s ma y no t hav e occurre d unti l Neogene time .

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Timing of the Cenozoic uplift of the South Swedish Dome We propos e a ne w theor y fo r th e Cenozoi c development o f southern Scandinavi a base d o n a combination o f thes e observation s fro m th e exposed basemen t i n souther n Swede n an d th e geological evidence from th e sedimentar y cove r in th e adjacen t area s i n easter n Denmark , e.g . Sjaslland, wher e Paleocen e sediment s ar e preserved an d a Cenozoic cove r o f c . 500 m ha s been remove d (Figs 1 and 13) . Neogene uplif t of the whol e o f souther n Scandinavia , centre d around th e Sout h Swedis h Dome , ma y accoun t for th e erosio n o f the missin g Cenozoi c cove r i n eastern Denmark , an d b e i n agreemen t wit h th e geomorphological evidenc e i n Sweden . The formatio n o f th e Sout h Swedis h Dom e started i n Late Palaeozoic tim e with continuation in Mesozoic tim e (Cederbom 2001 ; Fig. 16) . The present for m aros e after the Neogene uplif t wit h subsequent formatio n o f a ne w erosio n surfac e (the Sout h Smalan d Peneplain ) an d re-exposur e of sub-Cambrian an d sub-Cretaceous surfaces . In Denmark, th e Neogen e uplif t o f th e Sout h Swedish Dom e le d t o remova l o f Cenozoi c cover rock s i n large part s o f the territory , wher e the Skagerrak-Kattega t Platfor m i n particula r suffered dee p erosio n (Fig s 1 and 13) . A Neogene uplif t o f the Sout h Swedis h Dom e is i n agreemen t wit h interpretatio n o f fission track data, whic h indicat e th e remova l o f abou t 650m o f Uppe r Cretaceous-Paleogen e sedi ments fro m S W Sweden an d of about 1000 m of such sediments from SE Sweden during Neogene time (Cederbo m 2002) . A n Uppe r Cretaceous Palaeogene cove r ove r larg e part s o f souther n Sweden i s in agreement wit h evidence fo r a vast Cretaceous cove r i n sout h Swede n (Lidmar Bergstrom 1982 , 1995 ) an d wit h th e repor t o f redeposited Eocen e diatom s i n lakes i n Smaland at abou t 200 m a.s.l . (Cleve-Eule r 1941) . Th e suggested Plio-Pleistocen e age of the weath ering product s foun d o n th e cresta l par t o f th e South Swedis h Dom e is consistent wit h Neogen e uplift o f th e Dom e (Lidmar-Bergstro m e l al 1997). The re-exposur e o f th e sub-Cretaceou s hill y relief a t altitude s belo w 125 m a.s.l . occurre d after th e formatio n o f th e Sout h Smalan d Peneplain, whe n th e remainin g Cretaceou s cover wa s eroded . Consequently , th e Neogen e rise o f th e Sout h Swedis h Dom e mus t hav e taken plac e i n tw o phases : a ?Miocen e an d Pliocene phas e woul d b e i n agreemen t wit h the erosiona l patter n i n Denmar k (se e th e section 'Timin g o f maximu m buria l an d subsequent erosion' .

A ris e o f th e Sout h Swedis h Dom e durin g Neogene tim e i s als o consisten t wit h th e clockwise shif t i n sedimen t transpor t direction s that i s observed i n the NE Nort h Se a Basin. The transport direction s change d fro m southward s in Oligocene-Early Miocen e tim e t o southwest wards in mid- an d lat e Miocen e tim e an d finally westwards i n Pliocen e tim e (Fig . 1 ) (Clause n et al . 1999) . Th e Oligocen e transpor t directio n from th e nort h i n th e easter n Nort h Se a Basi n corresponds t o a Lat e Paleogen e uplif t phas e o f the Souther n Scande s (Lidmar-Bergstro m e t al . 1997, 2000 ; Clause n et al. 2000) . The lat e Cenozoi c sediment s foun d i n S W Denmark ar e thus expected t o reflect th e erosio n of Palaeozoic-Paleogene cover rocks in Sweden as well as the re-exposure o f a Mesozoic surfac e on Precambria n rock s an d th e subsequen t formation o f th e Sout h Smalan d Peneplain . Several observation s agre e wit h thes e sugges tions, fo r example , th e occurrenc e o f silicifie d Lower Palaeozoi c fossil s i n Miocen e san d i n Jylland (Spjeldnae s 1975) , an d th e occurrenc e o f gibbsite in the Gram Formation, whic h indicates erosion o f re-expose d tropica l soi l durin g lat e Miocene tim e an d a shor t fluvia l transpor t (Rasmussen & Larsen 1989) . Larse n & Dinesen (1959) foun d a n increasing content of amphibole in Miocen e sediment s i n th e Danis h area , an d suggested tha t th e considerabl e amoun t o f immature weatherin g material wa s derive d fro m exposed basemen t an d cove r rock s o n th e Scandinavian Shield. Discussion and conclusions Using a combinatio n o f basi n modellin g an d analysis of sonic data from differen t stratigraphi c levels to estimate erosion allow s identification of anomalous values and extends the areal coverage relative t o tha t achieve d b y th e applicatio n o f a single method. We estimate erosion in the eastern North Se a Basi n t o b e o n averag e 200-60 0 m lower tha n suggeste d i n previou s studies . I n particular, w e fin d erosio n estimate s base d o n Chalk velocitie s t o b e i n goo d agreemen t wit h estimates base d o n basi n modelling , an d thi s i s interpreted a s bein g du e t o th e homogeneou s composition o f th e Chalk , the larg e thicknes s of the Chalk section over which the mean velocity is calculated i n mos t wells , an d th e stres s dependence o f Chal k compaction . Basi n model ling predict s lo w hea t flo w durin g Oligocene Pliocene time , t o accoun t fo r observe d stee p vitrinite reflectanc e gradient s an d present-da y temperatures i n central an d northern Jylland . We suggest thi s to be partly due t o transient thermal effects induce d b y depositio n followe d b y

NEOCENE UPLIFT AND EROSION OF SOUTHERN SCANDINAVI A

erosion, bu t furthe r studie s o f thi s matte r ar e needed. Comparison o f results fro m differen t method s indicates tha t erosio n i s overestimate d whe n based o n sonic data from Lowe r Jurassic shal e in NE Denmark ; thi s coul d b e du e t o variations i n lithology, bu t furthe r studie s ar e neede d t o full y understand thes e variations . I t i s conclude d tha t maximum buria l of the Mesozoi c successio n throughout th e regio n occurre d befor e Neogen e erosion wher e Paleogene-Neogen e strat a ar e preserved in large parts. A previous suggestion of deep erosio n in the Sorgenfrei-Tornquis t Zon e during the Late Cretaceous-Paleogene inversion is rejected. The thickness o f the missin g sectio n remove d by lat e Cenozoi c erosio n increase s fro m th e eastern Nort h Se a toward s th e Norwegia n an d Swedish coasts . W e estimat e th e erosio n t o b e c. 500 m i n a broa d zon e acros s Denmar k from N W t o SE . Thi s zon e largel y conform s to th e are a wher e Paleocene-Miocen e deposit s subcrop th e Quaternary , an d th e erode d sedi ments mus t thu s hav e bee n o f Paleocen e Miocene age . Erosio n decrease s toward s zero in the wester n an d souther n par t o f th e Danis h North Sea . Th e ag e o f th e remove d sediment s must be progressivel y younge r in this direction, as for th e ag e o f the Quaternar y subcrop. To the north, th e missin g uppe r Cretaceous-Dania n section reache s c . 1000 m o n an d alon g th e Skagerrak-Kattegat Platform . Th e deepe r ero sion o n an d alon g th e Skagerrak-Kattega t Platform relativ e t o th e Danis h Basin cannot be explained b y glacia l erosio n o r by a drop i n sea level, an d thi s provide s a furthe r argumen t fo r tectonic uplif t o f th e Platfor m durin g Neogen e time. Finally , w e hav e demonstrate d tha t th e predicted thicknes s o f th e missin g strat a corresponds t o a likel y Cenozoi c depositiona l history. The patter n o f lat e Cenozoi c erosio n i n Denmark agree s wit h a Neogene uplif t o f south Norway centre d aroun d th e Souther n Scande s and o f souther n Scandinavi a centred aroun d th e South Swedis h Dome . Th e formatio n o f th e South Swedis h Dom e starte d in Late Palaeozoi c time, bu t th e presen t for m aros e afte r Neogen e uplift wit h subsequen t formatio n o f th e Sout h Smaland Peneplai n an d re-exposur e o f sub Cambrian an d sub-Cretaceou s surfaces . The re exposure o f th e sub-Cretaceou s hill y relie f a t altitudes belo w 125 m a.s.l . occurre d afte r th e formation of the South Smaland Peneplain, when the remainin g Cretaceou s cove r wa s eroded . Consequently, th e Neogen e ris e o f th e Sout h Swedish Dom e mus t hav e take n plac e i n tw o phases. We suggest that these phases correspon d

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to the erosiona l episode s o f ?Miocen e an d PlioPleistocene ag e tha t affecte d th e sedimentar y cover i n Denmark . Th e area l exten t o f th e ?Miocene episod e appear s restricte d t o southern Sweden an d th e Skagerrak-Kattega t Platform , whereas th e effect s o f th e late r episod e reac h from th e Balti c an d fa r int o th e Nort h Sea . A Neogene ris e o f the Dome i s consistent wit h the clockwise shif t i n sedimen t transpor t direction s observed i n th e N E Nort h Se a Basin , changin g from southward s i n Oligocene-Earl y Miocen e time to southwestwards in mid- and late Miocen e time an d finall y westward s i n Pliocen e tim e (Clausen e t al. 1999) . Th e Oligocen e transpor t direction fro m th e nort h correspond s t o a Lat e Paleogene uplif t phas e o f the Souther n Scandes . Apart fro m erosio n induce d b y tectoni c uplift , glacial erosio n wa s also important, a s witnessed by th e larg e volume o f Pleistocene sediment s in the central North Sea Basin (e.g. Japsen 1998) . A drop in sea level durin g late Cenozoic tim e ma y also hav e increase d erosio n (e.g . Nielse n e t al . 2001). The Neogen e uplif t alon g th e easter n margi n of the North Sea Basin and of the South Swedish Dome fit s int o th e genera l patter n o f lat e Cenozoic vertica l movement s aroun d th e Nort h Atlantic (Japse n & Chalmer s 2000) . A mode l explaining these phenomena must be constrained by observation s o f th e magnitud e and timin g of uplift an d erosion based on independent methods as w e hav e aime d a t demonstratin g here, an d i t must thus separate the effects o f Paleogene uplif t of plat e boundarie s fro m thos e o f Neogen e intraplate uplift . Appendix: Normal velocity-depth trends for homogeneou s formation s A revised normal velocity-depth trend for the North Sea Chalk Japsen (1998 ) publishe d a norma l velocity depth tren d fo r th e Chal k Grou p base d o n a n analysis o f dat a fro m 84 5 well s throughou t the North Se a Basi n an d OD P data . Fo r th e shallowest part of the trend, no data representing normal compactio n wer e foun d fo r th e Chal k of the Nort h Se a Basin , s o soni c lo g dat a fro m Eocene to Recent ooze and chalk deposits from a stable platfor m wer e use d t o guid e th e tren d (Urmos e t al . 1993) . A t intermediat e depths , Japsen (1998 ) applie d qualitativ e argument s t o identify Nort h Se a dat a representin g norma l compaction along the lower bound for velocity depth dat a fo r whic h th e effec t o f overcompac tion a s a result of erosion i s minimal. At greate r

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depths, dat a representin g norma l compactio n were identifie d along th e upper bound wher e th e effect o f undercompactio n a s a resul t o f overpressuring i s a minimum. Later, however , Japse n (2000 ) foun d additional geologica l constraint s t o refin e th e identification o f referenc e dat a a t intermediat e depths wher e th e influenc e o f erosio n an d overpressuring i s difficul t t o ascertain . Becaus e the soni c metho d identifie s deviation s fro m maximum burial , post-erosiona l reburia l o f a formation wil l reduc e it s observabl e buria l anomaly, e.g . a pre-Quaternar y erosio n o f 500m wil l b e maske d b y a subsequen t Quaternary reburia l o f 500 m (Eq . 2) . Thi s implies that , wher e th e Quaternar y sequenc e i s thick, eve n mino r deviation s fro m maximu m burial a s a resul t o f Neogen e erosio n ma y correspond to a substantial missing section. Deep erosion is , however , no t likel y wher e th e base Quaternary hiatu s i s minor , fo r example , wher e the Quaternar y sequenc e i s underlai n b y Neogene sediments . Normally compacte d Chal k is thus likely to be found i n areas wher e th e Quaternary sequenc e i s thick, Neogene deposits are present, and pressure is hydrostatic . Consequently , th e norma l velocity-depth tren d fo r th e Nort h Se a Chal k should follow the upper bound for data from such areas, wherea s dat a representin g undercompac tion a s a resul t o f overpressurin g shoul d plo t below th e tren d (Fig . 4) . A revised baselin e wa s thus define d b y Japse n (2000 ) usin g suc h maximum velocit y dat a fo r 90 0 < z < 1700m. Th e revise d tren d line s u p wit h th e maximum velocit y dat a use d t o defin e th e original tren d fo r z > 200 0 m, an d wit h a velocity a t the surfac e o f 1550ms" 1 :

The fourt h o f th e abov e segment s i s unchanged fro m th e origina l trend , i n whic h the uppe r thre e segment s wer e expresse d a s 1600 + z , 50 0 + 2z , an d 937. 5 + I.15z (Japse n 1998). Th e revise d tren d i s shifte d toward s shallower depth s b y a mea n o f 160 m fo r th e velocity interva l affecte d b y th e revisio n an d where Nort h Se a data ar e found , 210 0 < V < 4875ms" 1 ; th e maximu m shif t i s 210 m fo r 2920 < V< 3920ms" 1 .

The shif t toward s highe r velocitie s fo r th e revised baselin e result s i n a reductio n i n estimates o f erosio n b y u p t o 210m , an d a n increase i n estimate s o f overpressur e b y u p t o 2 MPa fo r dat a point s tha t plo t abov e an d below th e line , respectivel y (overcompactio n as a resul t o f overpressure : dZ B/100 = 210/100MPa-2MPa; se e Japse n 1998) . Th e increased overpressur e tha t th e revise d mode l predicts i s a n improvemen t relativ e t o th e original model , whic h explaine d onl y 80 % o f the observe d overpressur e i n th e Chal k fo r 5 2 wells in the central North Sea located away from diapirs an d wher e th e overpressur e exceede d 4 MPa (Japsen 1998) . The corresponding percentage base d o n th e revise d baselin e is 91% . Thi s improvement i s particularl y clear fo r dat a fro m relatively shallo w dept h o r moderat e over pressure, e.g . th e Danis h Da n fiel d wher e overpressure i s 7. 3 MPa, an d fo r whic h th e overpressure predictio n fro m velocit y dat a ha s been increase d fro m 4. 3 t o 6.5 MPa. A baseline for marine shale of the Lower Jurassic F-I Member Japsen (2000) formulated a constrained baseline, V^h, fo r marin e shal e dominate d b y smectite illite based on velocity-depth data for the Lower Jurassic F- I Membe r fro m 3 1 Danis h well s o f which 2 8 have data for th e Chal k (Fig. 6):

The baselin e wa s reconstructe d b y correctin g present formatio n depth s fo r th e effec t o f lat e Cenozoic erosio n a s estimated from th e velocity of th e overlyin g Chalk i n thes e well s relative to the revise d Chal k tren d (Eq . 3) . Th e correcte d depths correspond t o th e buria l o f th e formation before erosio n whe n th e sediment s wer e a t maximum buria l a t mor e location s tha n today . The baselin e can thu s be traced mor e easil y i n a plot o f velocit y versus the correcte d depths , and is wel l define d a t grea t dept h wher e velocity depth dat a fo r normall y compacte d shal e a t maximum buria l ca n b e difficul t t o identif y (2.1 < z < 3.8km) . Thi s formulatio n i s a constrained, exponentia l transi t time-dept h model tha t fulfil s reasonabl e boundar y con ditions a t th e surfac e an d a t infinit e depth : V0 = 1550ms" 1 an d V^ = 5405ms" 1 ; maxi mum velocity-dept h gradien t 0.6ms"1 m"1 for z = 2.0km. Th e shal e tren d give n b y Eq . 4 corresponds closel y to baselines for marine shale found b y othe r worker s (Scherbau m 1982 ; Hansen 1996 ; se e discussion by Japse n 1999) .

NEOCENE UPLIFT AND EROSION O F SOUTHERN SCANDINAVI A

A baseline for the Lower Trias sic Bunter Shale Japsen (2000 ) formulate d a segmented , linea r baseline, V^ sh, fo r th e Lowe r Triassi c Bunte r Shale base d o n velocity-dept h dat a fro m 14 2 British an d Danis h well s o f whic h 9 1 hav e velocity-depth data for the Chalk (Fig . 7) :

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baselines fo r th e Bunte r Shal e an d th e Bunte r Sandstone ma y b e du e t o th e dominanc e o f quartz i n both formations, and , correspondingly , a hig h conten t o f quart z i n th e Fjerritsle v Formation clos e t o th e expose d Scandinavia n basement ma y explai n th e relativel y hig h velocity observe d i n well s i n tha t area . We wish to thank reviewers D. Issler an d J. Turner, as well a s U . Gregersen , A . Mathiesen , C . Pulvertaft , E. S . Rasmusse n an d O . Vejbae k (al l GEUS) fo r constructive comment s tha t improve d th e manuscrip t considerably.

References The trend indicates a pronounced variation of the velocity gradient with depth. The gradient is only 0.5 m s"1 m"1 in the upper part, and increases to 1.5 ms"1!!!"1 fo r depth s aroun d 2km , fro m where i t decrease s graduall y wit h dept h t o 0. 5 and the n 0.2 5 ms"1!!!"1. Th e declin e o f th e gradient wit h dept h reflect s tha t velocit y approaches a n upper limit . The Bunte r Shal e baselin e wa s reconstructe d by applying the same procedure as for the Lower Jurassic shal e b y correctin g presen t formatio n depths fo r th e effec t o f late Cenozoic erosio n a s estimated fro m Chal k velocities . Th e tren d wa s constructed t o predic t likel y value s nea r th e surface (V o — 1550ms"1), an d i s base d o n reference dat a wit h correcte d depth s fro m 160 0 to 5600m (Japse n 2000). Rather tha n proposin g a specifi c baselin e fo r the Lowe r Triassi c Bunte r Sandstone , Japse n (2000) found that the trend derived for the Bunter Shale wa s a reasonabl e approximatio n fo r a dataset fro m 13 3 Britis h an d Danis h well s o f which 87 have velocity-depth data for the Chalk (see Fig . 7) . Dat a fro m shal e ar e preferabl e t o those fro m sandston e i n studie s o f maximu m burial fo r severa l reasons : shal e porosit y i s les s affected b y diageneti c processes , shal e doe s no t act a s a n aquife r wit h th e consequen t porosit y variations, an d shale ma y be mor e unifor m with regard t o both grai n siz e an d mineralogy. Burial anomalies fo r the Bunte r Sandstone ca n thus be used t o plac e a n uppe r limi t o n estimate s o f erosion base d o n Bunter Shale data . The dominance of smectite-illite in the distal parts o f th e Fjerritsle v Formatio n (Lindgreen , pers. comm.) , and of kaoli n in the continenta l Bunter Shal e was suggeste d by Japsen (2000 ) t o be a possibl e explanatio n o f wh y baseline s fo r these tw o formations diverge, and why those fo r Bunter Shal e an d Bunte r Sandston e converg e a t depth. Alternatively , th e similarit y o f th e

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Cenozoic sequenc e developmen t i n th e easter n parts o f th e Nort h Sea . Journal o f th e Geological Society, London, 156 , 809-816 . CLAUSEN, O.R. , NIELSEN , O.B. , HUUSE , M . & MICHELSEN, O . 2000 . Geologica l indication s fo r Palaeogene uplif t i n th e easter n Nort h Se a Basin . Global an d Planetary Change, 24, 175-187 . CLEVE-EULER, A . 1941 . Alttertiar e Diatomee n un d Silicioglagellanten i m innere n Schwedens . Palaeontographica, 92A , 165-208 . FREDEN, C. 1994 . Geology. National Atlas of Sweden. SNA, Stockholm . GEMMER, L. , NIELSEN , S.B. , HUUSE , M . & LYKKE ANDERSEN, H . 2002 . Post-mid-Cretaceou s easter n North Se a evolutio n inferre d fro m 3- D dynami c modelling. Tectonophysics, 350 , 315-342 . HANSEN, S . 1996 . Quantificatio n o f ne t uplif t an d erosion o n the Norwegian Shel f south of 66°N fro m sonic transi t time s o f shale . Norsk Geologisk Tidsskrift, 76 , 245-252. HILLIS, R.R. 1995 . Quantificatio n of Tertiary exhumation i n th e Unite d Kingdo m souther n Nort h Se a using soni c velocit y data . AAPG Bulletin, 79 , 130-152. HOGBOM, A.G. 1910 . Precambrian geology of Sweden. Bulletin of the Geological Institution of the University o f Uppsala, 10 , 1-80 . HOGBOM, A.G . & AHLSTROM , N.G . 1924 . Ube r di e sub-kambrische Landflach e a m fuss e vo m Kinne kulle. Bulletin o f th e Geological Institution o f th e University o f Uppsala, 19 , 55-88. HUUSE, M., LYKKE-ANDERSEN , H . & MICHELSEN , O . 2001. Cenozoi c evolutio n o f th e easter n Danis h North Sea . Marine Geology, 111, 243-269. ILIFFE, I.E . & DAWSON , M.R . 1996 . Basi n modellin g history an d predictions . In : GLENNIE , K.W . & HURST, A . (eds ) A D 1995: NW Europe's Hydrocarbon Industry. Geologica l Society , London , 83-105. JAPSEN, P . 1993 . Influenc e o f litholog y an d Neogen e uplift o n seismi c velocitie s i n Denmark ; impli cations fo r dept h conversio n o f maps . AAPG Bulletin, 11, 194-211. JAPSEN, P . 1998 . Regiona l velocity-dept h anomalies . North Se a Chalk : a recor d o f overpressur e an d Neogene uplif t an d erosion . AAPG Bulletin, 82 , 2031-2074. JAPSEN, P . 1999 . Overpressure d Cenozoi c shal e mapped fro m velocit y anomalie s relativ e t o a baseline fo r marin e shale , Nort h Sea . Petroleum Geoscience, 5 , 321-336. JAPSEN, P . 2000. Investigatio n o f multi-phas e erosio n using reconstructe d shal e trend s base d o n soni c data, Sol e Pi t axis , Nort h Sea . Global an d Planetary Change, 24 , 189-210 . JAPSEN, P . & BIDSTRUP , T . 1999 . Quantificatio n of lat e Cenozoic erosio n i n Denmark base d o n soni c dat a and basi n modelling . Bulletin o f th e Geological Society o f Denmark, 46 , 79-99. JAPSEN, P . & CHALMERS , J.A . 2000 . Neogen e uplif t and tectonics aroun d the North Atlantic : overview. Global an d Planetary Change, 24 , 165-173 . JAPSEN, P . & LANGTOFTE , C . 199la . Geological Ma p of Denmark 1:400 000. The Danish Basin. 'Base

Chalk' and the Chalk Group, Two-way Traveltime and Depth, Thickness and Interval Velocity. Geological Surve y of Denmark Map Series , 29. JAPSEN, P . & LANGTOFTE , C . 1991b . Geological ma p of Denmark 1:400 000. The Danish Basin. 'Top Triassic' and the Jurassic-Lower Cretaceous, Two-way Traveltime and Depth, Thickness and Interval Velocity. Geologica l Surve y o f Denmar k Map Series , 30 . JENSEN, L.N . & SCHMIDT , B.J . 1992 . Lat e Tertiar y uplift an d erosion i n the Skagerrak area; magnitude and consequences . Norsk Geologisk Tidsskrift, 12 . 275-279. JENSEN, L.N . & SCHMIDT , B.J . 1993 . Neogen e uplif t and erosion offshore South Norway; magnitude and consequences fo r hydrocarbo n exploratio n i n th e Farsund Basin . In : SPENCER , A.M . (ed. ) Generation, Accumulation, and Production of Europe's Hydrocarbons; III. Springer , Berlin, 79-88. KNUDSEN, K.L . 1995 . Kvartaeret . In : NIELSEN , O.B . (ed.) Danmarks geologi fra Kridt ti l i dag. Aarhus Universitet, Aarhus, 247-269. KONRADI, P. 1995. Foraminiferal biostratigraphy of th e post mid-Miocen e in tw o borehole s i n th e Danis h North Sea. In: MICHELSEN , O . (ed.) Proceedings o f the 2nd Symposium o n Marine Geology. Danmarks Geologisice Unders0gels e Series C, 12 , 101-112 . KORNFALT, K.A . & LARSSON , K . 1987 . Geological Maps and Cross-sections of Southern Sweden. Svensk Karnbranslehanterin g (SKB ) Technica l Report, 87-24 . LARSEN, G . & DINESEN , A . 1959 . Vejle Fjord Formationen ved Brejning. Danmark s Geologisk e Unders0gelse II . Raekke , 82 . LAURSEN, G.V . 1992 . Foraminifer a o f th e easter n North Sea . In : LAURSEN , G.V. , HEILMANN CLAUSEN, C . & THOMSEN , E . (eds ) Cenozoic Biostratigraphy of the Eastern North Sea based on Foraminifera, Dinoflagellates, and Calcareous Nannofossils. Geologis k Institut , Aarhus . 1-68. LlBORIUSSEN, J. , ASHTON , P . & TYGESEN . T . 1987 .

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NEOCENE UPLIFT AND EROSION O F SOUTHERN SCANDINAVI A LIDMAR-BERGSTROM, K . 1999 . Uplif t historie s revealed b y landform s o f th e Scandinavia n domes. In : SMITH , B.J. , WHALLEY , W.B . & WARKE, P.A . (eds ) Uplift, Erosion an d Stability: Perspectives on Long-term Landscape Development. Geologica l Society , London , Specia l Publications, 162 , 85-91. LIDMAR-BERGSTROM, K. , OLLIER , C.D . & SULEBAK , J.C. 2000. Landforms and uplift history of southern Norway. Global an d Planetary Change, 24 , 211-231. LIDMAR-BERGSTROM, K. , OLSSON , S . & OLVMO , M . 1997. Palaeosurface s an d associate d saprolite s i n southern Sweden . In : WIDDOWSON , M . (ed. ) Palaeosurfaces; Recognition, Reconstruction and Palaeoenvironmental Interpretation. Geologica l Society, London , Specia l Publications , 120 , 95-124. LINDGREEN, H . 1991 . Elementa l an d structura l changes i n illite/smectit e mixed-laye r cla y minerals durin g diagenesi s i n Kimmeridgian-Volgia n (-Ryazanian) clay s i n th e Centra l Trough , Nort h Sea an d th e Norwegian-Danis h Basin . Bulletin o f the Geological Society o f Denmark, 39 , 1-82 . LYKKE-ANDERSEN, H . 1991 . Nogl e hovedetrae k a f Kattegats kvartaergeolgi—fore!0big e resultate r a f en seismis k undersjagels e 1988-1991 . Dansk Geologisk Forening, Arsskriftfor 1990-91, 57-65. MARIE, J.P.P . 1975 . Rotliegende s stratigraph y an d diagenesis. In : WOODLAND , A.W . (ed.) Petroleum and the Continental Shelf of North-west Europe. Applied Science , London , 205-211 . MICHELSEN, O . 1989 . Revision o f th e Jurassic Lithostratigraphy of the Danish Subbasin. Geological Surve y of Denmark Serie s A , 24. MICHELSEN, O . & NIELSEN , L.H . 1993 . Structura l development o f th e Fennoscandia n borde r zone , offshore Denmark . Marine and Petroleum Geology, 10, 124-134 . NIELSEN, L.H . & JAPSEN , P . 1991 . Deep Wells i n Denmark 1935-1990. Litho strati graphic Subdivision. Geologica l Surve y o f Denmar k Serie s A, 31. NIELSEN, S.B. , PAULSEN, G.E. , HANSEN, D.L . et al. 2002. Paleocen e initiatio n o f Cenozoi c uplif t i n Norway. In : DORE , A.G. , CARTWRIGHT , J.A. , STOKER, M.S. , TURNER, J.P . & WHITE , N . (eds) Exhumation of the North Sea Margin: Timing, Mechanisms and Implications for Petroleum Exploration. Geologica l Society , London , Specia l Publications, 196 , 45-65. PEDERSEN, G.K. 1985. Thin, fine-graine d stor m layers in a mudd y shel f sequence : a n exampl e fro m th e Lower Jurassi c i n th e Stenlill e 1 well, Denmark . Journal o f th e Geological Society, London, 142 , 357-374. PEULVAST, J.-P . 1985 . Post-orogenic morphotectoni c evolution o f th e Scandinavia n Caledonides durin g the Mesozoi c an d Cenozoic . In : GEE , D.G. & STURT, B.A . (eds ) Th e Caledonide Orogen — Scandinavia an d Related Areas. Wiley, Chichester, 979-996. RASMUSSEN, E.S. & LARSEN , O. 1989 . Mineralogi o g geokemi afdet 0vre Mioccene Gram ler. Danmarks Geologiske Unders0gelse , Seri e D, 7.

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CLOETINGH, S . 1995 . Meso-Cenozoi c morpho tectonic evolutio n o f souther n Norway : Neogen e domal uplif t inferre d fro m apatit e fissio n trac k thermochronology. Tectonics, 14, 700-714. RUDBERG, A . 1954 . Vasterbotten s berggrunds morfologi. Geographica, 25 , 1 -457. SCHERBAUM, F . 1982 . Seismic velocitie s i n sedimen tary rocks ; indicator s o f subsidenc e an d uplift . Geologische Rundschau, 71 , 519-536 . SHIPBOARD SCIENTIFI C PARTY , e t al . 1991 . Site 807. In: KROENKE , L.W. , BERGER, W.H . & JANACEK , T.R. (eds ) Proceedings o f th e Ocean Drilling Program, Initial Reports, 130. Ocea n Drillin g Program, Colleg e Station , TX, 369-493 . SPJELDN^S, N . 1975 . Palaeogeograph y an d facie s distribution i n th e Tertiar y o f Denmar k an d surrounding areas . Norges Geologiske Unders0gelse Bulletin, 316, 289-311 . URMOS, J. , WILKENS , R.H. , BASSINOT, F. , LYLE , M. , MARSTERS, J.C. , MAYER, L.A . & MOSHER , D.C. 1993. Laboratory and well-log velocity and density measurements fro m th e Onton g Java Plateau: ne w in-situ correction s t o laborator y dat a fo r pelagi c carbonates. In : BERGER , W.H. , KROENKE, L.W., MAYER, L.A . & JANECEK , T.R. (eds) Proceedings of the Ocean Drilling Program, Scientific Results, 130. Ocean Drillin g Program, College Station , TX, 607-622. VAN OLPHEN , H . 1966 . Collaps e o f potassiu m montmorillonite clay s upo n heating—'potassiu m fixation'. In : BAILEY , S.W . (ed.) Clays an d Clay Mineralogy. Pergamon , Oxford , 393-405 . VEJB^K, O.V . 1997 . Dyb e strukture r i dansk e sedimentaere bassiner. Geologisk Tidsskrift, 1997/4 , 1-31. VEJB^K, O.V. & BRITZE , P . 1994 . Geological Ma p o f Denmark 1:750 000 Top Pre-Zechstein (Two-way Traveltime an d Depth). Geologica l Surve y o f Denmark Map Series , 45 . VIK, E . & HERMANRUD , C . 1993 . Transient therma l effects o f rapi d subsidenc e i n th e Haltenbanke n area. In: DORE , A.G., AUGUSTSON, J.H., HERMAN RUD, C. , STEWART , D.J . & SYLTA , O . (eds ) Basin Modelling: Advances and Applications. Norwegia n Petroleum Societ y Specia l Publication , 3 , 107-117. YUKLER, M.A . 1978 . One-dimensiona l mode l t o simulate geologic , hydrodynami c an d thermo dynamic developmen t o f a sedimentar y basin . Geologische Rundschau, 67 , 960-979. ZAGWIJN, W.H . 1989 . Th e Netherland s durin g th e Tertiary and the Quaternary: a case story of coastal lowland erosion . Geologic e n Mijnbouw, 68 , 107-120.

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Cenozoic uplif t and denudation o f southern Norway : insights fro m the North Se a Basi n MADS HUUSE 1'2 Department of Earth Sciences, University of Aarhus, Aarhus, Denmark 2 Present address: Department of Earth Sciences, Cardiff University, Cardiff CF 10 3 YE, UK (e-mail: m.huuse@ abdn.ac.uk) 1

Abstract: Th e Cenozoi c evolutio n o f th e Nort h Se a Basi n i s described , drawin g o n subsurface dat a an d a serie s o f palaeogeographica l map s compile d fro m a variet y o f published studies , mainl y emphasizin g the developmen t o f the easter n par t o f the basin . A model that accounts for the sedimentation history o f the North Sea Basin an d the topograph y (including maximu m an d mea n surfac e elevation ) o f souther n Norwa y i s proposed . Th e model involve s regiona l plume-relate d uplif t o f a n initia l low-elevatio n peneplai n i n earl y Paleogene tim e followe d b y repeate d episode s o f climati c deterioratio n an d eustati c fall , most notably a t the Eocene-Oligocene transition, i n late Mid-Miocene time, and eventuall y culminating wit h th e developmen t o f ful l glacia l condition s i n souther n Norwa y i n Plio Pleistocene time. Thes e episodes correspond to periods of accelerated sediment suppl y fro m southern Norwa y tha t reflec t increased rates of incision (dissection ) o f the sourc e area. It is argued tha t th e present-da y elevatio n o f > 2 km o f mountai n peak s i n souther n Norwa y adjacent t o deep valleys an d fjords coul d hav e been cause d b y isostatic uplift in response to dissection o f a high-elevatio n peneplain . Henc e i t ma y no t b e necessar y t o invok e lat e Cenozoic tectonic uplift event s t o explain th e present-day topograph y o f souther n Norway .

This pape r describe s th e Cenozoi c infil l histor y of th e Nort h Se a Basi n an d it s implication s fo r understanding ho w th e topograph y o f souther n Norway wa s created . A t present , th e Nort h Se a Basin i s fille d t o it s brim . I t comprise s th e shallow Nort h Sea , th e low-relie f area s o f Denmark, norther n Poland , norther n Germany , the Netherlands and SE England. This low-relief area i s surrounde d b y th e topographicall y hig h areas of southern Norway to the NE, Scotland to the N W and the Centra l Europea n Massi f t o the south (Fig. 1) . During Cenozoic time the margins of the North Sea Basin became exhumed whereas the centr e o f the basin subside d more tha n 3 km (Fig. 2) . Th e subsidenc e histor y o f th e basi n i s fairly wel l constraine d b y th e preserve d sedi mentary record , althoug h th e details ar e stil l the subject of debate (McKenzie 1978 ; Nielsen et al. 1986; Vinke n 1988 ; Cloeting h e t al 1990 , 1992 ; Ziegler 1990 ; Jo y 1992 ; Gallowa y e t a l 1993 ; Jordt e t al 1995 ; Liu & Galloway 1997 ; Japse n 1998; Michelse n e t a l 1998 ; Huus e 2002 ; Nielsen e t al 2002) . I n contrast, the magnitude, mechanisms an d exac t timin g o f uplif t an d subsequent denudatio n of the basin margins and the hinterland mountains has remained enigmatic to thi s da y (Torsk e 1972 ; Dor e 1992 ; Jense n & Michelsen 1992 ; Jense n & Schmid t 1993 ; Jord t

et al 1995 ; Rohrma n e t al 1995 ; Hanse n 1996 ; Riis 1996 ; Solhei m e t a l 1996 ; Japse n 1998 ; Michelsen et al. 1998 ; Dor e et al 1999 ; Lidmar Bergstrom 1999 ; Chalmer s & Cloeting h 2000 ; N0ttvedt 2000 ; Huuse , 2002). I t i s ofte n argue d that uplif t an d denudatio n occurre d i n tw o phases, on e synrif t an d on e post-rif t (e.g . Rii s & Fjeldskaa r 1992 ; Eyle s 1996 ; Rii s 1996 ; Rohrman & va n de r Bee k 1996 ; Lidmar Bergstrom et al 2000) , althoug h the two phases can b e difficul t t o separat e (Japse n & Chalmer s 2000). The magnitud e of uplif t an d denudatio n ma y be estimate d fro m proxie s suc h a s geothermo metry (fissio n tracks , vitrinit e reflectance , etc.), compaction estimates , structura l trends , sedi mentary geometrie s an d geomorphology . Thes e methods are , however , al l associated wit h rather large uncertainties , a s indicate d b y th e discre pancies observe d whe n comparin g uplif t esti mates derive d usin g different method s (compar e Jensen & Schmid t 1993 ; Japse n & Bidstru p 1999; Huuse 2002). It is even more speculative to assess th e mechanism s an d th e exac t timin g of uplift, an d th e amoun t and timin g of denudation that followed . The uplif t an d denudation history of souther n Norwa y i s a n intriguin g problem i n its own right, but is also o f significan t interes t to

From: DORE , A.G. , CARTWRIGHT , J.A. , STOKER , M.S., TURNER , J. R & WHITE , N . 2002 . Exhumation o f th e North Atlantic Margin: Timing, Mechanisms an d Implications for Petroleum Exploration. Geologica l Society , London, Special Publications, 196, 209-233 . 0305-8719/027$ 15.00 © The Geological Societ y of London 2002 .

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Fig. 1 . Present-day topography and bathymetry o f NW Europe. The study areas of Jordt el al. (1995) (long dash) , and o f Michelsen e t al. (1998 ) (shor t dash ) an d location s o f Figs 2 , 4. 5 and 9 are show n fo r reference .

the oi l industry , a s uplif t an d denudatio n ma y have bot h positiv e an d negativ e effect s o n th e hydrocarbon potentia l o f th e N W Europea n margin (e.g. Sales 1992 ; Jensen & Schmidt 1993 ; Dore & Jensen 1996 ; Dor e e t al. 1997 , 1999) . Uplift an d denudation Definition When discussin g the uplif t histor y of any area of the Eart h i t i s extremel y importan t t o properl y

define what is actually meant by the term 'uplift' , i.e. i s i t regiona l surfac e uplif t o r loca l uplif t o f mountain peak s ('uplif t o f rocks') . I t i s equally important t o properl y defin e th e referenc e level to which uplift i s measured. In this paper the term 'surface uplift ' i s use d t o describ e uplif t o f th e Earth's surfac e with respect t o th e geoi d (mea n global se a level ) average d ove r a n are a o f c. 10 4km2, a s thi s i s th e relevan t scal e fo r studying vertica l movement s o f th e lithospher e (England & Molnar 1990) . The local inversion of former norma l faults i s thus not considered here.

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Fig. 2. Depth to the Upper Cretaceous-Danian limeston e (Ziegle r 1990 ; Japsen 1998) . The stud y areas of Jordt et al (1995 ) (lon g dash) , o f Michelse n e t al (1998 ) (shor t dash ) an d o f Huus e (1999 ) (continuou s line) an d locations o f Figs 4 , 5, 6 and 8 are show n for reference.

Surface uplif t ove r larg e area s require s upwar d displacement o f rocks wit h respect t o the geoid , i.e. wor k agains t gravity , an d thu s require s a n active tectoni c mechanis m o f larg e magnitud e (England & Molnar 1990; Summerfield & Brown 1998; Nielse n e t al., 2002) . Whe n discussin g relatively smal l amount s (som e hundred s o f metres) o f surfac e uplift , i t i s importan t t o consider eustati c change s (Englan d & Molna r 1990; Huuse , 2002). I n particular, when discussing Cenozoic uplift events, the long-term eustatic fall sinc e mid-Eocene tim e of c. 200m (e.g. Haq et al. 1987 ) corresponds t o (nontectonic) surfac e uplift o f th e sam e magnitude . Finally , i t i s important t o not e tha t 'surfac e uplift' , 'uplif t o f rocks' an d 'denudation ' ar e relate d i n th e

following wa y (Englan d & Molna r 1990) : surface uplif t = uplif t o f rock s — denudation. Hence, th e implici t assumptio n ofte n mad e that roc k uplif t i s equa l t o surfac e uplif t i s true only o n th e rar e occasion s whe n denudatio n i s zero (Englan d & Molna r 1990) . Th e ter m 'denudation' i s use d her e instea d o f 'exhuma tion' as the latter is generally used to describe the re-exposure o f a burie d lan d surfac e (se e Summerfield & Brow n 1998) . I t shoul d b e noted als o tha t surfac e uplif t i n itsel f doe s no t cause increase d denudation , a s i t i s th e loca l relief rathe r tha n averag e elevatio n tha t govern s denudation rate s (Ahner t 1970 ; Summerfiel d 1991). Moreover , denudatio n rate s wil l usually be les s tha n rates o f surfac e uplift, a s otherwise

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there would b e no mountains. Indeed , it has been shown tha t i n som e case s th e denudationa l response ma y la g severa l ten s o f millio n year s behind a n uplif t even t (Summerfiel d & Brow n 1998). Tectonic mechanisms A larg e numbe r o f activ e tectoni c mechanism s have bee n propose d t o explai n th e denudatio n record an d present-da y topograph y o f th e Nort h Atlantic margins , i n particula r th e Norwegia n and Scottis h highland s an d th e eas t Greenlan d margin (e.g . Dor e 1992 ; Japse n & Chalmer s 2000). Crusta l shortenin g i s a well-know n mechanism fo r generatin g surfac e uplift , bu t the amoun t o f shortenin g observe d i n N W Europe i s nowher e nea r enoug h t o explai n th e amount o f exhumatio n inferre d fro m vitrinit e reflectance profile s an d fission-trac k analyse s (Brodie & Whit e 1995) . Crusta l shortenin g i s observed i n th e shap e o f elongat e dome s alon g the N W Europea n Atlanti c margi n (Dor e & Lundin 1996 ; Boldree l & Anderse n 1998 ; Dor e et al. 1999) , bu t thes e dome s ar e significantl y smaller than , for example , th e sout h Norwegia n dome, whic h must be explained b y other means . Other activ e tectoni c mechanism s capabl e o f generating regiona l surfac e uplif t includ e increased buoyanc y b y heate d lithospher e (White 1992) , mantl e upwellin g (Rohrma n & van der Beek 1996 ; Nadi n et al 1997) , magmati c underplating of the lower crust (McKenzie 1984 ; White & Lovel l 1997) , an d delaminatio n o f th e lower lithospher e (Bir d 1979 ; Molna r e t al . 1993). These mechanisms coul d all be related t o the impingemen t o f th e Icelan d Plum e ont o th e North Atlanti c an d subsequen t Nort h Atlanti c rifting a t th e Paleocene-Eocen e transitio n (c. 5 4 Ma). Thu s the y ar e al l plausibl e agent s o f surface uplif t (se e discussio n b y Nielse n e t al . 2002). However, i t has been argued tha t no singl e mechanism seem s capabl e o f explainin g al l o f the observe d (an d inferred ) 'up s an d downs ' around and within the North Atlantic (Dore et al. 1999). In lat e Cenozoi c tim e th e N W Europea n margin wa s fa r remove d fro m th e Icelan d Hotspot (Lawve r & Mulle r 1994 ; Dor e e t al . 1999). Becaus e o f th e increase d distanc e t o th e Iceland Plum e an d th e lac k o f extensiv e crusta l shortening, activ e tectoni c surfac e uplif t o f lat e Cenozoic ag e i s difficul t t o invok e withou t jumping t o rathe r exoti c explanations . Th e various tectoni c mechanism s capabl e o f causin g early Paleogen e surfac e uplif t aroun d th e Nort h Atlantic ar e discusse d i n a companio n pape r (Nielsen e t al . 2002) , whic h argue s tha t earl y

Paleogene delamination o f the lithospher e i s the most plausibl e mechanism . Thi s mechanis m causes severa l hundre d metre s o f almos t instantaneous uplif t followe d b y deceleratin g uplift rate s throug h th e remainde r o f th e Cenozoic. To kee p th e mode l develope d i n th e presen t paper a s simpl e a s possibl e an d independen t o f the mechanis m of the tectonic uplift, i t is simply assumed tha t significan t (500-1000m ) surfac e uplift occurre d i n relatio n t o th e arriva l o f th e plume an d riftin g o f th e Nort h Atlanti c in early Paleogene time . Also , i t i s assume d tha t fo r th e case o f souther n Norwa y thi s uplif t wa s permanent, as opposed t o transient, uplift cause d by increase d buoyanc y of heated lithosphere. Isostatic uplift response to denudation Gravity dat a indicat e tha t th e mea n surfac e topography o f souther n Norwa y i s isostaticall y compensated a t depth (Balling 1980; Rohrman & van de r Bee k 1996) , suggestin g tha t Cenozoi c denudation wa s compensate d b y regiona l iso static uplif t o f th e crust . I t ha s bee n argue d that intense localize d denudatio n (dissection ) o f a n initially fla t topograph y o f hig h elevatio n could theoretically caus e mountai n peak s t o reac h elevations o f twic e o r mor e th e heigh t o f th e initial topograph y (see Molnar & England 1990 ; Gilchrest e t al . 1994) . I t shoul d be note d tha t in this cas e uplif t i s a consequenc e o f denudation and not vice versa. This is possible because of the regional isostatic response to dissection, which is governed b y th e isostati c respons e functio n (/) , which equals pc/pm, where pc is the density of the material erode d fro m th e to p o f th e crust , p m i s the densit y o f th e mantl e a t th e dept h o f compensation (p m ~ 3.3gem~~) , an d / i s th e amount o f isostatic uplift pe r uni t mea n dept h of dissection (Gilchres t e t al . 1994) . I n th e centra l parts o f souther n Norway , wher e th e remove d material is likely to have been mainly crystalline rock (p c ~ 2.7-3.0gem" 3 ), the compensation is relatively hig h (approachin g 0.8-0.9) . I n mar ginal areas , wher e a greate r proportio n o f th e rocks remove d ar e likel y t o hav e bee n o f sedimentary origi n (p c ~ 2.3-2.7gcm~~ ) i t i s somewhat lowe r (c . 0.7-0.8). Climatic and eustatic change The denudatio n histor y of souther n Norway an d the infil l pattern s o f th e Nort h Se a Basi n wer e modulated b y climati c an d eustati c change s (Spjeldnaes 1975 ; Dor e 1992 ; Jord t e t a l 1995 ; Eyles 1996 ; Solhei m et al. 1996 ; Michelsen et al.

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Fig. 3 . Cenozoic stratigraphy of the eastern North Sea Basin. Time scale according to Berggren et al. (1995). The composite 6 18 O curve is compiled fro m Mille r et al. (1987 , 1998) . Th e Mille r e t al (1987 ) curv e was shifte d - 0.25% c at 1 0 Ma and +0.20%c at 37 Ma to match values of the Miller et al. (1998) curve. Correlation of North Sea sequence s wit h th e Berggre n e t al . (1995 ) tim e scal e an d th e 8 18O curv e i s base d o n th e calcareou s nannofossil zonatio n o f Martin i (1971) . Th e Nort h Se a 5 18O curve o f Buchard t (1978) , adjuste d t o fi t earl y Eocene an d mid-Miocene peaks on the Miller curve, is shown for reference. *Episodes of marked increase in the supply of coarse clastic sediment s to the North Sea Basin. 1 , North Atlantic rifting an d volcanism; 2, major icesheet expansion (Lear et al. 2000). The mid-Paleocene event roughly coincides wit h the commencement of rift related uplif t o f area s borderin g th e Nort h Atlanti c rift . Th e lates t Eocene-earlies t Oligocene , th e lat e MidMiocene and the Plio-Pleistocene events all coincide with major coolin g events and eustatic lowerings as a result of increased continental ic e volume (Lea r et al. 2000). Foraminifera: NSP , Nort h Se a Planktic; NSB , Nort h Sea Benthic; NSA, Nort h Sea Agglutinated.

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1998; Clause n e t al 1999 , 2000; Huus e 2002 ; Nielsen e t al 2002) . Most researcher s agre e tha t eustasy peaked i n response t o relativel y larg e spreadin g rate s a t mid-ocean ridge s and greenhouse climate durin g mid-Cretaceous time . Th e long-ter m trend s o f global se a level (eustasy) may be estimated fro m spreading rates o f mid-ocean ridge s (e.g . Pitman 1978; Komin z 1984) , seismi c stratigraph y (Vail et a l 1977 ; Ha q e t a l 1987) , flexura l back stripping (Peka r & Mille r 1996 ; Steckler e t a l 1999) an d from stable isotopes (Fig. 3; 8 18O, Sr, Mg/Ca; Matthews 1984; Millers al 1987 , 1996 , 1998; Abre u & Anderson 1998 ; Lear et al 2000) . From thes e studie s a genera l consensu s ha s emerged tha t long-ter m globa l se a leve l ha s fallen som e 150-25 0 m durin g Cenozoi c time . This, of course, corresponds to a surface uplift of the sam e magnitud e relativ e t o sea level. A long-ter m tren d o f climati c coolin g starte d in lat e Mid-Eocen e time , bot h globall y (Savi n 1977; Wolf e 1978 ; Mille r e t a l 1987 , 1998; Zachos e t al 1992 ; Lear e t al 2000 ) an d in the North Sea region (Buchardt 1978; Collinson et al 1981). I n associatio n wit h th e chang e toward s a colder climate , th e continenta l ic e shee t o n Antarctica expanded , causin g a pronounce d eustatic lowering , culminatin g a t th e Eocene Oligocene transitio n (c . 3 4 Ma; Mille r e t a l 1998; Lea r e t a l 2000) . A simila r episod e o f cooling an d eustati c lowerin g occurre d onc e again i n lat e Mid-Miocen e time , followin g a n early t o mid-Miocen e war m perio d (Molna r & England 1990 ; Mille r e t a l 1998 ; Lear e t a l 2000), an d i n Plio-Pleistocen e time , eventuall y causing ful l glacia l condition s i n N W Europ e (Eyles 1996 ; Solhei m e t a l 1996 ; Lear e t a l 2000). With the exception of the Plio-Pleistocen e glaciations, the role of climate change has largely been neglecte d i n previou s studie s o f th e Nort h Atlantic margins. Thi s ma y be due t o the notion that climate a s a mechanism act s on a time scal e an orde r o f magnitud e shorte r tha n tha t o f tectonics (see Vail et al 1991 ) or simply because of lack of attention to Cenozoic climat e changes . However, i t i s importan t t o bea r i n min d tha t a major long-ter m chang e i n climat e ma y signifi cantly affec t denudatio n rate s (Summerfiel d & Brown 1998) . Major episode s of climatic cooling during the Cenozoic generall y corresponde d t o majo r icesheet expansions on Antarctica that caused major eustatic lowering s (Lea r e t a l 2000) . Thu s i t is possible tha t th e effect s o f majo r eustati c fall s and stepwis e climati c deterioratio n coul d caus e effects simila r t o thos e widel y attribute d t o regional surfac e uplift , i.e . accelerated denuda tion o f topograph y an d increase d sedimen t

supply t o adjacen t basin s (Donnell y 1982 ; Molnar & England 1990 ; Huuse 2002). Rationale From the above discussion it should be clear that the following factors should be considered when attempting t o accoun t fo r th e observe d topogra phy o f souther n Norwa y an d th e sedimentar y record o f the North Sea Basin: (1) plume-related surface uplif t alon g Atlanti c margin s i n earl y Paleogene time ; (2 ) episodic inversio n tectonic s driven by Alpine compression an d Atlantic ridge push; (3 ) stepwis e climati c deterioratio n sinc e mid-Eocene tim e a s documente d b y stabl e isotope record s (Fig . 3); (4 ) eustati c lowerin g of c. 250 m since mid-Cretaceous time (c. 200 m since mid-Eocen e time) ; (5 ) passiv e (flexural? ) isostatic respons e t o denudatio n and deposition. These ar e al l relativel y well-documente d phenomena, althoug h ther e ma y b e som e uncertainty abou t th e exac t mechanism s o f early Paleogen e uplif t an d th e magnitud e an d effects o f climati c deterioratio n an d eustati c fall (see Dor e e t al 1999 ; Nielsen et al 2002) . Another facto r tha t mus t b e addresse d i n a n account of the genesis of present-day topography is palaeo-topography , i.e. the topograph y o f th e so-called 'pre-uplif t peneplain ' (Stuevol d & Eldholm 1996 ) or 'palaei c surface ' o f Norwa y (Gjessing 1967 ; Lidmar-Bergstrom e t al 2000) . It is assumed here tha t the present-day elevation of th e palaeo-peneplai n roughl y coincide s wit h the summit envelope (see Dore 1992) . The purity of th e Lat e Cretaceous-Dania n chalk s o f th e North Se a Basi n indicate s tha t an y siliciclasti c source are a mus t have been clos e t o a peneplain by lat e Cretaceou s tim e (Hancoc k 1975) . However, th e presenc e o f uppe r Cretaceou s siliciclastic wedge s alon g th e Atlanti c margi n (Knott e t a l 1993 ; Dore e t a l 1999 ) demonstrates tha t ther e mus t hav e bee n som e topography abov e se a level . Thi s topograph y probably coincide d wit h th e Shetlan d Platfor m and the present-day Norwegian mainland . Som e topography wa s probabl y als o generate d alon g the lat e Cretaceou s an d earl y Paleogen e inver sion zone s (Ziegle r 1990 ; Gemmer e t al 2002) . The elevatio n an d relie f o f th e topograph y o f southern Norwa y an d Shetlan d i s difficul t t o constrain becaus e o f th e absenc e o f lat e Mesozoic an d Cenozoi c sediments , bu t a maximum elevatio n o f th e orde r o f a fe w hundred metre s above (palaeo-) se a level, gently sloping toward s th e shorelin e wit h onl y mino r local relie f seem s plausible . Th e shorelin e wa s probably clos e t o o r slightl y inboar d o f th e present shoreline s o f souther n Norway, whereas

UPLIFT AND DENUDATION O F SOUTHER N NORWA Y

all o f Denmar k an d souther n Swede n wa s submerged (Spjeldnae s 1975 ; Ziegle r 1990 ; Stuevold & Eldhol m 1996) . Thi s combinatio n of maximu m elevatio n an d shorelin e position s would correspond t o an average surfac e slope of about 0.1°, which does not seem unrealistic for a peneplain acros s a former mountai n range . Against thi s background , i t wil l b e assesse d whether th e interactio n o f th e above-mentione d mechanisms ca n accoun t fo r th e present-da y

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topography o f souther n Norwa y an d th e strati graphic record o f the adjacent basins. If this were the case, the n there would be no need t o invoke late Cenozoic tectoni c uplif t events . Conversely, if th e combinatio n o f the abov e factor s doe s no t support th e hypothesis , on e ma y begi n t o speculate abou t lat e Cenozoi c tectoni c events . The scenari o give n her e i s base d mainl y o n qualitative evidenc e an d simpl e isostati c calcu lations, an d shoul d therefor e b e regarde d a s a

Fig. 4 . Post-Dama n depocentre s i n th e easter n Nort h Se a superimpose d o n dept h contour s o f th e Uppe r Cretaceous-Danian limestone . Depocentre s wer e mappe d b y Bidstru p (1995 ) and Michelsen e t al. (1998).

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somewhat subjectiv e perception o f how present day topograph y an d sedimentar y thicknesse s o f the Nort h Se a are a ma y b e accounte d for . Th e importance o f globa l (allocyclic ) mechanism s (climate, eustasy ) is stressed, a s opposed t o local or regiona l (autocyclic ) mechanism s (e.g . tec tonics). O n th e basi s o f geodynami c modelling , Nielsen et al (2002 ) have provided a quantitative account o f a similar scenario . The ultimate test of any qualitative scenario of the development o f source area s and their basins would be to carry ou t mass-balanced palaeogeo graphical reconstruction s suc h as that performed for th e Mississippi catchmen t and the associate d depocentre i n th e Gul f o f Mexic o (Ha y e t al . 1989). Suc h a n approac h require s a detaile d integration o f regiona l denudatio n estimate s from onshor e area s wit h th e extensiv e offshor e database o f seismi c an d wel l dat a an d thus calls for internationa l co-operatio n an d integratio n of large regiona l databases. Cenozoic evolution of the North Sea Basin Database The Cenozoi c evolutio n o f th e Nort h Se a Basin has bee n piece d togethe r fro m a larg e databas e that cover s variou s part s o f th e basin . Seismi c data, wel l data , outcrops , structur e an d isopac h maps (Fig. 4; Nielsen et al 1986 ; Bidstrup 1995 ;

Jordt e t a l 1995 ; Michelse n e t a l 1995 , 1998 ; Joy 1996 ; S0rense n etal 1997 ; Huus e & Clausen 2001), seismi c an d sedimentary facies maps (Joy 1996; Mudg e & Buja k 1996 ; Danielse n e t a l 1997; S0rense n e t a l 1997) , pattern s o f clino form breakpoin t migratio n (Fig s 5 an d 6 ; S0rensen e t a l 1997 ; Clause n e t al 1999) , an d palaeogeographical compilation s (Graman n & Kockel 1988 ; Kocke l 1988 ; Ziegle r 1990 ) hav e been compile d t o yiel d palaeogeographica l (palaeobathymetric) map s o f th e entir e Nort h Sea Basi n (Fig . 7) . Huuse , 2002 ha s provide d a complete lis t o f references for eac h o f th e map s (Fig. 7b-g) an d an account of their compilation. Palaeogeographical development It i s generall y accepte d tha t th e overal l infil l o f the Nort h Se a Basin was dominated b y westerly source area s durin g Paleocene an d Eocene time , whereas easterl y sourc e area s dominate d during the remainde r o f Cenozoic tim e (e.g . Jordt et a l 1995, 2000 ; Jo y 1996 ; Mudg e & Buja k 1996 ; Michelsen e t a l 1998) . However , thi s pictur e may be severely biased if one merely looks at the preserved sediments . Fo r example , regiona l cross-sections o f th e norther n Nort h Se a (e.g . Jordt e t a l 1995 , Fig . 3) , sho w a completel y preserved Paleocene-Eocen e successio n pro grading fro m th e Shetlan d Platform , wherea s similar age progradational deposits ar e truncated

Fig. 5 . Spatial and temporal migratio n o f clinoform breakpoint s i n the eastern Nort h Se a Basin superimposed o n depth contours of the Upper Cretaceous-Danian limestone (clinofor m breakpoint s afte r Funnel l 1996 ; Clause n et al 1999) . The clockwise infil l fro m th e Oligocene tim e onwards should be noted. This caused the northern part of the are a t o be fille d t o base leve l som e 15—2 0 Ma befor e th e souther n part .

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Fig. 6 . ENE-WSW-oriented seismic profil e (RTD81-22 ) showing large-scale depositiona l geometrie s alon g the Norwegian-Danish secto r boundary . (For location , se e Figs 2 an d 7. ) I n the Centra l Grabe n are a a condense d upper Paleocene t o lower Middle Eocene successio n is overlain by thick upper Middle-Upper Eocene smectiti c clays with chaotic internal structure onlapping towards the east. The overlying succession o f Oligocene silty clays interbedded wit h thic k san d unit s exhibit s markedl y southwestwar d progradationa l geometries . Th e Miocen e deposits ar e les s markedl y progradationa l alon g thi s profil e an d consis t mainl y o f silt y cla y wit h thi n san d stringers. High-angle progradational geometries ar e also observed i n the upper Pliocene successio n to the WSW. The final phase of infill in Pleistocene tim e is characterized b y regional onlap of mainly shallow-water sediments supplied fro m th e SSE . Th e height o f the Oligocene-Miocene clinoforms indicates that palaeo-water depth s in the Centra l Grabe n are a wer e substantia l (500-1000m).

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Fig. 7 . Palaeogeographical developmen t of the eastern Nort h Se a Basin (modified afte r Huus e 2002); a t (a) midPaleocene, (b ) late Eocene, (c ) late Oligocene, (d ) mid-Miocene (e ) late Miocene, (f ) lat e Pliocene, and (g) midPleistocene time . Th e map s ar e base d o n observation s fro m th e easter n Nort h Se a Basi n integrate d wit h a n extensive number of published data. Se e Huus e (2002) fo r complete listin g o f data and reference s used .

UPLIFT AND DENUDATION O F SOUTHERN NORWA Y

at a hig h angl e toward s th e Norwegia n coast . These progradationa l wedge s ar e probabl y th e remnants o f a large r syste m o f progradationa l lower Paleogen e deposit s originatin g fro m southern Norway . It is thus likely tha t significant amounts of sediment were supplied from easterl y source area s durin g Paleocen e an d Eocene tim e (Jordt et al. 2000). Thi s is also indicated by sandprone depocentre s o f lat e Paleocen e t o earl y Eocene ag e S W o f Norwa y (Fig . 4) . Possibl e causes fo r th e differentia l preservatio n o f th e Paleocene-Eocene successio n o n eithe r sid e o f the northern North Se a will be discussed later in this paper . The large middl e t o upper Eocene depocentr e in th e Centra l Troug h (uni t 3 , Fig. 4 ; Michelse n et al . 1998 ) consist s mainl y o f smectite dominated cla y (Thyber g e t a l 2000) . Th e deposits contai n little , i f any , evidenc e fo r palaeo-transport direction s an d compris e a n abnormally thic k successio n of hemipelagi c clays deposite d fa r awa y fro m potentia l sourc e areas. In th e easter n Nort h Sea , th e Eocene Oligocene transitio n i s characterize d b y a massive increas e i n the amoun t of coarse clasti c sediments supplie d fro m souther n Norway , resulting i n a n almos t 1 km thic k depocentr e o f markedly progradationa l sand-pron e deltai c sediments o f Oligocen e ag e i n th e Norwegian Danish Basi n (Figs 4-6) . Thi s shif t fro m hemi pelagic clay s an d marl s t o silt y an d sand y clay s was probabl y a n effec t o f th e lat e Eocen e climatic deterioratio n (Buchardt 1978; Collinson et al . 1981) , whic h le d t o increase d seasonalit y and eustati c lowerin g (Ivan y e t al . 2000 ; Lea r et al . 2000) . Th e lowe r temperature s an d increased seasonalit y probabl y increase d th e amount o f precipitatio n an d cause d significan t changes i n vegetatio n (Spjeldnae s 1975 ; Collin son et al., 1981), thus increasing th e erosivity o f the geomorphologica l syste m (se e Summerfiel d & Brown 1998) . The clockwis e rotatio n o f th e directio n o f progradation (Fig . 5 ; Clausen et al. 1999 ) show s how the basin was filled by sediments prograding from th e NE , east , S E an d finall y sout h durin g Oligocene t o Pleistocen e time . A s th e mos t proximal part s o f th e basi n wer e fille d durin g Oligocene tim e (Fig . 7c ; Danielsen e t al. 1997) , sediments simpl y bypasse d th e Oligocen e depocentre durin g Miocen e time , fillin g u p th e easternmost par t of the basin coinciding wit h the central part s o f Denmar k (Fig s 4- 6 an d 7d) . Following th e Hodde transgressio n (Koch 1989) , sediments started prograding from the east across the norther n par t o f German y int o the relatively deep waters of the German Bight of the southern

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North Se a Basi n (Fig s 5 an d 7e ; Kocke l 1988 ; Jiirgens 1996) . Th e distal toes o f these lat e Mid to Uppe r Miocen e delta s o f th e souther n Nort h Sea obliquel y onla p th e Oligocene-lowe r Mid Miocene depocentre s farthe r nort h (se e Fig s 6 and 8) . In Pliocen e tim e th e easter n part s o f th e basin ha d bee n fille d an d th e patter n o f progradation wa s eas t t o west along mos t o f th e North Se a Basi n (Fig s 5 an d 7f ; S0rense n e t al . 1997). Th e fina l phas e o f infil l durin g lates t Pliocene-mid-Pleistocene tim e wa s mainl y sourced fro m th e SS E b y th e ancestor s t o th e large N W Europea n river s o f th e present-da y landscape (Figs 5 and 7g; Gibbard 1988; Zagwijn 1989), whereas sediments from souther n Norway made thei r way int o the northern Nort h Se a and the North Atlanti c (Jord t e t al 1995 , 2000 ; Rii s 1996; Evan s et al. 2000) . The general picture of infill is thus comparable with tha t o f passiv e margins , whic h generall y show a n overal l basinwar d progradatio n o f successive sedimentar y wedges . However , th e last phase of infill i s remarkable i n that the lowe r to middl e Pleistocen e sediment s ar e regionall y very extensive compare d wit h previous units. In fact, the y attai n thicknesse s o f u p t o mor e tha n 500m i n area s wher e accommodatio n ha d previously bee n fille d b y th e Miocen e an d Pliocene delta s (se e wester n par t o f Fig . 6) , thus indicatin g tha t additiona l accommodatio n was bein g create d i n earl y t o mid-Pleistocen e time. Apar t fro m th e centra l part s o f th e basin , the lowe r t o middl e Pleistocene sediment s ar e generally o f relativel y shallow-wate r origin , suggesting tha t sedimentatio n rate s kep t u p with th e increase d subsidenc e rate s (Huus e 2002). Correlation with regional tectonic events Apart fro m regiona l plume-relate d uplif t i n Paleocene-early Eocen e tim e ther e i s onl y minor sig n of tectonic activit y in Cenozoic tim e in the North Sea Basin. This is mainly in the form of local inversion of old fault system s such as the Sorgenfrei-Tornquist Zon e an d th e Centra l Graben (Vejba? k & Andersen 1987 , 2002; Ziegler 1990). Th e inversio n o f thes e structure s wa s probably cause d b y th e combine d effect s o f Atlantic ridg e pus h an d Alpin e compressio n (e.g.Vejbaek & Anderse n 2002 ) rathe r tha n b y Alpine compression alone . The effect o f Atlantic ridge push is also see n a s large inversio n dome s along th e N W Atlantic margi n (Dor e & Lundin 1996; Boldree l & Anderse n 1998 ; Dor e e t a l 1999). It i s unlikel y tha t intra-plat e compressio n could caus e significan t uplif t o f cratoni c sourc e

UPLIFT AND DENUDATION O F SOUTHERN NORWA Y

areas such as southern Norway (Rohrma n & van der Bee k 1996 ; Dor e e t al 1999) . O n the othe r hand, i t ma y b e possibl e tha t intra-plat e compression coul d creat e enoug h disturbanc e along ol d faul t zone s t o cause avulsion of majo r rivers an d thu s resul t i n majo r change s o f sediment inpu t direction s suc h a s observe d i n mid-Miocene tim e (se e Figs 5 and 7d and e). It ha s bee n suggeste d tha t abnormall y rapi d early Pleistocen e subsidenc e coul d hav e bee n caused b y intra-plat e compressio n (e.g . Cloe tingh et al 1990 , 1992) . However, the absence of evidence fo r earl y Pleistocen e compressiona l faulting in the North Sea area makes it less likely that extensiv e compressio n wa s th e caus e o f rapid earl y Pleistocene subsidence . Another effect t o consider is the load-induce d subsidence cause d b y th e las t phas e o f infil l o f the North Sea Basin. By late Pliocene time only a narrow seawa y o f relativel y grea t wate r dept h existed i n th e centra l Nort h Sea . Th e infil l o f a narrow (sa y 400 m) deep basi n remainin g a t the beginning of Pleistocene tim e could cause som e degree o f flexura l dow n warping o f th e margins. The exten t t o whic h th e sediment-loade d subsidence woul d b e distribute d laterall y is , however, strongl y dependen t o n th e flexura l strength o f th e lithosphere , an d numerica l modelling o f th e loadin g effec t need s t o b e carried ou t t o quantif y thi s effect .

Correlation with global climate and sea level The suppl y o f coars e clasti c sedimen t t o th e North Sea Basin accelerated severa l times during Cenozoic time , mos t notabl y i n lat e Paleocen e time, at the Eocene-Oligocene transition, in late Mid-Miocene tim e an d in Plio-Pleistocene tim e (Fig. 3) . I t appear s straightforwar d tha t th e abruptly increase d suppl y o f siliciclasti c sedi ments in late Paleocene tim e was caused by uplift

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of sourc e area s i n relatio n t o th e arriva l o f th e plume and rifting of the North Atlantic. Because of th e substantia l tim e la g (>20Ma) , direc t plume-related effect s canno t satisfactoril y explain th e increases i n sedimen t suppl y a t th e Eocene-Oligocene transition , an d i n lat e Mid Miocene an d Plio-Pleistocen e times . Hence , i t should b e investigate d whethe r thes e increase s could have been caused by other well-documented phenomena such as major long-term climatic and eustati c changes . Althoug h plume-relate d effects canno t accoun t fo r abrup t change s i n sediment supply, the effect o f delamination of the lithosphere ma y b e par t o f th e explanatio n b y providing decelerating bu t continue d uplif t afte r the initial uplift puls e (Nielsen et al. 2002) . Previous studie s hav e note d a conspicuou s correlation betwee n majo r Nort h Se a sequenc e boundaries and major 8 8 O increases (Jordt et al. 1995; Huus e & Clause n 2001 ; Huus e 2002) . Huuse (2002 ) als o note d tha t large-scal e sedimentation pattern s o f th e easter n Nort h Se a Basin ar e comparabl e wit h sedimentatio n pat terns observe d alon g man y continenta l margins (see Donnelly 1982 ; Bartek et al. 1991 ; Cameron et al . 1993 ; Mille r e t al . 1998 ; Serann e 1999 ; Huuse & Clause n 2001) . Moreover , th e period s of increase d sedimen t suppl y a t th e Eocene Oligocene transition , an d i n lat e Mid-Miocen e and Plio-Pleistocene times roughly correlate with episodes o f climati c deterioratio n an d majo r increases in continental ice volume (see Buchard t 1978; Molnar & England 1990; Lear et al. 2000). Other evidenc e o f significan t increase s i n denudation come s fro m geomorphologica l studies, which have indicated that rates of stream incision accelerate d i n lat e Cenozoi c time , both in souther n Norwa y an d i n th e easter n US A (Lidmar-Bergstrom e t al . 2000 ; Mill s 2000) . Such virtuall y contemporaneou s increase s i n denudation indicate a global rather than regional or local caus e (Molna r & England 1990) .

Fig. 8 . North-south-oriente d seismi c profil e (DA94-04 ) showin g a conformabl e uppe r Paleocen e t o Eocen e succession consistin g o f hemipelagic cla y an d marl (unit s 1—3 ) overlying the Uppe r Cretaceous-Dania n Chal k Group. Th e condense d Paleocene-Eocen e successio n i s overlai n b y a thic k progradationa l successio n o f Oligocene t o mid-Miocen e ag e (unit s 4-6) . Th e post-middl e Miocen e sequenc e (uni t 7 ) is characterize d b y a thick shallowing-upward s aggradationa l successio n o f marin e cla y an d silt , onlappin g th e mid-Miocen e unconformity. P-wave velocities fro m fou r wells located alon g the profile (S-l, R-l, Inez-1, F-l) indicate that the velocity o f the post-Chalk Grou p is very clos e t o 2kms~ 1 (i.e . 1 s two-way trave l tim e (TWT ) ~ 1 km). Thi s relationship is used to directly compare geometrie s on the seismic wit h inferred amount s of 'missin g section' or 'Neogene uplift ' base d o n sonic-derive d compactio n trend s o f th e Chal k Grou p (Japse n 1998 ) an d o f Jurassi c shales (Jense n & Schmidt 1993) , respectively . Estimate s o f 'missin g section' an d 'missin g overburden' base d on integration o f chal k compactio n an d vitrinit e reflectanc e dat a (Japse n & Bidstru p 1999 ) ar e als o shown . Th e amount o f uplif t estimate d fro m th e seismi c geometrie s i s significantl y lowe r tha n tha t base d o n compactio n trends. Q , Quaternary valley . (Fo r location , se e Figs 2 and 7.)

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Cenozoic uplift an d denudation of southern Norway Constraints on magnitude Apatite fission-trac k thermochronolog y (AFTT ) has become a popular prox y fo r inferring uplift . However, AFTT relates to cooling histories of the rocks analysed , i.e . t o denudatio n an d no t directly t o uplif t (Brow n 1991) . Thi s i s important, a s denudatio n ma y la g uplif t b y several ten s o f millio n year s (Summerfiel d & Brown 1998) . Moreover, AFTT is sensitive onl y to temperature s o f minimu m c. 60° C (depth s >2-3 km) and is thus generally no t sensitive to late Cenozoi c effect s (Gunnel l 2000) , althoug h inversion o f th e AFT T dat a ma y hin t a t lat e Cenozoic coolin g historie s (e.g . Rohrma n e l al. 1995). The latter indicate a maximum denudation of 2 ± 0.5k m i n th e inne r fjord s o f souther n Norway decreasin g outwar d t o < 0.5 km a t the present-day coastline , wherea s th e amoun t o f denudation at the highest mountain peaks suc h as Jotunheimen i s belo w th e resolutio n o f AFT T inversion (Rohrma n e t al 1995 ; Fig . 8) . Th e rather large uncertainties of denudation estimates based o n th e AFT T metho d cal l fo r additiona l constraints o n denudatio n o f souther n Norway . The most powerful method fo r producing reliabl e denudation estimate s i s probabl y (U-Th)-H e thermochronometry, couple d wit h investigations of cosmogeni c isotop e studie s t o dat e expose d landforms, an d i t i s recommende d tha t suc h investigations be carrie d ou t i n the future . The denudatio n estimate s cite d abov e ar e almost opposit e t o th e estimate s o f Rii s & Fjeldskaar (1992) , whic h sho w greates t averag e amounts of denudation at the coastline (thei r fig. 11). Thi s stud y relie s o n geomorphologica l characteristics of the landscape an d extrapolation of well-date d offshor e surface s ove r lan d area s (Dore 1992 ; Rii s & Fjeldskaar 1992 ; Rii s 1996) . The use of this method i s somewhat problemati c when databl e rock s ar e absen t i n th e area s o f maximum uplif t an d th e ag e o f th e surfac e onshore i s thus poorly constrained . It is possible, however, t o defin e relativel y coheren t palaeo surfaces i n eve n severel y denude d area s usin g geomorphological criteria , althoug h th e ag e o f the surfac e wil l b e conjectura l (Gjessin g 1967 ; Riis & Fjeldskaa r 1992 ; Rii s 1996 ; Lidmar Bergstrom e t al . 2000) . Th e us e o f palaeosur faces thu s ma y yiel d som e ide a abou t th e magnitude o f uplift , bu t wit h rathe r poo r constraints o n timing , an d may , i n som e cases , agree poorly wit h result s o f AFTT. A widely used method o f estimating uplif t an d denudation i s t o compar e regiona l compactio n

trends o f shale s an d chalk s relativ e t o a define d 'normal trend' (Jense n & Schmidt 1993 ; Hansen 1996; Japsen 1998) . The estimates from compac tion-trend method s ar e highl y variable , eve n using the same wells (Fig. 8 ) and it would appea r that th e regiona l trends o f over - (an d under- ) compaction ar e mor e usefu l tha n th e absolut e values. Th e estimate s o f 'missin g section ' ma y be further constraine d by integrating compaction analyses with vitrinite reflectance studie s (Fig. 8; Japsen & Bidstrup 1999). However, although this integration yield s even lowe r estimate s of uplif t and denudation , the y ar e stil l o f th e orde r o f 300 m abov e thos e base d o n seismi c geometrie s (Fig. 8 : compare with 'buria l anomaly' o f Japsen & Bidstru p 1999) . Japse n & Bidstru p inferre d that lat e Neogen e erosio n i s responsibl e fo r th e 'burial anomaly ' show n i n Fig . 8 . I f thi s i s th e case, the n a 300-400 m thick wedge of Pliocene sediments must have been deposited between the (complete) Middle-Uppe r Miocen e successio n and the Lower-Middle Pleistocene unit in Fig. 8. However, i n Pliocen e tim e th e easter n Danis h North Se a wa s mainl y bypasse d b y sediments , which wer e deposite d i n larg e delta s 100k m farther to the west and SW (Fig. 7e and f) - Hence, the notio n of rapid lat e Neogen e depositio n an d erosion i n the eastern Danish North Sea inferre d from compaction-base d exhumatio n estimate s does no t agre e wit h th e palaeogeographica l evolution o f the area . As demonstrate d b y th e abov e example , i n areas wher e the bulk of the Cenozoic successio n is preserved , i t i s possibl e t o us e large-scal e depositional geometrie s observe d i n seismi c profiles t o estimat e th e amoun t o f tiltin g an d denudation tha t ha s occurre d withi n th e basi n during Cenozoic time . It should be noted that it is the amoun t o f til t o f previousl y horizonta l surfaces tha t indicate s differentia l uplif t o r subsidence; th e volum e o f sedimen t o f an y given age reflects only the amount of denudation. Geometrical denudation estimates and 'uplift ' and/or denudatio n estimate s base d o n compac tion trends al l show simila r trends of denudation increasing northwards , bu t th e amplitude s var y widely, wit h the geometrica l estimat e being th e lowest (Fig . 8) . I t i s beyon d th e scop e o f thi s study to scrutinize the methods and assumptions behind compactio n analyses , bu t i f th e seismi c geometries ar e t o b e relied upon , it appear s that compaction method s overestimate the amount of uplift an d denudatio n o f th e easter n Nort h Sea . Moreover, i t i s remarkabl e tha t th e geometrica l estimate i s i n accor d wit h a n estimat e o f 'overburial' o f th e Chal k Grou p base d o n 3 D basin modellin g (S.B . Nielse n pers . comm . 2000). Th e modellin g stud y incorporate s a

UPLIFT AND DENUDATION O F SOUTHERN NORWA Y

thermally subsidin g North Se a Basin, subjec t to long-term eustati c fall , varyin g sedimen t inpu t and sedimentar y loading , withou t th e influenc e of late Cenozoic tectonics . Hence, it appears that the large-scal e strata l geometrie s observe d o n seismic dat a (Fig s 6 and 8 ) reflect th e infil l o f a thermally subsidin g basi n durin g generall y falling se a level. Constraints on timing Paleogene an d Neogen e sediment s ar e almos t exclusively locate d aroun d the fringe s of the uplifted area s i n Scandinavi a an d Britai n an d there ar e n o sediment s preserve d i n th e mos t elevated parts . Hence , unti l high-resolutio n fission-track dat a becom e available , th e timin g of uplif t an d denudatio n has t o be inferre d fro m the sedimentar y recor d o f th e adjacen t basins . This task requires regional dat a coverage to filter out loca l variation s i n sedimen t supply , whic h could caus e potentiall y misleadin g sedimen tation patterns . Whe n tryin g t o establis h uplif t and denudatio n historie s fro m th e sedimentar y record i t i s extremely importan t to bear i n mind that althoug h a volum e o f sedimen t i s directl y related t o denudatio n i t tells u s ver y little abou t uplift. Also , i t i s important t o bea r i n min d that isopach map s sho w only the present distribution of th e erosiona l products . Hence , i t i s possibl e that sediment s recordin g earl y denudatio n hav e now bee n remove d a s a result o f late r uplif t (o r base-level fall ) an d associated erosion . Thus, the present distributio n o f sediment s ma y b e dominated b y th e mos t recen t episode s o f denudation and the associated isostati c response . This i s especiall y tru e fo r th e earlies t proxima l sediments deposite d alon g th e margin s o f th e source area . It seem s likel y tha t th e apparen t lac k o f proximal sediments of Paleocene an d Eocene ag e off Norwa y i s du e t o cannibalizatio n an d redeposition. Thi s i s indicated b y th e truncation of thick, highly progradational sedimen t wedge s off th e wes t coas t o f Norwa y (se e Jord t e t al. 1995; Fig. 3) . The less marked truncation around the Shetlan d Platfor m coul d b e du e t o th e narrowness o f th e Shetlan d topograph y (abou t one-third th e widt h o f th e sout h Norwegia n dome) causing less uplif t a s a result of erosiona l unloading. However, variations in thermal uplif t and subsidenc e (wit h th e Shetlan d Platfor m experiencing th e larges t o f both ) ma y als o hav e influenced thi s pattern. From th e sedimentar y recor d i t appear s tha t denudation accelerate d at least fou r times durin g Cenozoic time : i n lat e Paleocen e time , a t th e Eocene-Oligocene transition , i n lat e Mid -

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Miocene tim e an d in Plio-Pleistocene time . Th e first episode wa s probably a response to regional uplift associate d wit h plume activit y an d riftin g in th e Nort h Atlanti c region . Th e remainin g episodes appea r t o correlat e wit h simila r responses o n continenta l margin s fa r fro m th e North Atlanti c domain , thu s indicatin g th e influence o f globa l factor s suc h a s climat e an d eustasy. Hence , i t seem s tha t th e latte r thre e episodes o f accelerate d denudatio n ma y reflec t climatic and eustatic changes rather than tectonic events. Isostatic response to localized denudation (dissection) The width of the south Norwegian dome is of the order o f 300-40 0 km alon g an y cross-sectio n (Figs 1 and 9). The area of the dome i s therefore c. 10 5km2 and remova l of materia l fro m the surface o f th e dom e i s thu s likel y t o b e isostatically compensate d b y additio n o f mantl e material a t depth (England & Molnar 1990) . This is in agreement wit h gravity data (Balling 1980) , which indicat e tha t th e topograph y o f souther n Norway i s isostaticall y compensate d a t depth . Because th e lithospher e ha s a certai n flexura l strength ther e ar e likel y t o b e flexura l effects , seen a s subdue d uplif t (o r subsidence ) respons e to erosion (o r sedimentation), alon g th e margins of th e dome . A schemati c illustratio n o f th e formatio n o f dissected topograph y fro m a n initia l low elevation peneplai n i s show n i n Fig . 1 0 an d described below , leavin g ou t flexura l effect s a t the margins of the rock column. The dimension s of the rock column are comparable wit h those of the centra l part s o f th e sout h Norwegia n dome . Removal o f crusta l materia l (p c ~ 2.7gcm~ 3 ) from th e to p o f th e 200k m wid e column s i s isostatically compensate d b y additio n o f mantle material (p m ~ 3.3gcm~ 3 ) a t depth . Th e iso static compensation (/ = p c/pm) for denudation is thus c . 0.8 . Hence , 1 km (mean ) denudatio n would decreas e th e surfac e elevatio n b y onl y 0.2 km as a result of the isostatic response, which would caus e 0.8k m uplif t o f th e entir e roc k column. Regional studie s o f activ e orogen s (Europea n Alps, Andes, Himalayas) indicat e that dissection is rarely fully develope d i n their central parts and that onl y abou t hal f o f th e pea k heigh t a t th e centre o f orogen s ca n b e explaine d b y th e isostatic respons e t o denudation (Gilchrest e t a l 1994). Looking a t a regional topographi c cross section o f souther n Norwa y (Fig . 9 ) i t appear s that denudatio n i s unevenl y distributed , wit h

Fig. 9 . Topographic cross-section s o f th e sout h Norwegia n dome . Locatio n i s shown i n Fig. 1. The uppe r profil e shows a hypothetica l cross-sectio n o f th e dom e a t midEocene tim e (c. 40 Ma). The lower profile shows the present topography and summit envelope (afte r Torske 1972 ) and the mean surfac e elevation averaged ove r 50-100 km. The mea n surfac e elevation is a qualitative estimate o f the isostaticall y compensated topograph y envelope . I n mid-Eocen e tim e (c. 40 Ma) a hilly relie f ha d develope d i n response t o weathering of a Mesozoic peneplai n uplifted t o 1 -1.5 km elevatio n in earliest Eocen e time. A warm climate , dense vegetatio n an d a low-gradien t local relie f probably cause d lo w rates o f denudation . It should be note d tha t sea leve l wa s c . 200 m highe r at 40 Ma tha n a t present . Th e present-da y deepl y incise d relie f probabl y developed i n response t o repeated episode s o f climatic deterioration an d eustati c lowering during mid - to lat e Cenozoic time , culminating wit h full glacia l conditions and extreme rates of incision i n Plio-Pleistocene time. Deep dissection of the former high-elevation peneplain caused mountai n peak s to rise to approximately twice their initial elevation. As a result of the eustatic fall o f c. 20()m sinc e mid-Eocene time , the mea n surfac e elevation (with respec t t o se a level ) has remaine d a t roughly the sam e leve l throughout.

UPLIFT AND DENUDATIO N O F SOUTHERN NORWA Y

maximum denudatio n o f 1-1. 5 km i n a zon e stretching c . 100k m inboar d o f th e coastline . Local denudation is of the order of 0.7-0.8 km in the area of highest mean elevation an d decrease s to 0.5 km farther t o the SE . Average denudation (defined a s summit envelope minus mean surfac e elevation) i s highest (c . 1 km) abov e th e centra l and northwestern parts of the dome (Fig. 9). This picture is , o f course , strongl y dependen t o n th e length scal e ove r whic h denudatio n is average d and on the 3D distribution o f valley incision, bu t it seem s t o indicat e tha t th e centra l part s o f th e dome have been affected b y the combined effect s of loca l denudatio n (causin g surfac e lowering ) and isostatic uplift in response to deep incision of neighbouring areas .

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caused accelerate d rate s o f strea m incisio n an d thus increased denudatio n rates. Finally, in PlioPleistocene time , ful l glacia l condition s i n th e high-lying part s o f souther n Norwa y cause d extreme rates o f incision, eventually carving out the dissecte d topograph y w e observ e toda y (Fig. 9) . In th e scenari o describe d here , th e onl y tectonic uplif t o f th e surfac e (an d of th e roc k column) occurred i n early Paleogene time . Dee p dissection o f the uplifted peneplain an d resultant isostatic uplif t o f th e roc k colum n cause d th e remaining par t o f th e roc k uplift , wherea s the mea n surfac e wa s lowere d b y c . 20 % o f th e amount o f mea n denudation . I t shoul d b e note d that th e averag e denudatio n o f 1 km (surfac e lowering of 0.2 km) is compensated by a eustatic lowering o f the sam e magnitude , thu s maintain ing th e mea n surfac e elevatio n (wit h respec t t o Early Cenozoic uplift and late Cenozoic sea level) a t c. 1 km. denudation; a hypothetical model The landscape evolutio n as depicted i n Figs 9 A regional cross-sectio n o f the south Norwegian and 1 0 i s i n agreemen t wit h denudatio n a s dome is shown in Fig. 9. The upper panel shows a recorded b y th e offshor e stratigraphi c recor d hypothetical low-relie f landscap e afte r regiona l (Figs 4-7 ) an d wit h geomorphologica l studies , uplift o f a peneplain, befor e th e developmen t o f which indicat e accelerate d rate s o f incisio n i n the deepl y incise d valley s an d fjord s tha t late Cenozoi c tim e (Lidmar-Bergstro m et al. characterize th e present-da y topography , illus - 2000). A lat e Cenozoi c increas e i n th e rate s of stream incisio n wa s als o foun d i n a regiona l trated b y the lower profile . A hypothetica l mode l o f Cenozoi c uplif t an d study o f the easter n US A (Mill s 2000) , support denudation o f th e centra l part s o f souther n ing the notion of an allogenic control on incision Norway i s show n i n Fig . 10 . A t th e en d o f rates, i.e. climate and eustasy rather than regional Mesozoic time , souther n Norwa y wa s probabl y tectonics. Th e Plio-Pleistocen e increas e i n th e worn dow n t o a peneplai n o f som e mea n rates o f incisio n i s als o reflecte d b y th e elevation abov e se a level . Th e exac t elevatio n occurrence o f thic k Plio-Pleistocen e sedimen t of the peneplain i s poorly constraine d an d here it wedges al l alon g th e N W Europea n Atlanti c is assume d tha t th e mea n heigh t o f th e centra l margin (Rii s & Fjeldskaa r 1992 ; Riis 1996 ; parts wa s close t o 200 m elevation, althoug h the Evans e t a l 2000 ) an d i n th e centra l Nort h Se a highest peak s i n souther n Norwa y ma y b e (Figs 4 and 6). The concept of mountain building by isostati c respons e t o dissection o f an uplifte d remnants o f inherite d topograph y (Rii s & peneplain i s als o backe d b y th e apatit e fissionFjeldskaar 1992) . Late Paleocene-early Eocene plume- and rift- track analyse s o f Rohrma n e t al . (1995) . Thes e related tectonic s uplifte d the peneplain t o c. 1 km results indicate rapid late Cenozoic denudatio n at elevation (includin g pre-uplif t elevation) , bu t the bas e o f the deepes t valley s (fjords ) an d only incision was initially minimal in the central parts minor denudatio n (belo w th e detectio n of the dome, as a result of the low local relief, and threshold) o n th e mountai n peaks . However , a war m humi d climat e favourin g dens e veg- acquisition o f high-resolutio n fission-trac k dat a etation an d limited runoff . Th e bul k of sediment is neede d t o furthe r constrai n th e Cenozoi c supplied t o th e basi n a t thi s tim e (40-6 0 Ma) denudation history of souther n Norway. The model propose d here is not in agreemen t probably derived fro m the margins of the uplifted with th e interpretation s b y Rii s (1996) , wh o dome (not shown in Fig. 10) . In late Eocene-early Oligocene time , climatic found tha t easter n Denmar k ha s suffere d mor e deterioration (increase d seasonality ) and eustatic than 1 km o f Plio-Pleistocen e denudation . Thi s lowering cause d increase d rate s o f strea m estimate wa s probabl y drive n b y a n attemp t t o incision. Climat e recovere d durin g lat e Oligo - honour th e patter n o f lat e Neogen e erosio n cene-early Mid-Miocen e time , thu s stabilizin g inferred fro m compactio n analyses . A s argue d above, suc h estimates probabl y overestimat e th e incision rates . In lat e Mid-Miocen e tim e anothe r phas e o f amount o f exhumatio n o f th e easter n Nort h Se a climatic deterioratio n an d eustati c lowerin g Basin b y severa l hundre d metres. Moreover , th e

Fig. 10. Schematic illustration of uplift and bisection of an initial low-elevation peneplain. The dimensions are comparable with those of the contral parts of the south Nonwegian dome. (see text for discussion)

UPLIFT AND DENUDATIO N O F SOUTHERN NORWA Y

study relied on extrapolation o f key surfaces over several hundre d kilometres , obviousl y a some what tricky disciplin e whe n Cenozoic sediment s are absent over most of the area. In particular, the occurrence o f lat e Earl y Eocen e diatom s i n northern Finlan d (Tynn i 1982 ; Fenne r 1988 ) ha s been use d t o infe r a Paleogene episod e o f low elevation peneplanation an d submergence (Rohrman et al 1995 ; Rii s 1996) . The locations of the diatom find s are , however , al l o f relativel y lo w elevation (<300m ) an d coul d thu s hav e bee n submerged durin g the early Eocene highstan d of sea level (Fig . 3) without the need for Paleogene peneplanation o f souther n Norway . Th e sub sequent emergence o f the diatom locations could easily b e explaine d b y th e combine d effect s o f post mid-Eocen e eustati c lowerin g o f c. 200 m and flexura l uplif t a s a resul t o f denudatio n (dissection) an d isostati c uplif t (o f rocks) o f th e adjacent norther n Norwegian highland . Modelling studie s (e.g . Rii s & Fjeldskaa r 1992; Stuevol d & Eldhol m 1996 ) generall y fal l some hundre d metre s shor t o f explainin g th e present topograph y o f southern Norway, leading the researcher s t o sugges t tha t activ e tectoni c mechanisms suc h a s mantl e phas e change s o r palaeo-topographic relie f ar e required to explain present-day topograph y of souther n Norway . However, thes e studie s assume d a constan t global se a leve l throug h Cenozoi c tim e a s opposed t o th e 200 m eustati c lowerin g tha t occurred sinc e mid-Eocene time . Neglecting this fall woul d introduc e 200 m surfac e uplif t o f presumed tectoni c origi n int o th e modellin g results. Differences between southern Norway and the Shetland Platform Differential preservatio n o f upper Paleocene and Eocene sediment s acros s th e northern Nort h Sea could b e du e to a number of factors suc h as: (1 ) different amount s o f transien t uplif t cause d b y heating o f th e lithospher e beneat h th e Shetlan d Platform an d souther n Norway ; (2 ) differen t amounts o f permanen t uplif t cause d b y under plating o r delaminatio n o r othe r non-transien t mechanisms; (3) presence v. absence of extrusive volcanic rock s i n Scotlan d v . Norway ; (4 ) different (flexural ) isostati c respons e t o denudation; (5 ) palaeobathymetric variations ; (6) width of th e topograph y (ful l isostati c o r flexura l isostatic behaviour). Plate tectoni c reconstruction s (Skogsei d e t al. 2000) indicat e tha t th e Shetlan d Platfor m wa s closer t o the Iceland Plum e a t the time of riftin g (c. 54 Ma) and thus experienced a larger thermal

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(transient) uplift tha n southern Norway. It should thus be expected that the depocentres adjacent to the Shetlan d Platfor m subsequentl y experience d greater therma l subsidenc e an d thus had greate r preservation potentia l tha n thos e a t th e sam e latitude of f souther n Norway. Possible mechanism s responsibl e fo r non transient uplif t o f souther n Norwa y an d othe r uplifted area s wer e discusse d b y Nielse n e t al . (2002), wh o invoke d delaminatio n o f th e litho sphere a s th e uplif t mechanis m fo r souther n Norway. Other areas, e.g. Scotland, may not have been affecte d t o th e sam e degre e o r ma y eve n have experience d underplating , leadin g t o a different uplif t an d subsidenc e history . Early Paleogen e extrusiv e volcani c rock s ar e abundant i n wester n Scotland , bu t absen t i n Norway. Extrusiv e volcani c rock s ar e readil y erodible an d volcani c activit y contribute s t o a high loca l relief , bot h o f whic h ma y hav e contributed t o a large r sedimen t suppl y fro m Scotland in early Paleogene time (see Hall 1991) . The south Norwegian dome has a much large r areal exten t tha n Scotlan d an d th e Shetland s (Fig. 1) . I t i s possibl e tha t thi s coul d caus e differences i n th e amoun t o f (flexural ) isostatic response to denudation (and deposition) betwee n the tw o areas . Th e relativel y larg e exten t o f the south Norwegian dome probabl y cause s most of the loadin g and unloadin g of materia l acros s southern Norwa y t o b e isostaticall y compen sated, excep t aroun d the fringe s of the dome , where flexural strength may have an effect. I n the case o f Scotlan d an d Shetlands , th e narrownes s of th e topograph y ma y caus e a relativel y larg e part o f th e loadin g an d unloadin g o f materia l across th e platfor m t o b e supporte d b y th e flexural strengt h o f th e lithosphere . Hence , th e amount o f roc k uplif t a s a resul t o f denudatio n may b e significantl y les s fo r Scotlan d an d th e Shetlands than for souther n Norway. The effect s of loadin g b y deposition adjacen t t o the uplifte d area woul d als o hav e a greate r effec t o n th e Shetlands b y depressin g an y isostati c uplif t response t o denudation. Palaeogeographical reconstruction s (Fig . 7a ; Ziegler 1990 , fig . 54 ) indicat e tha t sediment s deposited eas t o f Scotlan d an d th e Shetland s were fe d acros s a narro w shel f int o relativel y deep wate r wherea s th e sediment s o n th e Norwegian sid e wer e deposited o n a wide shelf . Such differences i n palaeobathymetry could have caused th e greate r thicknes s an d bette r preser vation o f depocentre s observe d eas t o f Scotlan d and the Shetlands . The 'self-perpetuating ' mode l of uplif t and denudation o f souther n Norwa y invoke d her e (Fig. 10 ) works only until maximum equilibrium

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of valle y slope s i s reached ; the n mountai n top s will commenc e erodin g an d regiona l down wearing will prevail (Ahnert 1984; Gilchrest et al. 1994). Becaus e o f thei r narrowness , th e Shet lands migh t alread y hav e reache d thi s stag e i n Eocene time whereas i t has not yet been reache d in th e central part s o f souther n Norway (Fig . 9) . This could explain why sediment supply from th e northern par t o f the Britis h Isles slowed dow n in Neogene time . Conclusions The Cenozoi c evolutio n o f th e Nort h Se a Basin has been summarize d base d o n a compilation of detailed studie s from the eastern North Sea Basin integrated wit h published regiona l studies . The first pulses of coarse clastic sediment into the basin occurred i n response to regional surfac e uplift coincidin g wit h increase d plum e activit y and th e onse t o f Nort h Atlanti c riftin g i n lat e Paleocene an d earlies t Eocen e time . Th e preserved uppe r Paleocen e an d Eocen e sedi ments mainl y originat e fro m th e west , wherea s progradational siliciclasti c wedge s o f simila r age hav e bee n truncate d toward s th e sout h Norwegian dome . In early Oligocene time, sediment supply from Scotland an d the Shetlands diminished, probably as a resul t of therma l subsidenc e and regiona l down-wearing o f topography in Scotland an d the Shetlands. Th e suppl y o f coarse clasti c materia l from souther n Norwa y accelerate d a t th e Eocene-Oligocene transition , an d fro m Oligo cene tim e onward s th e central , easter n an d southern Nort h Se a Basi n wa s fille d b y larg e deltas prograding in a clockwise fashio n from th e NE (Oligocene-earl y Miocene) , eas t (lat e Miocene-Pliocene), an d finall y SS E durin g Pleistocene tim e (Fig s 5 and 7). It i s argue d her e tha t seismi c geometrie s ca n be use d wit h confidenc e t o infe r uplif t an d denudation o f area s wher e Cenozoi c sediment s are preserved . Suc h geometri c uplif t estimate s have bee n compare d wit h estimate s base d o n compaction trend s alon g a profile i n th e easter n Danish Nort h Sea . Al l method s sho w simila r trends o f increasin g uplif t toward s souther n Norway, bu t th e compariso n indicate s tha t compaction-based estimate s ma y overestimat e the amoun t o f uplif t b y severa l hundre d metres . Recent uplif t estimate s integratin g chal k com paction an d vitrinit e reflectanc e dat a (Japse n & Bidstrup 1999 ) ar e lower tha n previous compac tion-based estimates , bu t stil l c . 300 m to o high compared wit h th e geometri c estimates . Th e geometric estimate s ar e simila r t o estimate s o f overburial base d o n 3 D basi n modellin g o f a

passively subsiding basin filled during falling sea level (S.B. Nielsen pers. comm. 2000), indicating that lat e Cenozoi c tectoni c event s ma y no t b e required to explain the sedimentation and erosion patterns of th e Nort h Se a Basin. A mode l describin g th e developmen t o f topography i n souther n Norwa y i s develope d (Fig. 10 ) based o n th e evidenc e provide d b y th e Cenozoic successio n of the North Se a Basin and the present-da y topograph y o f th e adjacen t land areas. Th e mode l involve s lat e Paleocene earliest Eocen e uplif t o f a low-elevatio n pene plain develope d durin g late Mesozoi c an d earl y Paleocene tim e i n a n are a roughl y coincidin g with present-da y souther n Norway . Th e uplif t episode wa s followe d b y c . 1 5 Ma o f war m climatic conditions favouring a dense vegetation and lo w amount s of runoff . I n combination with low loca l relie f an d hig h se a level , thi s cause d relatively low rates of denudation during Eocene time. Subsequently, repeated episodes of climatic deterioration an d eustati c fal l cause d increase d rates of incision, most markedl y a t the EoceneOligocene transition , and i n lat e Mid-Miocen e and Plio-Pleistocen e times . Th e developmen t of full glacia l condition s i n souther n Norwa y i n Plio-Pleistocene tim e was probably instrumental for carvin g out the main parts of the deep valleys and fjords o f southern Norway, i.e. dissecting the initial high-elevatio n peneplai n an d causin g uplift o f th e adjacen t mountai n peak s t o twic e the initia l elevatio n o f th e uplifte d peneplain . This mode l i s i n agreemen t wit h th e large-scal e sedimentation pattern s o f th e Nort h Se a Basin , with geomorphologica l studie s (Lidmar-Berg strom e t al . 2000 ) an d wit h apatit e fission-trac k modelling (Rohrma n e t al . 1995) . I t als o accounts fo r th e occurrenc e o f thic k Plio Pleistocene wedge s alon g th e N W Europea n Atlantic margi n (Riis 1996 ; Evan s e t al. 2000) . The rise of the mountain peaks is facilitated by the regional isostatic response t o dissection of an initially flat surface of some elevation (Molnar & England 1990 ; Gilchres t et al. 1994) , which will cause an uplift of the rock column c. 0.8 times the amount o f mea n denudation . Th e amoun t o f denudation average d ove r th e sout h Norwegia n dome approximate s 1 km (Rii s & Fjeldskaa r 1992). Thi s woul d correspon d t o c . 200 m o f mean surface lowering. However, the eustatic fall of c . 200 m sinc e mid-Eocen e tim e ha s maintained the mean surfac e elevation of central southern Norwa y a t approximatel y th e leve l o f the uplifted peneplain . To test the model advocated here and to furthe r constrain th e uplif t an d denudatio n histor y o f southern Norway and other uplifted area s there is a nee d fo r a n integrate d approach , leadin g t o

UPLIFT AN D DENUDATION O F SOUTHERN NORWAY

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mass-balanced palaeogeographica l reconstruc - BIRD, P . 1979 . Continenta l delaminatio n an d th e Colorado Plateau . Journal o f Geophysical tions. Thi s require s integratio n o f well-date d Research, 884,7561-7571. sediment volume s (mainl y offshore ) wit h detailed denudation estimates from both onshore BOLDREEL, L.O . & ANDERSEN , M.S . 1998. Tertiar y compressional structure s o n th e Faroe-Rockal l and offshor e regions . In particular , the onshore Plateau i n relation t o northeast Atlantic ridge-pus h regions are in need of more detailed studies, e.g. and Alpin e forelan d stresses . Tectonophysics, 300, based o n apatite (U-Th)-He thermochronome13-28. try, t o resolve the Cenozoic denudation history in BRODIE, J . & WHITE , N . 1995 . Th e lin k betwee n sufficient detail . Ideally , suc h a n integrate d sedimentary basi n inversio n an d igneou s under plating. In : BUCHANAN , J.G. & BUCHANAN , P.G. database woul d b e fe d int o 3 D geodynami c (eds) Basin Inversion. Geologica l Society , London, models incorporatin g isostati c an d flexura l Special Publications , 88 , 21-38. effects o f loading and unloading of the crust . The content s o f thi s pape r deriv e fro m a Ph D stud y carried ou t a t th e Departmen t o f Eart h Sciences , University o f Aarhus , co-sponsore d b y th e Danis h Natural Scienc e Researc h Counci l (Grant s 940116 1 and 9502760) and the Faculty of Science, University of Aarhus. Th e pape r wa s writte n whil e i n receip t o f funding b y EFP-200 0 (Projec t ENS-1313/00-0001) . Discussions wit h N. Balling, O . R. Clausen, P. Japsen and S . B . Nielse n ar e greatl y appreciated . I than k A. Hurst, G. E. Paulsen and the referees for comments, which helped improv e th e manuscript.

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Tectonic impact on sedimentary processes during Cenozoic evolution of the northern North Sea and surrounding areas JAN INGE FALEIDE 1, RUN E KYRKJEB0 2, TOMAS KJENNERUD 3, ROY H. GABRIELSEN 2, HENRI K JORDT 1'4, STEIN FANAVOLL 3'5 & MORTEN D . BJERKE 1'6 1 Department of Geology, University of Oslo, P.O. Box 1047, Blindern, N-0316 Oslo, Norway (e-mail: j.i.faleide@ geologi.uio.no) 2 Geological Institute, University of Bergen, Allegaten 41, N-5007 Bergen, Norway 3 SINTEF Petroleum Research, N-7465 Trondheim, Norway 4 Present address: Aarhus Amt, Lyseng Alle 1, DK-8270, H0jbjerg, Denmark 5 Present address: IPRES International Ltd., Nedre Vollgt. 4, N-0158 Oslo, Norway 6 Present address: PGS, Strandveien 4, N-1366 Lysaker, Norway Abstract: Thi s pape r focuse s o n th e Cenozoi c evolutio n o f th e norther n Nort h Se a an d surrounding areas, with emphasis o n sediment distribution, composition an d provenance, a s well a s o n timing , amplitud e an d wavelengt h o f differentia l vertica l movements . Quantitative informatio n abou t palaeo-wate r dept h an d tectoni c vertica l movement s ha s been integrate d wit h a seismi c stratigraphi c framewor k t o bette r constrai n th e Cenozoi c evolution. Th e data an d modelling result s suppor t a probable tectoni c contro l o n sedimen t supply an d o n the formatio n o f regional unconformities . Th e sedimentar y architectur e an d breaks ar e relate d t o tectoni c uplif t o f surroundin g clasti c sourc e areas , thu s the offshor e sedimentary recor d provide s th e bes t ag e constraint s o n Cenozoi c exhumatio n o f th e adjacent onshor e areas . Tectoni c subsidenc e accelerate d i n Paleocen e tim e throughou t the basin, with uplifted area s to the east and west sourcing prograding wedges, which resulted in large depocentres clos e to the basin margins. Subsidence rates outpaced sedimentatio n rates along th e basi n axis , an d wate r depth s i n exces s o f 60 0 m ar e indicated . I n Eocen e time s progradation fro m th e Eas t Shetlan d Platfor m wa s dominan t an d majo r depocentre s wer e constructed in the Viking Graben area, with deep water along the basin axis. At the EoceneOligocene transition , souther n Norwa y an d th e easter n basi n flan k becam e uplifted . Th e uplift, i n combinatio n wit h progradin g unit s fro m bot h th e eas t an d west , gav e ris e t o a shallow threshold in the northern North Sea, separating deeper water s to the south and north. The uplift an d shallowing continued into Miocene tim e when a widespread hiatu s formed in the norther n Nort h Sea , as indicate d by biostratigraphi c data . The Pliocen e basi n configuration was dominated by outbuilding of thick clastic wedges from the east and south. Considerable lat e Cenozoi c uplif t o f th e easter n basi n flan k i s documente d b y th e stron g angular relationshi p an d tiltin g o f th e complet e Tertiar y packag e belo w th e Pleistocen e unconformity. Cenozoi c exhumatio n is documented o n both side s o f the North Sea , but the timing is not well constrained. Two major uplift phases in early Paleogene an d late Neogen e times ar e relate d t o rifting , magmatis m an d break-u p i n th e N E Atlanti c an d isostati c response t o glacia l erosion , respectively . Additional uplif t event s may be related t o mantle processes an d the episodi c behaviou r of the Icelan d plume.

The norther n North Se a rif t basi n (Fig . 1 ) has & va n Hoor n 1989 ; Gabrielse n e t al 1990 ; been affecte d b y tw o majo r episode s o f riftin g Robert s et al. 1990;Ziegler 1990, 1992; Yielding since Devonian time. The two events took place et al . 1992 ; Milto n 1993 ; Ratte y & Ha y ward in Permia n t o earlies t Triassi c an d lat e Mid - 1993 ; N0ttved t e t a l 1995 ; Faerset h 1996 ; Jurassic t o earlies t Cretaceou s times , eac h Faerset h e t al . 1997) . Mos t o f th e post-rif t followed b y period s o f post-rif t therma l relax - tectoni c subsidenc e related t o th e Lat e Jurassic ation and subsidence (Badley et al. 1988;Ziegle r rif t phas e had cease d a t th e en d o f Cretaceous From: DORE , A.G., CARTWRIGHT, J.A. , STOKER, M.S. , TURNER, J.P. & WHITE , N . 2002 . Exhumation o f th e North Atlantic Margin: Timing, Mechanisms and Implications for Petroleum Exploration. Geological Society, London, Special Publications, 196, 235-269 . 0305-8719/02/$15.0 0 © The Geological Society of London 2002.

J. I . FALEIDE ETAL.

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V

^

Fig. 1 . Regional settin g an d locatio n o f stud y are a in the norther n Nort h Sea .

time an d th e basi n ha d becom e a wid e are a o f deposition wit h low relief. Henc e i t is suggeste d that therma l equilibriu m wa s reache d i n lates t Cretaceous tim e (Gabrielse n e t al. 2001). The Cenozoic norther n Nort h Se a basin form s a wid e sagge d depocentr e containin g u p t o c. 2500 m o f sediments . Shift s i n depocentr e locations, outbuildin g direction s an d sedimen t composition hav e bee n relate d t o differentia l vertical movement s o f th e basi n floo r an d surrounding clasti c sourc e area s (Nielse n e t al . 1986; Rundber g 1989 ; Rundber g & Smalle y 1989; Gallowa y e t a l 1993 ; Jord t e t a l 1995 , 2000; Michelsen e t al 1995 , 1998 ; Clause n et al 1999b; Martinse n e t a l 1999 ; Thyber g e t a l 2000). Basi n modellin g (bot h backwar d an d forward) ha s bee n applie d t o quantif y th e syn and post-rift evolution o f the northern Nort h Se a basin (e.g . Jo y 1992 , 1993 ; Robert s e t al 1993 ; White & Latin 1993 ; Hal l & White 1994 ; Nadi n & Kusznir 1995, 1996) . However, an d as pointed out b y thes e workers , th e inpu t parameters , an d especially th e assesse d palaeo-wate r depth , ar e

associated wit h uncertaint y an d ma y hav e resulted i n erroneous subsidenc e histories . This pape r focuse s o n the Cenozoic evolutio n of the northern North Sea and surrounding areas, with emphasi s o n th e timing , amplitud e an d wavelength o f differentia l vertica l movements . Quantitative informatio n abou t palaeo-wate r depth an d tectoni c vertica l movement s ha s been integrated wit h a seismi c stratigraphi c frame work t o better constrai n the Cenozoic evolution. We will also briefly discuss the main mechanisms responsible for the tectonic evolution of the area. The main study area of the northern North Sea is restricte d t o 1°W-5° E an d 58-62° N (Fig s 1 and 2) . However , when discussin g th e Cenozoi c sediment distribution and provenance we have to include th e surroundin g lan d area s (Fig . 1) . Cenozoic exhumatio n i s documente d o n bot h sides of the North Sea, mainly from geomorpho logical an d apatit e fission-trac k studie s (Gree n 1986, 1989 ; Bra y et al 1992 ; Lewis et al 1992 ; Holliday 1993 ; Rohrma n et al 1995 ; Rii s 1996 ; Lidmar-Bergstrom e t a l 2000 . A t leas t tw o significant episode s o f Cenozoi c exhumatio n have bee n suggested , bu t th e timin g is no t wel l constrained an d i n man y place s i t i s difficul t t o separate th e tw o (Japse n 1997 ; Japse n & Chalmers 2000) . Th e Cenozoi c developmen t of the stud y area i s also linke d to the plate tectonic evolution o f th e Nort h Atlanti c (Talwan i & Eldholm 1977 ; Eldhol m e t al 1990 ; Dor e e t a l 1999). The wor k ha s bee n carrie d ou t a s par t o f th e project Tectoni c impac t o n sedimentar y pro cesses in the post-rift phase—improve d models' , which focuse d o n th e Cretaceous-Cenozoi c succession filling in the structural relief resulting from Lat e Jurassic-earliest Cretaceous riftin g i n the norther n Nort h Se a (Fig . 3) . Thi s pape r summarizes th e Cenozoi c par t o f thi s study , an d Gabrielsen e t al (2001 ) have given a summary of the Cretaceou s post-rif t developmen t i n th e northern Nort h Sea . Seismic mapping In thi s stud y w e hav e interprete d a gri d o f regional high-qualit y seismi c reflectio n profile s tied t o ke y wells . The mai n databas e comprise d eight regional seismi c reflectio n surveys , providing the best dat a coverage i n the Norwegian part of th e norther n Nort h Sea . I n addition , fou r regional dee p seismi c reflection lines were used . Regional crusta l transect s (Fig . 2 ) wer e con structed alon g thes e line s b y combinin g the m with conventiona l seismi c line s an d gravit y an d magnetic data . Th e fou r transect s wer e depth converted using velocity information from wells ,

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Fig. 2 . Structural map of stud y are a in the northern North Sea and location of regional transects and key wells. ESB, East Shetland Basin; ESP, Eas t Shetland Platform; HP , Horda Platform; LT , Lomre Terrace; MFB, Mal0 y Fault Blocks; MgB, Magnus Basin; MrB, Marulk Basin; SB, Stord Basin; SG, Sogn Graben; TS, Tampen Spur; UH, Utsira High; UT, Uer Terrace; VG, Viking Graben; WG, Witchground Grabe n ;AG,Asta Graben.

interval velocitie s fro m stackin g velocitie s an d velocities fro m dee p seismic refractio n data. The transects wer e used in the modelling . Twelve well s (Fig . 2 ) wer e selecte d t o b e principal sources for analysis on the basis of their location a t or close t o the regiona l transect s an d as representative s o f differen t structura l pos itions. The key well s wer e analyse d to obtai n information abou t ag e an d lithologica l compo sition o f th e mai n Cenozoi c sequences , t o

recognize stratigraphica l breaks , an d to estimate palaeo-water depths. In addition, biostratigraphic data and well logs from mor e than 60 wells were used to calibrate th e interpretation of the seismi c data. The Cenozoic seismic stratigraphi c framework is based o n the work byJordt et al (1995 , 2000) . The Cenozoi c successio n i s subdivide d int o 1 0 seismic sequence s (CSS- 1 t o CSS-10 ) (Fig . 4). The sequence s hav e bee n date d usin g biostrati-

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Fig. 3 . Regiona l seismi c lin e NVGTI-92-105 . (Se e Fig . 2 fo r lin e locatio n an d Fig . 4 fo r Cenozoi c seismi c stratigraphy.) BT , Base Tertiary; MC, mid-Cretaceous ; BC , Bas e Cretaceous ; twt . two-wa y trave l time .

Fig. 4 . Cenozoi c seismi c stratigraphi c framework , based o n Jordt e t al (1995 , 2000) .

graphic dat a fro m ke y well s publishe d b y Steurbaut e t a l (1991) , Eidvi n & Rii s (1992) , van Vee n e t al . (1994) , Gradstei n & Backstro m (1996), Martinsen e t al (1999 ) and Eidvin et a l (1999, 2000) , i n additio n t o th e 1 2 ke y well s o f the TecSe d projec t (Fig . 2) . Th e seismi c stratigraphic framewor k wa s relate d t o th e tim e scale o f Gradstei n & Og g (1996 ) an d forma l lithostratigraphy o f th e norther n Nort h Se a (Isaksen & Tonsta d 1989 ; Kno x & Hollowa y 1992). The Paleogen e successio n comprise s fou r seismic sequence s (CSS-1 , CSS-2 , CSS- 3 an d CSS-4). Th e CSS- 1 sequenc e i s o f Lat e Paleocene-earliest Eocen e ag e an d it s to p corresponds t o th e to p o f th e Balde r tuffs . Th e CSS-2 sequenc e i s o f Eocen e age . Th e seismi c boundary a t the top o f CSS-2 correlate s wit h the Eocene-Oligocene transition , whic h i s associ ated wit h a hiatus , i n particula r alon g th e basi n flanks. Th e CSS- 3 sequenc e cover s a narro w period i n Earl y Oligocen e time . A n upwar d change i n seismic signatur e from progradation t o marked aggradatio n an d onlap occurs agains t the top o f CSS- 3 i n the norther n Nort h Sea . A mid Early Oligocene ag e is inferred for this sequence boundary. The to p of sequence CSS- 4 i s of latest Oligocene age .

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Fig. 5 . Regional profiles 1- 5 across the northern North Sea. (See Fig. 2 for profile location s and Fig. 4 for colour codes of the Cenozoic seismi c sequence s CSS- 1 to CSS-10.)

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The Miocen e successio n i s divided int o thre e seismic sequence s (CSS-5 , CSS- 6 an d CSS-7) , and the boundaries separatin g thes e are dated i n the centra l Nort h Se a (Jord t e t al 1995; Michelsen e t al . 1995) . Miocen e sediment s ar e also presen t i n th e norther n Nort h Sea , bu t i t i s difficult t o mak e thi s subdivisio n her e becaus e some o f th e sequence s thi n belo w seismi c resolution. Mos t worker s agre e o n a prominen t mid-Miocene hiatus ; however , th e biostratigra phy is equivocal (Steurbaut et al. 1991 ; Eidvin & Riis 1992 ; Gradstein & Backstro m 1996; Martinsen e t al . 1999 ; Eidvi n e t al . 2000 ; se e Jordt e t al. 2000 for discussion) . The Plio-Pleistocen e sequence s (CSS-8 , CSS-9 an d CSS-10 ) consis t mostl y o f glacia l sediments. The CSS- 8 sequenc e is of Pliocen e age (Eidvin & Riis 1992 ; Gradstein & Backstrom 1996; Eidvi n et al. 2000) whereas the CSS-9 and CSS-10 sequence s ar e of Pleistocene age . The seismi c interpretatio n focuse d o n identi fication o f sequenc e geometries , locatio n o f

depocentres, outbuildin g directions , recognitio n of tectoni c influenc e an d establishmen t o f a general Cenozoic framewor k for further analyses and basi n modelling . Seismi c sequenc e geome tries an d outbuildin g direction s provid e infor mation abou t change s i n th e basi n topography , and the y ar e relate d t o underlyin g structures, shifts i n provenance area an d changes in relative sea leve l an d sedimen t accumulatio n rates. Fo r each seismi c sequenc e w e hav e constructe d time-thickness maps.

Palaeo-water depth Temporal an d spatia l variation s i n palaeo-wate r depth are crucial parameters i n basin analysis, as changes i n palaeobathymetr y detai l th e amoun t of sedimen t underfil l durin g basi n evolutio n (Gradstein & Backstro m 1996) . I n forwar d o r backward basin modelling, palaeo-water depth is an importan t inpu t paramete r tha t control s th e measured subsidenc e o r uplif t an d thu s th e

Fig. 6 . Palaeo-water depth summary for the Cenozoic sequence of well 30/10-6 (Viking Graben). Modified fro m Kyrkjeb0 e t al. (2001) . Dominant lithologies: shales-mudstone s i n green; sandstones i n yellow. Sedimentatio n rates in mm ka" 1 ar e also shown. (Se e Fig. 4 for Cenozoic seismic stratigraphy. ) K-6 , Maastrichtian.

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1

Fig. 7 . Cenozoi c palaeo-wate r dept h fo r selecte d Vikin g Graben wells , modifie d fro m Kyrkjeb 0 e t al . (2001) . Long-term eustati c curv e fro m Ha q e t al . (1987 , 1988 ) i s show n i n magent a fo r comparison . (Se e Fig . 2 fo r location o f wells.)

Fig. 8 . Tectoni c modelling and subsidenc e analysi s along Transec t 1 (see Figs 2 an d 5) . Tectoni c subsidenc e curves fo r selecte d position s ar e show n abov e th e profile . Relative vertical movement s indicated by arrows (red, uplift; blue , subsidence). MgB, Magnus Basin; ESB, East Shetland Basin; TS, Tampen Spur ; VG, Viking Graben; LT, Lomre Terrace ; HP , Horda Platform. Modified fro m Kyrkjeb 0 et til. (1999) .

CENOZOIC EVOLUTION OF THE NORTHERN NORTH SEA recognition of tectonic events. However, palaeowater depth is difficult t o determine, and depends largely o n the qualit y an d quantit y of data fro m boreholes an d seismi c dat a (Bertra m & Milto n 1989; Joy 1992 , 1993 , 1996 ; Hall & White 1994 ; Jones & Milto n 1994 ; Nadi n & Kuszni r 1995 , 1996). By carefull y integratin g seismic stratigraphi c observations wit h palaeo-wate r dept h estimate s from structura l restoratio n an d micropalaeonto logical data , change s i n accommodatio n spac e throughout Cretaceous-Tertiar y time s ca n b e documented o n a regiona l scal e i n th e norther n North Se a (Fig s 6 an d 7 ) (Kyrkjeb 0 e t al 2001). The palaeorelie f wa s restore d alon g th e regional transect s (Fig s 2 an d 5 ) usin g th e depositional geometries , indication s o f zer o o r near-zero wate r dept h (suc h a s subaeria l unconformities an d coals ) an d faul t restoratio n (Kjennerud e t al . 2001) . Th e metho d provide s estimates o f palaeorelie f alon g th e transect s a t the base of each seismic unit rather than absolute palaeo-water depth . I n mos t case s thi s basi n relief could be characterized as a minimum water depth. Subaeria l topograph y canno t b e deter mined b y thi s method . Progradin g sequences , which are characteristic for parts of the Cenozoic development i n th e norther n Nort h Sea , ar e associated wit h th e leas t uncertaint y i n th e restorations. The 1 2 key well s fro m th e Norwegia n part of the norther n Nort h Se a (Fig . 2 ) wer e studie d i n great detai l t o asses s palaeo-wate r depths . Th e micropalaeontological analysi s (Gillmor e e t al . 2001) gav e a rang e estimat e o f palaeo-wate r depth an d i n some case s additiona l maxima and minima ar e included (Fig . 6). Next, a most likely trend through time was determined for each well by integratin g th e result s fro m th e structura l restoration an d th e micropalaeontologica l anal ysis (Fig . 6) . A s i t i s no t possibl e t o determin e exactly the palaeo-water depth, we have focused on determinin g most likel y dept h intervals , and identifying th e principal shallowing and deepening trends. In the seismic interpretation i t was not possible t o differentiat e betwee n th e Miocen e seismic unit s CSS-5 , CSS- 6 an d CSS- 7 i n th e northern Nort h Sea . Therefore , palaeo-wate r depth estimate s wer e no t obtainabl e fo r eac h Miocene uni t by structura l restoration . Furthermore, onl y a few sample s o f Miocen e ag e have been availabl e fo r micropalaeontologica l anal ysis. The sensitivity for shallowing o r deepenin g trends i s considere d t o b e n o bette r tha n 100 m (Kyrkjeb0 e t al 2001) . The inferre d Cenozoi c deepenin g o r shallow ing trend s fro m th e investigate d well s ar e

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generally i n good agreemen t wit h each othe r on a regiona l scale , especiall y whe n th e tectoni c position withi n th e basi n i s take n int o accoun t (Fig. 7). The inferred general trends in Cenozoic time are (Kyrkjeb0 et al. 2001): (1) deepening in Early t o Lat e Paleocen e time ; (2 ) shallowin g from Lat e Eocen e t o Lat e Miocen e time ; (3 ) deepening fro m Lat e Miocen e to Early Pliocen e time; (4) shallowing during Pliocene time . Most worker s agre e on falling global eustati c sea leve l throughou t Cenozoi c time s (Pitma n 1978; Watt s & Steckle r 1979 ; Ha q et al . 1987 , 1988). However , th e suggeste d amplitude s vary between abou t 10 0 an d 300m , whic h ar e les s than th e suggeste d amplitude s of th e deepenin g or shallowing trend (Fig . 7). The Late Eocene to Late Miocen e shallowin g correlate s wit h th e long-term eustati c curv e o f Ha q e t al . (1987 , 1988), bu t th e genera l shallowin g tren d wa s probably amplifie d b y tectono-therma l effects . The deepening events in Early to Late Paleocen e and Lat e Miocen e t o Pliocen e time s canno t b e explained b y th e eustati c sea-leve l curve , an d must therefor e b e explaine d b y purel y tectono thermal event s (Kyrkjeb 0 e t a l 2001) . Fro m Mid-Eocene time , ther e i s a reasonabl e corre lation betwee n eustati c curve s derive d fro m sequence stratigraphi c studie s (e.g . Ha q e t a l 1987, 1988 ) an d th e composit e oxyge n isotop e record reflecting variations in global climate and glaciations (Abreu & Anderson 1998) . Modelling The Cenozoi c deepenin g o r shallowin g trend s summarized above clearly point towards tectonothermal event s affectin g Cenozoi c basi n evol ution. On e o f th e mai n objective s o f ou r stud y was to constrain the amplitude and wavelength of differential vertica l movements . Togethe r wit h timing, these are critical factors to a discussion of possible mechanism s fo r Cenozoi c vertica l movements. To asses s tectoni c subsidenc e an d uplif t i n a more quantitativ e way , w e hav e carrie d ou t modelling alon g th e fou r regiona l transect s (Figs 2 , 5 an d 8 ) (Kyrkjeb 0 e t a l 1999) . Conventional backstrippin g technique s (e.g . Steckler & Watts 1978 ; Watt s et al 1982 ; Allen & Allen 1990 ; Robert s e t al 1993 , 1998 ; Kuznir et a l 1995 ; Nadi n & Kusznir 1995, 1996 ) were applied to the transects after key parameters such as age , lithology , porosity-dept h an d palaeo water depth for each seismi c sequenc e had bee n constrained from analysi s of the seismic and well data. Th e lithospheri c respons e t o loadin g o r unloading of sediment and water was assumed to be compensate d fo r by local Air y isostasy .

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Vertical movement s differen t fro m thos e predicted b y therma l contractio n an d sedimen t and wate r loadin g wer e regarde d a s potentia l tectonic signals . Kyrkjeb 0 et al (1999 ) regarde d events to be truly tectonic whe n the amplitude of the vertical movement s exceed s th e uncertainties related t o changes in palaeo-water depth adjusted for eustati c sea-leve l change s accordin g t o Ha q et al. (1987, 1988) . Figur e 9 shows how change s in palaeo-water dept h ar e crucial for detection of tectonic event s othe r tha n thermal cooling . The tectoni c subsidenc e analysi s o f th e northern Nort h Se a point s t o th e followin g Cenozoic trend s (Kyrkjeb 0 e t al . 1999 ) (Fig s 8 and 9) : (1 ) accelerate d tectoni c subsidenc e i n Paleocene time; (2 ) standstill to uplift i n Eocen e time followe d b y uplif t i n Oligocen e time ; (3 ) tectonic subsidenc e durin g Miocen e tim e fol lowed b y uplift i n Pliocene time . Kyrkjeb 0 et a l (1999) emphasize d tha t these trend s ar e general , and that local variation s occur . The genera l lac k of palaeo-wate r dept h informatio n fo r th e Miocene perio d affect s th e resolutio n i n thes e

Fig. 9. Typica l palaeo-wate r dept h and tectoni c subsidence curve s i n the northern Nort h Se a compare d with standar d rif t mode l (post-rif t therma l cooling) .

results. On the basis of evidence discussed below , we believ e tha t th e chang e fro m uplif t t o subsidence occurre d i n Mid-Miocen e tim e an d the subsequen t change fro m subsidenc e to uplif t probably too k plac e i n Late Pliocen e time .

Cenozoic evolutio n Here w e integrat e th e result s o f th e studie s described abov e an d summariz e th e Cenozoi c evolution o f th e norther n Nort h Se a an d surrounding area s usin g seismic sections , time thickness maps an d palaeo-water dept h maps. Paleocene time The CSS- 1 sequenc e (Uppe r Paleocene-lower most Eocen e units ) i s typifie d b y progradin g wedges tha t buil t ou t fro m th e Eas t Shetlan d Platform an d from souther n Norway. Lowermos t CSS-1, markin g th e beginnin g o f th e Tertiar y sequence, was probably characterized by shallow water dept h an d littl e relie f (Fig . lOa) . Th e lowermost Paleocen e brea k (66-6 2 Ma) i s interpreted to represent both an erosional vacuit y and a hiatus . Par t o f th e brea k i s probabl y a marine condensatio n o f sedimentatio n befor e progradation o f th e Paleocen e depositiona l systems fro m th e eas t an d th e wes t (Martinse n et al. 1999) . The prominen t depocentre s alon g th e wester n basin margi n (Fig . lOc ) wer e mainl y source d from th e uplifted Eas t Shetlan d Platform an d the Scottish Highlands . Th e progradin g shelf-slop e system i s i n par t fairl y san d ric h an d severa l phases of sand deposition were related to tectonic uplift an d erosio n o f th e sourc e area . Th e sequence geometrie s reflec t tha t som e o f th e Mesozoic grabe n faults , particularl y alon g th e western margin of the East Shetlan d Basin, were reactivated an d that differentia l compactio n too k place ove r deeper-seate d Mesozoi c faul t blocks . The syndepositiona l faulting wa s not related to a new phas e o f rifting , bu t probabl y t o regiona l subsidence. Th e shap e o f th e depocentr e wa s defined b y a combinatio n o f increase d sub sidence towards the basin centre, and differentia l compaction o f th e Mesozoi c sediment s i n th e graben relativ e t o thos e o n th e platfor m area s (Milton e t al. 1990) . A depocentr e i n th e northeaster n par t o f th e North Sea (Fig. lOc ) was probably source d fro m mainland Norway . I t ha s a progradationa l stacking patter n tha t thicken s pronouncedl y eastward althoug h i t ha s bee n subjecte d t o late r erosion (Fig . 11) . Thi s uni t i s characterize d b y more fine-graine d lithologies . Th e cla y minera l distribution show s a n increasin g kaolinite -

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Fig. 10. (a) Palaeo-water depth map for earliest Paleocene time (lowermost CSS-1); (b) palaeo-water depth map for earlies t Eocene time (lowermos t CSS-2); (c ) time-thickness map fo r CSS-1 (Uppe r Paleocene-lowermost Eocene sequence) ; (d) time-thickness map for CSS-2 (Eocen e sequence).

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Fig. 11 . Seismic line NVGTI-92-105 showing CSS- 1 building ou t from the east i n northernmost Nort h Sea . (Se e Fig. lO c fo r line location and Fig . 4 for Cenozoi c seismi c stratigraphy. )

smectite t o illit e rati o toward s th e depocentr e (Rundberg 1989 ) indicatin g tha t these sediment s were derived fro m the east (Thyberg et al 2000) . A relativ e sea-leve l ris e i n latest Paleocen e earliest Eocen e tim e i s indicated b y aggradatio n and even sedimen t thicknes s i n th e uppermos t part o f CSS-1 , whic h include s th e Balde r Formation, comprisin g interbedde d shale s an d tuff layers . Explosiv e volcanis m i n th e Britis h and th e Faeroe-Greenlan d Tertiar y volcani c provinces has been regarded a s the source for the widespread volcani c ashe s foun d i n th e Nort h Sea (Mal m e t a l 1984 ; Kno x & Morto n 1988 ) and onshor e Denmar k (Spjeldnae s 1975 ; Nielsen & Heilmann-Clause n 1988) . The highe r smecti tic conten t i n th e uppe r par t o f th e CSS- 1 sequence indicate s weatherin g o f a basalti c source rock . Th e mos t likely provenanc e are a for thes e sediment s i s the Late Paleocene-Earl y Eocene basal t province alon g the North Atlantic rift zon e (Fig . 12 ) in additio n t o th e widesprea d tuff deposits . Uppermos t CSS- 1 probabl y indi cates a condense d sectio n relate d t o period s of low clasti c sedimen t suppl y an d a starve d depositional environment. The sequenc e geometrie s an d thicknes s distribution o f CSS- 1 indicat e tha t dee p marin e conditions existe d alon g th e Vikin g Grabe n an d

towards th e continenta l margi n i n th e north . A deepening during Paleocene time , reaching water depths of about 800 m in the deepest part s of the basin, ha s bee n inferre d fro m biostratigraphi c data (Gradstei n e t al . 1994 ; Gradstei n & Backstrom 1996 ; Gillmor e et al 2001 ; Kyrkjeb0 et al 2001 ) and structural restoration (Kjenneru d et a l 2001 ) (Fig s 7 an d lOb) . A s th e inherite d post-Cretaceous bathymetry along the North Sea basin margin areas was not significant (Fig . lOa) , the accommodation and development of the thick Upper Paleocen e wedges probably resulte d fro m subsidence o f th e basi n floo r combine d wit h increased sedimen t suppl y relate d t o tectoni c uplift and subsequent erosion o f the adjacent land areas. Shallo w marin e condition s probabl y prevailed alon g th e basi n margin s i n th e wes t and east. A delta succession characterized by the presence o f coa l an d i n situ lignites , a s wel l a s freshwater floras and faunas, has been reported in the Mora y Firt h (Andrew s et al 1990) . Uplift abov e se a leve l o f th e Hebrides Shetland axi s t o th e N W of the Nort h Se a basi n was accompanie d b y th e outbrea k o f extensiv e volcanicity i n th e area s surroundin g the presen t NE Atlanti c Ocean . Th e uplif t establishe d a wholly ne w geograph y (Fig . 12) . An easterl y to southeasterly flowin g drainag e syste m becam e

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Fig. 12. Regional setting in Late Paleocene-earliest Eocene time. 1 , early Tertiary flood basalt province; 2, areas of uplift; 3 , intrabasinal highs; 4, main depocentres; 5 , plate boundary-line o f breakup; 6, V0ring and FaeroeShetland escarpments; 7, outbuilding directions; 8, Sorgenfrei-Tornquist Zone-Fennoscandia n Border Zone.

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established o n th e Orkney-Shetlan d Platfor m and Scottis h Highlands , resultin g i n th e rework ing o f th e sedimentar y cove r ove r th e newl y emergent terrains , an d dispersa l o f thei r detritu s into th e Nort h Se a basi n (Johnso n e t al 1993 ; Jones & Milton 1994) . The northwester n corne r o f souther n Norwa y (Fig. 12) , o n tren d wit h th e Hebrides-Shetlan d axis, wa s als o uplifted . Th e uplifte d are a wa s bounded b y the 0ygarden an d M0re-Tr0ndelag fault complexe s toward s th e wes t an d NW , respectively. I t extended southward s to about the Sognefjorden area . Northeastward s th e uplifte d area wa s probabl y bounde d b y th e landwar d extension o f th e Ja n Maye n Lineament . Earl y Tertiary volcanis m (c . 5 6 Ma) i s documente d within thi s are a jus t of f th e coas t o f Norwa y (Vestbrona Formation; Bugg e et al. 1980 ; Torsk e & Prestvik 1991 ; Prestvik et al. 1999) . Clast s of Danian chal k hav e bee n foun d i n Pliocen e sediments buildin g out from Norway i n this are a (Thyberg e t al . 2000) , indicatin g that n o clasti c source are a existe d her e befor e th e Lat e Paleocene uplift . Most o f th e sediment s derive d fro m th e uplifted are a o n th e Norwegia n sid e wer e deposited i n th e northeaster n Nort h Se a an d southeastern margin of the M0re Basi n (Fig. 12) . However, som e sediment s ma y hav e bee n transported southwards , wher e onl y a smal l portion reached the North Sea in the NorwegianDanish Basin. Here, a Late Paleocene depocentr e (Fig. 12 ) reflects outbuilding fro m th e NE . Th e source are a fo r th e CSS- 1 sequenc e i n thi s are a may also have include d th e uplifted Sorgenfrei Tornquist Zon e (STZ)-Fennoscandia n Borde r Zone (FBZ) , whic h wa s inverte d i n respons e t o Alpine compression . Th e uplifte d STZ-FB Z must hav e bee n a barrie r fo r an y sediment s coming fro m souther n Scandinavia . Erosiona l valleys an d kars t topograph y sugges t tha t th e uppermost chal k wa s subaeriall y expose d i n the eastern Nort h Se a i n mid-Paleocen e time s (Huuse 1999 ; Clause n & Huus e 1999) . Th e fossil conten t i n th e Danis h Fu r Formatio n indicates tha t th e lowermos t Eocen e sediment s were deposite d relativel y clos e to a coastlin e i n the Norwegian-Danish Basin . Erosional product s (sands ) wer e also trans ported northwestwar d fro m th e uplifte d Heb rides-Shetland are a int o th e Faeroe-Shetlan d Basin (Fig . 12) . Renewe d tectoni c movemen t allowed depositio n adjacen t t o majo r faults , bu t Late Paleocen e an d Earl y Eocen e event s wer e dominated b y regiona l uplif t an d massiv e volcanism associated wit h North Atlantic rifting. Towards th e en d o f Paleocen e times , transfe r faulting ha d largely ceased, an d the northwestern

margin o f th e Faeroe-Shetlan d Basi n wa s inundated b y subaeriall y emplace d lava s o f th e Faeroes lowe r series , an d t o th e N E th e Erlen d and West Erlend volcanoes were erupting (Stoker etal 1993) . Late Paleocene-earlies t Eocen e depocentre s in th e wester n M0r e an d V0rin g basin s wer e sourced fro m a n uplifte d are a i n th e wes t alon g the incipien t plat e boundar y (Fig . 12) . O n th e M0re Margina l High , Paleocen e deposit s ar e thin, typicall y 200-350m , an d lithologie s consist mainl y o f lava s an d shallow-wate r sediments. A thic k wedg e i n th e M0r e Basi n was source d fro m th e west. Aggradation of sequence CSS-1.2 (Jordt et al. 1995) indicate s tha t sedimen t suppl y fro m Norway wa s significantl y reduce d i n earlies t Eocene time , and tha t th e lan d areas furthe r eas t were subjecte d t o a marin e transgression . Reduced runof f cause d b y climati c change , changes i n lan d vegetation , tectonic quiescence or erosiona l levellin g o f souther n Norway , ar e possible mechanisms that could explain a sudden reduction in sediment supply from th e east in the very earlies t Eocene tim e (Jordt et al 2000) . Eocene time There i s a composite brea k i n the Earl y Eocen e deposition correspondin g t o th e CSS-l-CSS- 2 boundary. Th e biostratigraph y indicates that the break spans the period 55-52 Ma. The lower part of the break (c. 55-54 Ma) is laterally extensive. The uppe r par t o f th e brea k (c . 54-5 2 Ma) seemingly ha s a limited latera l extent , and in the North Se a basi n i s accompanie d b y a shif t o f sedimentation i n th e basinward s direction , an d subsequent onla p ont o th e basi n margin . Thus , the brea k increase s i n exten t toward s th e east . This Earl y Eocen e brea k i s interprete d t o represent bot h a n erosiona l vacuit y (althoug h evidence is limited) and a hiatus (Martinsen et al. 1999). The Eocen e (CSS-2 ) depocentre s (Fig . lOd ) are mainl y locate d i n th e centra l an d wester n parts of the northern North Sea basin and indicate outbuilding int o the basin fro m the uplifted East Shetland Platform . CSS- 2 infill s an d drape s th e topography o f to p CSS- 1 an d th e depocentr e i s located basinward s o f th e forme r CSS- 1 shel f edge. Eocene depositio n was dominated by deepwater an d slop e processe s includin g turbidit y currents. Dee p water existe d throughou t mos t of Eocene time s (Fig s lO b an d 13a) . Water depth s of abou t 1000 m hav e bee n estimate d fo r th e Lower Eocen e Frig g Formation , whic h wa s deposited a s a submarin e fa n i n a deep-se a environment (Heritie r et al . 1979 , 1981) .

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Fig. 13 . (a) Palaeo-water dept h map for latest Eocene tim e (lowermos t CSS-3); (b) palaeo-water depth map for Early Oligocen e tim e (lowermos t CSS-4); (c) time-thickness map for CSS-3 (Lowe r Oligocene sequence) ; (d) time-thickness map for CSS-2 (Uppe r Oligocen e sequence) .

250

J. I. FALEIDE ETAL.

CSS-2 consist s o f stacked , thickl y bedde d immature sandstone s wit h interbeds o f clay- and siltstone a t th e margi n o f th e Eas t Shetlan d Platform an d i n th e Mora y Firt h Basi n an d generally of claystones in the Viking Grabe n an d Central Grabe n area s (Mudg e & Buja k 1994) . Lignite layer s penetrated i n well 8/27a-l poin t to the existenc e o f swamp y coasta l lake s landwar d of th e inne r shel f (Veeke n 1997) . The Eocen e successio n thin s eastwards an d is almost absen t i n som e section s clos e t o th e eastern basi n margin . However , a mino r Earl y Eocene (CSS-2.1 of Jordt et al 1995 ) depocentr e off Sognefjorden , i n th e sam e are a a s fo r th e underlying CSS- 1 sequence , comprise s a pro gradational uni t tha t downlap s ont o th e upper most Balde r surface . Th e outbuildin g o f th e lower part o f sequence CSS- 2 i n the northeastern North Se a an d depositio n o f san d alon g th e eastern basi n margi n indicat e renewe d tectoni c uplift an d sedimen t suppl y fro m th e eas t (Rundberg 1989 ; Jord t e t al 1995 , 2000) . From mid-Eocen e tim e onwards , Tertiar y sedimentation i n th e centra l area s o f th e Nort h Sea basi n wa s dominate d b y monotonou s sequences o f marin e mud s an d silts , althoug h sands wer e deposite d a t the periphery. Th e finegrained CSS- 2 sediment s of low rate s of non volcanic clasti c sedimen t suppl y indicat e a relatively dee p marin e facies . Mor e o r les s starved sedimentatio n conditions , as a result of a relative rise in sea level, favoured the enrichmen t in bot h volcaniclasti c an d organi c matte r (Thyberg e t al . 2000) . Th e unifor m litholog y and compositio n o f th e Eocen e sediment s (smectitic mudstones ) i n th e northeaster n Nort h Sea (Rundber g 1989 ; Thyber g e t al 2000) , and the velocit y distributio n o f sequenc e CSS- 2 (Jordt e t a l 2000) , suggest s a rathe r unifor m source an d sedimen t suppl y fro m a n are a dominated b y basalti c rocks , i.e . th e Nort h Atlantic basal t provinc e t o th e N W (Thyber g et al 2000) . The dominance of fine-grained Middle-Upper Eocene deposit s i n easter n Nort h Se a well s (Rundberg 1989 ; Steurbau t et al 1991 ; Mudge & Bujak 1994 ; Michelse n e t a l 1995) , an d th e regional thinnin g o f CSS-2. 2 o n th e Hord a Platform an d furthe r south , suggest s tha t th e Scandinavian topograph y wa s to o limite d t o supply sufficien t clasti c materia l an d t o allo w progradation fro m th e east, and that large parts of Scandinavia, eas t o f th e stud y area , ma y hav e been submerge d i n Mid-Lat e Eocen e time s (Jordt e t al 1995) . During Lat e Eocen e time , a lowerin g o f relative se a leve l (eustati c fal l and/o r uplift? ) promoted th e advanc e o f delt a system s int o th e

North Se a fro m th e west , wit h depositio n o f successions characterize d b y th e presenc e o f i n situ lignite s an d freshwate r flora s an d faunas . Biostratigraphical evidenc e indicate s tha t Uppe r Eocene strata ar e absent i n much of the norther n North Se a an d alon g th e basi n flank s furthe r south. In these areas, Middle Eocene deposit s are overlain unconformabl y b y Lowe r Oligocen e shales (e.g . Gradstei n et al 1994) . The boundary between th e Lowe r an d Middle Eocene sequence s i n th e Norwegian-Danis h Basin i s associate d wit h a chang e i n sedimen t outbuilding directio n fro m th e N E (Norway ) t o the wes t (Britis h Isle s an d Shetlan d Platform ) (Michelsen et al 1995) . A complete deep marine Eocene successio n i s preserve d i n th e Norwe gian-Danish Basi n an d i n th e Centra l Grabe n (Heilmann-Clausen e t a l 1985) . Here , th e Eocene successio n appear s unaffecte d b y erosion. In th e Faeroe-Shetlan d Basin , riftin g an d volcanism waned through Early Eocene tim e and was followe d b y subsidence . Middl e t o Uppe r Eocene sediment s of the Faeroe-Shetland Basin are generall y fine-graine d claystone s an d silt stones wit h occasiona l limestone s an d sand stones. Thes e reflec t th e retur n t o full y marin e conditions afte r a regressiv e phas e i n Lat e Paleocene an d Early Eocen e time s (Stoke r e t a l 1993). In th e M0r e Basin , subsidence durin g Eocen e time allowe d a thic k sequenc e t o accumulate . Wells drille d o n th e flan k o f th e Margarit a Spu r penetrated nearl y 1000 m o f Lowe r t o Middl e Eocene sediments , overlai n unconformabl y b y the Oligocene sequenc e (Stoke r et al 1993) . The thickness increas e o f CSS-2 toward s th e Atlantic continental margi n indicate s sedimen t suppl y from tha t direction . A t th e V0rin g margi n of f mid-Norway, uplif t create d th e mai n wester n source are a fo r th e Paleogen e V0rin g Basi n sediments (Skogsei d & Eldhol m 1989 : Stuevold etal 1992 ; Stuevold & Eldholm 1996 : Hjelstue n etal 1999) . Oligocene time Biostratigraphical dat a indicate a hiatus between the Eocen e (CSS-2 ) an d Oligocen e (CSS-3 CSS-4) sequence s (va n Vee n e t a l 1994 ; Gradstein & Backstro m 1996) . Martinse n e t a l (1999) place d a widesprea d composit e brea k i n Early Oligocen e tim e (33-2 7 Ma), o f variabl e extent dependin g o n location. During Oligocene time (CSS-3 and CSS-4) the basin configuratio n changed . Th e thickes t Oligocene strat a i n th e norther n Nort h Se a basin ar e locate d ove r th e Norwegia n secto r o f

CENOZOIC EVOLUTION OF THE NORTHERN NORTH SEA the Nort h Vikin g Graben , wher e u p t o 900 m accumulated. Isopach s o f Oligocen e deposit s indicate a genera l north-sout h trend , wit h thinning both t o the west an d east (Fig . 13 c and d). The concentratio n o f CSS-3 sediment s in the centre o f th e norther n Nort h Se a basi n an d contemporaneous erosio n o n th e flank s indicat e an increase d topographi c relief , whic h wa s probably cause d b y uplif t alon g th e easter n an d the wester n basi n margin s i n Earl y Oligocen e time. The sequenc e geometrie s reflec t uplif t o f th e eastern basin flank (Fig. 14) . Minor progradation occurred a s a consequenc e o f tectoni c uplif t activity i n th e eas t initiate d i n Earl y Oligocen e time. A n increase d rat e o f inpu t o f coarser grained materia l occurre d i n Earl y Oligocen e (CSS-3) time , probabl y i n response to uplift an d erosion o f soft Eocen e an d Paleocene sediments . However, th e volcani c inpu t i s stil l distinctiv e (Thyberg et al 2000) . Ne w ashe s fro m Icelan d and reworke d Eocen e sediment s fro m Norwa y may hav e contribute d t o th e hig h smectit e content. Parts o f the basi n wer e uplifted together wit h Norway s o tha t CSS- 4 (Uppe r Oligocen e sequence) locall y overlie s CSS- 2 (Eocen e sequence) (Fig s 13 b an d 14) . A mino r dro p i n sea level , leadin g t o truncatio n o f th e Eocene Lower Oligocene sequence(s) , was followed by a renewed ris e tha t resulte d i n onla p o f CSS- 4 deposits. Lat e Oligocen e depositio n wa s characterized b y basina l accumulatio n o f aggradin g silts an d clays . Locall y CSS- 4 build s ou t fro m uplifted sourc e area(s ) withi n th e basin . Th e CSS-4 depocentre s ar e mainly located alon g the eastern basi n margin (Fig . 13d) . Clausen et al. (1999a) studied intraformational faults withi n th e Uppe r Oligocen e sequenc e (CSS-4) i n the Trol l are a o n the norther n Hord a Platform. Th e fault s hav e a strongl y dominan t NW-SE tren d indicatin g SW-N E extension , and forme d abov e a n intra-Oligocen e unconformity tha t was establishe d a s a result o f uplif t o f Fennoscandia an d a sea-level fall . Outbuilding continued from th e East Shetland Platform durin g bot h CSS- 3 an d CSS- 4 times , resulting i n a thic k depocentr e i n th e Vikin g Graben. I n the transgressive phas e followin g th e intra-Oligocene lowstan d o f se a level , a majo r shelf-deltaic forese t uni t u p t o 500 m thic k prograded fro m th e Eas t Shetlan d Platform into the central par t o f the norther n Nort h Se a basin. The Earl y Oligocen e glacio-eustati c sea-leve l fall tha t was caused by expansion of ice sheets in Antarctica (Miller et al. 1987; Zachosef a/. 1992 ; Abreu & Anderso n 1998 ) wa s overprinte d b y local vertica l tectoni c movement s i n th e Nort h

251

Sea area . Palaeo-wate r dept h dat a indicat e shallowing (Fig . 7 ) an d a t th e en d o f Oligocen e time the northern North Sea formed a narrow and shallow basi n separatin g deepe r basin s t o th e south and north (Fig. 13b ) (Thyberg et al. 1999) . In th e Norwegian-Danis h Basin , th e marke d shift fro m Lat e Eocen e dista l deep-wate r sedimentation t o shor e progradatio n wit h sedi ment transpor t fro m nort h an d N E i n earlies t Oligocene tim e indicate s rapid tectoni c uplif t o f southern Norwa y (Fig . 15 ) (Jord t e t al . 1995 , 2000; Michelse n e t al . 1995 ; Danielse n e t al . 1997; Clause n e t al . 2000) . Th e geometr y an d thickness o f CSS-3 indicat e that water depths in the Central Nort h Sea may have exceeded 60 0 m in Earl y Oligocen e tim e (Jord t e t al . 1995 ; Michelsen et al. 1995). The presence of reworked Paleocene and Eocene nannofossils in Oligocene sediments is interprete d to be the resul t of denudation o f Paleocene-Eocen e sediment s exposed a t the basin margin to the east (Clausen et al. 2000) . The Lowe r Oligocen e sequenc e form s a n extensive depositiona l uni t that extend s paralle l to th e mid-Norwegia n coas t fro m M0r e t o Lofoten (Rokoenge n et al. 1995) . The sediments are interprete d as deltai c and coasta l deposits , probably forme d i n a wave-dominate d environ ment wit h extensiv e longshor e drift . Seismi c lines acros s thi s uni t show tha t i t i s represente d by an interval of steeply dipping reflectors, which are interpreted to be a set of prograding foresets. The bas e o f th e successio n i s seismicall y identified a s a downla p surface . Th e wester n boundary is a morphological ram p formed by the distal foresets (Eidvi n et al. 1998) . Outbuilding fro m Norwa y observe d fro m th e southeastern Nort h Se a t o Lofote n (Fig . 15 ) reflects th e onse t o f regiona l tectoni c uplif t o f Scandinavia at the Eocene-Oligocene transition. Southern Norwa y wa s uplifted , erosio n acceler ated an d sedimen t transpor t toward s th e wes t became mor e important. Early Oligocen e outbuildin g is also observe d at th e edg e o f th e Hebride s Shel f (Stoke r e t al . 1993). I n th e Faeroe-Shetlan d Channel , th e Oligocene sequenc e is thin. During Earl y Oligocen e time , th e Green land-Svalbard ga p bega n opening , allowin g greater communicatio n betwee n th e Atlanti c and Arcti c oceans , an d causin g a n influ x o f colder wate r int o the Nort h Sea . Thi s chang e i s evidenced b y th e replacemen t o f warm-wate r fish by borea l forms , an d a dro p o f 12° C i n th e North Se a botto m temperature s (Buchard t 1978). Thi s chang e wa s coinciden t wit h shallowing o f th e basi n a s a resul t o f sedimen t infilling an d tectoni c uplift .

252 SG8043-201

NE

1.5-

(a) 5 km

SG8043-303

CSS-3

(b) 25km

Fig. 14 . Seismic line s showing pinchout of CSS-3 towards area at the eastern basin flank uplifted a t the EoceneOligocene transition , (a) Seismic line SG8043-201; (b) seismic line SG8043-303. (Se e Fig. 13 b for line locations and Fig. 4 for Cenozoic seismi c stratigraphy.)

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Fig. 15 . Regional setting in Early Oligocene time . 1 , early Tertiary flood basalt province; 2, areas of regional uplif t (Rohrman et al 1995 ; Japsen 1997) ; 3, domes or anticlines (Dore & Lundin 1996); 4, main depocentre; 5 , extinct spreading axis ; 6, V0ring an d Faeroe-Shetland escarpments; 7 , outbuilding directions .

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J. I . FALEIDE ETAL.

Miocene time A prominen t hiatu s i s presen t i n th e Miocen e sequence bu t th e biostratigraph y i s equivoca l (Steurbaut e t al 1991 ; Eidvi n & Rii s 1992 ; Gradstein & Backstro m 1996 ; Martinse n e t al 1999; Eidvi n e t al . 2000) . Accordin g t o Martinsen e t al. (1999) , thi s break extend s fro m latest Oligocene (c . 25 Ma) to Late Miocene tim e (c. 8-9Ma), interrupted b y sedimentar y units of Late Oligocen e an d Earl y t o Mid-Miocen e age . The Miocen e brea k i s locall y characterize d b y erosion o f the underlying Oligocen e sequence . During Miocen e tim e th e norther n Nort h Se a constituted a shallo w marin e basin , whic h connected th e deepe r centra l Nort h Se a an d th e Norwegian-Greenland Sea (Fig. 16a) . The mai n Miocene depocentre s (CSS-5 , CSS-6 and CSS-7) are locate d sout h o f 60° N (Fig . 16b) . Farthe r north some o f the sequence s ar e thin or missing, so tha t i t i s no t possibl e t o ti e thei r boundarie s within the seismi c grid . The CSS- 5 sequenc e (Lowe r Miocen e sequence) buil t ou t int o th e norther n Nort h Se a from a basin margi n t o th e west . I t comprise s a seismically well-define d sequenc e reachin g about 200 m i n thicknes s i n th e basi n centre . The bas e i s marke d b y onla p ont o th e irregula r uppermost Oligocen e sequenc e boundary ,

whereas th e to p i s identifie d a s a prominen t reflector representin g th e bas e o f th e sand y Utsira Formation . T o th e north , th e Lowe r Miocene sequenc e pinche s ou t a t abou t 61°30'N (Eidvi n et al. 2000) . Depositio n during Early-Mid-Miocene tim e wa s dominate d b y low accumulatio n rates. The overlyin g sequenc e (CSS-6 ) thin s north ward an d ma y b e absen t i n larg e part s o f th e study are a nort h of 60°N. A mid-Miocene fall i n glacio-eustatic se a level (Miller et al. 1987 ; Haq et al . 1987 , 1988 ; Abre u & Anderso n 1998 ) i n combination wit h regiona l uplif t gav e ris e t o erosion an d formatio n of a prominent unconfor mity, in particular along the eastern basin margin. A Miocen e progradin g wedg e observe d i n th e M0re Basi n (Martinse n e t al . 1999 ) probabl y represents outbuildin g from uplifte d an d erode d areas in southern Norway and the northern North Sea (Fig . 17) . A Middl e Miocen e sequenc e boundary i n th e Norwegian-Danis h Basi n represents a shif t fro m a progradin g reflectio n pattern belo w t o a n aggradin g patter n abov e (Michelsen e t al 1995) . In quadran t 3 5 o f th e Norwegia n norther n North Se a ther e i s evidenc e o f incisio n int o Oligocene strata (Fig. 18 ) (Rundberg et al 1995 : Gregersen 1998 ; Martinse n e t a l 1999) . Th e formation o f th e incise d valley s (submarin e or

Fig. 16 . (a) Palaeo-water dept h ma p for Late Oligocene time (lowermost CSS-5); (b ) time-thickness map for the Miocene units (CSS- 5 to CSS-7).

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Fig. 17. Regional setting in mid-Miocene time reflecting a hiatus in the northern North Sea. 1 , early Tertiary flood basalt province ; 2, areas o f regional uplift (Rohrma n e t al. 1995 ; Japsen 1997) ; 3 , domes or anticlines (Dore & Lundin 1996) ; 4, mai n depocentres ; 5 , extinct spreadin g axis ; 6 , V0ring an d Faeroe-Shetland escarpments ; 7, outbuilding directions.

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J. I . FALEIDE ETAL.

Fig. 18 . Neogene incised valleys i n the northern Nort h Sea. (a) Seismic exampl e acros s incised valley ; (b) semi regional map showing distribution of incised valleys (modified fro m Martinse n et al. (1999); (c ) earliest Miocene time map, Block 34/ 6 (Martinsen et al. 1999) ; (d ) earliest Miocene tim e map, Bloc k 35/4 (Gregersen 1998) ; (e) earliest Miocen e tim e map, Block 35/8.

subaerial?) an d th e timin g o r ag e ar e critica l questions i n constrainin g th e histor y o f vertica l movements. Martinse n e t al . (1999 ) interprete d the incise d feature s a s subaerial , incise d valley s because o f th e morphologica l patterns . A

subaerial origi n fo r th e incise d feature s implies considerable tectoni c uplif t followe d b y Lat e Miocene-Pliocene subsidence . Th e incisio n probably too k plac e durin g formatio n o f th e mid-Miocene unconformity. However, according

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257

Fig. 19 . (a) Palaeo-water dept h map for earliest Pliocen e tim e (lowermos t CSS-8) ; (b ) time-thickness map for CSS-8; (c ) time-thickness map for CSS-9; (d ) time-thickness map for CSS-10.

to Eidvin et al. (2000) th e incision is Pliocene i n mid-Miocen e unconformit y i n th e norther n age (see discussion below). Nort h Sea . Th e formatio n accumulate d i n tw o The Utsir a Formatio n correspond s t o par t o f mai n depocentre s alon g th e basi n centre , sout h CSS-7 an d wa s deposite d o n to p o f th e an d nort h o f 60°N , respectively . Th e Utsir a

258

J. I . FALEIDE ETAL.

Formation i s interprete d a s a shallow-marin e deposit, she d mainl y fro m th e easter n basi n margin durin g lat e Mid-Lat e Miocen e risin g relative se a leve l (Isakse n & Tonsta d 1989 ; Rundberg 1989 ; Gallowa y et al 1993 ; Gregerse n et a l 1997 ; Martinse n e t a l 1999) . Th e Utsir a Formation i s san d rich , bu t toward s th e wester n basin margin it fines considerably. The sediment s were mainly derived from Scandinavia (Rundberg 1989), bu t sedimen t influ x fro m th e Shetlan d Platform o r Scotland i s also proposed (Gregerse n et al 1997) . Althoug h the Utsira sandstone s ar e likely t o hav e bee n supplie d fro m th e basi n margin, the y wer e probabl y significantl y reworked b y basina l current s (Gallowa y e t a l 1993). Th e basi n i s unlikely t o hav e bee n mor e than 150-200 m deep, s o that such currents could have existed . Th e wate r depth increase d an d culminated i n earlies t Pliocen e tim e wit h condensed deposit s o n whic h th e succeedin g Pliocene deposit s downlapped . Late Miocen e uplif t o f sout h centra l Norwa y promoted furthe r tiltin g o f th e Hord a Platfor m area, an d hence , erosio n o f Miocen e sediment s (Clausen e t a l 1999a) . Th e continue d uplif t o f Scandinavia force d th e Lat e Miocen e an d Pliocene deposit s t o prograd e westward , down lapping ont o the mid-Miocene unconformit y and the underlying Lower-Middle Miocene deposit s and fillin g i n th e mor e centra l part s o f th e northern Nort h Sea . Deeper-wate r starve d

conditions existe d i n th e Norwegian-Danis h Basin an d the M0re Basi n i n Late Miocen e tim e contemporaneous wit h depositio n o f th e Utsir a Formation i n the norther n Nort h Sea . The Mid-Miocen e an d younger sedimentation in the souther n Nort h Se a was dominated b y the expansion of massive delta systems on its eastern seaboard associate d with former Baltic rivers that drained fro m th e Fennoscandian Shield (Bijlsm a 1981; Gibbar d 1988 ; Camero n e t al 1993) . Plio-Pleistocene time A majo r depocentr e o f mainl y Uppe r Pliocen e (CSS-8) glacia l sediment s derived fro m uplifte d Norway is located i n the northernmost North Sea (Fig. 19) . Sequence geometrie s characterize d by westward progradin g clinoform s (Fig . 20 ) sho w that th e dominan t sedimen t transpor t direction s are from th e Norwegian shelf margin towards the west an d NW , bu t occasiona l input s fro m th e Shetland Platfor m ar e als o observed . Th e regional downla p surfac e a t th e bas e o f th e Pliocene sequenc e i s interprete d a s reflectin g starved sedimentatio n probabl y cause d b y a relative sea-leve l ris e (Gregerse n e t a l 1997) . The clinoform s withi n sequence CSS- 8 indicat e deep wate r in excess of 500m (Figs 19 a and 20). An ag e o f 2.75 Ma i s taken as the maximum age of the Upper Pliocene section , corresponding to a large increase in the supply of ice-rafted material

Fig. 20 . Seismi c lin e NVGTI-92-10 8 showin g Pliocen e progradatio n fro m th e eas t an d flat-lyin g Pleistocen e sediments abov e angula r unconformity . (Se e Fig . 19 b fo r lin e locatio n an d Fig . 4 fo r Cenozoi c seismi c stratigraphy.)

CENOZOIC EVOLUTION OF THE NORTHERN NORTH SEA related t o a marke d expansio n o f norther n European glacier s (Eidvin et al 2000) . A perio d o f transgressio n i n Earl y Pliocen e time resulte d i n strongl y reduce d rate s o f deposition, an d sediment s o f thi s ag e ar e preserved mainl y in the centra l Nort h Sea (Eidvin et al 2000) . Most of the Lower Pliocen e sediments, whic h ma y hav e existed , wer e probably erode d i n the subsequent period during an extensiv e relativ e fal l i n th e globa l se a leve l (4.1-2.9 Ma). A period o f regression i n earlies t Late Pliocen e tim e probably resulted i n erosio n of most of the Norwegia n continental shelf with the exceptio n o f th e deepe r area s o f th e Centra l and Vikin g grabens . I f th e incise d valley s described abov e (Fig . 18) wer e forme d a t thi s time (earlies t Lat e Pliocen e time) , an d the y formed abov e se a level , i t implie s extremel y rapid Lat e Pliocen e subsidenc e t o creat e th e water depth inferred from th e clinoform geometries within the CSS-8 sequenc e (>500m) . This period wa s immediatel y followed , i n th e late r part of Late Pliocene time, by rapid deposition of glacially derive d sediment s prograding along the entire shelf . In general , th e Pleistocen e developmen t i s a continuation o f th e Lat e Pliocen e evolution , but is marked by more extensive erosion of the inner shelf (Eidvi n et al. 2000). Flat-lyin g Pleistocene beds lie with an angular unconformity on more or less progradationa l Uppe r Pliocen e deposit s (Fig. 20) . The lowe r par t o f th e Pleistocen e sequence an d uppermos t par t o f th e Uppe r Pliocene sequence ar e eroded ove r large area s of the continental shelf. The base of the Pleistocene section i s date d t o 1. 2 Ma (Sejru p e t a l 1995) . This dating coincides wit h a marked intensifica tion of glacial activity, as is observed in the deepsea recor d (Ruddima n e t a l 1986 ; Berger & Jansen 1994) . Repeated glaciation s (Sejru p et al 1995) erode d th e margina l part s o f the norther n North Sea and the associated glacio-eustati c sea level falls an d the overall lowstand have resulted in majo r channe l cut s o r incise d valley s i n th e marginal part s o f th e Nort h Se a basi n (Sejru p etal 1991) . Differences i n Pliocen e an d Pleistocen e depositional pattern s ar e probabl y th e resul t o f changes i n glaciatio n cycle s tha t occurre d a t c. 1. 1 Ma. Durin g th e perio d befor e this , th e Fennoscandian ice cap probably extended only to the presen t coastlin e (Janse n & Sj0hol m 1991) . Subsequent t o c . 1. 1 Ma, glacier s periodicall y extended ove r th e continenta l shel f an d trans ported sediment s ove r greate r distance s (Sejru p et al 1995 , 2000 ; King et al 1996) . Large-scale Pliocen e progradatio n i s als o observed offshor e mid-Norwa y (Fig . 21 )

259

(Stuevold & Eldhol m 1996 ; Hjelstuen e t a l 1999; Eidvi n etal 2000) . The narrow-elongate d depocentr e alon g th e basin axis in the central North Sea shows that the basin flank s wer e uplifte d o r exhume d togethe r with mainlan d Norwa y an d th e Britis h Isle s (Fig. 21 ) (Hilli s 1995a , 1995b ; Hanse n 1996; Japsen 1998 , 1999 , 2000) . Th e Grea t Europea n delta, fe d mainl y fro m th e eas t an d south , continued to expand northwards into the southern and centra l North Se a (Zagwij n 1989 ; Camero n etal 1993;Funnel l 1996) . The regional uplif t o f Scandinavi a is believed to be the main control on sediment suppl y to the prograding Pliocen e sequences . Oscillation s o f the eustatic se a level punctuate d the tectonicall y controlled progradatio n and affected variation s in the accommodatio n space , an d thu s create d th e high-frequency sequence s (S0rensen etal 1997) . An upwar d increasin g numbe r o f Pliocen e sequences i n th e centra l Nort h Se a ha s bee n related t o a n increasin g Neogen e uplif t o f Scandinavia an d adjacen t areas , combine d wit h a general lowered glacio-eustati c se a level during Pliocene time . Biostratigraphi c studie s b y Seidenkrantz (1992 ) suppor t a shallowin g upward and gradually colder environment during Late Pliocen e time . Th e gradua l coolin g (Buchardt 1978 ) an d ice-ca p growt h cause d a general glacio-eustati c sea-leve l lowerin g (Ha q et a l 1988 ) through Pliocen e time , which , together wit h th e uplif t durin g Neogen e tim e (Ghazi 1992 ; Jensen & Schmid t 1992 , 1993 ; Hansen 1996) , led to erosion o f the exposed shel f areas an d increased sedimen t influx . Possible mechanisms causing vertical movements The main mechanisms suggeste d i n the literature to cause Cenozoic vertica l movement s alon g the NE Atlanti c margin an d adjacen t area s include: (1) arriva l o f th e Icelan d plum e an d th e subsequent latera l spreadin g of th e plum e head; (2) mantle processes and thermal regime, and the episodic behaviou r o f th e Icelan d plume ; (3 ) emplacement o f magma in and at the base of the crust (magmati c underplating ) leadin g t o isostatic uplift ; (4 ) metamorphi c phas e change s i n the mantle; (5) intra-plate stress related to ridgepush from the North Atlantic plate boundary and/ or t o Alpin e compression ; (6 ) erosio n an d isostatic rebound; (7) flexura l effects . It i s beyond th e scop e o f this paper to test all these mechanism s quantitatively . Here, som e o f them will be briefly discussed in light of our new constraints on timing, amplitude and wavelength

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Fig. 21. Regional settin g in Late Pliocene time . 1 , early Tertiary flood basalt province ; 2, areas of regional uplif t (Rohrman e t at. 1995 ; Japse n 1997) ; 4 , mai n depocentres ; 5 , extinc t spreadin g axis ; 6 , V0rin g an d Faeroe Shetland escarpments ; 7 , outbuilding directions; 8 , major rive r system (Gibbard 1988) .

CENOZOIC EVOLUTION OF THE NORTHERN NORTH SEA of Cenozoic differentia l vertica l movements . For each o f th e mai n regiona l event s w e wil l summarize th e mai n observation s o r fact s tha t any proposed mechanism ha s to explain .

Late Paleocene-Early Eocene vertical movements In simpl e terms , th e Lat e Paleocene-Earl y Eocene uplif t i s referre d t o a s 'margina l uplift ' related t o th e riftin g an d break-u p o f th e N E Atlantic. However , ther e ar e larg e variation s i n the wavelengt h an d amplitud e o f th e uplif t observed alon g th e N E Atlanti c margins , an d possibly als o i n timing. Important observation s tha t hav e t o b e explained include : (1 ) uplif t o f th e Hebrides Shetland axi s an d northwester n corne r o f southern Norway ; (2 ) uplif t o f th e are a alon g the incipien t plat e boundary, with subaeria l sea floor spreading; (3) accelerated subsidence in the North Sea . The accelerate d tectoni c subsidenc e i n th e North Se a basin i n Paleocene tim e an d uplif t o f surrounding area s wa s probabl y associate d wit h arrival o f th e Icelandi c plum e an d crusta l break-up in the North Atlantic . Ther e appea r to have bee n tw o majo r phase s o r pulse s o f Earl y Tertiary magmatis m withi n th e Nort h Atlanti c igneous province , wit h a pea k a t 5 9 Ma an d a t 55 Ma (Whit e & Lovel l 1997 ; Ritchi e e t al 1999). Several uplif t mechanism s ca n b e attribute d to thi s regiona l tectonomagmati c event : (1 ) transient therma l uplif t generate d b y buoyanc y of lithospher e heate d b y ho t plum e materia l (Sleep 1990 ; Clif t & Turne r 1998) ; (2 ) transient dynami c uplif t generate d b y mantl e fluid flo w drive n b y th e ascendin g ho t plum e material (Nadi n e t al . 1995 , 1997) ; (3 ) permanent, isostaticall y compensate d uplif t generated b y crusta l thickenin g cause d b y igneous underplatin g associate d wit h mantl e plume activit y (Brodi e & Whit e 1994 , 1995 ; White & Lovel l 1997 ; Clif t & Turne r 1998 ; Clift 1999) . It i s likel y tha t on e o r mor e o f thes e mechanisms contribute d t o th e regiona l uplif t of th e norther n Britis h Isle s an d adjacen t area s (surrounding th e Britis h Tertiar y igneou s pro vince). Th e mai n developmen t o f Paleocen e sandstone reservoir s alon g th e axi s o f th e Faeroe-Shetland Basi n appear s t o hav e bee n synchronous wit h th e phase s o f therma l uplif t along th e basi n margi n an d pulse d volcanis m (White & Lovel l 1997 ; Naylo r e t al 1999) .

261

A simila r relatio n betwee n outbuildin g o f sandy deposit s an d th e episodi c natur e o f th e plume-related uplif t ha s bee n suggeste d fo r th e southern Nort h Se a area (Kno x 1996) . However, i t i s mor e difficul t t o se e how the y can hav e contribute d t o th e formatio n o f th e regional dom e i n souther n Norway . We sugges t that th e uplifte d are a relate d t o th e Paleocene Early Eocen e phas e o n th e Norwegia n sid e i s much narrower, covering the northwestern corner of souther n Norway , includin g th e onl y are a where early Tertiary volcanism has been reported (Vestbrona Formation , Bugg e e t al . 1980 ; Prestvik e t al . 1999) . Magmati c underplatin g added to relatively thic k continental crust and/or buoyancy fro m heate d lithospher e ca n bot h explain the uplift . The Late Paleocene-Early Eocene accelerate d subsidence in large parts of the North Sea basin is also difficul t t o explain . Whit e & Lati n (1993 ) and Hall & White (1994) claimed that a period of crustal extension occurred i n early Tertiary time . Minor norma l faultin g is observable in the basin to support this hypothesis. Skogseid e t al . (2000 ) relate d Paleocen e vertical movement s i n th e Nort h Atlanti c region t o th e arriva l o f th e Icelan d mantl e plume t o lithospheri c level s an d th e subsequen t lateral spreadin g o f th e plum e head . Th e thickness o f th e plum e bod y i s bot h a functio n of th e distanc e fro m th e plum e centr e an d th e structure o f th e overlyin g lithosphere . Th e ho t plume materia l preferentiall y fill s trap s a t th e base o f th e lithospher e create d eithe r b y previous plat e deformation s o r b y continuin g lithospheric thinnin g (Slee p 1996 , 1997) . I n their dynami c mode l th e centra l rif t zon e wit h the thinnes t lithospher e an d th e thickes t plum e body i s uplifte d an d eroded . I n region s wher e the lithospher e i s thic k wit h respec t t o th e thickness o f th e plum e bod y (e.g . th e Nort h Sea) significan t an d rapi d subsidenc e i s expected durin g th e plum e emplacement . Thi s subsidence shoul d theoreticall y b e followe d b y a rapi d rebound , resultin g i n ne t uplif t a s th e plume movemen t cease s (Skogsei d e t al . 2000). Rapid Eocene subsidenc e has been related to a decrease i n dynamic uplift cause d by a reduction in plum e activit y (Nadi n e t al . 1997) . Rapi d decay o f uplif t i s attribute d t o a temperatur e decrease o f th e plum e i n Earl y Eocen e tim e (c. 5 5 Ma), resultin g i n a decreas e i n dynami c uplift. Thi s even t coincide d wit h a decreas e o f the plum e activit y i n th e Britis h Tertiar y an d Greenland igneou s provinces , an d th e initiation of sea-floo r spreadin g betwee n Greenlan d an d NW Europe .

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Oligocene-Miocene domal uplift of southern Norway Important observations tha t have to be explaine d include: (1) Early Oligocene onse t of outbuilding from Norway-Denmar k t o Lofoten ; (2 ) con temporaneos uplif t o f th e easter n basi n flan k (offshore areas) ; (3 ) a break i n sedimentatio n a t the Eocene-Oligocen e transition ; (4 ) low velocity o r hotte r uppe r mantl e beneat h th e domes i n souther n and northern Norway . The doma l uplif t o f th e Norwegia n mainland probably starte d i n Lat e Eocen e t o Earl y Oligocene time , an d als o contribute d t o uplif t of th e shel f areas . Th e uplif t an d shallowin g i n the norther n Nort h Se a basi n continue d int o Miocene time . Severa l mechanism s hav e bee n suggested t o explain th e Neogene doma l uplif t of southern Norwa y an d simila r uplifte d area s around th e North Atlantic. Rohrman & va n de r Bee k (1996 ) propose d a model based on simple fluid dynamics and linked this to the elevated Nort h Atlanti c upper-mantl e thermal regim e surroundin g the Iceland hotspot . In thei r model , a n anomalou s hot , buoyan t asthenospheric layer or lens is present a t the base of the lithosphere. O n interaction of the hot (lowviscosity) asthenospher e laye r wit h th e col d (higher-viscosity) shiel d lithosphere , diapir s wil l start to form as a result of convective instabilities. The diapiris m wa s probabl y triggere d b y litho spheric stres s pattern s tha t cause d th e plat e reorganization a t c . 3 0 Ma. Th e mode l implie s that uplif t i s transien t an d wil l chang e t o subsidence whe n th e therma l anomal y decays . However, diapir s penetratin g th e lithospher e could generat e partia l meltin g an d produc e underplating, hereb y generatin g permanen t uplift. Th e area s o f uplif t ar e associate d wit h strong negativ e Bougue r gravit y anomalie s an d reduced lithospheri c P - an d S-wav e velocities , suggesting a n anomalou s mantl e structur e an d temperature underneat h th e dome s (Banniste r etal. 1991) . Plio-Pleistocene glacial erosion and uplift Important observation s tha t have to be explaine d include: (1) Upper Pliocene and older sequences tilted awa y fro m th e dom e i n sout h Norway; (2) Pleistocene sediments, nearl y flat-lying above an angular unconformity ; (3 ) the angular unconfor mity separatin g Pleistocen e an d Pliocen e sedi ments becomin g les s pronounce d toward s bot h the central North Se a and the shelf offshore midNorway; (4 ) accelerate d subsidenc e o f basi n centres adjacen t t o the uplifted landmasses .

The seismi c line s sho w tha t th e post-Eocen e sequences see m to be uniformly tilted away fro m the Scandinavia n dome , indicatin g tha t a significant par t o f th e Neogen e uplif t o f sout h Norway occurre d i n lates t Pliocen e an d earlies t Pleistocene times . Isostati c response t o unloading cause d b y glacia l erosio n contribute d significantly t o th e Neogen e uplif t (Rii s & Fjeldskaar 1992) . However, unloading cannot be the onl y operatin g mechanis m fo r th e sout h Norway Neogen e uplif t (Rii s 1996) .

Intra-plate stress: compressional deformation Tertiary epeirogen y i s ofte n attribute d t o compression tha t i s assume d t o b e relate d i n a general sens e t o Alpin e mountai n building . However, t o remov e c . 3k m o f sedimentar y rock fro m a basi n c . 100k m wid e require s > 15k m o f shortenin g (Brodi e & White 1994) . Minor Tertiary compressio n i s observed al l ove r the continenta l shelf, but nowher e is it sufficien t to accoun t for th e require d amoun t of uplif t an d erosion. I n addition , exhumatio n dramaticall y increases fro m sout h t o north , wherea s th e observed compressio n decreases markedl y in the same direction. Compressional structure s o f Cenozoi c age , including simpl e dome s o r anticlines , revers e faults an d broad-scale inversion , are widesprea d in th e Nort h Se a an d alon g th e N E Atlanti c margin (Fig s 1 5 an d 17 ) (Dor e & Lundi n 1996). A multiphas e growt h histor y ha s bee n reported fo r som e o f th e structure s a t th e Norwegian margi n (Mid-Eocen e t o Earl y Oligocene an d Miocen e times ; Dor e & Lundin 1996) an d i n th e Faero e regio n (Lat e Eocen e t o Early Oligocen e an d mid-Miocen e times ; Boldreel & Anderse n 1998 ; Anderse n e t al 2000). Vagne s e t a l (1998) , o n th e othe r hand , reported a surprisingl y constan t growt h rat e fo r the Orme n Lang e Dom e fro m earlies t Eocen e time t o th e present . Thes e studie s al l relat e th e Cenozoic compressiona l deformatio n t o tw o main sources : (1 ) far-fiel d effect s reflectin g episodes o f deformatio n i n th e Alpin e Orogen y and (2 ) ridge-pus h fro m th e Nort h Atlanti c mid-ocean ridg e system . Late Miocen e an d Pliocen e differentia l vertical movement s i n th e Nort h Se a basi n (increased subsidenc e rate s a t basi n centr e an d relative uplif t alon g basi n edges ) hav e als o been relate d t o change s i n intra-plat e stres s correlated wit h plate tectoni c adjustments in the Alpine hinterlan d an d th e Atlanti c spreadin g system (Cloeting h e t al . 1990 ; Gallowa y e t al . 1993).

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Summary and conclusions

References

This pape r ha s focuse d o n th e Cenozoi c evolution o f th e norther n Nort h Se a an d surrounding areas , wit h emphasi s o n sedimen t distribution, composition an d provenance as well as o n timing , amplitud e an d wavelengt h o f differential vertica l movements . Th e dat a an d modelling result s suppor t a probabl e tectoni c control on sediment suppl y and on the formation of the regional unconformities . The sedimentar y architecture an d break s ar e relate d t o tectoni c uplift o f surroundin g clasti c sourc e areas ; thus , the offshore sedimentary recor d provide s the best age constraint s o n th e Cenozoi c exhumatio n of the adjacen t onshor e areas . However , climati c and glacio-eustati c change s als o playe d a rol e and fo r som e event s it is difficul t t o separat e th e effects o f tectonics an d eustasy . Cenozoic exhumatio n i s documente d o n bot h sides of the North Sea , but the timing is not well constrained. Two major uplif t phase s ar e clearly reflected b y th e availabl e data : (1 ) a n earl y Paleogene (Late Paleocene-Early Eocene) phase and (2 ) a late Neogene (Plio-Pleistocene ) phase. The firs t i s relate d t o rifting , magmatis m an d break-up i n th e N E Atlanti c associated wit h the arrival o f the Icelan d plum e an d th e subsequent lateral spreadin g of the plume head. The latter is related t o th e isostati c respons e t o unloadin g caused b y glacial erosio n durin g th e widesprea d Northern Hemispher e glaciations . However , other uplif t event s interpose d betwee n thes e episodes, i n particula r o n th e Scandinavia n side during Oligocen e an d Miocen e times . Som e o f these may also be related to mantle processes and the episodi c behaviou r o f th e Icelan d plume . Intra-plate stres s relate d t o ridge-pus h fro m th e North Atlanti c plat e boundar y and/o r t o Alpin e compression als o contribute d t o differentia l vertical movements , bu t canno t b e th e mai n uplift mechanis m for the long-wavelength domal uplifts.

ABREU, V.S . & ANDERSON , J.B. 1998. Glacial eustasy during th e Cenozoic ; sequenc e stratigraphi c implications. AAPG Bulletin, 82 , 1385-1400. ALLEN, PA . & ALLEN , J.R . 1990 . Basin Analysis; Principles an d Applications. Blackwel l Scientific , Oxford. ANDERSEN, M.S. , NIELSEN , T. , NIELSEN , T. ,

The pape r i s base d o n wor k withi n th e projec t 'Tectonic impact o n sedimentary processe s in the postrift phase—improve d models' . Th e projec t wa s supported b y the Research Counci l of Norway through grant 32842/211 . The author s would like to thank the companies participatin g in the project (Amoc o Norway Oil Company , de n Norsk e Stat s Oljeselska p (Statoil) , Mobil Exploratio n Norwa y Inc. , Nors k Agi p A/S , Norsk Hydr o ASA , Phillip s Petroleu m Compan y Norway, Sag a Petroleu m ASA) . Th e seismi c dat a were kindl y made availabl e by TGS-NOPEC . W e are also gratefu l t o D . Mudg e an d K . G . R0sslan d fo r reviewing the paper an d for suggesting improvements.

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Scotland's denudational history: an integrated view of erosion and sedimentation at an uplifted passive margin ADRIAN HALL 1'2 & PAUL BISHOP3 Department of Geography, University of Edinburgh, Drummond Street, Edinburgh EH8 9XP, UK 2 Fettes College, Carrington Road, Edinburgh EH4 1QX, UK (e-mail: am.hall@ fettes.com) 3 The CRUST Project, Department of Geography and Topographic Science, University of Glasgow, Glasgow G12 8QQ, UK 1

Abstract: Denudationa l histor y i s commonly reconstructe d fro m basi n sediment s derive d from th e denude d sourc e area , an d les s frequentl y fro m th e sourc e are a itself . Norther n Britain i s a n importan t sourc e are a fo r th e surroundin g sedimentar y basin s an d thi s pape r reviews the erosional histor y o f Scotland fro m Devonian time to the present using evidenc e both fro m onshor e geolog y an d geomorpholog y an d fro m pattern s o f sedimentatio n i n surrounding basins. Cover rocks were extensive in Scotland durin g late Palaeozoic tim e but the persistenc e of sedimen t sourc e area s withi n the uplan d area s of Scotlan d make s it unlikely tha t basemen t high s wer e eve r completel y buried , an d depth s o f post-Devonia n erosion of basement have been correspondingly modest (< 1 -2 km) . During Mesozoic time, Scotland experience d severa l majo r erosional cycles , beginnin g with uplift, reactivatio n of relief an d strippin g o f cove r rocks , followe d b y progressiv e reductio n o f relie f throug h etchplanation an d culminatin g i n extensiv e marin e transgression s i n Lat e Triassic , Lat e Jurassic an d Late Cretaceou s time . Mid-Paleocene pulses of coarse sedimen t t o the Mora y Firth Basi n coincide d wit h majo r uplift. Thi s uplif t wa s associate d wit h major differentia l tectonics withi n th e Highlands, wit h warping an d faulting alon g th e margins o f the Minc h and th e inne r Mora y Firt h Basins . Tectoni c activit y was renewe d o n a lesse r scal e i n lat e Oligocene tim e an d continued into Late Neogene time . Differential weathering an d erosion under the warm to temperate humid climates of Neogene time created th e major elements of the preglacial relief, wit h formation o f valleys, basins, scarps and inselbergs, feature s ofte n closely adjuste d to lithostructura l control s and , i n som e cases , wit h precursors tha t can b e traced bac k t o Devonia n time . Th e histor y tha t ca n b e 'read ' fro m th e onshor e regio n complements th e source area history interpreted from sedimentar y basins derived from thes e

areas

The nature and rate of deposition of sedimentary determin e the exten t t o whic h potential energy basin sequence s depen d o n man y factors , provide d by uplif t ca n b e converte d to kineti c including rate s o f basi n subsidence , sea-leve l energy , namely, th e geomorphologica l character history an d sourc e area characteristics , such a s o f th e sourc e area, an d th e exten t t o whic h th e climate history, lithology and uplift rates . Source sourc e are a 'knows ' abou t th e uplif t event . A area uplif t i s generall y interprete d t o b e high-elevation , upliftin g sourc e are a tha t i s associated wit h a n essentiall y instantaneou s efficientl y connecte d to base leve l will generate 'basin' signa l of high rates of flu x o f sediment s hig h volume s of sedimen t via a combination of that hav e experienced limited chemica l weath - rive r incision , mass movement from stee p valley ering. Althoug h i n som e cases suc h a respons e side s and efficien t evacuatio n o f sediment . The may be demonstrable (e.g. Copeland & Harrison southward s drainin g river s o f th e Himalaya s 1990), th e degre e t o whic h majo r uplif t i s offe r excellen t examples of suc h system s (e.g. signalled by a sedimentary pulse depends on the Burban k e t al 1996 ; Hancoc k e t al 1998) , extent t o whic h the potentia l energ y associated whic h contras t markedl y wit h th e system s with hig h elevation ca n b e converte d int o th e drainin g northwards fro m th e Himalaya s to th e kinetic energ y (and henc e th e strea m power ) high-elevation , bu t low-energ y an d internall y necessary t o detach , entrain and transpor t high drained , Tibeta n platea u (Summerfiel d & volumes o f sediment . Two interrelate d factors Brow n 1998) . From: DORE , A.G., CARTWRIGHT, J.A., STOKER, M.S. , TURNER, J.P . & WHITE , N . 2002 . Exhumation of th e North Atlantic Margin: Timing, Mechanisms and Implications for Petroleum Exploration. Geological Society, London, Special Publications, 196, 271-290. 0305-8719/027$ 15.00 © The Geological Society of London 2002.

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The mos t rapidl y upliftin g areas, suc h a s th e Southern Alp s (Hoviu s 2000 ; Tippett & Hoviu s 2000), Himalaya s (Fieldin g 2000) , Japa n (Ohmori 2000 ) an d Taiwa n (Li n 2000) , ar e associated wit h plate convergenc e zones . O n the other hand , passive margi n uplands , suc h a s the southern Africa n upland s (Brow n e t al 2000) , the Wester n Ghat s i n Indi a (Gunne l & Fleitou t 2000), th e S E Australia n highland s (Bisho p & Goldrick 2000) , an d th e Scottis h Highland s portion o f the western European Atlantic margin, are commonl y associate d wit h lo w t o ver y lo w rates o f denudatio n an d sedimen t flux . Flemin g et a l (1999 ) an d Cockbur n e t a l (2000) , fo r example, hav e reporte d ver y lo w rate s o f denudation fro m th e souther n Africa n passiv e margin highlands , o f th e sam e orde r o f magnitude a s thos e reporte d fro m th e S E Australian upland s from mas s balance , geomorphological an d thermochronologica l studie s (Bishop 1985 ; Bisho p & Goldrick 2000). Bishop & Goldrick (2000 ) also reported widesprea d and persistent disequilibri a i n S E Australia n rive r long profiles, reflecting the low gradients and low stream powe r o f this margin's drainag e systems . These long profil e disequilibria ca n be attribute d to passiv e denudationa l reboun d (Bisho p & Brown 1992 ; Bishop & Goldrick 2000) , wit h the margin evidentl y no t havin g experience d activ e tectonic uplif t durin g Cenozoi c tim e apar t fro m temporary uplif t event s relate d t o transien t thermal effect s a t th e centra l volcanoe s tha t mark easter n Australia' s Cenozoi c passag e ove r mantle hotspot s (Wellma n & McDougal l 1974 ; Wellman 1986 ; McDougall & Duncan 1988; Sun etal 1989) . The tectoni c characte r an d historie s o f mos t of th e passiv e margin s descibe d abov e hav e been reconstructe d largel y fro m subaeria l terrestrial data , wit h relativel y littl e relianc e on th e sedimentar y basi n record . Th e sedimen tary recor d ha s bee n use d mainl y fo r mas s balance studie s o f thes e margin s t o determin e rates o f sourc e are a subaeria l denudatio n (e.g . Bishop 1985 ) o r a s a guid e t o th e evolutio n o f the rive r system s (e.g . Rus t & Summerfiel d 1990). Rate s o f sourc e are a denudatio n ma y also b e determine d mor e directl y fro m th e source are a itsel f usin g geomorphologica l studies (e.g . Bisho p 1985 ; Not t e t a l 1996) , cosmogenic isotop e analysi s (Fleming e t a l 1999; Cockbur n e t a l 2000) , an d low temperature thermochronologica l techniques , such a s apatit e fission-trac k analysi s (Gleado w & Brow n 2000) . Conflicts , whic h ar e no t ye t fully resolved , ar e ofte n apparent , however , between geomorphologica l interpretation s o f source are a histor y an d thermochronologica l

approaches t o sourc e are a denudatio n (Koh n & Bishop 1999) . Reconstruction of the evolution of the Scottish Highlands (Fig . 1 ) o n th e Wester n Europea n continental margi n ha s relie d o n bot h offshor e and onshor e data , wit h ofte n muc h greate r emphasis o n th e offshor e record , n o doub t because o f th e wealt h an d qualit y of thes e dat a (see Jones e t al 2002) . Ther e i s a corresponding wealth o f dat a fro m onshor e areas , an d i n thi s paper w e re-examin e th e post-Palaeozoi c geo morphological histor y of th e Scottis h Highlands as a sourc e are a fo r surroundin g basins , especially th e mai n sedimen t receivin g area, th e North Se a Basin . W e hav e tw o aims : (1 ) t o summarize criticall y th e onshor e dat a o n th e evolution o f th e Scottis h Highlands , fo r a readership tha t migh t no t b e full y awar e o f th e literature on this topic; (2) to assess the extent to which sourc e are a uplift , denudatio n an d geomorphological development , an d landscap e antiquity, ca n b e 'read ' fro m th e sourc e are a itself, thereb y complementin g th e offshor e record.

An outline of the history of the Highlands source regio n from Palaeozoic time to the present The dispositio n an d provenanc e o f (ofte n thin ) remnants of Devonian sediments show that many key morphotectoni c element s o f th e curren t Highlands relie f wer e alread y establishe d b y the end of Devonian time (Fig. 2). These includ e the main Grampia n watershed, the linear depressio n of th e Grea t Glen , th e larg e basin s o f N E Scotland an d majo r valle y system s draining NE towards th e Mora y Firt h alon g th e Caledonia n fracture zones . Th e Caledonia n mountain s ha d been eroded , exposin g man y lat e Caledonia n Newer Granites , togethe r wit h som e olde r intrusions (Watso n 1985) . Th e Caledonia n granites wer e intrude d int o alread y stabilize d crust o r thei r intrusio n complete d th e stabiliz ation proces s (Leak e & Cobbin g 1993) . Recon struction o f th e sub-Devonia n relie f aroun d th e inner Moray Firth implies surfaces of high relief, with fault-bounde d half-basins an d fault-guided valleys partl y infille d wit h conglomerate s an d sandstones. Thes e Devonia n fill s wer e largel y removed betwee n lat e Palaeozoi c tim e an d th e present s o tha t th e curren t leve l o f erosio n lie s close t o tha t a t th e en d o f Devonia n tim e (se e Leake & Cobbin g 1993) . A simila r equivalenc e exists i n th e N W Highlands , wher e th e presen t terrain lie s a t th e sam e genera l elevatio n a s th e base of the Torridonian sequence (Watso n 1985) .

SCOTLAND'S DENUDATIONA L HISTOR Y

N

Fig. 1. Scotland, showing sites and features mentioned in the text.

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Fig. 2 . Indicators o f post-Devonian depth s o f denudation (after Hal l 1991) .

There i s little direc t evidenc e o f event s i n th e Highlands durin g th e Carboniferou s period , a s rocks o f thi s ag e ar e restricte d t o th e margina l basins o f th e Midlan d Valle y an d th e Mora y Firth. Th e considerabl e thicknesse s o f Carbon iferous sediments, wit h up to 4 km in the Midlan d Valley (Franci s 1991 ) an d 1.5k m i n th e oute r Moray Firt h (Andrew s et al. 1990) , wer e largel y sourced fro m th e Highlands . Th e Westphalia n outliers i n Morver n restin g o n th e Moin e sequence impl y a former cove r o f Carboniferous rocks i n part s o f th e S W Highland s (Franci s 1991). A reworke d Carboniferou s microflor a i s present i n Jurassi c rock s a s fa r wes t a s th e onshore outcrop s i n the inner Mora y Firt h Basi n (Andrews e f al . 1990) . The Highland s ar e generall y show n a s a n emergent an d exposed basemen t are a i n palaeogeographical map s o f th e Carboniferou s perio d (Guion e t al . 2000 ) bu t th e exten t an d dept h o f Carboniferous denudatio n i n th e Highland s remain unclear . Recen t apatit e fission-trac k studies impl y tha t a s muc h a s 3k m o f lat e

Palaeozoic cove r rocks hav e been remove d fro m the Highland s are a (Thomso n e t al . 1999 ) bu t deep Lat e Palaeozoi c erosio n appear s incompa tible with the widespread surviva l of near-surface volcanic an d intrusiv e rock s o f lat e Carboni ferous ag e (Watso n 1985 ; Hal l 1991) . Volcani c activity continue d throughout Permian tim e an d is represente d i n th e Highland s b y th e campto nite-monchiquite dyk e swarm s o f Orkne y an d the Wester n Highland s (Franci s 1991) . Associ ated uplif t wa s probabl y limite d a s th e tota l volume o f magm a wa s smal l (Watso n 1985) . This accord s wit h the surviva l of Carboniferous deep weatherin g mantle s i n norther n Scotlan d that acted a s a major source o f kaolinitic detritu s for Jurassi c sediment s i n th e inne r Mora y Firt h (Hurst 1985a) . Alternating period s o f moderat e uplift , reduction o f relie f an d marin e transgressio n affected th e Highland s durin g Mesozoi c tim e (Hall 1991) . Triassi c sediment s ar e u p to 500m thick agains t th e Grea t Gle n Faul t bu t thi n t o 150m around Elgin (Frostick et al. 1988) . By the

SCOTLAND'S DENUDATIONA L HISTOR Y

end o f th e Triassi c perio d uplif t ha d cease d an d relief wa s considerabl y reduced , an d th e oute r Moray Firt h forme d par t o f a n extensiv e continental plai n o f lo w relie f (Andrew s e t al 1990). Calcrete s an d silcrete s forme d an d th e Rhaetic Se a transgresse d clos e t o th e presen t margins o f th e inne r Mora y Firth . Continue d transgression i n Earl y Jurassi c tim e sa w th e deposition o f fluviatil e san d aroun d th e margin s of th e Mora y Firth . San d an d cla y mineralog y suggests derivatio n dominantl y fro m Devonia n and Carboniferou s cove r rock s t o th e nort h (Hurst 1985a ) an d fro m Moinia n chloriti c metasediments t o th e sout h (Hurs t 1985b) . Thermal domin g i n Mid-Jurassi c tim e i n th e Moray Firt h Basi n cause d dee p truncatio n o f Early Jurassi c and older sediment s an d sediment transfer t o th e Vikin g Grabe n an d inne r Mora y Firth. Crustal collapse i n the central North Sea in Callovian tim e wa s accompanie d b y rapi d sedimentation i n th e inne r Mora y Firt h an d synsedimentary movements along the Helmsdal e Fault (Anderto n e t al . 1979) . Margina l marin e sands oversteppe d the current basi n margin s an d may hav e covere d th e axi s o f th e Grea t Gle n (Hallam & Sellwood 1976 : Wignal l & Pickering 1993). Th e fault s controllin g sedimentatio n i n the inne r Mora y Fort h appea r t o have als o bee n active o n th e adjacen t lan d are a (Robert s & Holdsworth 1999) . Tectonic activit y was renewed a t the JurassicCretaceous boundar y i n th e Mora y Firt h Basin . Uplift o f th e Halibu t Hors t le d t o erosio n o f Carboniferous sandstones . Fault scarps alon g the northern margi n o f th e inne r Mora y Firt h generated coars e mas s flo w deposit s (Anderto n et al . 1979) . Earl y Cretaceou s sediment s late r overstepped th e Helmsdal e Faul t nort h o f Helmsdale t o overli e Jurassi c an d Devonia n sediments (Cheshe r & Lawso n 1983) . I n Morvern, Cretaceous greensand s rest on Moinian schists (Georg e 1966) . A smal l outlie r o f lat e Hauterivian-early Barremia n glauconiti c sand stone rest s o n Devonia n an d basemen t rock s i n eastern Bucha n (Hal l & Jarvis 1994) . By Lat e Cretaceou s tim e th e Highland s ha d been reduce d t o a n are a o f relativel y lo w relief . Cretaceous sequence s alon g th e easter n margi n of th e Hebride s basi n ar e thin, implyin g limite d sediment supply , an d li e clos e t o se a level , implying tectoni c stabilit y (Hancoc k 2000) . Terrigenous sedimentatio n cease d i n th e Mora y Firth wit h th e depositio n o f thic k chal k sequences. O n land , th e sub-Cenomania n sur face, before transgression, carried dee p kaolinitic weathering mantles , late r reworke d t o for m th e highly quartzose sands of Lochaline (Humphrie s 1961) an d th e kaoliniti c Paleocen e sand s an d

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muds of the inner Moray Firth (Carman & Young 1981). The extent of marine transgression i n Late Cretaceous tim e i s unclear bu t depositio n o f th e chalk i n depth s o f severa l hundre d metre s o f water (Hancoc k 1975 ) suggest s that only a small area o f th e Highland s ca n hav e escape d submergence. Around 6 0 Ma, th e passag e o f th e Icelan d plume wa s accompanie d b y majo r magmati c activity in western Scotland (Bell & Jolley 1997) . Magmatism involve d emplacemen t o f igneou s centres, extrusio n o f floo d basalt s wel l beyon d the presen t outcro p an d injectio n o f regiona l dyke form s t o for m th e Tertiar y Igneou s Province. Th e perio d o f magmatis m wa s brief , concentrated betwee n 6 1 an d 5 5 Ma (Jolle y 1997), an d in individual igneous centre s volcanism wa s largel y confine d t o singl e palaeomag netic polarit y interval s o f 0.4- 3 Ma (Musse t 1984). Accelerate d san d accumulatio n i n th e Moray Firth Basin (Liu & Galloway 1997 ) can be linked via sediment routeways and provenance to erosion o f uplifted sourc e area s o n the OrkneyShetland Platfor m an d in the Highlands (Jone s & Milton 1994) . Sediment flux reached a maximum in Late Paleocene tim e and declined int o Eocene time (Jo y 1993 ; Whit e & Lovell 1997) . Small outlier s o f thi n Cretaceou s sequence s occur o n bot h th e wester n (Hancoc k 2000 ) an d the easter n (Hal l & Jarvis 1994 ) margin s o f th e Highlands. A s an y emergen t area s o f th e Highlands ha d bee n reduce d t o lo w relie f b y the end o f Cretaceous tim e (Hal l 1991) , patterns of Tertiar y uplif t ca n b e reconstructe d usin g th e present summi t topograph y o f th e Highland s (Fig. 3) . Th e distributio n o f summit s abov e 800 m defines a zone of maximum uplift; terrai n that no w form s th e mai n watershed s o f th e N W Highlands an d o f th e Grampia n Mountains . I n Northern Scotland , hig h summit s overloo k th e sedimentary basin s o f Th e Minc h an d th e innermost Mora y Firth , basin s wit h margin s that retain attenuate d and localized sequence s of Mesozoic sediment s (Fig . 4) . Majo r differentia l tectonics i s implie d betwee n th e Highland s an d the surroundin g basins . Earl y Tertiar y reactiva tion o f th e Helmsdal e Faul t produce d a majo r fault scarp , no w marke d b y th e lin e o f hill s between Be n Wyvi s an d Helmsdale . Anothe r major escarpmen t existe d i n Early Tertiar y tim e on th e wes t coas t o f th e Norther n Highlands , stretching fro m th e Cuillins to Cape Wrath. Thi s escarpment i s no w dissecte d int o a chai n o f isolated hill s an d hil l groups , includin g th e inselbergs o f Suilve n an d Quinag . It s alignment runs parallel to the edge of the Minch Basin but it is not fault controlled , implyin g that uplift o f the NW Highland s wa s associated wit h significan t

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Fig. 3 . Patterns o f Tertiary uplif t implie d b y Highlan d summi t height s (i n hundreds of metres) . Fin e line s give locations o f diagrammatic cross-section s i n Fig. 4 (longer section lines ) and Fig. 5 (shorter sectio n lines).

warping o n it s wester n margin . Sout h o f th e Great Glen, ther e is also evidence o f differentia l uplift. I n Buchan the preservatio n o f Cretaceou s chalk flint s an d greensan d demonstrate s modes t Tertiary uplif t ye t the Cairngorms, onl y 50 km to the west , eve n toda y reac h 1300 m (Fig . 5) . Differential movement s ar e required , wit h possible downwarpin g toward s th e eas t i n th e Dalradian bel t betwee n Ballate r an d Keit h (Ringrose & Migo n 1997 ) an d dislocatio n a t the easter n edg e o f th e Mount h an d th e Hil l o f Fare (Hall 1987) . O n th e southwester n edg e o f the Wester n Grampian s lie s a zon e o f lowe r summits, centre d o n Cowal , wher e th e presenc e of vesicula r dyke s (Gun n e t al. 1897 ) suggest s relative proximit y t o th e Earl y Tertiar y lan d surface (Fig . 4) . Th e preservatio n o n th e Lom e Plateau o f a smal l Carboniferou s outlie r a t Bridge o f Awe , restin g o n Devonia n lava s (Johnstone 1966) , i s noteworthy , a s ar e th e fragments o f sub-Triassi c surface s (Godar d 1965) foun d i n Morvern . Thes e occurrence s together impl y tha t post-Caledonia n vertica l movements o f the Cowal peninsula and adjacent areas hav e bee n modes t whe n compare d wit h

those that have affected th e main area of the S W Grampians. In earl y Eocen e tim e th e Highlan d are a foundered a s i t move d awa y fro m th e Icelan d plume (Nadi n & Kuszni r 1995 ) an d sedimen tation rate s droppe d i n th e Nort h Se a (Li u & Galloway 1997) . This coincided with the onset of a period , c . 2 0 Ma i n duration , o f humi d an d initially subtropica l conditions , an d apparentl y limited uplift. Dee p kaolinitic weathering covers probably develope d widel y i n associatio n wit h extensive erosion surface s (Hall 1991) . Tectonic activity wa s resume d throughout NW Europ e i n Late Oligocen e time , wit h th e onse t o f majo r uplift o f Fennoscandia (Rohrman et al. 1995 ) and basin developmen t throughou t wester n Britain , including th e Hebridea n regio n (Fyf e e t al . 1993). Th e presenc e o f depositiona l hiatuse s west o f th e Shetland s (Rid d 1981) , deltai c an d lignitic sands east of the Shetlands (Johnson et al. 1993) an d unconformitie s i n th e centra l Nort h Sea (Gatliff e t al. 1994 ) indicate significant uplif t in th e Scottis h are a an d associate d erosio n an d enhanced sedimen t suppl y (Li u & Gallowa y 1997).

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Fig. 4 . Schematic cross-section s illustratin g major morphotectoni c units along two traverses across the Scottis h Highlands.

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Fig. 5 . Schemati c cross-section s illustratin g differentia l tectonic s an d deundatio n i n th e Tertiar y Igneou s Province: Sky e an d Mull (fro m variou s dat a sources) .

Early t o mid-Miocen e tim e wa s th e secon d long perio d o f relativ e tectoni c stabilit y i n th e Tertiary period, lasting for over 1 0 Ma (Le Coeur 1999). Kaoliniti c weatherin g resume d unde r humid, war m t o temperat e condition s (Bersta d & Dypvik 1982) and it is possible that this was an important period for the formation, extension and remodelling o f erosio n surface s b y etc h processes. Miocen e marin e clays , u p t o 8 m thick , together wit h Cretaceous sandstone , ar e reporte d from Leavad , Caithness, as part of a large glacial erratic transporte d fro m th e Mora y Firt h (Crampton & Carruthers 1914) . The occurrenc e appears t o demonstrat e Miocen e marin e sedimentation i n the inne r Moray Firth . A marke d chang e occur s i n th e cla y miner alogy o f Nort h Se a sediment s i n Lat e Miocen e time. A n increasin g conten t o f feldspar, chlorit e and illite (Karllson et al. 1979; Bersta d & Dypvik 1982) reflect s climatic coolin g an d th e inpu t of immature terrigenous material. This material was increasingly sourced first from Fennoscandi a an d later fro m th e Rhin e an d Balti c rive r systems , reflecting uplif t o f source areas. The significance of Neogen e uplif t i n th e morphogenesi s o f northern Britai n ha s lon g bee n recognize d (George 1966) , althoug h it s scal e an d patter n are a s ye t uncertain . Exhumatio n o f th e Chal k from beneat h c . 1 km i n th e inne r Mora y Firt h appears t o hav e bee n achieve d largel y i n Neogene time (Japsen 1997 ; Japsen & Chalmer s 2000) and implies contemporaneou s uplif t o f the Scottish Highlands . Geomorphological evidenc e for Lat e Tertiar y tectonic s i s provide d b y th e apparent warpin g o f mid-Tertiar y erosio n surfaces i n norther n Scotlan d (Godar d 1965) , th e uplift, warpin g an d dislocatio n o f Lat e Tertiar y

surfaces i n wester n Scotlan d (L e Coeu r 1988, 1999) an d th e widesprea d evidenc e o f valle y incision and deepening of topographic basins set into mid-Tertiary erosion surface s throughout the Highlands a t this time (Hall 1991) . The influ x o f ice-rafte d materia l t o th e Hebridean margi n a t 2. 5 Ma mark s th e onse t of mid-latitude climati c deterioration (Stoke r el al. 1994). Episodi c mountai n glaciation i s likel y t o have occurre d thereafte r (Clapperto n 1997 ) but the firs t ic e sheet s reache d th e Nort h Se a Basin only afte r 1 Ma (Andrew s et al. 1990) . Multiple glaciation o f th e Highland s an d th e adjacen t shelves during the Quaternary period brought the transfer o f sedimen t fro m curren t lan d an d nearshore are a to the axia l area of the North Se a Basin an d t o th e continenta l shelve s (Clayto n 1996). On land, each glaciation tended to remove the deposit s o f it s predecesso r s o tha t onl y th e deposits of the last (Late Devensian) ice sheet are usually preserved . The impac t o f th e Quaternar y glaciation s has varie d i n tim e an d space . Th e volum e o f sediment deposite d i n th e centra l Nort h Se a indicates tha t sedimentation , an d hence denuda tion o f th e adjacen t lan d masse s an d shelves , doubled betwee n th e dominantl y non-glacia l conditions of Pliocene and early Pleistocene time and th e condition s o f episodi c ic e shee t glaciation ove r the las t 1 Ma. Th e las t ic e sheet s during Lat e Quaternar y tim e wer e als o les s effective agent s o f erosio n an d transportatio n than those of mid-Quaternary time, reflecting th e earlier remova l o f pre-glacia l weathere d rock , the progressiv e adaptatio n o f th e glacie r be d t o the efficien t evacuatio n o f ice , and th e greate r thickness an d extent of the Elsterian an d Saalia n

SCOTLAND'S DENUDATIONAL HISTOR Y

ice sheet s (Glasse r & Hal l 1997) . Th e Scottis h Highlands als o exhibi t a wid e rang e o f glacia l landscapes, fro m th e deepl y dissecte d terrai n of the wester n Highland s t o th e zone s o f selectiv e linear erosio n o f the Cairngorms an d the limite d erosion o f th e Bucha n lowland s (Linto n 1959 ; Clayton 1974) . Th e averag e dept h o f glacia l erosion acros s Britai n is estimated a t 76 m, with 175m i n mountainou s zone s o f intens e erosio n and a s little a s 1 5 m in zones o f slow-movin g or cold-based ic e (Clayton 1996). This is equivalent to a volum e les s tha n th e tota l amoun t o f Quaternary sedimen t o n the shelve s surrounding Britain, implyin g tha t a significan t componen t has bee n derive d fro m th e dee p erosio n o f material fro m th e inne r shelve s (Clayto n 1996) , including the inner Moray Firth an d The Minch . Mass transfe r o n thi s scal e mus t hav e cause d isostatic uplift i n the glaciated mountai n areas of western and northern Britain. The depth of som e of th e wes t coas t fjord s ma y reflec t glacia l incision int o th e still-risin g edg e o f th e N W Highlands. Former cover rocks in the Scottish Highlands Palaeogeographical map s o f the Highland s hav e tended t o sho w th e are a a s a persisten t topographic hig h (e.g . Anderto n e t al. 1979 ; Ziegler 1981) , bu t th e exten t an d thicknes s o f former cove r rock s i n th e Scottis h Highland s remain controversial . Th e regio n i s routinel y seen a s a sourc e are a fo r sediment s tha t hav e accumulated i n th e surroundin g basins . Thi s seems consisten t wit h evidenc e o f relativel y modest depth s (
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Current model s o f Highlan d sourc e are a evolution base d o n AFT T (Thomso n e t al . 1999) sugges t that : (1 ) lat e Palaeozoi c cove r rocks u p t o 3k m thic k formerl y covere d th e basement rock s o f th e Highlands ; (2 ) Tertiar y erosion has removed 1- 2 km of rock from abov e the current topography. There i s littl e doub t tha t lat e Palaeozoi c rocks onc e covere d considerabl y large r area s of the Highland s massi f tha n a t present . Devonia n outliers occu r widel y aroun d th e coasta l ri m o f the Mora y Firt h an d reac h thicknesse s o f several hundre d metres i n fault-bounde d basins. Carboniferous rock s formerl y extende d acros s part o f th e S W Highland s (Georg e 1960 ) an d reach a thicknes s o f ove r 1 km i n th e inne r Moray Firt h (Thomso n e t a l 1999) . Ye t i t i s unlikely tha t lat e Palaeozoi c rock s onc e covered al l o f th e Highland s an d henc e pre Mesozoic overburde n thicknesse s o f 2- 3 km seem unreasonable . Th e Lowe r an d Middle Ol d Red Sandstone s aroun d th e Orcadia n Basi n were largel y derive d fro m mountain s site d i n the are a o f th e curren t Easter n Grampian s an d Western Highland s (Mykur a 1983 ) and ther e i s little obviou s sig n tha t thes e mountai n area s were eventuall y wor n down sufficientl y t o b e buried. Typically , a regiona l unconformit y separates th e Uppe r Ol d Re d Sandston e fro m older sediment s an d thi s i s ascribe d t o regiona l tectonics (Mykur a 1983) . I n Morayshire, o n th e margin o f th e Orcadia n Basi n wher e uplif t might be expecte d t o be limited , th e Uppe r Old Red Sandston e rest s i n place s directly o n th e Moine sequenc e (Hom e 1923) , indicatin g removal o f th e Lowe r an d Middl e Ol d Re d Sandstone before deposition . A simila r situation occurs i n th e Midlan d Valley , where th e Uppe r Old Re d Sandstone , includin g conglomerate s with pebble s o f metamorphi c rock s derive d from th e Souther n Highlands, rests with marked angular unconformit y o n th e Lowe r Ol d Re d Sandstone (Franci s et al 1970) . B y implication , substantial remova l o f th e Lowe r an d Middl e Old Re d Sandston e ha d bee n achieve d befor e the en d o f Devonia n time . The Highland s als o acte d a s a sourc e regio n for Carboniferou s sediment s i n th e Midlan d Valley (Franci s 1991 ; Guio n e t al, 2000 ) an d Namurian-Westphalian sandstone s i n th e Stir ling distric t contai n heav y mineral s ultimatel y derived from low-grade metamorphic rocks north of th e Highlan d Boundar y Faul t (Franci s e t a l 1970). Alon g th e easter n margi n o f th e Minc h Basin numerou s smal l outlier s o f Permo-Tria s occur. Only at Inninmore on the Sound of Mull is Carboniferous sedimen t foun d underlyin g Tria s units an d here th e Carboniferou s rock s ar e only

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100-160m thick . Elsewher e th e Permo-Tria s sequence rest s o n olde r rock s (Johnston e & Mykura 1989) . B y implication , th e Carbonifer ous rock s wer e remove d befor e th e star t o f th e Trias perio d o r wer e neve r lai d dow n t o an y significant thicknes s alon g th e wester n edg e o f the Northern Highlands . The thic k sequence s o f Mesozoi c clasti c sediments i n th e centra l Nort h Sea , whic h might b e take n t o indicat e dee p erosio n o f th e Highlands, deriv e onl y i n part fro m th e Scottis h area. A majo r contributio n o f materia l fro m Fennoscandia occurre d durin g Triassic an d early Jurassic time (Ziegle r 1981) . Erosio n o f intra basinal high s als o provide d sedimen t (Andrew s et al. 1990) . Ye t th e presenc e o f igneou s an d metamorphic debri s i n sediment s a t the margins of the Highland s implies continuin g erosion an d thus exposur e o f th e basemen t o f th e Highland s area (Hudso n 1964 ; Hurst 1985a , 1985b) . Th e volume o f debri s als o implie s a significant , but unknown depth of erosion. I n the Minch Basin, it is possibl e to quantif y depth s of Permia n to Jurassic erosio n o f th e terrai n surroundin g th e Sea of the Hebrides Trough . Preserve d sedimen t volumes ar e clos e t o th e originall y deposite d volumes (Stee l 1978 : Fyf e e t a l 1993 ) an d suggest tha t c. 280 m o f roc k wa s remove d i n Permo-Trias tim e an d c . 210 m i n Jurassi c tim e from th e mai n contributin g are a o f th e souther n Outer Hebride s Platform . Thes e relativel y lo w values ar e consisten t wit h th e remova l o f onl y thin cover from the Highlands . Depths o f Tertiar y erosio n fro m th e sourc e area o f th e Mora y Firt h Basi n i n th e Highland s can be quantified using sedimen t volume s in the North Sea . The Nort h Se a acte d a s a sedimen t trap throughou t Tertiar y tim e (Li u & Gallowa y 1997), wit h onl y a narro w an d shallo w connec tion t o th e Norwegian-Greenlan d Se a (Nielse n et al. 1986) . In early Tertiar y tim e the dominant sediment sourc e wa s the Scottis h Highlands an d the Orkney-Shetlan d Platfor m bu t throughou t Neogene tim e sediment was increasingly derived from Scandinavi a an d th e grea t rive r system s o f NW Europe . Roug h calculation s indicat e th e removal o f 600-800 m o f roc k fro m th e contributing are a eas t o f th e mai n Scottis h watershed durin g Tertiar y tim e (Hal l 1991) . A s only 25% of the Scottish lan d area now lies above 300m (Haynes 1983) , i t appears that the summit envelope surfac e o f th e Highland s a t 900 1200m (Fig . 3) lie s onl y a fe w hundre d metre s below th e uplifte d sub-Cretaceou s lan d surface . On these estimates, there appears to be a good fit between th e volume o f rock remove d b y erosio n from the Highlands and that received i n the North Sea Basin . I n th e inne r Mora y Firth , soni c

velocities i n th e Kimmeridg e Cla y indicat e removal o f aroun d 1 km o f overburde n (Hilli s et a l 1994 ; Thomson & Hilli s 1995 ) and thi s seems generally compatible with the existence of hills, suc h a s Be n Rinnes , a t u p t o 800 m elevation o n bot h th e norther n an d souther n margins o f the basin. In contrast , the remova l of over 1. 1 km o f roc k fro m abov e th e presen t terrain o f Caithnes s an d Sutherlan d (Thomso n et al 1999 ) seems excessive. On Morvern, AFTA data indicate c. 1.7k m of Tertiary erosio n o n the Strontian Granodiorit e (Thomso n e t a l 1999) , yet th e curren t lan d surfac e nearb y retain s thi n sequences o f Mesozoic sediment s buried beneath Early Tertiary lavas (Johnstone & Mykura 1989). Further judgement on thes e issue s must await continuing wor k usin g AFT T an d (U-Th)/H e thermochronology. Non e th e less , onshor e geological an d geomorphologica l dat a fro m th e passive margin s of souther n Afric a (BJ . Bluck , pers. comm) an d southeastern Australi a (Kohn & Bishop 1999 ; Bishop & Goldrick, 2000) see m to suggest tha t AFTT ma y overestimat e the depths of forme r cove r rock s an d o f denudation . Of course, th e regiona l patter n o f AFT T data , an d the denudatio n tha t the y imply , cannot b e take n to provid e a detaile d histor y o f an y particula r locality. Th e lava s o f 10 0 Ma ag e o n th e sout h coast o f Ne w Sout h Wales , fo r example , ar e inconsistent wit h dee p denudatio n havin g occurred alon g an d acros s th e whol e o f th e coastal stri p belo w th e S E Australia n escarp ment, a n interpretatio n tha t ha s bee n implici t i n much o f th e AFT T discussion . Likewise , ther e are man y area s throughou t th e S E Australia n highlands, a s i n th e Scottis h Highlands , wher e ancient landscap e element s (datin g from Meso zoic tim e i n th e S E Australia n case ) hav e bee n identified (Youn g 1981 ; Bir d & Chiva s 1989 : Twidale 1994 ; Twidale & Campbel l 1995 ; Hill 1999). Hil l (1999 ) ha s argue d tha t variation s in relief i n th e S E Australia n highlands and acros s the coasta l stri p belo w th e escarpment , an d differential preservatio n o f Mesozoi c landscap e elements, ma y explai n th e apparen t conundru m in S E Australia of kilometre-scale denudatio n in Late Mesozoi c tim e an d th e preservatio n o f Mesozoic landscap e elements . Wher e local , detailed AFTT data are available in SE Australia, particularly i n areas wher e the regional structur e is dominate d b y individua l faul t blocks , differ ential faul t bloc k movemen t i s clearl y indicated by th e AFTT dat a (Koh n & Bisho p 1999 ; Kohn et a l 1999) . Elevate d geotherma l gradient s would als o assis t i n minimizin g th e amount s o f denudation require d fo r Lat e Mesozoi c fission track age s t o cro p ou t a t th e presen t groun d surface i n S E Australia , but ther e i s currentl y a

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Fig. 6 . Early Tertiary differentia l tectonic s an d sediment transport (from variou s data sources).

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strong consensu s amon g AFT T researcher s tha t geothermal gradient s alon g th e S E Australia n continental margi n wer e probabl y no t signifi cantly greate r tha n 25-30°Ckm" 1 i n Lat e Mesozoic tim e (Dumitr u e t al. 1991 ; Brow n et aL 1994) . These matter s wil l almost certainl y emerge a s importan t fo r th e Scottis h Highland s as more detai l become s available .

Cenozoic uplif t and denudation o f the Scottish Highland s The timing , amoun t an d patter n o f uplif t i n th e Highlands throug h Cenozoi c tim e ar e graduall y becoming clearer . Uplif t appear s t o hav e bee n intermittent, wit h th e mai n phase s centre d o n Late Paleocen e an d Lat e Oligocen e tim e an d a later phas e o f continuin g uplif t startin g i n Lat e Miocene tim e (Hal l 1991 ; Li u & Gallowa y 1997). In terms o f sediment suppl y to the Mora y Firth, th e firs t o f thes e phase s i s th e mos t important (Nielse n e t al 1986 ; Li u & Gallowa y 1997), implyin g tha t thi s wa s th e perio d o f maximum Tertiar y uplif t i n th e Highlands . Th e timings of these vertica l movement s matc h thos e in wester n Norwa y (Rii s 1996) . Indeed, th e fou r major unconformitie s datin g fro m mid-Paleo cene, mid - t o lat e Oligocene , mid-Miocen e an d latest Pliocen e tim e ar e remarkably consisten t in development acros s th e centra l an d norther n North Sea and the Norwegian Shel f (Huuse et al., 2001). I t appear s therefor e tha t th e Scottish Norwegian sectio n of the NE Atlantic margin has reacted contemporaneousl y t o crustal break-up in Paleocene time , intraplat e deformatio n i n Neo gene tim e (Stuevol d & Eldholm 1996 ) an d to the onset o f regiona l glaciation . Th e magnitud e o f these events , however , ma y wel l hav e varie d spatially. Th e maximu m uplif t o f souther n Fennoscandia appear s t o hav e occurre d fro m late Oligocen e t o Pliocen e tim e (Dor e e t al . 1998; Japse n & Chalmer s 2000 ) an d thu s postdates th e mai n uplif t phas e i n th e Scottis h Highlands. Th e reconstructe d buria l histor y o f the Chalk in the western Nort h Sea (Japsen 1997 ) implies tha t part s o f easter n Scotlan d ma y hav e experienced maximu m Tertiar y uplif t i n Neo gene time . Major uplif t occurre d a t th e Paleocene Eocene boundar y vi a a combinatio n o f permanent an d transien t (dynamic ) uplif t (Jones e t al . 2002). Thi s uplif t even t i s consistent wit h th e evidenc e o f rapi d loca l denudation i n th e Hebridea n Igneou s Province . As muc h a s 2k m o f roo f an d cove r rock s ar e missing fro m abov e th e igneou s centre s o f Skye, Mul l an d Arra n (Georg e 1966) , bu t thi s unroofing i s likel y t o hav e bee n spatiall y

variable (Fig . 6 ) (se e foregoin g discussio n o f SE Australia) . Holness (1999) , fo r example , ha s argued tha t th e metamorphi c grad e o f th e Torridonian arkos e intrude d b y th e Rhu m Igneous Comple x implie s 500-55 0 m o f over burden, whic h i s les s tha n th e curren t relie f o n Rhum, an d 'point s t o a topograph y a t th e tim e of meltin g [i.e . metamorphism ] a s ver y simila r to tha t o f today ' (p . 538) . Thi s erosio n o f th e Hebridean Igneou s Provinc e too k plac e remark ably quickly . Som e 2 km o f basal t wa s remove d from Mul l betwee n 5 8 an d 5 6 Ma an d th e unroofing o f th e Wester n Granophyr e o f Rhu m took plac e withi n 3 Ma (Emeleu s 1983) . Sedimentary interbed s i n th e Rhu m Centra l Complex demonstrat e tha t th e comple x wa s already unroofe d an d wa s undergoin g activ e erosion a s th e lava s wer e bein g extrude d (Emeleus & Forste r 1979 ; Holnes s 1999) . Palynological studie s o f forme r interbasalti c vegetation sugges t tha t durin g th e 0.2 4 Ma o f existence o f th e Sky e Lav a Fiel d ther e wer e two majo r subsidenc e phase s an d on e uplif t phase. Therma l domin g relate d t o th e emplace ment o f th e Cuilli n centr e cause d elevatio n t o altitudes i n exces s o f 1200 m a.s.l . (abov e se a level; Jolle y 1997) . Lav a erupte d mainl y fro m fissures spread ou t over a n area o f c. 40 000km 2 stretching fro m Harri s t o th e Firt h o f Clyd e over th e axi s o f th e dyk e swarm s (Presto n 1982). B y th e en d o f th e magmati c phase many igneous centre s wer e reduce d t o clos e t o present levels , a s show n b y th e lat e lava s resting o n th e Wester n Granophyr e o f Rhu m (Emeleus 1983 ) an d th e deepl y denude d basalts lying beneat h th e Sgur r o f Eig g pitchston e (Dickin & Jone s 1983) . Thi s dee p erosio n implies tha t th e lav a field s wer e als o largel y stripped i n earl y Tertiar y time , togethe r wit h any underlyin g Palaeozoi c an d Mesozoi c sediments. A s Permia n sandstone s o n Lewi s and Jurassi c sandstone s o n Sky e ar e associate d with maximu m burial depth s o f onl y 1.5-2k m (Carter e t al . 1995) , i t i s likel y tha t th e Paleogene uplif t even t wa s a ke y phas e i n th e stripping o f Paleozoi c an d Mesozoi c cove r rocks throughou t th e Inne r Hebrides . The Leava d clay s probabl y relat e t o earl y Miocene marin e transgressio n (Crampto n & Carruthers 1914) , althoug h a moder n fauna l study o f thes e intriguin g deposits i s required . If correct, a n earl y Miocen e ag e implie s tha t dee p erosion o f th e Jurassi c an d lowe r Cretaceou s sequence i n th e inne r Mora y Firt h (Hilli s et al . 1994) wa s largel y complete d betwee n Lat e Paleocene an d Late Oligocene time . Jones e t a l (2002 ) interprete d gravit y dat a from th e N E Atlanti c i n terms o f Neogene uplif t

SCOTLAND'S DENUDATIONAL HISTOR Y

but because of the absence of later Tertiary rock s from muc h of the Highlands area it is not easy to assess amount s of subsequent uplift an d denudation. Th e smal l fault-bounde d Lat e Oligocen e basins o f The Minc h an d the NW Scottis h shel f provide important evidence of a phase of tectonic activity tha t ca n b e trace d throughou t western Britain an d ma y b e relate d t o mino r plat e reorganization (Evan s et al 1991 , 1997) . It may be significan t tha t the Late Oligocen e floodplai n and swamp deposits in these basins do not rest on Eocene sediment s bu t o n kaolinize d basement , implying prolonge d weatherin g unde r humi d conditions in Eocene time. O n the NW Scottish shelf, th e Oligocene sediment s ar e succeeded b y Miocene shallo w marin e sand s (Evan s e t al . 1997). The sequence implies that despite Eocene crustal collaps e Th e Minc h an d th e adjacen t shelf remaine d abov e se a leve l unti l marin e transgression i n Miocen e time . Crustal movemen t resume d i n Lat e Miocen e time (Stoke r e t al . 1994) . Japse n (1997 ) considered tha t th e amoun t o f exhumatio n during Neogen e uplif t wa s equivalen t i n Britai n to tha t achieve d i n Lat e Paleocene-Earl y Eocene tim e (se e als o Japse n & Chalmers , 2000). Thi s seem s unlikely , a s Neogen e sediments i n th e Nort h Se a ar e largel y source d from Scandinavi a an d N W Europe , rathe r tha n Britain (Jord t e t al . 1995) . Moreover , th e larg e volumes o f coars e clasti c sediment s o f Lat e Paleocene an d earl y Eocen e ag e i n th e Mora y Firth contras t markedly with the more restricte d sequences o f Neogen e mud s (Li u & Gallowa y 1997). Th e increasin g resolutio n o f th e Neogene successio n i n th e centra l an d norther n North Se a an d o n th e Norwegia n shel f help s t o constrain th e timin g o f thi s lat e uplift , wit h pulses o f terrigenou s sediment s eviden t i n Lat e Miocene tim e and , coinciden t wit h th e firs t Scandinavian ic e sheets , i n Lat e Pliocen e tim e (Eidvin e t a l 2000) . The thick Neogen e sequence s i n the Faeroe Shetland basi n syste m (Stoke r e t al . 1993 ) ar e more difficul t t o accoun t for . Likewise , Jone s et al . (2002 ) reporte d tha t th e volum e o f Paleocene sedimen t i n th e Faeroe-Shetlan d basins canno t b e accounte d fo r b y denudatio n of a source area in NW Scotland. They suggeste d that Faeroe-Shetland basin sediment s ma y also have been derive d fro m sourc e area s othe r tha n NW Scotland , such as the Faeroe Islands regio n or more widely in northern and western Scotland, and a similar explanation may have to be invoked to reconcil e th e thicknes s o f th e Neogen e sequence i n the Faeroe-Shetland basi n system s and th e apparentl y mino r Neogen e uplif t o f Scotland.

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The signatur e o f thes e Neogen e vertica l movements i s recorde d b y element s o f th e Highland topography. In the Northern Highlands, Godard (1965 ) recognize d thre e majo r erosio n surfaces: a surfac e betwee n 40 0 an d 550m , which cut s acros s part s o f th e Tertiar y igneou s centres an d s o i s post-Paleocen e an d possibl y Eocene i n age ; a n intermediat e an d extensiv e Scottish Surfac e a t c . 300 m o f possibl e Oligocene age ; th e lo w coasta l platea u o f th e Niveau Pliocene a t c. 100 m. Assuming that these erosion surface s originate d clos e t o se a level , i t has bee n estimate d tha t uplif t i n th e N W Highlands i s o f th e orde r o f 400 m i n th e las t 40Ma(LeCoeur 1999) . The long-ter m tren d ha s bee n toward s th e progressive tiltin g or downwarping o f the High lands toward s th e inne r Mora y Firt h Basin . Despite a degre e o f glacia l diversio n an d disruption th e mai n drainag e route s continu e to flow toward s th e Mora y Firth , jus t a s i n Earl y Tertiary an d Lat e Devonia n time . Th e mos t extensive erosion surfac e recognized south of the Moray Firt h i s th e Easter n Grampian s Surface , comprising th e ramp-lik e interfluve s o f th e Monadliath, th e Dee-Do n watershe d an d th e Mounth. Thes e surface s slop e fro m thei r inne r margins around the Cairngorms a t around 80 0 m towards the inner Moray Firth and the North Sea, dropping t o elevation s o f c . 500 m a t th e oute r margins (Hall 1991) . Recent morphometri c analysi s ha s confirme d that regiona l tiltin g wa s fa r fro m uniform . Erosion surface s between 20 0 and 600 m on the Monadliath are strongly influenced b y the Grea t Glen Faul t an d th e Ericht-Laido n Fault . Thi s may imply a significant tectoni c event in mid- to late Tertiary time in which an extensive mediumlevel erosio n surfac e wa s disrupte d b y bloc k movement (Ringros e & Migon 1997) . Ringros e & Migo n als o identifie d a possibl e zon e o f flexure located i n the Dalradian belt between the high tops of the Cairngorms an d the lowlands of Buchan. Lat e Neogen e faultin g ha s als o bee n proposed alon g the S E coasts o f Rhum an d Col l along th e Camasunary-Skerryvor e Faul t (L e Coeur 1988 ) an d i n th e Elgi n are a (Hal l 1991 ) and ma y represen t par t o f a continuing patterns of neotectoni c activit y (Muir Wood 1989) . Denudation and landscape evolution The mai n morphotectoni c unit s i n th e Scottis h area were alread y i n existence b y th e en d o f th e Palaeozoic era . The Orkney-Shetland Platform, the Highland s an d Souther n Upland s massif s have remained abov e se a level for mos t o f postPalaeozoic tim e an d hav e she d sedimen t t o

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surrounding basins . Th e persistenc e o f differential movement s betwee n basemen t massif s an d sedimentary basin s ha s allowe d conservation , during surfac e lowering , o f majo r topographi c features fro m a s lon g ag o a s Devonia n time , including watershe d zones , drainag e pattern s and, especially , th e scarp s a t th e margin s o f th e buoyant basement massifs . Th e Helmsdale Fault , for example , ha s bee n intermittentl y activ e throughout Mesozoi c an d Cenozoi c tim e (Andrews e t al 1990) . Superimposed o n thi s physiographi c frame work ar e th e mai n Tertiar y landforms . Th e magnitude o f Paleocen e uplif t an d denudatio n means tha t thes e Tertiar y landform s ar e o f mid to lat e Tertiar y age . Onl y i n area s o f minima l displacement, suc h a s Bucha n an d Caithness , i s the preservatio n o f extensiv e Mesozoi c land forms feasible . Exhumed terrain s ar e o f restricte d exten t i n Scotland. The y includ e th e rugge d sub Torridonian surfac e i n N W Scotlan d (Godar d 1957; Stewar t 1972 ) an d th e equall y irregula r sub-Devonian surfac e expose d beneat h th e out liers around the margins of the Moray Firt h Basi n (Godard 1965) . Asid e fro m area s suc h a s these , the oldes t landform s recognize d i n Scotlan d occur i n the lowland s o f Buchan. This i s an are a of long-ter m relativ e stabilit y wher e typica l Highland rock s ar e associate d wit h surface s o f low relie f (Clayto n & Shamoo n 1999) . I t is als o an are a o f ver y limite d glacia l erosion . Th e antiquity of the landscape i s demonstrated b y the Cretaceous residue s i n th e for m o f chal k flint s within th e quartzite - an d kaolin-ric h Bucha n Gravels (Hal l 1987 ) an d a n outlie r o f Lowe r Cretaceous greensan d (Hal l & Jarvi s 1994) . These residue s res t o n weathere d igneou s an d metamorphic rocks , know n i n borehole s t o extend t o depth s o f man y ten s o f metre s (Hal l 1986). Largely b y correlation wit h clay mineral s in Nort h Se a sediments , th e spatiall y restricte d kaolin-rich weatherin g profile s hav e bee n assigned t o pre-Pliocen e tim e an d th e les s mature, bu t stil l dee p sand y weatherin g profile s to Pliocene and Pleistocene time (Hall 1985 ; Hall et al . 1989) . Give n th e apparen t stabilit y o f th e area, i t is conceivable tha t som e o f the kaolinitic weathered material s ar e older , survivin g fro m Paleogene tim e o r exhume d fro m beneat h Lat e Cretaceous cover rock s (Hal l 1993) . Deep weatherin g i s o f fundamenta l import ance in understanding the nature and evolution of the pre-Quaternar y relie f throughou t Scotlan d and N W Europ e (Godar d 1965 ; Thoma s 1989 ; Migon & Lidmar-Bergstrom 2001) . Its presenc e has bee n use d widel y a s a n indicato r o f th e preservation o r limited modification of preglacial

forms (Linto n 1951 ; Godar d 1961 ; Hal l & Sugden 1987) . Ther e i s often a close correspon dence between pre-glacia l morphology, level s o f rock resistanc e t o chemica l weatherin g (Godar d 1962) an d deep weatherin g pattern s (Hal l 1986) . The Tertiar y perio d i n Scotlan d wa s a tim e o f warm to temperate humid climates when most, if not all , o f th e countr y stoo d abov e bas e level , conditions favourin g th e dee p penetratio n o f weathering. The sustaine d etching out o f contrasts i n roc k resistance b y chemica l processe s le d t o th e formation o f majo r landform s o f differentia l weathering an d erosion . Thes e includ e valle y systems, such as the pre-glacial headwaters of the Spey, Do n an d Dee , wher e ther e i s pervasiv e litho-structural control s o n valle y alignmen t (Threlfall 1981) . Dee p topographi c basins , some floore d b y Devonia n sediments , are strung out alon g man y o f th e valley s tha t draine d eas t from th e mai n watershed s i n th e N W Highland s and Grampia n Highland s toward s th e Mora y Firth (Fig . 4) . Th e larges t example s includ e the Rannoch, Atholl and Naver basins, with areas of more tha n 500km 2 (Linto n 1951) . I n N E Scotland, th e basi n floor s ar e preferentiall y located o n biotite-bearin g granit e an d gabbr o and borehole s sho w widesprea d dee p weath ering, reachin g depths o f a s muc h a s 5 0 m (Hall 1986, 1991) . Thes e susceptibl e rock s hav e provided th e foc i fo r weatherin g an d erosio n a s the surroundin g terrain wa s uplifted . In counter point stan d inselbergs . Thes e isolate d hill s include thos e o f resistance , notabl y quartzit e hills suc h a s Schiehallion . Inselbergs of position also occu r (Godar d 1965) , where the isolation of the hil l mas s appear s t o b e a resul t o f back wearing o f slopes . A fe w exhume d hill s occur , notably th e sub-Devonia n inselber g o f Scarabe n in Caithness (Crampton & Carruthers 1914) . The quartzite inselber g o f Mormon d Hil l i n Bucha n may b e a Mesozoic relic , as i t is associated with deep kaolinization and lies close to or at the level of th e sub-Cenomania n surfac e (Hal l 1987) . The landform s o f differentia l weatherin g an d erosion occu r a s mesoscal e feature s a s par t o f major erosio n surfaces . Th e Bucha n Surface , a t an elevatio n o f c . 100m , include s mos t o f th e lowlands of NE Scotland, apart from th e glacially modified coasta l strip. It is an etch surface, where subtle difference s i n roc k resistanc e giv e ris e t o hills an d depression s an d dee p weatherin g i s widespread. Th e origin s of the eastern part of the surface dat e back to late Mesozoic time , for Late Cretaceous greensan d an d chal k wer e deposite d on it s surface , bu t i t ha s a lon g histor y an d it s various element s ar e o f differen t age . Th e preservation o f unwor n flint s a t th e bas e o f th e

SCOTLAND'S DENUDATIONA L HISTORY

flint gravels in Buchan testifies t o th e proximity of the sub-Cenomania n surfac e (Bridgland et al. 1997; Merritt et al. 2002). O n the high ground of central Bucha n a t 100-150 m highl y kaoliniti c saprolites an d th e flin t gravel s themselve s hav e been ascribe d a pre-Pliocen e ag e (Hal l 1985) . The lowe r tier s o f th e terrai n suppor t dee p sandy weatherin g cover s an d appea r t o b e o f Plio-Pleistocene age . It is likely tha t high-level erosio n surface s ar e also polycycli c an d spatiall y diachronous . Th e surfaces a t 800 m i n th e Gaic k Fores t (Hal l & Mellor 1988 ) an d a t hig h elevation s i n th e Cairngorms (Hal l 1996 ) retai n pocket s o f dee p sandy weatherin g datin g fro m th e lates t Plio Pleistocene phas e o f etching . Withi n th e Cairngorms mountain s ar e a rang e o f majo r paleic form s whic h hav e a longe r history , including high-leve l basin s an d ope n valley s (Hall 1996 ) an d th e majo r depressio n o f th e Upper Avo n Embaymen t (Linto n 1950) . Th e precursors o f th e majo r Tertiar y river s o f N E Scotland, th e De e an d th e Don , wer e alread y established in Paleocene tim e and fed material to the Gannet Fan in the North Sea (Morton 1979) . The Cairngorms and the Eastern Grampians have been a n are a o f positive relie f sinc e a t least th e start of Tertiary time and these headwater erosion surfaces have evolved far above sea level. Phase s of valley incision, indicated b y benches o n valley and basin sides, ar e likely to have been driven by changes i n loca l bas e level s rathe r tha n i n response t o regional uplift .

Discussion an d conclusion s Variations in th e characte r and rate o f suppl y of sediments t o th e offshor e regio n ha s generall y been interprete d i n term s o f sourc e are a responses to Cenozoic regional tectonics. Indeed, Liu & Galloway (1997 ) wer e explici t abou t the link between uplif t an d erosion i n the North Se a Basin: Tectoni c uplift o f source area s exercise s the commandin g rol e i n modulatin g lon g ter m sediment suppl y to the North Sea' (p . 1506). The precise proces s linkage s betwee n uplif t an d enhanced denudatio n are , however , no t alway s as clea r a s i s implie d b y th e assume d lin k between uplif t an d enhance d sedimen t flux (see Introduction). White & Lovel l (1997 ) acknowledge d tha t a simple proces s lin k betwee n a n increas e i n offshore sedimentatio n rat e an d th e uplif t o f Scotland canno t be assumed , an d suggested tha t uplift ma y be linked closel y i n time to increase s in offshore sedimentation via an uplift-driven fall in relativ e se a leve l an d a n associate d mobiliz ation o f sediment s tha t wer e i n storag e i n th e

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nearshore are a an d o n th e continenta l shelf . I t seems clear, none the less, that parts of Scotland, for example , the Tertiary Igneous Province, have responded rapidly to presumed uplift events. This is probabl y becaus e thes e wester n area s ar e composed o f smal l catchment s tha t ar e wel l connected to base level and that, therefore, would be expected t o respond rapidly to a fall in relative sea level. The preservation o f ancient landscap e elements throughou t Scotlan d demonstrates , however, tha t suc h rapi d respons e canno t b e assumed fo r al l areas , an d tha t no t al l area s ar e equally sensitive to base-level changes. Variations in lithology will also be expected to be associated with variation s i n landscap e sensitivit y (se e Brunsden & Thorne s 1979) ; lithologica l vari ations in the respons e to Cenozoi c weatherin g regimes i n th e Scottis h Highland s poin t t o a variant o f suc h sensitivity . I n th e cas e o f th e Tertiary Igneou s Province, i t is also worth noting that constructio n o f th e Province' s centra l volcanic edifices and the extrusion of the regional lava fields must have caused a relative elevation of the lan d surface , i n tur n triggerin g incisio n an d enhanced denudation. The relative magnitudes of this latter effect an d uplift remain to be quantified. Many problems remain in understanding longterm landscape development in Scotland and yet there are many promising lines of enquiry. There is a n urgen t nee d fo r detaile d morphometri c analysis o f th e relie f comparabl e wit h tha t available i n souther n Fennoscandi a (e.g . Lid mar-Bergstrom e t al . 2000) . Th e scattere d outliers o f sedimentar y rock s o f Devonia n t o possible Miocen e ag e tha t occu r o n lan d i n Scotland ar e importan t archive s o f information regarding th e evolutio n o f th e terrain , bot h vi a provenance studies and by investigation of burial histories. Th e dyk e system s o f lat e Caledonian , Carboniferous, Permia n an d Paleogen e ag e that criss-cross muc h o f Scotlan d hav e ye t t o b e examined i n detai l wit h regar d t o depth s o f emplacement beneat h contemporar y lan d sur faces. Saprolite s occu r throughou t Scotland , including burie d an d formerl y burie d pre Cenozoic weatherin g mantle s an d Cenozoi c deep weatherin g profiles. Thes e saprolite s await detailed mineralogica l stud y an d datin g (Hal l 1993), using techniques such as K-Ar datin g of mica clays (Sturt et al. 1979 ) and D-H analysi s of kaolin s (Gil g 2000) . Finally , ther e ar e opportunities t o tie in the uplift an d denudational histories of regional sourc e areas within Scotland with sub-basin s an d fan s i n th e offshor e area , such as the Tertiary Barra Fan NW of The Minch (Stoker et al 1993 ) and the Early Tertiary Gannet Fan fed by rivers draining the Eastern Grampians (Gatliffetai 1994) .

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The preservatio n o f ancien t landscap e remnants o f varyin g spatia l exten t i n th e Scottish Highland s point s t o th e type s o f spatial variabilit y o f denudatio n tha t hav e been identifie d i n mor e detaile d studie s i n southeastern Australia . Jone s e t al (2002 ) ha s called fo r improve d quantitativ e understanding of th e Cenozoi c denudatio n of norther n Britain and i t i s therefor e timel y fo r thermochronolo gically base d studie s o f th e denudatio n o f Scotland: (1 ) t o mov e beyon d regiona l denudational studie s t o mor e focuse d AFT T studie s and (U-Th)/H e analysis in apatite (Zeitler et al. 1987) t o identif y spatia l variabilit y i n denuda tion; (2 ) t o us e AFT T an d (U-Th)/H e analysi s to asses s th e exten t o f differentia l movement s of crusta l blocks ; (3 ) t o us e th e (U-Th)/H e system t o asses s th e age(s ) o f majo r landscap e elements feature s (e.g . Hous e e t al . 1998) ; (4 ) to us e cosmogeni c isotop e analysi s t o asses s i n more detai l th e spatia l an d tempora l variation s in Quaternar y denudation . Despite th e adven t o f thes e variou s recentl y developed techniques , th e majo r challenge s o f dating landscape s an d landforms , an d o f determining th e timing o f uplift an d denudation, remain. S o far, the excellent tempora l resolutio n available i n th e offshor e recor d elude s u s i n studying the onshore, source area record, but this shortcoming doe s no t justify , fo r example , oversimplified assumption s concernin g th e relationships betwee n uplift an d denudation. A.M.H. thank s th e Carnegi e Trus t fo r th e Univer sities o f Scotlan d fo r financia l suppor t o f fieldwork . P.B. gratefull y acknowledge s th e suppor t o f th e Australian Researc h Council . Th e CRUS T projec t (Constraining Regiona l Uplift , Sedimentatio n an d Thermochronology) i s supporte d b y a Scottis h Higher Educatio n Counci l Researc h Developmen t Grant. W e thank K . Lidmar-Bergstrom, J . Cartwrigh t and th e Edito r fo r thei r review s an d suggestion s t o improve thi s contribution , an d M . Shan d fo r preparing th e figures . References ANDERTON, R. , BRIDGES , PH. , LEEDER , M.R . & SELLWOOD, B.W . 1979. A Dynamic Stratigraphy o f the British Isles. Alle n an d Unwin, London. ANDREWS, I.J., LONG, D., RICHARDS , PC., THOMSON , A.R., BROWN , S. , CHESHER , J.A . & MCCORMAC , M. 1990 . The Geology o f th e Moray Firth. Britis h Geological Survey , Keyworth . BELL, B.R . & JOLLEY , D.W . 1997 . Applicatio n o f palynological dat a t o th e chronolog y o f th e Palaeogene lav a field s o f th e Britis h Province : implications fo r magneti c stratigraphy . Journal o f the Geological Society, London, 154, 701-708.

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morphological surface s with offshore data. Global and Planetary Change, 12, 331-357. RINGROSE, P.S . & MIGON , P . 1997 . surfaces . In : WIDDOWSON, M . (ed. ) Palaeosurfaces: Recognition, Reconstruction and Palaeoenvironmental Interpretation. Geologica l Society , London , 25-36. ROBERTS, A.M . & HOLDSWORTH , R.E . 1999 . Linkin g onshore an d offshor e structures : Mesozoi c exten sion i n th e Scottis h Highlands . Journal o f th e Geological Society, London, 156, 1061-1064 . ROHRMAN, M. , BEEK , P.v.d. , ANDRIESSEN , P . & CLOETINGH, S . 1995 . Meso-Cenozoi c morphotec tonic evolutio n o f souther n Norway : Neogen e domal uplif t inferre d fro m apatit e fissio n trac k thermochronology. Tectonics, 14, 704-718. RUST, DJ. & SUMMERFIELD , M.A. 1990 . Isopac h an d borehole dat a a s indicator s o f rifte d margi n evolution i n southwester n Africa . Marine an d Petroleum Geology, 1, 277-287. SMITH, K. , GATLIFF , R.W . & SMITH , N.J.P . 1994 . Discussion o n the amount of Tertiary erosion in the UK usin g soni c velocit y analysis . Journal o f th e Geological Society, London, 151 , 1041-1044 . STEEL, R.J . 1978 . Triassi c rif t basin s o f Northwes t Scotland—their configuration , infilling and devel opment. In : FINSTAD , K.G . & SELLEY , R.C . (eds ) Mesozoic Northern North Sea Symposium. Norwegian Petroleu m Society , Stavanger , Paper 7 . STEWART, A.D . 1972 . Precambria n landscape s i n northwest Scotland . Geological Journal, 8 , 111-124. STOKER, M.S. , KITCHEN , K . & GRAHAM , C.C . 1993 . United Kingdom Offshore Regional Report: the Geology of the Hebrides and the West Shetland Shelves, an d Adjacent Deep Water Areas. HMSO , London. STOKER, M.S., LESLIE, A.B. , SCOTT, W.D. & 6 OTHERS 1994. A recor d o f lat e Cenozoi c stratigraphy , sedimentation an d climat e chang e fro m th e Hebrides Slope , N E Atlanti c Ocean . Journal o f the Geological Society, London, 151 , 235-249 . STUEVOLD, L.M . & ELDHOLM , O . 1996 . Cenozoi c uplift o f Fennoscandia inferre d from a study of the mid-Norwegian margin . Global an d Planetary Change, 12, 359-386. STURT, B.A. , DALLAND , A . & MITCHELL , J.L . 1979 . The ag e o f th e sub-Jurassi c tropica l weatherin g profiles o f And0ya , norther n Norway , an d th e implications fo r th e Lat e Palaeozoi c paleogeogra phy o f th e Nort h Atlanti c region . Geologische Rundschau, 68 , 523-542. SUMMERFIELD, M.A . & BROWN , R.W . 1998 . Geomorphic factor s i n th e interpretatio n o f fission-track data . In : VA N DEN HAUTE, P . & D E CORTE, F . (eds ) Advances i n Fission-Track Geochronology. Kluwer , Dordrecht, 269-284. SUN, S.-S. , MCDONOUGH , W. E & EWART , A . 1989 . Four componen t mode l fo r Australian basalts . In: JOHNSON, R.W . (ed. ) Intraplate Volcanism i n Eastern Australia an d Ne w Zealand. Cambridg e University Press , Cambridge , 333-347 .

SUTTER, A.A . 1980 . Palaeogene sediments from th e U.K. sector o f th e central North Sea. Ph D thesis , University o f Aberdeen. THOMAS, M.F . 1989 . Th e rol e o f etc h processe s in landfor m development . Zeitschrift fu r Geomorphologie, NF , 33 , 129-142 . THOMSON, K . & HILLIS , R.R . 1995 . Tertiar y structuration an d erosio n o f th e inne r Mora y Firth. In : SCRUTTON , R.A. , STOKER , M.S. , SHIMMIELD, G.B . & TUDHOPE , A.W . (eds ) Th e Tectonics, Sedimentation and Palaeooceanography o f th e North Atlantic Region. Geologica l Society, London , Specia l Publications , 90 , 249-269. THOMSON, K. , UNDERHILL , J.R., GREEN , P.P. , BRAY , R.J. & GIBSON , H.J . 1999 . Evidenc e fro m apatit e fission track analysi s fo r the post-Devonia n buria l and exhumation history o f the northern Highlands , Scotland. Marine an d Petroleum Geology, 16 , 27-39. THRELFALL, W.F . 1981 . Structura l framewor k o f th e central and norther n Nort h Sea. In: ILLING , L.V. & HOB SON, G . (eds ) Petroleum Geology o f th e Continental Shelf o f NW Europe. Heyden, London , 98-103. TIPPETT, J.M . & Hovius , N . 2000 . Geodynami c processes in the Souther n Alps. In: SUMMERFIELD, M.A. (ed. ) Geomorphology an d Global Tectonics. Wiley, Chichester, 109-134 . TWIDALE, C.R . 1994 . Gondwana n (Late Jurassi c an d Cretaceous) palaeosurface s o f th e Australia n craton. Palaeogeography, Palaeoclimatology, Palaeoecology, 112 , 157-186 . TWIDALE, C.R / & CAMPBELL , E.M . 1995 . Pre Quaternary landform s i n th e lo w latitud e context: the exampl e o f Australia . Geomorphology, 12 , 17-35. WATSON, J . 1985 . Scotlan d a s a n Atlantic-Nort h Se a divide. Journal o f th e Geological Society, London, 142,221-243. WELLMAN, P . 1986 . Intrusion s beneath larg e alkalin e intraplate volcanoes . Exploration Geophysics, 17 , 135-139. WELLMAN, P . & MCDOUGALL , I . 1974 . Cainozoi c ingenous activit y i n easter n Australia . Tectonophysics, 23, 49-65. WHITE, N . & LOVELL, B. 1997. Measuring the pulse of a plume with the sedimentary record. Nature, 387 , 888-891. WIGNALL, P.B . & PICKERING , K.T . 1993 . Palaeo ecology an d sedimentolog y acros s a Jurassic faul t scarp, N E Scotland . Journal o f th e Geological Society, London, 150 , 323-340 . YOUNG, R.W . 1981 . Denudationa l histor y o f th e south-central upland s o f Ne w Sout h Wales . Australian Geographer, 15 , 7 7 — 88. ZEITLER, P.K. , HERCZIG , A.L. , MCDOUGALL , I . & HONDA, M . 1987 . U-Th-H e datin g o f apatite : a potential thermochronometer . Geochimica e t Cosmochimica Acta, 51, 2865—2868. ZIEGLER, P.A. 1981. Evolution o f Sedimentary Basins of North-West Europe. Heyden, London.

Cenozoic evolution of the Faroe Platform: comparing denudation and deposition MORTEN SPARR E ANDERSEN 1, AAGE BAC H S0RENSEN 1, LARS OL E BOLDREEL 2 & TOVE NIELSEN 1 1

Geological Survey of Denmark and Greenland (GEUS), Thoravej 8, DK-2400 Copenhagen NV, Denmark (e-mail: msa@ geus.dk)

University of Copenhagen, Department of Geology, 0ster Voldgade 10, DK-1350, Copenhagen K, Denmark Abstract: Throughou t Paleocen e an d Eocen e tim e th e Faroe-Shetlan d Channe l an d th e eastern par t o f the Faroe Platform wa s a subsiding marin e basin . I n Early Paleocen e time , basin-floor fan s o f a British provenance wer e deposite d i n the eastern par t o f the basin. I n Late Paleocen e time , c . 6 km o f basalt entered th e basi n fro m th e wes t and north , an d th e basin wa s constricte d b y th e larg e volume s o f basal t tha t entere d th e basin , creatin g th e Faroe-Shetland Escarpment . I n Eocen e tim e subsidenc e continue d i n th e basina l areas . Again, sediments o f a dominantly easter n provenanc e were deposited. Throughou t Eocen e time, erosio n product s fro m th e Faroe Platform wer e possibly deposite d i n the Faroe Ban k Channel an d the Norwegian Sea Basin, but only to a limited degre e i n the Faroe-Shetland Channel. The oldest sediment s of documented western provenance on the eastern margin of the Faroe Platfor m are of Early Oligocen e age . During a compressional phas e commencing in Mid-Late Miocene time some basinal areas emerged and erosion took place on the top of emerged anticlines . However , denudation throughout Late Miocene an d Early Pliocene time was apparentl y rather limited compared wit h a Late Pliocene phas e of denudation. During this phase of denudation, a large progradational wedge was deposited on the eastern margin of th e Faro e Platform . O n th e basi s o f a structura l analysi s o f th e Faro e Platform , th e amount o f basal t remove d fro m i t durin g Cenozoic tim e i s estimated t o be c . 46 000 km3 (131 100 X 1012kg). Usin g 2900kgnT ^ a s th e densit y o f basal t an d 2300kgnT 3 a s sediment densit y th e estimate d amount o f remove d basal t i s i n fai r agreemen t wit h th e estimate o f th e volum e o f sediment s derive d fro m th e platfor m (c . 56000km 3, 114800 X 1012kg). The greates t depositio n rate s o n the eastern Faro e Platform an d in th e Faroe-Shetland Channe l apparentl y occurre d afte r tw o distinc t inversion o r compressio n events in Mid-Eocene an d Mid-Late Miocene time . However, uplift o f the Faroe Platfor m could hav e been force d b y denudation rather tha n endogenous processes .

In this paper w e present a preliminary attemp t to documente d Mesozoi c riftin g i n N W Europ e quantify th e denudationa l an d depositiona l (Fig . 2). history o f th e Faro e regio n durin g Cenozoi c Th e mos t importan t Cenozoi c structura l time (Fig . 1) . The stratigraphi c subdivisio n an d element s withi n th e Faro e regio n ar e show n i n discussion o f th e depositiona l histor y is mainl y Fig . 1 . Most of the structural elements are known based o n publishe d wor k (Nielse n & Va n fro m previou s work (e.g . Boldree l & Andersen Weering 1998 ; Anderse n e t al 2000 ; Stoke r 1994 ; Lamer s & Carmichae l 1999) . Th e et al. 2002). Foinave n Basi n i s th e Paleocene-Eocene depo centre in the southern part of the Faroe-Shetland Channel. Geographicall y i t i s equivalen t to th e „, , j • i • Foinave n Sub-basi n o f Lamer s & Carmichae l Structural elements an d regional setting (1999 ) Th e Fugloy ^^intr oduced here , is a In a broader context the Faroes ar e located close broa d Eocene-Oligocen e depocentr e i n th e to th e lin e o f th e Lat e Paleocen e break-u p norther n par t o f th e Faroe-Shetlan d Channel , between Europ e an d Greenland, centrall y i n the Uppe r Eocene-Oligocen e sediments , whic h ar e area affecte d b y th e Icelandi c mantl e plum e thi n o r absen t i n th e Foinave n basin , thicke n during Paleocen e time , an d o n th e margi n o f northwar d i n th e Fuglo y Basi n (Fig . 3) . From: DORE , A.G., CARTWRIGHT, J.A. , STOKER , M.S. , TURNER, J.P. & WHITE , N . 2002 . Exhumation o f th e North Atlantic Margin: Timing, Mechanisms an d Implications for Petroleum Exploration. Geologica l Society, London, Special Publications, 196, 291-311. 0305-8719/027$ 15.00 © The Geological Society of London 2002.

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Fig. 1 . Structural element s o f the stud y area. Interprete d an d inferre d fault s ar e show n schematically . Important transfer zone s ar e show n wit h shading . Fuglo y Basi n i s introduce d a s a ne w ter m representin g th e norther n Eocene-Oligocene depocentre i n the Faroe-Shetland Channel. I t overlies th e Paleocene Flet t Basin , th e Coron a Ridge and a poorly describe d depocentre west of Corona Ridge . Foinave n Basi n is a depocentre in the souther n part o f th e Faroe-Shetlan d Channe l throughou t Paleocene-Eocene time . Othe r element s hav e bee n describe d elsewhere (e.g . Andersen e t al. 2000) . FBK , Faroe Ban k Knoll ; DVC , Darwin volcani c centre ; WR , Westra y Ridge. Geographically, th e Fuglo y Basi n overlie s th e Paleocene Flet t Basi n (Kitchen & Ritchie 1987) , the Corona Ridg e (e.g . Rumph et al. 1993 ) and a relatively poorl y constraine d Paleocen e depo centre wes t o f th e Coron a Ridge . Th e Westra y Transfer Zone (e.g . Lamers & Carmichael 1999) , including th e anticlina l structur e th e Westra y Splay, separate s th e Foinave n Basi n fro m th e Flett an d Fugloy basin s an d the Corona Ridge .

During Paleocen e time , norma l faultin g occurred alon g approximatel y NNE-SSW- an d east-west-trending fault s producin g subsidence of the Foinaven and Flett basins (e.g. Dean et al. 1999; Lamer s & Carmichae l 1999) . T o som e extent, thi s ma y reflec t compactio n o f a thic k underlying successio n o f Cretaceou s sediment s (Ebdon e t al . 1995) . However , th e Paleocen e depocentres ar e offse t an d slightl y rotate d

CENOZOIC EVOLUTIO N OF THE FAROE PLATFORM

293

Fig. 2 . Th e N E Atlanti c regio n i n earlies t Eocen e tim e (magneti c chro n C23) . Th e approximat e exten t o f Paleocene pre-break-u p an d syn-break-u p basalt s i s shown . Th e approximat e locatio n o f th e centr e o f th e Icelandic plume and its extent is shown according to White & McKenzie (1989). FSC, Faroe-Shetland Channel.

Fig. 3 . Seismi c sectio n fro m Foinave n Basi n int o Fuglo y Basin . Th e seismi c characte r o f FPC- B change s dramatically a s th e sectio n crosse s th e Westra y Splay . Locatio n o f profil e i s show n in Fig . 7 . B y courtes y o f Veritas.

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relative t o th e Lat e Cretaceou s depocentre s (Dean e t al 1999) . I n addition , a strike-sli p component alon g som e east-wes t faul t plane s may possibl y b e indicate d b y abrup t chang e o f the thro w alon g strik e (e.g . Jud d Fault) . Thi s suggests a n importan t Paleocen e rif t even t controlled b y a stres s syste m rotate d relativ e t o the mid-Cretaceou s syste m (Anderse n 1999 ; Dean e t al . 1999) . Cenozoi c subsidenc e curve s calculated fo r well s i n th e Wes t Shetlan d are a could suppor t th e concept o f Paleocene riftin g i n the Faroe-Shetlan d area . However , Paleocene Recent subsidenc e i n basina l area s o f th e northern Nort h Se a an d Paleocen e denudatio n in larg e area s o f the Britis h Isle s ar e considere d arguments fo r th e existenc e o f a larg e mantl e plume, th e Icelan d plume , i n th e are a durin g Paleocene tim e (Hal l & Whit e 1994) . Th e Iceland plum e woul d hav e eve n stronge r influence i n th e Faroe-Shetlan d are a tha n i n the North Sea , but in the Faroe regio n separatio n of rift-induce d tectoni c movement s fro m plume induced tectoni c movement s ha s no t bee n possible. Following the Paleocene rif t episod e (o r late in the rift period) th e Faroe Platfor m was , within a short time span (59-55.5 Ma), covered b y a thick succession o f basalt s (stratigraphi c thicknes s exceeds 5km : Hal d & Waagstei n 1984) . Sinc e Eocene time , basina l area s o f th e Faro e regio n subsided abou t 2500 m (e.g . Boldree l & Andersen 1993 ; Ritchi e e t a l 1999) . Durin g

this perio d bot h th e Shetlan d regio n an d th e Faroe Platfor m supplie d sediment s t o basina l areas (e.g . Andersen e t al. 2000) . Stratigraphic correlation Work toward s a unifie d mid-lat e Cenozoi c stratigraphy acros s th e Faroe-Shetland Channe l is currently in progress unde r the auspice s of the EU-funded STRATAGE M projec t (Evan s 2000) . This wor k shoul d provid e a significantl y improved stratigraphi c mode l fo r th e Faro e Platform an d it s margins . However , thi s strati graphic mode l i s no t finished , an d th e strati graphic subdivision used in this paper (Table 1 ) is adapted fro m Boldree l & Anderse n (1995) , Nielsen & Va n Weerin g (1998 ) an d Anderse n et al . (2000) . Revise d ag e assignment s an d correlation o f th e Neogen e sectio n acros s th e Faroe-Shetland Channe l are primarily based o n Cloke et al. (2000) an d Stoker et al (2002) . The stratigraphic correlation s i n Tabl e 1 ar e con sidered a n improvemen t relativ e t o thos e presented b y Anderse n e t al . (2000) . However , the chronostratigraphi c correlation s ar e stil l mostly tentative . Th e followin g change s hav e been introduce d relativ e t o Anderse n e t al . (2000). A progradationa l wedg e i n uni t FPC-D. 2 o n the Faro e Platform' s easter n margi n (Fig . 4 ) (Andersen e t al . 2000 ) i s simila r t o a Lat e Pliocene-early Pleistocen e progradationa l

Table 1 . Seismic units mapped o n th e Faroe Platform an d it s margins Bounded units on Faroe Platform Eastern margin

Northern margin

FPC-D.3

Unit 9

| Lithostratigraphy West Shetland BGS N3

mid-Pleistocene FPC-D.2

Seismic horizons Faroes East and South

CN 050

GU

CN O'lO

INU

Sequence 3

early Pliocene FPC-D.1

Sequence 2

FPC-C

Sequencel

PM n°n

Mid-Miocene/Late Miocene latest Oligocene/early Miocene

CN 010 FPC-B.4

Mid-Eocene/Late Eocene

[UrTitA

Nordland Group

Approximate age

N2

N1

LOEMU Westray Group

CP-100

FPC-B.3 Early Eocene/Mid-Eocene FPC-B.2

Sequence 0

intra-Ypresian

CP-060

Stronsay Group

CP-030 FPC-B.1

base Eocene

CP-010 FPC-A

Moray Group

Basalt basement

The seismic subdivisio n used on the eastern margi n o f the Platform (FPC-A-FPC-D) has also been recognized o n the S W margi n o f th e platfor m an d i n th e Faro e Ban k Basin . Th e correlatio n wit h th e Neogen e seismi c stratigraphy o n the West Shetlan d margi n (Stoke r 1999 ; Stoke r et al. 2002) is discussed in the text. The table also shows a tentative correlation wit h the seismic stratigraph y on the northern margi n o f the platform (Nielse n & Van Weering 1998) . Furthe r discussio n o f this correlatio n ha s bee n give n by Anderse n e t al. (2000) .

CENOZOIC EVOLUTION OF THE FAROE PLATFORM

295

Fig. 4. Seismic section across the NW part of the Foinaven Basin and the SE margin of the Faroe Platform. Th e distal par t o f FPC-B. l continue s fro m th e basi n ont o the margin . Location of profil e i s show n i n Fig . 7 . B y courtesy o f Western Geophysical.

wedge between the 'Glacial Unconformity' (GU ) and the 'Intr a Neogene Unconformity ' (INU ) on the Wes t Shetlan d an d Hebride s margi n (Stoke r 1999) an d to a Late Pliocene-Pleistocene wedge off Mid-Norwa y (Henrikse n & Vorre n 1996) . Therefore th e bas e o f FPC-D. 2 i s tentativel y correlated wit h the INU (Stoker et al 2002) , and the bas e o f FPC-D.3 is considered equivalen t to the G U (Stoker et al 2002) .

Recent biostratigraphi c wor k i n exploratio n wells offshor e West Shetlan d ha s indicate d tha t Neogene basi n inversio n an d contractio n i n th e UK secto r o f th e Faroe-Shetlan d Channe l commenced approximatel y a t th e boundar y between Mid - an d Lat e Miocen e tim e (Clok e et al 1999 , 2000 ; Cloke, pers. comm.). We have adopted thi s ag e fo r th e bas e o f FPC-D.l . However, distinc t along-strik e variation s i n th e

Fig. 5 . Seismi c section across the Fugloy Basin. Onla p from wes t ont o FPC-B.4 should b e noted. Location of profile i s show n i n Fig. 7. By courtesy o f Western Geophysical.

296

M.S. ANDERSE N ETAL.

timing o f Neogen e inversio n an d contractio n i n the Faroe-Shetlan d Channe l wer e recorde d b y Cloke et al (2000) . The bas e o f FPC- C i s tentativel y correlate d with th e 'Lat e Oligocene-Earl y Miocen e Unconformity' (LOEMU ) (Stoke r 1999 ; Stoke r et al . 2002) , an d th e base o f FPC- C i s thu s reassigned a Lat e Oligocene-Earl y Miocen e age. Th e correlatio n betwee n th e base o f FPC- C and the LOEM U is supporte d by tie s to well s 214/28-1 an d 204/28- 1 o n th e Wes t Shetlan d margin an d correlatio n wit h seismi c section s presented b y Stoke r (1999) . The Cenozoi c sequenc e i s thu s divide d int o four mai n unit s separate d b y thre e distinc t seismic marke r horizon s (Anderse n e t al. 2000): (1) FPC-A , which on the Faroe margi n an d in the Faro e Ban k Basin , consist s o f Paleocen e plateau basalt s (Fig . 4) . Furthe r eas t i n th e Foinaven an d Fuglo y basin s th e T36-T4 5 sequences (Ebdo n e t al . 1995 ) ar e considere d

time equivalent s o f FPC-A . O n th e norther n margin of the Faroe Platfor m FPC-A consists of a wedge o f seaward-dippin g basalt . A n oute r high characterized b y chaoti c reflection s i s foun d below the lower slope, basin wards from seaward dipping basalts . Accordin g t o th e emplacemen t model o f Plank e e t al . (2000) , th e oute r high , located seaward s of the seaward-dipping basalts , may represen t th e submergenc e o f th e volcani c centre afte r th e break-u p betwee n Faroe s an d Greenland. Furthe r nort h FPC- A consist s o f typical oceani c basemen t forme d fro m sub aqueously erupte d basalt. (2) FPC- B consist s o f a thic k successio n o f Eocene-Oligocene sediments , whic h i n th e Faroe-Shetland Channe l ar e mostl y o f easter n provenance. I n th e Foinave n Basi n FPC- B i s subdivided int o fou r subunit s FPC-B . 1-FPCB.4, al l characterize d b y fairl y goo d reflectio n continuity (Fig . 4) . Thi s subdivisio n i s no t recognized o n seismi c dat a fro m th e Fuglo y

Fig. 6 . Correlation char t o f Maastrichtian-Paleocene sedimentologica l an d volcanic event s in the N E Atlantic region. Fou r possible regiona l event s ar e indicated. Modifie d fro m Whit e & Lovell (1997) , Da m e t al. (1998) , Storey e t al. (1998 ) an d Larsen e t al. (1999) .

CENOZOIC EVOLUTION O F THE FAROE PLATFOR M

Basin, wher e FPC- B generall y i s characterize d by poor reflectio n continuity . However, the shif t from sediment s o f dominantl y easter n prove nance t o sediment s o f mixe d o r dominantl y western provenanc e i s marke d i n th e Fuglo y Basin by a seismic horizon, which approximately correlates wit h the base o f FPC-B.4 (Fig s 3 and 5). I n th e Foinave n Basin , th e thre e lowe r (Eocene) subunit s o f FPC- B ar e o f easter n

297

provenance but extend westward on to the Faro e Platform (Fig . 4) . I n Lat e Eocene-Oligocen e time sediments of a western provenance becom e abundant, an d i n th e uppe r par t o f FPC-B. 4 sediments of a western provenance are dominant. In th e Faro e Ban k Basi n FPC- B consist s o f a seismically homogeneou s successio n o f sedi ments originatin g mainl y fro m th e Faro e Platform. Sequenc e 0 (Nielse n & Va n Weering

Fig. 7 . Distributio n o f Paleocen e basal t i n th e Faro e region . Contour s o f th e dept h t o th e basal t ar e base d o n seismic interpretation . Contou r interva l 50 0 m. The approximat e exten t of erosional truncatio n o f the Paleocene basalt i s also shown .

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M.S. ANDERSEN ETAL.

1998), a thick succession of sediments onlapping the basalti c basemen t o n the norther n margi n o f the Faro e Platform , ma y b e equivalen t t o al l o r part of unit FPC-B. (3) Th e Lower-Middle Miocene uni t FPC- C is foun d o n th e easter n margi n o f th e Faro e Platform an d i n th e Fuglo y Basin , wher e i t thickens northwards . I n the Faroe Ban k Channe l FPC-C is fairly thin (<200m). The up to 500m thick Sequenc e 1 (Nielsen & Van Weering 1998 ) on th e norther n margi n o f th e platfor m ma y b e the equivalen t o f FPC-C (Andersen e t al 2000) . (4) FPC- D (of Late Miocene-Recen t age), in the Faroe-Shetlan d Channe l an d Faro e Ban k Channel, i s mostl y characterize d b y goo d reflection continuit y (Fig . 5) . Th e thicknes s i s variable, probabl y reflectin g comple x inter action o f down-slop e an d along-slop e sedimen t

Fig. 8. Log signatures of four basalt layers in well 2057 9-1. Th e constan t lo w value s o f th e gamma-ra y lo g (GR) through each of the fou r basal t layers shoul d be noted. The top of all four basalt layers is characterized by a porous 'scori a zone ' wit h downwar d decreasin g neutron porosit y (NPHI ) an d increasin g velocit y (decreasing D T (delt a time)). Th e bas e o f th e thre e upper layer s i s shar p o n thre e logs . The lowe r layer , Bl, is atypical for subaerially emplaced basalts, as the lower halves of the logs are a mirror image of the upper half. Thi s may indicat e rapid coolin g from belo w a s well a s fro m abov e durin g an d immediatel y afte r emplacement. Emplacement in water or wet sediments (a bulldozin g intrusion , Plank e e t a l 2000 ) coul d explain the log signatur e o f layer B4.

transport i n the deeper part o f the basins. FPC- D is divided int o thre e subunits : FPC-D. 1 (of Late Miocene-Early Pliocene age), FPC-D.2 (of Late Pliocene-early Pleistocen e age ) an d FPC-D. 3 (of late Pleistocene-Recen t age) .

Cenozoic tectonics, denudation and deposition in the Faroe region Paleocene time Several generation s of slope an d basin-floor fans were deposite d i n th e Faroe-Shetlan d Channe l (Ebdon e t al. 1995 ; Lamers & Carmichael 1999 ) (Fig. 6) . O n th e basi s o f wel l dat a an d seismi c interpretation i t appears tha t most if not all of the reasonably well-define d fa n deposit s i n th e eastern an d centra l par t o f th e Faroe-Shetlan d Channel wer e derive d fro m sedimen t source s i n the area aroun d th e Shetland an d Orkney Island s or furthe r sout h (Ebdo n e t al . 1995 ; Lamer s & Carmichael 1999) . Furthe r wes t i n th e channe l the basal t cove r prohibit s sufficien t seismi c imaging qualit y t o identif y fa n systems. N o direct evidenc e o f Earl y Paleocen e vertica l movements o n th e Faro e Platfor m i s availabl e (e.g. Ki0rbo e 1999 ; Naylo r et al. 1999) . In Lat e Paleocen e tim e a thic k successio n o f basalts wa s emplace d i n the Faro e are a (Fig . 7) . On th e Faro e Island s th e expose d remnan t o f this successio n i s represente d b y a c . 3000 m thick sequenc e o f parallel-bedded platea u basalt (Rasmussen & Noe-Nygaar d 1970) . A n additional c . 2000 m o f platea u basal t wa s penetrated i n Lopra-1 without reaching th e bas e of th e Uppe r Paleocen e basal t successio n (Hal d & Waagstei n 1984) . Parallel-bedde d platea u basalt ha s bee n recognize d o n seismi c reflectio n profiles a s a uni t o f parallel-bedde d reflector s below th e top of the basalt throughou t the Faro e Platform, th e Faro e Ban k Channe l an d th e western par t o f th e Faroe-Shetlan d Channe l (Andersen 1988 ; Boldreel & Andersen 1993) . On seismic section s a strong reflection, representing the top of the plateau basalt , ca n be traced a s far east a s wel l 205/9 1 i n th e U K secto r o f th e Faroe-Shetland Channe l (e.g . Naylo r e t al . 1999; Ritchie et al. 1999) . In this well the plateau basalts ar e represente d b y fou r distinc t basal t layers intercalate d wit h sediment s containin g terrestrial t o margina l marin e palynoflora s (e.g . Naylor e t al . 1999) . Thre e o f th e fou r basal t layers hav e lo g signature s comparabl e wit h th e log signatur e o f a typica l subaeriall y emplace d basalt flow from the Lopra-1 well (Fig. 8). North and west of the Faroe Platform a wedge of basalts characterized b y seaward-dippin g reflector s i s found (Smyth e e t al. 1983) . I n the norther n par t

CENOZOIC EVOLUTION O F THE FAROE PLATFORM

of the Faroe-Shetlan d Channel , a uni t of prograding reflector s i s interpreted a s the latera l continuation of the parallel-bedded platea u basalt (Gatliff etal 1984 ; Andersen 1988 ; Ritchie etal 1999). Th e eastwar d terminatio n o f th e progra dational uni t i s define d b y th e Faroe-Shetlan d Escarpment, an d i t ha s bee n suggeste d tha t th e break of the escarpment formed a Late Paleocen e shoreline (e.g . Smyth e 1983 ; Boldree l & Andersen 1994 ; Ritchi e e t al 1999) . On th e basi s o f observe d latera l thicknes s changes, w e estimate tha t the surfac e di p during emplacement o f th e platea u basal t wa s c . 0.2°. This value is considered reasonabl e a s the largest depositional dip s i n th e platea u basalts , c . 0.5°, have bee n observe d o n th e flank s o f individua l small shiel d volcanoe s (o f the scutulu m type) in the uppe r basal t serie s (Noe-Nygaar d 1968) . Volcanic feeder system s alon g the line of break up wer e emphasize d i n a recen t emplacemen t model fo r volcani c rifte d margin s (Plank e e t al . 2000). Wit h 0.2 ° monotonous di p awa y fro m feeder system s alon g th e lin e o f break-up , th e Faroe Platfor m woul d culminate c. 900 m abov e sea leve l clos e t o th e lin e o f break-u p (Fig . 9). Observations of feeder system s alon g the strait s between th e Faro e Island s ma y indicat e tha t the feede r system s wer e sprea d ou t ove r mos t of th e Faro e Platfor m (Rasmusse n & NoeNygaard 1970) . The Late Paleocen e topograph y of th e platfor m coul d thu s hav e bee n lower , culminating c. 500m above se a level (Fig. 9) .

299

Eocene-Recent denudation of the Paleocene plateau basalts Since the basalt was emplaced i n Late Paleocen e time considerabl e erosio n ha s take n place , an d throughout most of the Faroe Platfor m the top of the basal t i s a n erosiona l unconformit y (Fig . 7 ) (Rasmussen & Noe-Nygaar d 1970 ; Boldreel & Andersen 1993 ; Andersen e t al 2000) . T o estimate th e amoun t o f erosion , w e hav e reconstructed th e missin g section , takin g int o consideration tha t the basalts are parallel bedde d in mos t o f the stud y are a (Fig . 7) . The principl e of th e reconstructio n an d th e resultin g contou r map showing the top of the missing basalt section are show n i n Fig. 10. To derive a reasonable fi t between onshore and offshore geology, we had to assume tha t th e middl e an d uppe r basal t serie s thinned fro m mor e tha n 200 0 m on the Northern Islands t o abou t 1000 m o n Suduroy . Thi s i s i n agreement wit h detailed observation s i n th e upper an d middl e serie s (e.g . Waagstein 1988) . This reconstructio n als o implie s tha t o n th e northern par t o f th e Faro e Platfor m c . 70 0 m of basalt wa s emplace d abov e th e youngest known basalt o f the upper series . A detailed correlatio n between th e Faroes basal t serie s an d the Nansen Fjord an d Miln e Lan d formation s o n Eas t Greenland ha s bee n establishe d (Larse n e t al . 1999). A few dykes are found o n the Faroes wit h compositions simila r t o the basalts of the Geikie Plateau Formation , whic h overlie s th e Miln e

Fig. 9 . Tw o alternative surfac e topographie s o f th e Faro e basal t platea u i n Lat e Paleocen e time . NW F i s th e topography expected if the volcanic feeders wer e located in a narrow zone around the line of break-up. DF is the topography expected in the case of distributed feeder systems . (Se e text fo r further discussion. ) The present-day surface topography , Eocene-Recent sediments in the Faroe-Shetland Channel and the estimated missing section of basalt ar e show n for comparison . Location of profile i s show n i n Fig. 7.

300

M.S. ANDERSE N ETAL.

Fig. 10 . Reconstruction o f the structur e of the to p of the Faroes Basalt Plateau. The to p of the missing section is contoured wit h 500 m interval . The principle s o f th e reconstructio n ar e show n i n th e lowe r par t o f th e figure . Between A and E the top of the basalt i s an erosional unconformity . The to p of the missing section is drawn as a smooth curv e joining the top of the basalt a t A and E. The curve is constructed so that its tangents are parallel to bedding plane s a t points B , C an d D . The remove d par t o f the seaward-dippin g reflectors i s reconstructed by a smooth curv e joining th e las t poin t withi n the parallel-bedde d basalts , D, t o th e poin t o f truncation , E. A t an y location alon g this part of the curve, the dip of the curve is less than the dip of apparent bedding planes within the seaward-dipping reflectors .

CENOZOIC EVOLUTIO N O F THE FAROE PLATFOR M

301

Fig. 11. Estimated denudatio n on the Cenozoic Faroe Platform. Contour interval 500 m. The map was constructed by subtracting the present top of the basalt (Fig. 7) from th e reconstructed structure of the top of the basalt plateau shown as Fig. 10 .

Land Formation i n East Greenland (Larsen et al 1999). I t i s thu s possibl e tha t th e 700 m thic k missing sectio n i s equivalen t t o th e c . 1500 m thick Geiki e Platea u Formatio n i n Eas t Green land. An estimate of Eocene-Recent denudation of th e Faroe Platfor m is obtained by subtracting the actual to p of the basalt show n in Fig. 7 fro m the contou r map i n Fig. 10 . A map o f estimated Eocene-Recent denudatio n i s presente d i n Fig. 11 . On the basis of this map, c . 46000km 3 of basal t ha s bee n remove d fro m th e Faro e Platform sinc e earlies t Eocen e time . Studies of zeolite zonation in the Lopra-1 and Vestmanna-1 well s provid e independen t esti mates o f th e magnitud e o f denudatio n a t thes e two location s (J0rgense n 1984) . I n Lopra- 1 th e denudation estimat e fro m zeolit e zonatio n obtained b y J0rgense n (1984 ) i s comparabl e

with ou r estimat e base d o n simpl e structura l reconstruction. However , i n th e 650 m dee p Vestmanna-1 wel l J0rgense n (1984 ) foun d a complex crystallizatio n sequenc e characterize d by fossi l therma l gradient s i n th e rang e 67-100°C km" 1. Extrapolatio n o f thes e gradi ents t o th e presen t mea n temperatur e o f 7° C indicates tha t th e origina l to p o f th e basalt s should b e foun d c . 600-800 m abov e se a level. This implies that the surface of the basalt plateau should coincid e approximatel y wit h th e C horizon, betwee n th e middl e an d uppe r basalt . However, durin g th e regiona l mappin g o f th e Faroes neithe r a n angula r unconformit y a t th e base o f th e uppe r basal t serie s no r westwar d thinning of the upper basalt serie s wa s observe d (Rasmussen & Noe-Nygaar d 1970) . A re evaluation o f th e zeolit e mineralizatio n i n th e

302

M.S. ANDERSE N E T AL.

Fig. 12 . (a) Detai l of Fig. 1 1 showing denudatio n i n the central par t o f the Faroe Platform. Locatio n o f sample s taken for measurement of vitrinite reflectance and zeolite fission-trac k age s from Kou l et al. (1983) are shown on the map. (b) Measurements of vitrinite reflectance values of surface samples from th e Faroes, (c) Zeolite fissiontrack age s ar e plotte d agains t extrusio n ag e o f th e stratigraphi c leve l (M . Store y cite d b y Larse n e t al . 1999) . (d) Vitrinit e reflectance values plotted agains t estimated depth of denudation at the sampl e localities.

Vestmanna-1 well is in progress (J0rgensen, pers. comm.). Vitrinit e reflectanc e measurement s o f coal sample s fro m 1 4 localitie s o n th e Faroes , including cutting s fro m 1364 m dept h i n th e Lopra-1 well , yiel d value s o f R 0 i n th e rang e from 0.3 6 t o 0.48% (Fig . 12) . However, w e have not bee n abl e t o fin d systemati c variation of th e reflectance values that could be used to constrain the estimate o f denudation on the Faroe Platform. Few dat a ar e availabl e tha t giv e direc t indication o f th e timin g o f denudatio n o n th e

Faroe Platform . Becaus e o f th e absenc e o f apatites o f a suitabl e size , n o successfu l apatit e fission-track stud y o f the platform has eve r been completed. Th e onl y publishe d fission-trac k study o f sample s fro m th e Faroe s i s a reconnaissance stud y o f fissio n track s i n zeolite mineral s (Kou l e t al . 1983) . Al l zeolit e fission-track age s fro m th e Faroe s ar e younge r than curren t estimate s o f th e age s o f th e Faroe s basalts (Fig . 12) . There i s a considerable spread of th e fission-trac k age s withi n bot h th e middle

CENOZOIC EVOLUTION O F THE FAROE PLATFOR M

and uppe r basal t series . Th e annealin g charac teristics o f zeolite s ar e poorl y constraine d an d any conclusion s base d o n zeolit e fissio n track s are thu s speculativ e (H . Andriessen , pers . comm.). However , there i s n o apparen t conflic t between th e zeolit e fission-trac k age s obtaine d by Koul et al (1983 ) and the conclusions i n this and following sections. It appear s tha t th e zeolit e assemblag e i n Lopra-1 is the only independent estimat e o f th e thickness o f remove d basal t o n th e Faro e Platform, which can be used as a firm constraint on the map of estimated denudation of the Faroe Platform show n i n Fig . 11 . Th e calculate d volume o f remove d basal t (46000km 3) shoul d thus be considere d a first estimate tha t ma y be refined whe n additiona l constraint s becom e available. Eocene deposition Sediment transpor t from th e sout h and east int o the Foinaven Basin continued throughout Eocene time, and sediments were deposited as units FPCB.1-FPC-B.3. Durin g this perio d n o sediment s of an y significanc e entered th e Faroe-Shetlan d Channel fro m th e wes t (e.g . Anderse n e t al .

303

2000). Also , farthe r nort h i n th e Fuglo y Basin , sediments o f eastern provenanc e wer e dominan t throughout Eocene time. However, the thickness of the Eocene sequence is considerably less than in the Foinaven Basin and the three lower units of FPC-B ar e no t identifie d a s separat e unit s o n seismic data . Th e Westra y Transfe r Zon e thu s may have been a significant tectonic boundary in Eocene time (Fig . 1) . In the West Shetland area Early Eocene tuff , whic h presumably originated in th e now-subside d volcani c zon e aroun d th e line o f break-up , wa s intersperse d wit h th e mostly paralic sediments of the Balder Formation (e.g. Kno x et al. 1997). Equivalents of the Balder tuff ar e als o foun d o n th e Faro e Platfor m (Waagstein & Heilmann-Clausen 1995) . O n the Faroe Platfor m Ypresian-Lutetia n limeston e follows th e 'Balde r tuf f (Waagstei n & Heil mann-Clausen 1995) . Thi s limeston e contain s tuffaceous fragment s tha t appea r t o hav e bee n laid dow n directl y i n th e limestone . However , transported clasti c fragments , whic h woul d indicate denudatio n an d eastwar d drainag e o f the Faroe Platform, are not seen. It was therefore concluded tha t th e Faro e Platfor m ha d a fairl y low topography throughout most of Eocene tim e and that northward or westward drainage from a

Fig. 13 . Seismic sectio n throug h basin-floo r moun d i n souther n Fuglo y Basin . Locatio n o f profile i s show n i n Fig. 7 . By courtesy o f Western Geophysical .

304

M.S. ANDERSEN ETAL.

significant par t o f th e platea u wa s possibl y established i n Earl y Eocen e tim e an d prevaile d without significant modifications into Oligocen e time (Anderse n e t al 2000) . Oligocene time No Upper Eocene sediment s have been encountered o n th e easter n margi n o f th e Faro e Platform. Lowe r Oligocen e volcaniclasti c sandstones ar e th e oldes t know n sediment s originating fro m denudatio n o f th e platfor m t o the west (Waagstein & Heilmann-Clausen 1995) . Andersen et al. (2000) suggested that the Lower Oligocene volcaniclasti c sandstone s correlat e with seismi c uni t FPC-B. 4 o n th e oute r par t o f the margin an d in the Faroe-Shetland Channel. Abundant reworked dinoflagellate s o f Early an d Mid-Eocene ag e ar e foun d i n mos t o f th e investigated samples , an d reworke d dinoflagellates of Mid-Late Eocene ag e were found in two samples (Waagstei n & Heilmann-Clausen 1995) . Sediments o f Mid-Lat e Eocen e ag e ma y thu s have been present on the platform (Waagstein & Heilmann-Clausen 1995) . I t thu s appear s tha t uplift o f the easter n Faro e Platfor m occurre d i n Late Eocene-Earl y Oligocen e tim e afte r a period o f burial i n Earl y an d Mid-Eocene time . On th e basi s o f th e seismi c interpretation , inversion o f th e Foinave n Basi n commence d i n Mid-Eocene time. Therefore, the inversion of the Foinaven Basi n apparentl y starte d befor e th e Late Eocene-Earl y Oligocen e denudatio n (an d uplift) o f large part s o f the Faroe Platform . A mounded basin-floor body, low in unit FPCB.4, i s see n o n seismi c dat a fro m th e south eastern part of Fugloy Basin (Fig. 13) . This body is characterize d b y continuou s high-amplitud e reflections o n it s wes t flank . Th e centra l an d thickest par t o f th e bod y i s characterize d b y a fairly chaoti c reflectio n configuration . O n th e eastern flan k a progradationa l syste m o n th e lower West Shetland slope merges into or onlaps the basin-floor body (Fig . 13) . W e interpret thi s body as a basin-floor fan system, which was fed by sediments from the south and east. The central part ma y represen t along-slop e transpor t chan nels wherea s th e continuou s high-amplitud e reflections o n th e flank s ma y represen t mor e widespread 'overban k deposits' . Th e tectoni c event that caused inversion o f the Foinaven Basi n and movement s alon g Westra y Splay ma y hav e caused uplif t i n th e area s tha t supplie d th e sediments fo r th e basi n floor. Further nort h and west and above the basin-floor body, FPC-B.4 i s characterized b y a successio n o f parallel reflections cu t b y numerou s smal l fault s tha t mostl y terminate downwar d withi n uni t FPC-B.4 .

Upwards, th e fault s mostl y terminat e i n th e upper part o f FPC-B o r in FPC-C. Between th e faults, th e continuit y o f th e reflection s i s high , whereas regiona l continuit y i s relativel y poo r (Fig. 5) . Westwards, the densely faulte d paralle l reflectors terminat e b y downla p ont o th e to p of FPC-B.4 o r ont o reflector s in th e lowe r par t of FPC-B.4. W e sugges t tha t th e densel y faulte d parallel reflectors represent slop e and basin-floor sediments tha t were transported westward s fro m the Wes t Shetlan d area . Thes e sediment s ar e characterized b y syndepositional an d early postdepositional deformation . Thes e fault s ar e arranged i n a polygona l patter n suggestin g volumetric contractio n durin g buria l (Davie s et al . 1999 ) o r gravitationa l instabilit y o f th e thick succession o f sediments . The Faroe Ban k Knoll (Fig . 1 ) was emplace d as a centra l magmati c comple x presumabl y during early Eocene time. A hyaloclastic foreset breccia i s interprete d fro m seismi c data . Th e breccia terminate s i n a volcani c escarpmen t o n the eastern flank, suggesting that the Faroe Bank Knoll wa s emplaced i n a shallow sea . Through out mos t of Eocene time , th e Faro e Ban k Knoll was a barrie r fo r sediment s o f easter n prove nance. In the Munken Basin, eas t an d NE of the Faroe Ban k Knoll , th e Eocene-Oligocen e evolution wa s almos t identica l t o th e evolutio n in th e Foinave n Basin . Seismi c interpretatio n indicates that the Eocene an d Oligocene deposit s in the Faroe Ban k Basin ar e up to 1800 m thick. Most o f th e sediment s ar e presumabl y derive d from th e Faroe Platform . On th e norther n margi n th e equivalen t o f FPC-B i s a transgressiv e uni t (Nielse n & Van Weering 1998) , an d i s represente d b y marin e clay- and siltstones a t Deep Se a Drilling Projec t (DSDP) Sit e 336 (Talwani et al. 1976). Andersen et al. (2000) have suggested that deposition coul d have already started in Eocene time. However, no well tie s ar e available tha t constrain th e star t of deposition o n th e norther n margi n o f th e Faro e Platform. Early and Mid-Miocene time Lower and Middle Miocene sediment s are foun d on the shelf an d margin of the Faroe Platfor m as seismic uni t FPC-C . Thes e deposit s wer e previously considere d olde r (Uppe r Oligocene Lower Miocene : Anderse n e t al . 2000) . Th e revised ag e rang e o f thi s uni t i s base d o n th e tentative correlatio n o f th e bas e o f uni t FPC- C with th e Lat e Oligocene-Earl y Miocen e unconformity offshor e Wes t Shetlan d (Stoke r et al . 2002) . O n th e basi s o f th e geometr y and t o a lesse r degre e th e interna l reflectio n

CENOZOIC EVOLUTION OF THE FAROE PLATFORM

configuration, we suggest that unit FPC-C on the eastern margi n o f th e Faro e Platfor m wa s fe d almost exclusively by sediments derived from the platform. I n th e basina l settin g sediment s o f pelagic an d hemipelagi c origi n ma y hav e bee n dominant. On the northern margin of the Faroe Platform , down-slope an d eas t o f DSD P Sit e 336 , the equivalent of FPC-C wa s interpreted as a type 2 sequence b y Nielse n & Va n Weerin g (1998) . The uni t reflect s depositio n afte r a significan t drop i n relativ e se a leve l followin g continuou s subsidence throughou t Eocene-Oligocen e times. Late Miocene-Early Pliocene time Deposition o f FPC- C wa s followe d b y mil d folding (Anderse n e t al 2000) . Recen t biostratigraphic studie s i n exploratio n well s i n th e British secto r o f th e Faroe-Shetlan d Channe l show tha t deformatio n occurre d clos e t o th e Mid-Late Miocen e boundar y (Clok e e t al . 2000). Thi s tectoni c phas e wa s followe d b y deposition o f uni t FPC-D. l i n Lat e Miocene Early Pliocen e tim e (Fig . 14). On th e easter n margin of the Faroe Platform , sediment s o f unit FPC-D.l wer e deposite d i n smal l synclina l basins. Th e sediment s wer e apparentl y derive d

305

from th e anticlinal crests on the Faroe Platform , some o f whic h sho w indicatio n o f erosio n (Andersen e t al . 2000) . Along-slop e botto m currents apparentl y influence d Uni t FPC-D. l along th e slop e an d i n th e Faroe-Shetlan d Channel. Thi s i s especiall y eviden t alon g th e northern par t o f th e slop e eas t o f th e Faro e Platform. Late Pliocene time A larg e progradationa l wedge , comprising 1900km3 of undifferentiated sediment s originat ing on the Faroe Platform, was deposited eas t of the Faroe s durin g Pliocen e an d possibl y earl y Pleistocene tim e as unit FPC-D2 (Fig s 4 and 14) (Andersen e t al . 2000) . A simila r bu t muc h smaller wedg e is found S W of the Faroes i n the Skeivi Bank area. The total volume of sediments in th e Skeiv i Ban k Wedg e i s unknown . B y analogy wit h simila r wedge s offshor e Norwa y and th e UK , Anderse n e t al . (2000 ) hav e suggested tha t th e wedge s eas t an d S W of th e Faroes wer e mostl y deposite d i n Lat e Pliocen e time. A second depocentre within unit FPC-D. 2 is foun d i n th e norther n par t o f th e Faroe Shetland Channel . Thi s depocentr e partiall y coincides wit h the depocentr e o f the underlying FPC-D. 1. In the eastern depocentre the dominant

Fig. 14 . Seismic section from the eastern margin o f the Faroe Platform. Th e change of reflection characte r fro m FPC-B t o FPC-C an d erosional unconformit y o n top of FPC-C shoul d be noted. Locatio n o f profile is shown in Fig. 7 . By courtesy o f Western Geophysical .

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M.S. ANDERSE N ETAL.

trough-mouth settin g associate d wit h th e transverse valley s mentione d abov e (Kuijper s 2000).

seismic characteristic s ar e subparallel , continu ous and often drapin g reflectors (Fig. 5) . Onlaps onto the underlying unit are common. Therefore , we sugges t tha t th e sediment s i n th e easter n depocentre ar e intercalation s o f pelagi c o r hemipelagic sediment s an d turbidites . I n th e southern par t o f th e easter n depocentr e down slope sedimentation , representin g a basinwar d continuation o f th e progradationa l wedge , appears to be dominant.

Faroe Platform mass balance We have made a preliminary calculatio n o f th e volume o f sediment s originatin g fro m denuda tion o f th e Faro e Platfor m (Tabl e 2) . Th e volumetric calculation s ar e base d o n seismi c interpretations. Th e provenanc e o f sedimen t bodies wa s determine d base d o n th e geometr y and interna l reflecto r configuration . Thi s approach is rather crude, as sediment originatin g outside th e Faro e Platfor m coul d b e hidde n within th e units , which, on th e basi s o f seismi c expression, ar e interpreted a s originating o n th e platform. Fo r instance , pelagi c sediment s an d hemipelagic sediment s derive d fro m othe r denudation area s ar e probabl y include d i n th e calculations. Sediment s derive d fro m th e Faro e Platform ma y also have been carried outsid e the three depositiona l centre s considered . Seismi c units that appea r t o be derived fro m on e sid e of the Faroe-Shetlan d Channe l ma y contai n veneers derive d fro m th e opposit e side . I n addition t o th e error s mentione d above , th e magnitudes o f whic h ar e currentl y unknown , we estimat e 10-15 % erro r o n th e volume s

Pleistocene time The Faroe s wer e glaciate d durin g Pleistocen e time. Stud y o f th e glacia l landform s o n th e islands indicate s tha t during th e las t glaciatio n the ice cover was limited t o a small area around the islands, possibl y a s three smal l separat e ic e caps (Humlu m & Christiansen 1998) . Al l of the Faroe Platfor m experience d erosion . Transvers e valleys were cut into the sediments on the outer shelf, marginal valleys that follow the limit of the basalt outcro p aroun d th e Faroe s wer e cu t int o Pliocene an d olde r sediments , an d (late ) Pleistocene sediment s wer e deposite d i n thes e depositional lows . I n lat e Quaternar y tim e mass-flows occurre d o n th e norther n an d eastern margin s o f the platform. O n the easter n margin th e mass-flow s hav e bee n relate d t o a

Table 2. Estimates o f denudation o n the Faroe Platform an d deposition o f sediments derived from th e platform

FNM

FSC

FP

FBC

Sum

3

Volume (km ) Removed Paleocene basalt Eocene-Oligocene (FPC-B) 9900 Lower-Middle Miocene (FPC-C) 445 Upper Miocene-Lower Pliocene (FPC-D. 1) 225 Upper Pliocene-Recent (FPC-D.2-3) 450

0 0 0

Total 2110

0 25000

46000 -

ab

'

C

9900

C

-4600

46000

0 1000

0

0 -13110

0 4 1 2 3

0 -1630

0

12 d

Mass balance (mas s in kg X 10 ) Removed Paleocene basalt -13110 Oligocene (FPC-B) 2029 Lower-Middle Miocene (FPC-C) 912 Upper Miocene-Lower Pliocene (FPC-D.l) 461 Upper Pliocene-Recent (FPC-D.2-3) 922

5 5386 3 2421 3 1224 5 2448

Total 4325

5 5125

0 2029

5 -13110

FSC, Faroe-Shetlan d Channel; FNM, norther n margin of Faroe Platform; FBC , Faro e Bank Channel; FP, Faroe Platform. Positiv e numbers indicate deposition; negative numbers indicate denudation. a Excluding Eocene-Oligocene sediments that are removed later. b Including some Upper Eocene sediments. cBulk volume, al l Eocene-Recent sediments included. d Assuming dry density o f basalt is 2850 kg m (porosity c . 10%) , an d of sediments is 2050 kg m (porosit y c. 35%) .

CENOZOIC EVOLUTION O F THE FAROE PLATFORM

307

Fig. 15 . Estimated Eocene-Recen t sedimen t yield fro m th e Faroe Platform . A peak in sediment production in Late Pliocen e tim e i s seen c . 8 Ma after th e star t of the Neogene compressio n phase . A less pronounce d pea k partly overlaps the Late Eocene-Oligocene deformation event.

Fig. 16 . Cross-section s showin g aspect s o f th e structura l evolutio n o f th e Faro e Platform , (a ) Present-da y configuration o f the platform; removed basalt i s indicated b y grey shading, (b ) Load effect s o f Eocene-Recent sediments an d the removed basal t sectio n hav e been remove d assumin g Air y isostasy an d a mantle densit y of 3100kg m"3. (c ) A possible model fo r th e configuratio n of th e platfor m in lates t Paleocen e time . Th e mode l presented assumes volcanic feeder systems close to the line of incipient break-up. Location of profile is shown in Fig. 7.

308

M.S. ANDERSEN ETAL.

calculated for both removed basalt and deposite d sediments. Basalt is generally less porous than sediments. We hav e thu s recalculate d th e volume s o f denudation and deposition t o dry mass assuming the porosity of basalt i s c . 10 % and the porosity of volcaniclasti c sediment s i s c . 30% . The estimated tota l amoun t o f sediment s deposite d in th e basins (56000km 3, 11480 0 X lO^kg ) is reasonably clos e t o th e estimate d amoun t o f basalt remove d fro m th e Faro e Platfor m (46000km3, 13 1 100 X 10 12kg) (Table 1) . Overall the mean specific sediment yield based on the total Cenozoic denudation of the Faroe Platfor m and deposition i n the Faroe-Shetland Channel is approximately 5 5 X 10 3kg km~ 2 a" 1. Thi s i s within the normal rang e o f sedimen t yiel d fro m large river s (Hoviu s 1998) . O n th e basi s o f th e depositional dat a from th e easter n margin of the Faroe Platform and the Faroe-Shetland Channel it can be seen that denudation rates changed over time (Fig. 15). It appears that the denudation rate in Late Eocene-Oligocene time was rather hig h (c. 70xl0 3 kg km~ 2 a" 1) possibl y reflectin g uplift o f th e Faro e Platform . I n Early-Mid Miocene time it was decreased t o c. 40 X 10 3kg km~ 2 a" 1, an d i n Lat e Miocen e tim e th e denudation rat e wa s c . 3 5 X 10 3kg km~ 2 a" 1. This wa s followe d b y a sudde n increas e i n denudation rat e i n mid-Pliocen e tim e t o c. 175Xl0 3 kg km" 2 a" 1. A mor e detaile d analysis o f th e sedimen t successio n aroun d th e Faroe Platfor m woul d presumably revea l signifi cant fluctuation s o f denudationa l an d deposi tional rates. Peaks in depositional rate s occurred after th e tw o distinc t inversio n o r compressio n events during Late Eocene-Early Oligocene and Mid-Late Miocene time. However, only the Late Pliocene depositiona l peak is well constrained in time, an d i t occurre d c . 8 Ma after th e midNeogene deformatio n event . Discussion Fig. 17 . (a ) Th e denudatio n dept h o f a 20k m thic k crust a s a functio n o f upper-crusta l densit y fo r thre e compressional strain s (2% , 3% an d 5 % shortening) . Erosion to base level is assumed, (b) and (c) show two conceptual models, which could explain the uplift an d denudation o f the Faroe Platform , (b ) A small part of the basal t platea u i s lifte d abov e bas e leve l durin g a compressional event . Th e plateau i s then erode d unti l the top of the plateau again is at base level, (c) Basalt is emplaced abov e bas e level . Th e crus t will subsid e to retain isostatic equilibrium, and only a small part of the basalt will be above base level. As a result of thermal subsidence durin g erosion o f the plateau , th e erosio n surface wil l reac h bas e leve l befor e th e platea u i s completely removed.

A simplifie d profil e acros s th e Faroe-Shetlan d Channel fro m wel l 205/9- 1 in th e U K secto r i s shown in Fig. 16a. It has been suggeste d tha t the uplift an d denudatio n o f th e Faro e Platfor m t o some extent could be attributed to compressiona l deformation (Boldreel etal 1998 ; Andersen etal. 2000). Simpl e consideration s concernin g isostatic balanc e indicat e tha t denudatio n induced by tectoni c uplif t depend s o n th e densit y o f th e removed rock (Fig. 17). Unless the removed rock is relatively dense , small initial uplif t o f tectonic origin woul d giv e ris e t o onl y relativel y insignificant additiona l uplif t an d denudation . However, remova l o f heav y materia l suc h a s

CENOZOIC EVOLUTION OF THE FAROE PLATFORM

basalt wil l giv e ris e t o significan t denudatio n even in the case of small tectonic strain . As can be see n i n Fig . 17 , abou t 2-3 % crusta l shortening coul d accoun t fo r th e averag e denudation o f th e Faro e Platform . Crusta l shortening o f the Faro e Platform , a s a result o f compression, i s indicate d bot h b y dip-sli p analysis o f faul t plane s o n th e Faro e Island s (Geoffrey 1993 , 1994 ) an d b y interpretatio n o f offshore seismi c data (e.g. Boldreel et al. 1998) . However, horizo n balancin g base d o n th e regional profil e i n Fig . 16 a indicate s tha t th e accumulated crusta l shortenin g o f th e platfor m since th e emplacemen t o f th e Lat e Paleocen e basalts presumably is les s tha n 1 % and ma y b e insignificant. White & Lovell (1997 ) propose d tha t fluctuating intensit y o f th e Icelandi c plum e wa s responsible fo r episodic Paleocen e uplif t o n th e British Isles . Therma l fluctuation s o f the plum e are evidence d b y V-shape d ridge s arrange d symmetrically aroun d th e Reykjane s Ridg e (White e t a l 1995) . However , th e V-shape d ridges wer e formed where thermal pulses of the plume intersected th e spreading ridge, an d there is apparently no documentation for off-axis uplif t connected t o th e Neogen e therma l pulses . I t i s thus unlikely that thermal pulses are responsible for th e Neogen e uplif t an d denudatio n o f th e Faroe Platform . Assuming Air y isostasy , w e ma y remove th e approximate effec t o f th e loa d o f sediment s deposited in the Faroe-Shetland Channel and of basalts remove d fro m th e Faro e Platfor m sinc e Late Paleocen e tim e an d obtai n th e profil e i n Fig. 16b . Comparing the profile in Fig. 16 b with an estimate of the Late Paleocene topography , we obtain a n approximatio n o f th e accumulate d tectonic vertica l movement s alon g th e profil e (Fig. 16c) . Th e larges t tectoni c subsidenc e (c. 1600m ) occurre d i n th e centr e o f th e Faroe-Shetland Channel . O n th e Faero e Plat form th e tectoni c subsidenc e i s 400-800m . Northwest o f th e Faro e Platfor m th e tectoni c subsidence is greater than 800 m. This distributio n o f estimate d tectoni c sub sidence ma y b e explaine d partl y b y therma l contraction o f th e hea d o f the Icelandi c plume , which would have been present below the entire Faroe regio n (e.g . Whit e & McKenzi e 1989) . Reduced dynami c support from th e plum e may also b e involve d (e.g . Nadi n e t al . 1997) . However, th e large r tectoni c subsidenc e i n th e Faroe-Shetland Channel and NW of the platform indicates a n additiona l componen t o f tectoni c subsidence cause d b y post-rif t therma l contraction. Mid-Cretaceou s riftin g o f th e Faero e Shetland Channel with j8 ~ 2 has been proposed

309

(e.g. Nadi n e t a l 1997) . Les s pronounce d Paleocene riftin g (e.g . Dea n e t a l 1999 ) ma y also be involved. On the Faroe Platform, tectonic subsidence has been mor e than compensated by isostatic uplif t cause d b y th e remova l o f Paleocene basalt . Th e lo w tectoni c subsidenc e around the Faroe Islands (260-300 km in Fig. 16) is presumably an artefact introduced because we assumed tha t th e lithospher e ha s n o flexura l strength. Suitable 2D flexural modelling, including a continuation o f th e profil e o n th e uplifted Shetland Platform , S E o f th e Faroe-Shetlan d Channel, shoul d modif y th e subsidenc e values , especially aroun d the Faroe Islands. Flexural strength of the lithosphere is ignored in th e discussio n above . However , i t i s stil l evident from Fig. 16 that Eocene-Recent erosion of th e Faro e Platfor m coul d b e th e dominan t cause fo r th e Eocene-Recen t uplif t o f th e platform. Therefore , w e sugges t tha t th e Eocene-Recent uplif t o f th e Faro e Platfor m is primarily cause d b y isostati c compensatio n fo r the erosio n o f th e platfor m (Fig . 17c) . I n thi s context, th e widesprea d Lat e Pliocene-Pleisto cene denudatio n tha t provide d th e sediment s found i n th e larg e Pliocen e wedge s aroun d th e Faroes is most readily understood in association with the climatic-eustactic changes that affecte d the are a durin g th e Norther n Hemispher e glaciation. We woul d lik e t o than k W.G . Harra r an d tw o anonymous referee s fo r constructiv e revie w o f th e paper. Th e vitrinit e reflectanc e measurement s wer e made b y C . Guva d an d supporte d financiall y b y th e Atlantic Margi n Grou p (Mobi l Nort h Se a Limited , Enterprise Oi l pi c an d Statoi l U K Limited) . W e ar e grateful t o Western Geophysical and Veritas DGC fo r their contribution . W e wis h t o than k th e F0roy a Jar5fr05isavn an d the Geologica l Surve y o f Denmark and Greenland for permission to publish thi s work .

References ANDERSEN, M.S . 1988 . Lat e Cretaceou s an d earl y Tertiary extensio n an d volcanis m aroun d th e Faeroe Islands . In : MORTON , A.C . & PARSON , L.M. (eds ) Early Tertiary Volcanism an d th e Opening o f th e NE Atlantic. Geologica l Society , London, Special Publications, 39, 112-122. ANDERSEN, M.S . 1999 . Structura l evolutio n o f th e Faroe-Rockall regio n i n th e Cenozoicu m (abstract). Geonytt, Norsk Geologisk Vinterm0te, 6-8 January 1999. Nors k Geologis k Forening , Stavanger. ANDERSEN, M.S. , NIELSEN , T. , SORENSEN , A.B. , BOLDREEL, L.O . & KUIJPERS , A . 2000 . Cenozoi c sediment distributio n an d tectoni c movement s i n the Faroe region. Global an d Planetary Change, 24 (3-4), 239-259.

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BOLDREEL, L.O . & ANDERSEN , M.S . 1993 . Lat e Paleocene t o Miocene compressio n i n the Faeroe Rockall area . In : PARKER , J.R . (ed. ) Petroleum Geology of Northwest Europe: Proceedings of the 4th Conference. Geologica l Society , London , 1025-1034. BOLDREEL, L.O . & ANDERSEN , M.S . 1994 . Tertiar y development o f the Faeroe-Rockall Plateau base d on reflectio n seismi c data . Bulletin o f th e Geological Society of Denmark, 41, 162-180 . BOLDREEL, L.O . & ANDERSEN , M.S . 1995 . Th e relationship betwee n th e distributio n o f Tertiar y sediments, tectoni c processe s an d deep-wate r circulation aroun d th e Faero e Islands . In : ScRUT TON, R.A. , STOKER , M.S. , SHIMMIELD , G.B . & TUDHOPE, A.W. (eds) The Tectonics, Sedimentation and Palaeoceanography of the North Atlantic Region. Geologica l Society , London , Specia l Publications, 90 , 145-158. BOLDREEL, L.O. , ANDERSEN , M.S . & KUIJPERS , A . 1998. Neogene seismic facies and deep gateways in the Faroe Bank area, NE Atlantic. Marine Geology, 152, 129-140 . CLOKE, L , DAVIES , R.J. , FERRERO , C. , LINE , C . & BINGHAM, J . 2000. Inversion: th e critica l elemen t in determining th e success of the petroleum syste m of th e Faroe-Shetlan d Channel . Exhumation o f Circum-Atlantic Margins: Timing, Mechanisms and Implications for Hydrocarbon Exploration, London, 13-14 June 2000. CLOKE, I. , DAVIES, R.J. , LINE , C., HORNAFIUS, S . an d MCLACHLAN, K . 1999 . Petroleu m syste m o f th e Faroe-Shetland Basin . AAPG International Conference and Exhibition, Birmingham, 12-15 September 1999, pp. 119-123. DAM, G. , LARSEN , M . & S0NDERHOLM , M . 1998 . Sedimentary respons e t o mantl e plumes : impli cations fro m Paleocene onshore successions , wes t and east Greenland. Geology, 26, 207-210. DAVIES, R. , CARTWRIGHT , J. & RANA , J . 1999 . Gian t hummocks i n deep-wate r marin e sediments : evidence fo r large-scal e differentia l compactio n and density inversion durin g early burial. Geology, 27 (10), 907-910. DEAN, K. , MCLACHLAN , K . & CHAMBERS , K . 1999 . Rifting an d developmen t o f th e Faeroe-Shetlan d Basin. In : FLEET , AJ . & BOLDY , S.A.R . (eds ) Petroleum Geology of Northwest Europe: Proceedings o f th e 5t h Conference. Geologica l Society , London, 533-544. EBDON, C.C. , GRANGER , P.J. , JOHNSON , H.D . & EVANS, A.M . 1995 . Earl y Tertiar y evolutio n an d sequence stratigraphy o f th e Faeroe-Shetlan d Basin: implication s fo r hydrocarbon prospectivity . In: SCRUTTON , R.A. , STOKER , M.S. , SHIMMIELD , G.B. & TUDHOPE , A.W . (eds ) Th e Tectonics, Sedimentation and Palaeoceanography of the North Atlantic Region. Geologica l Society , London, Special Publications , 90 , 51-69. EVANS, D . 2000 . E U projec t o n developmen t o f th e glaciated Europea n Margin . EO S Transactions, American Geophysical Union, 81, 423-425. GATLIFF, R.W. , KITCHEN , K. , RITCHIE , J.D . & SMYTHE, D.K . 1984 . Interna l structur e o f th e

Erlend Tertiar y volcani c complex , nort h o f Shet land, revealed b y seismic reflection . Journal of th e Geological Society, London, 141, 555-562 . GEOFFROY, L. , ANGELIER , J . & GERGERAT , F . 1993 . Sur 1'evolution tectonique cassante tertiaire des lies Feroe, Atlantiqu e Nord ; l a compressio n feringi enne. Comptes Rendus de I'Academic de s Sciences, 316, 975-982 . GEOFFROY, L. , BERGERAT , F . & ANGELIER , J . 1994 . Tectonic evolutio n o f th e Greenland-Scotlan d ridge durin g th e Paleogene : ne w constraints . Geology, 22, 653-656. HALD, N. & WAAGSTEIN, R. 1984. Lithology of a 2-km sequence o f Lower Tertiar y tholeiiti c lavas drille d on th e Suduro y Faero e Island s (Lopra-1) . In : BERTHELSEN, O. , NOE-NYGAARD , A . & RASMUS SEN, J. (eds) The Deep Drilling Project 1980-1981 in th e Faeroe Islands. F0roy a Frodskaparfelag , Torshavn, 15-37 . HALL, B.D. & WHITE, N. 1994 . Origi n o f anomalou s Tertiary subsidenc e adjacen t t o Nort h Atlanti c continental margins . Marine an d Petroleum Geology, 11 , 702-714. HENRIKSEN, S . & VORREN , T.O . 1996 . Late Cenozoi c sedimentation an d uplif t histor y o n th e mid Norwegian continental shelf. Global and Planetary Change, 12, 171-199. KITCHEN, K. & RITCHIE, J.D. 1987 . Geological review of th e Wes t Shetlan d area . In : BROOKS , J.D . & GLENNIE, K.W . (eds) Petroleum Geology of North West Europe, Procedings of the Third Conference. Graham & Trotman, London, 737-749. Hovius, N . 1998 . Control s o n sedimen t suppl y b y large rivers . In : SHANLEY , K.W . & MCCABE , P.W . (eds) Relative Role o f Eustasy, Climate, an d Tectonism i n Continental Rocks. Societ y o f Economic Paleontologist s an d Mineralogists , Special Publications , 59, 3-16. HUMLUM, O. & CHRISTIANSEN , H.H. 1998 . Mountain climate an d periglacial phenomen a i n th e Faero e Islands. Permafrost an d Periglacial Processes, 9, 189-211. J0RGENSEN, O . 1984 . Zeolit e zone s i n th e basalti c lavas of the Faroe Islands. In: NOE-NYGAARD, O. & RASMUSSEN, J . (eds ) Th e Deep Drilling Project 1980-1981 i n th e Faeroe Islands. F0roy a Frods kaparfelag, Torshavn , 71-91. KI0RBOE, L . 1999 . Stratigraphi c relationship s o f th e Lower Tertiary of the Faeroe Basal t Plateau and the Faeroe-Shetland Basin. In: FLEET , A.J. & BOLDY , S.A.R. (eds ) Petroleum Geology o f Northwest Europe: Proceedings of the 5th Conference. Geological Society, London , 559-572 . KNOX, R.W.O.B. , HOLLOWAY , S. , KIRBY , G.A . & BAILY, H.E. 1997 . Early Paleogene Lithostratigraphy and Sequence Stratigraphy. Stratigraphic Nomenclature of the UK North West Margin, 2. British Geologica l Survey , Keyworth. KOUL, S.L. , CHADDERTON , L.T . & BROOKS , C.K . 1983. Eas t Greenlan d an d th e Faero e Islands : a fission trac k study . Kongelige Danske Videnskabernes Selskab, Matematisk-fysiske Meddelelser, 40, 7-34 .

CENOZOIC EVOLUTION OF THE FAROE PLATFORM KUIJPERS, A . 2000. Late Quaternary slop e instabilit y on th e Faero e margin : mas s flo w feature s an d timming of events. Geo-Marine Letters, 149-159. LAMERS, E . & CARMICHAEL , M.M . 1999 . Th e Paleocene deepwate r sandston e pla y Wes t o f Shetland. In: FLEET , AJ. & BOLDY , S.A.R . (eds) Petroleum Geology of Northwest Europe: Proceedings o f th e 5t h Conference. Geologica l Society , London, 645-659 . LARSEN, L.M. , WAAGSTEIN , R. , PEDERSEN , A . & STOREY, M . 1999 . Trans-Atlanti c correlatio n o f Palaeogene volcani c succession s i n th e Faero e Islands an d Eas t Greenland . Journal o f th e Geological Society, London, 156, 1081-1095 . NADIN, P.A. , KUZNIR, N.J . & CHEADLE , MJ . 1997 . Early Tertiary plume uplift of the North Sea and the Faeroe-Shetland Basins . Earth an d Planetary Science Letters, 148, 109-127 . NAYLOR, PH., BELL, B.R., JOLLEY, D.W., DURNALL, P. & FREDSTED , R . 1999 . Palaeogene magmatis m in the Faeroe-Shetlan d Basin : influence s o n uplif t history an d sedimentation . In : FLEET , AJ . & BOLDY, S.A.R. (eds) Petroleum Geology o f Northwest Europe: Proceedings of the 5th Conference. Geological Society, London , 545-558. NIELSEN, T . & VA N WEERING, T.C.E . 1998 . Seismic stratigraphy an d sedimentar y processe s a t th e Norwegian Se a margi n northeas t o f th e Faero e Islands. Marine Geology, 152, 141-157. NOE-NYGAARD, A . 1968 . O n extrusio n form s i n plateau basalts . Shiel d volcanoe s o f 'Scutulum ' type. In: FREDERIKSON, S. (ed.) Science in Iceland. Visindafelag Icelandinga, Reykjavik, 10-13. PLANKE, S. , SYMONDS , P.A. , ALVESTAD , E . & SKOGSEID, J . 2000 . Seismi c volcanostratigraph y of large-volum e basalti c extrusiv e complexe s o n rifted margins . Journal o f Geophysical Research, 105 (B8), 19335-19351. RASMUSSEN, J . & NOE-NYGAARD , A . 1970 . Geology of the Faroe Islands (Pre-Quaternary}. Danmarks Geologiske Unders0gelse I . Series, 25 . Reitzel s Forlag, Copenhagen . RITCHIE, J.D., GATLIFF, R.W. & RICHARDS, PC. 1999 . Early Tertiary magmatism in the offshore N W UK margin an d surrounds . In: FLEET , AJ . & BOLDY , S.A.R. (eds ) Petroleum Geology o f Northwest Europe: Proceedings of the 5th Conference. Geological Society, London , 573-584 . RUMPH, B. , REAVES , C.M. , ORANGE , V.G . & ROBINSON, D.L . 1993 . Structurin g an d transfe r zones i n th e Faero e Basi n i n a regiona l tectoni c context. In : PARKER, J.R. (ed.) Petroleum Geology of Northwest Europe: Proceedings of the 4th Conference. Geologica l Society , London , 999-1009.

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SMYTHE, D.K . 1983 . Faeroe-Shetlan d Escarpmen t and continenta l margi n nort h o f th e Faroes . In : MORTON, A.C . & PARSON , L.M . (eds ) Early Tertiary Volcanism and the Opening of the NE Atlantic. Geologica l Society , London , Specia l Publications, 39, 109-119. SMYTHE, D.K. , CHALMERS , J.A. , SKUCE , A.G. , DOBINSON, A . & MOULD , A.S . 1983 . Earl y opening history of the North Atlantic—I. Structure and origi n o f th e Faeroe-Shetlan d Escarpment . Geophysical Journal of the Royal Astronomical Society, 72, 373-398 . STOKER, M.S . 1999. Mid- to Late Cenozoic Stratigraphy. Stratigraphic Nomenclature of the UK North West Margin. Britis h Geologica l Survey , Keyworth. STOKER, M.S., NIELSEN, T. , VA N WEERING, T.C.E . & KUIJPERS, A . 2002 . Towards a n understandin g o f the Neogene tectonostratigraphic framewor k of the NE Atlantic Margin between Ireland an d the Faroe Islands. Marine Geology, in-press. STOREY, M. , DUNCAN , R.A. , PEDERSEN , A.K. , LARSEN, L.M . & LARSEN , H.C . 1998. 40 Ar/39Ar geochronology o f th e Wes t Greenlan d Tertiar y volcanic province . Earth an d Planetary Science Letters, 160 (3/4), 569-586. TALWANI, M., UDINTSEV, G. et al. 1976. Initial Reports of th e Deep Se a Drilling Project, 38 . U S Government Printing Office , Washington , DC. WAAGSTEIN, R. 1988 . Structure, composition an d ag e of the Faeroe basal t plateau. In: MORTON , A.C. & PARSON, L.M. (eds) Early Tertiary Volcanism an d the Opening o f the NE Atlantic. Geological Society, London, Specia l Publication, 39, 225-238. WAAGSTEIN, R . & HEILMANN-CLAUSEN , C . 1995 . Petrography an d biostratigraph y o f Paleogen e volcaniclastic sediments dredged from th e Faeroe s shelf. In : SCRUTTON , R.A. , STOKER , M.S. , SHIMMIELD, G.B . & TUDHOPE , A.W . (eds ) Th e Tectonics, Sedimentation and Palaeoceanography of th e North Atlantic Region. Geologica l Society , London, Special Publications , 90 , 179-197. WHITE, N. & LOVELL, B. 1997. Measuring the pulse of a plume with the sedimentary record . Nature, 387 (6636), 888-891 . WHITE, R.S. & MCKENZIE, D. 1989. Magmatism at rift zones: th e generatio n o f volcani c continenta l margins and floo d basalts. Journal of Geophysical Research, 94B, 7685-7729. WHITE, R.S., BOWN, J.W. & SMALLWOOD , J.R. 1995. The temperature of the Iceland plume and origin of outward propagatin g V-shape d ridges . Journal of the Geological Society, London, 152, 1039-1045.

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Late Neogene development of the UK Atlantic margin M. S. STOKE R British Geological Survey, Murchison House, West Mains Road, Edinburgh EH9 SLA, UK (e-mail: mss@ bgs.ac.uk) Abstract: Th e late Neogen e (Pliocene-Holocene ) interva l witnesse d a significant chang e in sedimentatio n styl e acros s th e U K Atlanti c margi n tha t culminate d i n it s presen t morphological expression . Th e onse t o f chang e i s marke d b y th e creatio n o f a regional, angular, erosiona l unconformit y tha t ca n b e trace d fro m th e Hebride s an d West Shetlan d margins int o th e adjacen t deep-wate r basin s o f the Rockal l Troug h an d Faeroe-Shetland Channel. I n th e Rockal l Trough , th e unconformit y i s a submarin e erosio n surfac e tha t i s dated a s earl y Pliocen e i n age , betwee n c . 3.8 5 an d 4. 5 Ma. O n th e Hebride s an d Wes t Shetland margins, th e dating o f the unconformity i s slightly less well constrained and spans the lates t Miocene(?)-earl y Pliocen e (c . 5.5-3. 8 Ma) interval . Th e formatio n o f thi s unconformity may have resulted fro m the seaward tilting and subsidence o f the shelf margin , which ma y hav e furthe r modifie d th e oceanographi c circulatio n patter n i n th e adjacen t basins. The sedimentary response to this event was relatively quick in the deep-water basins, which preserv e a recor d o f earl y Pliocen e (pos t 3.8 5 Ma) t o Holocen e sediment-drif t accumulation, albei t wit h a shif t i n th e focu s o f sedimentatio n relativ e t o th e underlyin g strata. I n contrast , majo r progradin g wedges , whic h hav e contribute d extensivel y t o th e construction o f the Hebrides and West Shetlan d margins an d to a lesser extent the Rockall Bank, dat e essentially fro m late Pliocene time and largely correlate with th e influx of ice rafted materia l t o th e margin . However , indication s fo r a restricte d lowe r Pliocen e component t o th e shelf-margi n successio n sugges t tha t thi s apparen t dela y o r la g i n sedimentation, relativ e to the basins, ma y be a natural respons e t o the rate o f denudation o f the adjacent landmasses . Th e regional observation s of f NW Britain suppor t th e concep t o f Neogene tectonic uplift .

It i s becomin g increasingl y apparent tha t high rates o f denudation in the lat e Neogene interval have strongl y influence d th e constructio n o f continental margins surrounding the NE Atlantic Ocean. Ther e ar e indication s tha t this phenomenon resulte d fro m th e uplif t o f adjacen t landmasses, althoug h th e mechanis m drivin g such a large-scale event remains open to question (Japsen & Chalmer s 2000 , an d reference s therein). However, there is no doubting that on e of th e mai n sedimentar y response s t o th e widespread denudatio n is preserved in th e for m of majo r progradin g shelf-margi n wedges . Substantial wedge s hav e bee n describe d fro m around th e N E Atlanti c Ocean , including : th e NW Britis h margi n (Stoke r 1999 ; Stoke r e t al 2001, 2002), the east Faeroes margin (Andersen et a l 2000) , th e northern North Sea (Gregersen et al. 1997) , the S W Norwegian margin (Sejru p et al . 1996) , th e mid-Norwegia n margi n (Rokoengen e t al . 1995 ; Henrikse n & Vorre n 1996; Evans et al. 2000), the Barents Sea margin (Vorren & Laberg 1997) , an d th e eas t an d west

Greenland margin s (Solhei m e t al . 1998 ; Chalmers 2000). Although ther e i s genera l agreemen t tha t circum-NE Atlantic shelf-margin progradation is of Plio-Pleistocen e age , a n improve d strati graphic resolutio n o f th e offshor e sequence s i s necessary to determine whether or not sedimentation i s linke d directl y t o th e causa l (?uplift ) mechanism o r i s a delaye d response . There ar e indications, fro m bot h offshor e Norwa y (Eidvin et al . 2000 ) an d N W Britai n (Stoke r e t al . 1994), tha t th e bul k o f th e sedimen t preserved in th e progradin g wedge s is o f lat e Pliocene Pleistocene age . However , th e evidenc e fro m offshore N W Britain , includin g th e deep-wate r oceanographic an d sedimentar y response , suggests a n earl y Pliocene ag e fo r th e onse t of change a s marke d b y th e lowe r boundin g unconformity. The ai m o f thi s pape r i s t o summariz e th e stratigraphic recor d an d styl e o f sedimentatio n for th e uppe r Neogen e (Pliocene-Holocene ) succession of f N W Britain , extendin g fro m th e

From: DORE , A.G., CARTWRIGHT , J.A., STOKER , M.S. , TURNER , J.P. & WHITE , N. 2002 . Exhumation of the North Atlantic Margin: Timing, Mechanisms and Implications for Petroleum Exploration. Geological Society, London, Special Publications, 196, 313-329 . 0305-8719/02/$15.00 © The Geological Society of London 2002.

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Hebrides an d Wes t Shetlan d shelve s int o th e adjacent deep-wate r basin s o f th e Rockal l Trough an d Faeroe-Shetlan d Channe l (Fig . 1) . The pape r utilize s an d build s upo n regiona l stratigraphic studies undertaken in this region, by the autho r (e.g. Stoke r 1999 ; Stoker et al 2001 , 2002), by focusing in more detail on the nature, distribution an d timin g o f formatio n o f th e regionally significan t intra-Pliocen e unconfor mity, an d th e subsequen t developmen t o f th e overlying shelf-margi n an d deep-wate r basina l successions. I n vie w o f it s margin-wid e occur rence, the event responsible fo r the development

of thi s unconformit y implies a regiona l contro l on continental margin evolution in this region.

Neogene setting off NW Britain The Neogene successio n preserve d o n th e continental margi n of f N W Britai n ha s bee n divided int o tw o megasequence s o f predomi nantly Miocene (including lower Pliocene i n the basins) an d Pliocene-Holocen e age , whic h ar e separated b y a regiona l unconformit y o f earl y Pliocene ag e (Stoke r e t a l 2002 ) (Fig . 2) . A basinal successio n o f sedimen t drift s an d

Fig. 1 . Bathymetric setting of the UK Atlantic margin (contours in metres), showing locations of BGS boreholes (•), ODP and commercial borehole s (O ) and major depocentre s referre d to in the text, the geoseismic section s illustrated in Fig. 2, and the seismic profiles in Figs 4-8. Depocentres : DF , Donegal Fan; BF, Barra Fan; ERW, East Rockall wedge; SSF, Sula Sgeir Fan; WSW, West Shetland wedge; EFW, East Faeroes wedge; NSF, North Sea Fan. Sediment drifts (shaded) : F, Feni Ridge; H, Hatton Drift; Bj , Bjorn Drift; G , Gardar Drift. Geographica l locations: WSS , Wes t Shetlan d Shelf ; FB , Faero e Bank ; BBB , Bil l Bailey' s Bank ; WTR , Wyville-Thomso n Ridge; LB , Lous y Bank ; RB, Rosemar y Bank ; GBB, George Blig h Bank ; AD, Anto n Dohrn Seamount ; HT, Hebrides Terrac e Seamount ; RKB , Rockall Bank ; HRB , Hatton Rockal l Basin ; HB , Hatton Bank ; FS , Faeroe Shelf; HS, Hebrides Shelf; MS, Malin Shelf; IS, Irish Shelf; PB, Porcupine Bank; PSB, Porcupine Seabight; PAP, Porcupine Abyssa l Plain.

Fig. 2. Interpreted geoseismic sections focusing on the upper Cenozoic stratigraphy. Section (a) based on profile OF94-29 (Faeroes margin) and BGS profiles 83/04-29 and 30 (West Shetland margin); section (b) based on profiles DG95-8 and -10 (courtesy of Conoco and Digicon) and BGS profile 83/04-6; section (c) based on BGS profiles 92/ 01-38 and -39; section (d) based on Mobil line M89-WB-2 acquired by the 'Rockall Continental Margin Consortium' (see Acknowledgements); section (e) based on profile WRM96-107 (courtesy of Fugro Geoteam). The variable early Pliocene reflecto r notation is based on several stratigraphic studies: CIO, Stoker et al. (2001); INU, Stoker (1999); CN-040, Andersen et al. (2000). Lotations of sections are shown in Fig. 1. The boxes indicate the approximate regional context of the seismic profiles illustrated in Figs 4, 5 and 7. TWTT, two-way travel time.

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contourites, locall y u p t o 400-500m s thic k (two-way trave l time , TWTT) , form s th e Miocene megasequence . Thes e deposit s charac teristically displa y extensiv e onla p an d upslop e accretion o n bot h flank s o f th e Rockal l Troug h and o n th e Wes t Shetlan d Slop e (Stoke r 1999 ; Stoker e t al. 2001). Transgressive shallow-wate r glauconitic sandstone s ar e locall y preserve d o n the Hebride s an d West Shetland shelves (Stoke r et al . 1993 , 1994 ; Stoke r 1999) . Thi s mega sequence developed i n response to the establishment o f th e moder n deep-wate r exchang e between th e Arcti c an d Nort h Atlanti c oceans , with depositio n o f sedimen t drift s occurrin g i n the Rockal l Troug h an d Faeroe-Shetlan d Channel as the deep-water currents stabilized. The indicatio n o f a significan t chang e i n th e development o f th e continenta l margi n i s preserved a t th e to p o f thi s megasequence , which ha s bee n extensivel y erode d bot h o n th e shelves an d i n th e adjacen t deep-wate r basins . The overlyin g Pliocene-Holocen e megase quence i s interpreted a s a third-order composit e lowstand system s trac t tha t rest s wit h angula r discordance o n th e Miocen e megasequenc e (Stoker 1999 ; Stoke r e t a l 2001) . Plio Pleistocene progradin g wedge s ar e th e mos t prominent indicato r o f a change i n styl e o f lat e Neogene sedimentatio n an d margin construction ; however, a les s well-documente d bu t equall y significant chang e i s th e modificatio n t o th e oceanographic circulatio n tha t substantiall y altered th e patter n o f deep-water sedimentation . It is the development o f the Pliocene-Holocene megasequence that forms the basis of the present study. Late Neogene stratigraphy and style of sedimentation The geometry and generalized distribution of the Pliocene-Holocene megasequenc e of f N W Britain i s depicte d i n Fig s 2 an d 3 . Th e mos t distinctive featur e o f th e megasequenc e i s highlighted b y th e shelf-margi n depocentre s that ar e th e prograding wedges . However , thes e wedges for m onl y on e componen t o f a mixe d depositional syste m tha t involve s bot h down slope an d alongslop e processes . Sedimen t drift s and contourite s dominat e th e basina l strat a an d where the y overlap wit h the prograding wedge s an intercalate d successio n o f alongslop e an d downslope sediment s i s generally developed . I n terms o f lat e Neogen e developmen t o f th e U K Atlantic margin , th e deep-wate r sedimentar y response i s inextricabl y linke d t o th e formation of th e progradin g wedges . Thes e feature s

together wit h the chronological evidenc e fo r the onset o f chang e (th e basa l unconformity ) ar e described belo w an d summarize d i n term s o f a late Neogene even t stratigraphy. The basal unconformity Nature o f th e boundary. Th e bas e o f th e Pliocene-Holocene megasequenc e i s marke d by a widesprea d unconformity . I n th e Rockal l Trough, the CI O reflector of Stoker et al. (2001 ) represents thi s boundar y (Fig . 2b-e) , whic h forms an erosional , angula r unconformit y truncating strata both on the flanks and in the axis of the Troug h (Fig s 4-6) . Adjacen t t o th e Hebridean margin , thi s reflecto r form s th e bas e of th e progradin g shelf-margi n successio n tha t includes th e Barr a an d Sul a Sgei r fan s (Fig . 2 b and d). Figure 4 illustrates the erosive characte r of th e basa l unconformit y a s trace d fro m th e basin floo r int o th e Barr a Fa n o n th e lowe r Hebrides Slope . Strata l termination s i n th e Miocene (t o lowe r Pliocene ) basina l deposit s are clearl y truncate d b y th e unconformity . Although some of this erosion ma y be associated with th e depositio n o f th e basa l debris-flo w package i n th e Fan , th e remova l o f a significan t section o f th e Miocen e successio n beyon d th e limit of the Fan (Fig. 4 (inset a) , and Figs 5 and 6a) implie s submarin e erosio n associate d wit h deep-water bottom currents. On the western flank of th e Rockal l Trough , thi s reflecto r is , itself , commonly erode d an d truncated by the present day sea-bed surfac e (Figs 2d-e an d 5); a furthe r consequence o f persisten t bottom-curren t activity (se e below) . O n the Hebridea n margin , the unconformit y ca n b e trace d fro m th e basi n floor onto the slope (Fig. 4). On the upper slope, it forms a seaward-tilted erosio n surfac e that has cut int o an d locall y remove d shallow-wate r sediments o f mid- to late Miocene age and older (Fig. 6b) (Stoker et al. 1993 , 1994) . This linkage between th e deep-wate r basi n an d th e shel f margin demonstrates the regional significanc e of this reflector . O n th e eas t Rockal l Bank , th e nature o f the base o f the Eas t Rockal l wedg e i s less clear , althoug h there i s downla p ont o olde r Miocene strat a (Fig . 2e) . In the Faeroe-Shetland region, the INU (intraNeogene unconformity) of Stoker (1999) and the CN-040 reflecto r o f Anderse n e t al . (2000 ) denote th e basa l unconformit y on , respectively , the Wes t Shetlan d an d Eas t Faeroe s margin s (Fig. 2a) . Wes t of Shetland , th e basa l unconfor mity can be traced fro m th e West Shetland Shelf into the Faeroe-Shetland Channel. O n the shelf margin, i t i s a planar , seaward-tilte d erosio n surface tha t truncate s th e underlyin g uppe r

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Fig. 3 . Map showing the generalized distributio n and gross depositional environmen t of the Pliocene-Holocene megasequence alon g the Atlantic margin between Irelan d an d the Faeroe-Shetland region. Abbreviation s a s in Fig. 1 , and FSC, Faeroe-Shetland Channel . Present-day contours (as in Fig. 1 ) are superimposed o n the map to act as an approximate guid e to the palaeomorphology of the continental margin . The present-day bottom-circulatio n pattern is taken from Stoke r (1998) .

Miocene and older strata (Fig . 7) . It is envisaged that the unconformity here formed a s a flat-lying erosion surfac e o n th e oute r shelf , whic h wa s tilted seawar d befor e th e depositio n o f th e overlying progradin g wedge . I n th e Faeroe -

Shetland Channel , th e unconformit y i s a planar to irregular , angula r erosio n surface . A t th e narrow S W end of the Channel, the eroded top of a sequenc e o f Miocen e sediment-drif t deposit s infilling a palaeo-erosiona l dee p i s clearl y

Fig. 4. BGS airgun profile 92/01-56 extends across the lower part of the Barra Fan and into the Rockall Trough and shows the early Pliocene unconformity (CIO) truncating the underlying Miocene (t o lower Pliocene) strata , and the lateral relationship between the Pliocene-Holocene shelf-margin and basinal deposits. Both insets (a) and (b) highlight the erosive nature of CIO; inset (b) also illustrates the interdigitating nature of the debris-flow and deep-marine deposits, and indicates that the lowest debris-flow package is underlain by a thin section of basinal sedimen t at its distal edge. Cenozoic reflector notation, C10-C30, is from Stoker et al. (2001). TWTT, two-way travel time; SBM, sea-bed multiple. Location of profile i s show n in Fig. 1 (see also Fig. 2).

Fig. 5. (a) Interpreted geoscismic section of BGS airgum profile 92/01-24 across the western flank of the Rockall Trough, Inset expanded in (b) shows detail of erosional sea

bed, the relict Miocene (to lower Pliocene) sediment drift, an d locally exposed Eocene strata , (c) The BGS sparker profile 92/01-35 , locate d S W of profile 24 , shows the early Pliocen e unconformit y (CIO ) preserve d beneath a drape o f Pliocene-Holocene sediments , togethe r wit h the locatio n o f BG S borehole 94.1 . Cenozoi c reflecto r notation, C10-C30, from Stoke r et al (2001) . Locations of profiles ar e shown in Fig. 1 (see also Fig. 2 for (a)). TWTT, two-way travel time.

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Fig. 6 . Seismic-stratigraphic setting of the middle to upper Cenozoic succession in the Hebrides-Rockall region. (a) Fugro-Geotea m commercia l seismi c profil e WRM96-11 5 fro m th e wester n flan k o f th e Rockal l Troug h showing the basinal successio n calibrate d t o ODP site 981 (modified from Jansen et al. 1996; Stoker et al. 2001). (b) BG S sparke r profil e 84/06-1 7 an d interprete d lin e drawin g acros s th e Hebride s Slop e showin g th e shelf margin successio n calibrated t o BGS borehole 88/7,7 A (modifie d fro m Stoke r et al 1994 , 2001) . Location s of profiles ar e shown in Fig. 1. TWTT, two-wa y travel time; GE, Geikie Escarpment .

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Fig. 7. BG S airgu n profil e 83/04-2 9 an d interprete d lin e drawin g fro m th e Wes t Shetlan d Shel f showin g th e seaward-tilted, shelf-margi n expressio n o f the early Pliocene unconformit y (INU) an d the progradational build out o f the overlying Pliocene-Holocen e megasequence. Th e latter display s severa l lowstan d (LST)-highstan d (HST) couplet s withi n a n overal l third-orde r composit e lowstan d system s tract . Reflecto r notatio n an d stratigraphic data based on Stoker (1999). Location o f profile is shown in Fig. 1 (see also Fig. 2). TWTT, two-way travel time . observed i n Fig . 8 . Moreover , th e reduce d thickness o f Miocene strat a on the flank s o f this deep i s als o partl y attributabl e t o erosio n associated wit h the unconformity. Farther t o the NE, Davies et al (2001 ) have described a similar erosional relationship . A t th e N W en d o f th e

profile i n Fig . 8 , thi s unconformit y i s locall y truncated by the se a bed, an d the erosional dee p (part o f the Judd Deeps) tha t here remains ope n represents a composit e erosio n surfac e tha t ha s been eroding into the underlying Paleogene strata throughout th e Neogen e interva l (Stoke r 1999) .

Fig. 8. BGS profile 83/04-32 from the West Shetland Slop e int o the Faeroe-Shetland Channel showin g the early Pliocene unconformit y (INU ) truncating th e underlyin g Miocen e sediment-drif t deposits , bu t bein g truncate d itself b y th e present-da y se a be d tha t form s a composit e erosio n surfac e i n th e are a o f th e Jud d Deeps . Stratigraphic dat a and notation are from Stoke r (1999) : LOEMU, lates t Oligocene-early Miocene unconformity . Location o f profile is show n in Fig. 1. TWTT, two-wa y travel time .

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East o f th e Faero e Islands , significan t erosio n also appear s t o have occurred a t the bas e o f th e East Faeroes wedg e (Andersen et al. 2000) . Age o f formation. I n th e Rockal l Trough , th e CIO reflecto r correlate s wit h reflecto r R 3 o f Jansen e t al (1996 ) a t Ocea n Drillin g Progra m (ODP) site 981 (Stoker et al 2001 ) (Fig . 6a). At this site , th e unconformit y occur s a t c . 271 m below se a bed (bsb) (Janse n et al 1996) . Several important planktoni c foraminife r datu m level s (first occurrence , FO ; las t occurrence , LO ) bracket thi s depth . Abov e th e unconformity , between 207. 8 an d 215.63 m bsb , th e L O o f Globorotalia cf . crassula (3. 3 Ma) i s recorded , and th e F O o f Globorotalia puncticulata (4.5 Ma) occur s jus t belo w th e unconformit y between 277. 7 an d 279.2 m bs b (Flowe r 1999) . These datum levels fall within the early Pliocen e Gr. puncticulata Zone , c. 4.2-3.2 Ma, of Weaver & Clement (1986) . Additionally , the CNll-12a (Okada & Bukr y 1980 ) o r NN15-1 6 (Martin i 1971) calcareous nannoplankto n biozone boundary, date d at 3.8 Ma by Berggren et al (1995) , is recorded a t about 220 m bsb, and the FOs of the nannofossil Pseudoemiliania lacunosa (3. 7 Ma) and the diatom Thalassiosira convexa (3.8 5 Ma) are recorde d a t depth s o f abou t 244 m bs b an d 244-272 m bsb , respectivel y (Jansen e t a l 1996). Thes e dat a sugges t a n early Pliocene ag e for th e unconformity , between 3.8 5 an d 4.5 Ma. This i s consisten t wit h biostratigraphi c infor mation fro m Britis h Geologica l Surve y (BGS ) borehole 94/ 1 tha t penetrate d reflecto r CI O farther nort h (Fig . 5c) , an d prove d a mid Miocene t o early Pliocen e (NN10-1 2 to NN15, about 9- 4 Ma) ag e rang e fo r th e sediment s immediately underlyin g th e unconformit y (Stoker e t a l 2001) . I t i s als o supporte d b y biostratigraphic dat a from well 214/4-1 in the NE Faeroe-Shetland Channel , whic h indicat e a n intra-early Pliocen e ag e fo r th e IN U reflecto r (Davies et al 2001) . On the Hebrides Slope , BGS borehole 88/7,7 A penetrated th e entir e preserve d Pliocen e shelf margin sectio n o n th e uppe r Hebride s Slop e (Fig. 6b), between th e Barra and Sula Sgeir fan s (Stoker e t a l 1994) . Althoug h th e bul k o f th e Pliocene section , betwee n 80. 5 an d 89.0 m bsb , is o f lat e Pliocene age , th e basa l 0.25 m (89.0-89.25 m) comprised a glauconite-rich lag gravel interprete d t o overli e th e unconformity , and date d by Foraminifera t o the earl y Pliocen e Gr. puncticulata Zone . Additionally , the benthonic species Uvigerina venusta saxonica was also present in this section; its LO in the North Sea is at the top o f the benthonic zon e NSB13 of King

(1989) date d a t c . 3. 8 Ma. Th e sediment s immediately belo w th e unconformity a t this sit e are dated a s late Miocene (abou t 9.5 Ma) in age; however, th e seismi c dat a indicat e tha t younger strata subcroppin g th e unconformit y ar e pre served furthe r upslop e fro m th e sit e (Fig . 6b) . On th e Wes t Shetlan d Shelf , th e ag e o f th e unconformity is less well constrained an d a more general lates t Miocene-earl y Pliocen e ag e ha s been inferre d fro m a regiona l stud y o f BG S boreholes an d commercia l well s (Stoke r 1999) . However, evidenc e fro m BG S borehol e 77/ 9 (Fig. 1) , which penetrate d th e proxima l par t o f the Wes t Shetlan d wedge , doe s indicat e th e presence o f th e earl y Pliocen e dinoflagellat e cysts Amiculosphaera umbracula an d Spiniferites splendidus an d th e planktoni c foraminife r Globorotalia crassaformis a t th e bas e (61.5 m bsb) o f the Pliocen e section . Thes e sugges t that the sediment s immediatel y abov e th e unconformity ar e unlikel y t o b e younge r tha n biozon e NN13-14, whic h implie s a n upper ag e limi t of c. 4.2-4. 0 Ma (Berggre n e t a l 1995) . Th e sediments immediatel y belo w th e unconformit y are o f lat e Miocen e age , date d o n th e basi s o f 87 Sr/86Sr between 5. 5 and 8 Ma (Stoker 1999) . Clearly, th e resolution o f the availabl e dat a is variable an d to som e exten t dependen t upo n the geological setting . The most precise dating is the intra-early Pliocen e ag e fro m th e Rockal l Trough, supporte d b y th e Faeroe-Shetlan d Channel data , wherea s i n th e generall y mor e erosional settin g o f th e shel f margi n a lates t Miocene-early Pliocen e ag e i s a t bes t defined . On th e basi s o f thes e data , th e unconformit y is here assigne d a n earl y Pliocen e age . Althoug h the possibility of some diachroneity spanning the latest Miocene-earl y Pliocen e (c . 5.5-3.8Ma ) interval canno t b e discounted , i t shoul d b e recognized that the cored sediment s subcroppin g the unconformit y d o no t necessaril y represen t the younges t pre-unconformit y strat a o n th e Hebrides an d Wes t Shetlan d shelves . Never theless, th e seawar d tiltin g o f th e unconformit y on the shelf margins must have occurred between latest Miocene an d early Pliocen e time . Prograding wedges Distribution, thickness and architecture. Alon g the easter n margi n o f th e Rockal l Trough , th e discrete depocentre s o f th e Barra-Donega l an d Sula Sgeir fans (Fig. 2b and d) form the largest of the progradin g sequence s of f N W Britain , ranging i n thicknes s fro m 30 0 t o 800m s TWTT. Wes t o f Shetland , th e Wes t Shetlan d wedge (Fig . 2a ) forms a more linea r depocentr e up t o abou t 30 0 ms TWTT thick , borderin g th e

LATE NEOCENE , UK ATLANTIC MARGIN

eastern margin of the Faeroe-Shetland Channel. Through thei r development , th e shel f brea k of f NW Britain has advanced locally by up to 50km throughout the late Neogene interval (Stoker et al 1993). Farther north , th e North Sea Fan (Fig. 3 ) encroaches into the UK sector but is more related to th e developmen t o f th e Norwegia n Channe l (see Sejru p e t al . 1996 ) an d i s no t considere d further i n this study. In genera l terms , th e progradin g wedge s ca n be subdivided into two sequences separated by an irregular, shelf-wide , erosiona l unconformity , which represent s th e glacia l unconformit y (Figs 2 , 6 b an d 7 ) (Stoke r e t a l 1994 ; Stoke r 1995). The development o f this boundary marks the onse t o f extensiv e shel f glaciatio n of f N W Britain (see below fo r age). The seismic-stratigraphic architecture of these wedges indicate s tha t thei r growt h patter n wa s initially restricte d t o th e oute r shel f an d uppe r slope, utilizing the accommodation spac e create d by th e seawar d tilting o f the oute r shel f (Fig s 2 and 7) . Locally , th e wedge s preserv e bot h a prograding an d an aggrading component , repre senting higher-orde r lowstand-highstan d coup lets (Fig . 7 ) (Stoke r 1999) . Althoug h th e predominant tren d i s towar d progradation , an d the developmen t o f a composit e lowstan d prograding wedge , th e occurrenc e o f thes e couplets doe s impl y som e degre e o f higher order relative sea-leve l fluctuation . Significantly , the preservation of these couplets is indicative of continued subsidenc e o f th e margi n throughou t the lat e Neogen e interval . A s th e shel f margi n developed, th e prograding wedge s extended into deep water , downlappin g olde r Miocen e (t o lower Pliocene ) sediment-drif t deposits , an d becoming intercalate d wit h th e basina l strat a (Figs 2 , 4 an d 8) . Whereas th e Sul a Sgei r Fa n clearly overlap s slightl y olde r Pliocen e basina l strata (Fig . 2b) , contemporar y basina l strat a i n the are a of the Barr a Fan may have been partl y eroded away , a s th e bul k o f th e fa n deposit s appear to rest on the early Pliocene unconformity (Figs 2 d and 4). Clos e inspectio n of the seismi c profile i n Fig. 4 (inset b) indicates that the distal edge o f th e basa l debris-flo w packag e overlie s post-unconformity basina l deposits . However , farther i n toward s th e Hebridea n Margin , an y basinal sedimentatio n i s likel y t o hav e bee n overwhelmed by th e shee r volum e o f the mass flow sedimentation. Th e upper part of the shelf margin successio n has locally been modified by slope failure , suc h a s tha t associate d wit h th e Peach Slid e o n th e Barr a Fa n (Holme s e t al . 1998). Shelf-margin progradatio n wa s no t exclusiv e to the Hebridean an d West Shetland margins . A

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significant progradin g wedge , th e Eas t Faeroe s wedge, i s preserve d of f th e Faero e Island s (Andersen et al. 2000) (Fig . 2a) , which togethe r with th e Wes t Shetlan d wedg e indicate s a symmetry t o margin constructio n i n this region. Farther south , Plio-Pleistocen e progradatio n occurs o n th e easter n flan k o f Rockal l Bank , where th e Eas t Rockal l wedg e i s develope d downlapping ont o Miocen e sediment-drif t deposits (Stoker 2002) (Figs 2e and 3). However, the Eas t Rockal l wedg e i s relativel y smal l i n comparison with the eastern flank of the Rockall Trough. Age o f th e wedges. Th e Plio-Pleistocen e succession ha s bee n teste d b y numerou s BG S boreholes an d commercial well s on the Hebride s and Wes t Shetlan d margin s (Stoke r e t al. 1993 ; Stoker 1999) , but relatively fe w of these provide the necessary detai l and resolution t o rigorousl y date th e sediments . BG S borehol e 88/7,7 A penetrated the entire Plio-Pleistocene successio n (89.25 m thick) preserve d o n the upper Hebride s Slope betwee n th e Barr a an d Sul a Sgei r fan s (Fig. 6b) . Althoug h lowe r Pliocen e strat a wer e proved i n th e basa l 0.25 m (describe d above) , biostratigraphic an d magnetostratigraphi c dat a indicate tha t th e bul k o f th e overlyin g 89.0 m section (t o se a bed ) i s o f lat e Pliocen e t o Pleistocene age , probably n o older tha n c. 3 Ma (Stoker e t al. 1994) . It was suggested by Stoke r et al . (1994 ) tha t th e lower-uppe r Pliocen e boundary may be lost in a core gap between 89. 0 and 88.4m bsb, thus the nature of this boundary remains uncertain. The gap in the ages recorde d by th e earl y an d late Pliocene strat a a t this sit e suggests that here the boundary may represent a hiatus of up to 0.8 Ma. The Pliocene-Pleistocene boundary (1.6 4 Ma) occur s a t 80.5 m bsb in th e borehole, wit h th e glacia l unconformit y pene trated a t 67.82 m bs b an d date d t o earl y mid Pleistocene (abou t 0.44 Ma) age. It is interesting to note that the influx of ice-rafted detritus (IRD) to th e Hebride s Slop e occur s a t 86.0 m bsb , coinciding mor e o r les s wit h th e Gauss Matuyama polarit y transitio n (2.4 8 Ma) a t 85.8m bsb (Stoker e t al. 1994) . On th e Wes t Shetlan d Shelf , BG S borehol e 77/9 proved lower Pliocene sediment s at the base of th e borehol e (describe d above) . However , BGS borehole s 84/ 1 an d 84/ 2 (Fig . 1) , whic h similarly penetrated th e proximal (landward ) part of th e wedge , sugges t tha t th e bul k of the West Shetland wedg e i s of lat e Pliocene-Pleistocen e age (Stoke r 1999) . I n borehol e 84/2 , th e sediments recovere d betwee n th e basa l uncon formity an d th e glacia l unconformit y ar e

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assigned a lat e Pliocene-earl y Pleistocen e ag e on th e basi s o f th e dinoflagellat e cys t species , Amiculosphaera umbracula an d Habibacysta tectata (Harland, pers. comm.). This is supported by borehol e 84/1 , wher e th e basa l par t o f th e wedge correlate s wit h th e benthi c foraminife r Cibicides grossa zone, dated to latest Plioceneearliest Pleistocen e tim e (Wilkinso n & Harlan d 1999). These dat a suggest that the main onset of deposition of the prograding wedge s occurred in late Pliocen e time . Borehol e 88/7J A furthe r suggests tha t thi s correlate s approximatel y wit h the influx o f ice-rafted materia l a t c. 2.48 Ma. Deep-water sedimentary response Style and focus of sedimentation. Contour following botto m current s hav e strongl y influ enced basina l sedimentatio n i n th e Rockal l Trough an d Faeroe-Shetland Channe l through out th e Neogen e period , an d a variet y o f sediment-drift an d associate d bedform s hav e been describe d (Kid d & Hill 1986 ; How e e t al 1994; How e 1996 ; Stoke r 1998 ; Stoke r e t a l 1998; Masson et al 2002) . However , the base of the Pliocene-Holocene megasequence indicate s that th e deep-wate r sedimentar y regim e wa s modified i n tw o ways : (1 ) b y extensiv e submarine erosio n tha t truncate d olde r drif t deposits, formin g th e basa l (earl y Pliocene ) unconformity; (2 ) by a major shif t in the focus of late Neogen e sedimentatio n relativ e t o th e underlying Miocen e deposits . Example s o f these changes ar e described below . In the Rockall Trough , the Miocene (t o lower Pliocene) sediment s characteristicall y onla p th e flanks o f th e basins , wit h elongat e mounde d sediment drift s displayin g significan t upslop e accretion (Stoke r 1998 ; Stoke r e t a l 1998 ) (Fig. 5) . Followin g earl y Pliocen e erosio n (e.g . Figs 4-6a) , sedimen t drift s continue d t o accumulate in the basin (e.g . Fig . 6a ) and alon g the easter n flan k o f th e Troug h (partl y overlapping wit h th e progradin g wedges) , bu t th e western flan k o f th e basi n (nort h o f abou t 56°30/N) has continued to be subject to erosion to the presen t da y (Stoker e t al 2001 ) (Fig s 2c- e and 3) . This intens e bottom-curren t activit y ha s restricted Pliocene-Holocen e accumulatio n i n this area to a thin drape (e.g. Fig. 5c) or a sea-bed veneer, commonl y marke d b y a gravel-la g contourite (How e e t a l 2001) , an d ha s erode d the underlyin g Miocen e drif t deposits , locall y exposing Paleogene strat a at the sea bed (Fig. 5a and b). The longevity of the erosion is confirme d where th e bas e o f th e Pliocene-Holocen e megasequence, th e CI O reflector , i s truncate d by th e se a be d formin g a composit e sea-be d

erosion surfac e (Fig . 5) . I n compariso n wit h Miocene (t o lower Pliocene) sedimentation , thi s represents a major eastwar d (basinward ) shift i n the accumulatio n o f sediment-drif t deposit s o n the wes t an d N W flank of th e norther n Rockal l Trough durin g Plio-Pleistocen e time . Farthe r south, i n contrast , basina l sedimentatio n alon g the wester n flan k o f th e centra l an d souther n Rockall Troug h prevaile d followin g earl y Plio cene erosion , an d sedimen t drifts , suc h a s th e giant elongat e Fen i Drif t (Fig s 1 an d 3) , continued to aggrade (Stoke r et al 2001) . In compariso n wit h th e Rockal l Trough , th e Faeroe-Shetland Channe l ha s generall y bee n more an area of sediment export, resulting in the preservation o f a thinne r Neogen e successio n (Stoker et al 1998 ; Stoke r 1999) . Nevertheless , the earl y Pliocen e unconformit y i s strongl y expressed, especiall y a t th e S W en d o f th e Channel, wher e formerl y extensive , thick , Miocene sedimen t drift s hav e bee n erode d an d locally remove d (Fig s 2 a an d 8) . Sinc e earl y Pliocene time , thi s par t o f th e channe l ha s remained largel y a n are a o f prolonge d erosio n (Fig. 3), characterized by a thin to locally absent sediment cover . Th e mai n focu s o f Pliocene Holocene sediment-drif t accumulatio n i n th e Faeroe-Shetland Channel occur s north of about 60°307N; here th e basina l sediment s increas e i n thickness an d onlap onto the lower-middle part of th e Wes t Shetlan d Slop e wher e the y overla p with th e Wes t Shetlan d wedge . Anderse n e t a l (2000) reporte d a simila r relationshi p fro m th e East Faeroes Slope . Rate o f response. O n th e basi s o f biostrati graphic data and sedimentation rates, Jansen et al (1996) conclude d tha t OD P sit e 98 1 doe s no t contain an y significan t hiatuses . However , th e FOs of T. convexa (3.85 Ma) and Gr. puncticulata (4.5 Ma) wer e recorde d b y Janse n e t a l (1996 ) immediately abov e an d below , respectively , th e unconformity, whic h doe s sugges t a hiatu s o f about 0.7 5 Ma, althoug h th e proble m o f bios tratigraphic resolutio n i s highlighte d b y thei r placing o f the CNll-12a nannofossi l boundary (3.8 Ma) about 50 m above the unconformity (see above). Nevertheless , thes e dat a sugges t tha t there i s a relativel y continuou s recor d o f lat e early Pliocen e t o Holocen e (pos t 3.8 5 Ma) sedimentation abov e th e unconformity . Thi s represents a n earlie r sedimentar y respons e than tha t prove d fo r th e progradin g wedge s (3 Ma o r younger). Moreover, th e influ x o f IRD at OD P site 981 is recorded a t a depth o f abou t 138m bsb , whic h i s jus t abov e th e L O o f th e planktonic foraminife r Neogloboquadrina

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atlantica (2.4 1 Ma), an d abou t 133 m abov e th e unconformity (Janse n et al 1996) . The thickness of th e pre-IR D post-unconformity sediment s clearly contrast s wit h th e progradin g wedges , which preserve a fa r mor e condense d recor d o f equivalent strata.

(1) The early Pliocen e event . This resulte d i n the formation of the early Pliocene unconformity between abou t 3. 8 an d 4. 5 Ma, althoug h som e diachroneity fro m lates t Miocen e tim e (abou t 5.5 Ma) cannot be discounted o n the basis of the shelf-margin record . (2) The early Pliocene (pos t 3.8 Ma) sedimen tary response. This is marked by the continuatio n Late Neogene development of the UK Atlantic of sediment-drift accumulation in the deep-water basins, albei t wit h a shif t i n th e focu s o f margin: a summary event stratigraphy sedimentation i n th e norther n Rockal l Troug h From the information presented in the preceding and S W Faeroe-Shetlan d Channel . A mor e sections, th e ke y element s o f th e lat e Neogen e restricted record o f early Pliocene sedimentatio n event stratigraph y of f N W Britai n ar e sum- is preserve d o n th e shel f margin s at th e bas e o f marized i n Fig. 9 , and further encapsulate d a s a the prograding wedges . (3) Th e lat e Pliocene-earl y mid-Pleistocen e gross palaeo-environmenta l reconstructio n i n Fig. 3 . I n genera l terms , a four-stage histor y of sedimentary response . Thi s interva l wa s domi development can be established that is applicable nated by the deposition o f the majo r progradin g wedges betwee n abou t 3 an d 0.4 4 Ma. I n th e to the entire U K Atlantic margin , as follows .

Fig. 9. Schematic diagram summarizing the late Neogene event stratigraphy and development of the UK Atlantic margin (see text for details).

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adjacent basins , deep-wate r sedimentatio n pre vailed an d i n area s wher e downslop e an d alongslope processe s interacte d overlappin g slope apron-sediment-drift system s developed . (4) Mid - t o lat e Pleistocen e shelf-wid e glaciation. Ic e sheet s reache d th e edg e o f th e Hebrides an d Wes t Shetlan d shelve s betwee n 0.44 an d c . 0.018 Ma an d furthe r contribute d t o shelf-margin progradatio n (Stoke r 1995) . Deep water sedimentatio n prevaile d i n th e adjacen t basins, and bottom-current processe s continu e to be active at the present day (Howe 1996; Masson et al 2002) . Discussion The result s presente d fro m th e U K Atlanti c margin rais e severa l interestin g aspect s o f broader relevanc e concernin g th e timin g an d mechanism o f chang e throughou t th e N E Atlantic region . Timing of change Perhaps on e o f th e mos t significan t aspect s arising fro m thi s stud y concerns the questio n of cause and effect an d its bearing o n the timing of change within late Neogene time. In general, the sedimentary record o f shelf-margin progradatio n around th e N E Atlanti c region , includin g N W Britain, appear s t o indicat e tha t i t essentiall y dates fro m lat e Pliocen e time . Fro m a compre hensive stud y o f th e Norwegia n continenta l margin, Eidvin et al. (2000) concluded that there was a lin k betwee n shelf-margi n progradatio n and th e intensificatio n o f norther n latitud e glaciation. Althoug h thi s climaticall y relate d thesis is supported to some extent by the present study, th e questio n remain s a s t o whethe r thi s linkage i s simpl y a response t o (a n effec t of ) a more fundamenta l causal event , e.g . uplift . Th e prograding wedge s effectivel y preserv e th e record o f denudation , b e tha t cause d b y glacia l processes o r otherwise , an d d o no t necessaril y represent a n immediat e respons e t o change . A s denudation rat e i s largel y a functio n o f loca l relief, whic h i n tur n i s relate d t o th e degre e o f fluvial dissection , ther e i s a n inheren t tendency for th e rat e o f denudatio n t o b e initiall y slowe r than the rate of uplift (Summerfiel d 1991) . On the basis of the results of the present study, it is here suggested that the late Neogene change was initiated befor e the deposition o f the bulk of the prograding wedges. The indicator o f change is manifes t b y th e regiona l earl y Pliocen e unconformity, whic h ca n b e trace d unambigu ously fro m th e shel f margi n into the deep-wate r basins (Fig . 2) . Biostratigraphi c dat a ar e

consistent fro m bot h setting s an d indicate lowe r Pliocene sediment s immediatel y overlyin g th e unconformity (Fig . 9) . A major question that cannot be resolved fro m the available dat a is whether or not there is some degree o f diachroneit y i n th e developmen t o f the unconformit y betwee n th e shel f margi n an d deep-water basin . O n th e shel f margin , th e youngest core d strat a fro m belo w th e unconfor mity are of latest Miocene (c. 5.5 Ma) age, which contrasts wit h the basinal sediments , wher e earl y Pliocene (4. 5 Ma) deposit s mak e u p th e sub unconformity section (Fig . 9). This disparity may suggest diachroneit y o f th e orde r o f 1 Ma, possibly extendin g th e onse t o f chang e o n th e shelf margi n int o lates t Miocen e time , befor e culminating i n th e margin-wid e earl y Pliocen e event. Furthe r samplin g i s necessar y t o bette r constrain th e timin g o f chang e o n th e shel f margin. Mechanics of change The mechanism of change in late Neogene tim e is a proble m o f ver y genera l interes t a s i t ha s affected a larg e proportio n o f th e landmasse s around th e Nort h Atlantic . I n a recen t revie w paper, Japse n & Chalmer s (2000 ) summarize d the curren t thinkin g o n Neogen e evolutio n around th e Nort h Atlanti c tha t favour s som e kind of tectonic event , which resulted i n uplift of basin margin s an d subsidenc e o f basina l area s adjacent to the uplifted landmasses. However , the mechanism responsible fo r uplift remain s uncertain, and there is a school of thought that suggests that th e so-calle d 'Neogen e uplift ' ma y i n fac t have begu n earlie r i n Paleogen e tim e (Clause n et a l 2000) . Wha t i s clea r i s tha t an y genera l model invoke d t o tr y an d explai n Cenozoi c uplift, includin g bot h Paleogen e an d Neogen e uplift, mus t be constrained b y observations fro m the entire NE Atlantic region . On th e U K Atlanti c margin , regiona l obser vations tha t mus t b e take n int o consideratio n include (1 ) the seaward tilting an d subsidence of the outer part of the Hebrides an d West Shetland margins a t som e tim e betwee n lates t Miocen e and earl y Pliocen e time , (2 ) th e earl y Pliocen e change i n the oceanographic circulatio n pattern , and (3 ) the predominantly late Pliocene onse t of shelf-margin progradation . Th e regiona l distri bution o f th e progradin g wedge s i s a furthe r consideration, a s west to NW progradation o f the NW Britis h margi n wa s accompanie d b y S E progradation fro m th e Faeroe Islands (Anderse n et a l 2000 ) an d the Rockall Ban k (Fig. 3) . One might argue whether o r not uplift i s necessary t o achieve thi s distributio n o f wedges ; instead , i t

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may b e achieve d b y th e eustatic lowering o f sea level, fo r whic h ther e i s evidenc e o f significant lowering fro m c . 4. 1 Ma (Vai l & Hardenbo l 1979). Presumably, thi s would also have affected the oceanographi c circulatio n pattern. However , eustacy alone does not explain either the seaward tilting of the continental margin off NW Britain, nor th e timin g o f shelf-margi n progradation . Instead, th e continuit y o f th e earl y Pliocen e unconformity fro m th e shel f margi n int o dee p water suggest s tha t th e modificatio n t o th e oceanographic circulatio n patter n wa s inextric ably linke d t o th e proces s responsibl e fo r th e tilting o f the margin . There i s evidenc e fo r Neogen e uplif t o f th e NW Britis h hinterlan d fro m onshor e studie s i n Britain an d Ireland (e.g . Japsen 1997 ; Galewsky et al 1998 ) an d the Irish Se a Basin (Green e t a l 2001). Uplift has also been invoked to explain the prograding wedg e developmen t o n the Faeroes e margin (Anderse n e t al . 2000) . Th e conjugat e development o f wedge s i n th e Faeroe-Shetland region implie s tha t th e intervenin g basi n mus t have subsided. Similarly, the development of the East Rockal l wedge , facin g th e Hebridea n margin, suggest s tha t no t onl y ha s th e Rockal l Bank undergon e som e degre e o f uplif t con comitant wit h that o f the Hebridean margin , bu t that th e Rockal l Troug h ha s probabl y compen sated fo r thi s wit h subsidence . Th e ensuin g change in the shape and palaeobathymetry of the continental margin is also likely to have modified the water circulation pattern. The adjustment and breaching o f sill s o r othe r barrier s an d th e development o r modificatio n o f deep-wate r gateways controlling bottom-current circulation, sedimentation an d erosio n i n th e Nort h Atlantic Ocean ha s bee n linked b y severa l worker s (e.g . Tucholke & Mountai n 1986 ; Eldhol m 1990 ; Thiede & Mhyre 1996 ; Wrigh t & Miller 1996 ) to regional tectonic events. Indeed, the triggering of early Neogene deep-wate r erosion , a n event that had margin-wide implications similar to those of the earl y Pliocen e event , an d th e subsequen t widespread accumulatio n o f Miocen e (t o lowe r Pliocene) sediment-drif t deposits of f NW Britain are mos t probabl y linke d t o th e plate-tectoni c development o f th e Norwegian-Greenlan d Se a (Eldholm 1990 ; Janse n & Raym o 1996 ; Stoke r 1998). Uplift i n combinatio n wit h climati c coolin g may have ultimately bee n a trigger for the onse t of glaciatio n (Raym o & Ruddima n 1992 ; Eyle s 1996) i n lat e Pliocen e time . Glacia l processe s superimposed o n uplift ma y be the reason for the enhanced denudatio n tha t resulte d i n th e prograding wedges , thu s accounting , i n part, fo r the delaye d shelf-margi n sedimentar y respons e

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to change . Th e presenc e o f IR D withi n th e wedges support s th e concep t tha t th e influ x o f much o f thi s materia l ma y b e linke d t o th e climatic evolutio n o f th e area . Th e subsequen t isostatic respons e t o denudation , includin g climatically enhance d denudatio n a s glaciatio n intensified, an d sedimen t loading in the adjacen t basin ma y hav e continue d t o driv e vertica l movements. Although thi s explanation of th e observation s is consistent wit h the Japsen & Chalmers (2000 ) view o f landmas s an d basin-margi n uplif t an d basin-centre subsidence , it doe s no t explai n the nature o f the mechanis m responsible for th e late Neogene event . However , i t doe s highligh t the observations tha t must be incorporate d int o any future genera l mode l o f lat e Neogen e evolution of th e N E Atlanti c region. Conclusions The patter n o f observation s o n an d aroun d th e UK Atlanti c margin , includin g th e evidenc e o f tectonic tiltin g an d subsidenc e o f th e shel f margins, th e modifie d deep-wate r curren t sys tem, an d th e distributio n o f th e progradin g wedges, support the idea of a late Neogene uplif t event. Athoug h th e mechanis m remain s unknown, i t i s suggeste d tha t thi s even t wa s initiated i n earl y Pliocen e (possibl y eve n lates t Miocene) tim e and has been a major influenc e in the subsequen t development of th e U K Atlantic margin. On this basis, it is proposed tha t the late Neogene even t stratigraph y reflect s a varie d sedimentary respons e t o uplift . Th e deep-wate r basins, bein g ope n t o th e N E Atlantic , wer e sensitive t o change s i n oceanographi c circula tion, an d th e earl y Pliocen e even t ha d a significant, an d possibl y relativel y sudde n (o f the orde r o f 10-100ka) , effec t o n pattern s o f deep-marine sedimentatio n an d erosion . I n contrast, i t i s suggeste d tha t th e shelf-margi n record reflect s som e dela y i n th e sedimentar y response, a s the prograding wedges do not begin to full y develo p unti l perhap s 1- 2 Ma afte r th e initiation o f uplift , albei t enhance d b y glacia l processes. I would like to thank D. Evans, K . Hitchen, and the two referees, P . Knut z an d J . Cartwright , fo r thei r constructive review s o f th e paper . I a m gratefu l t o I. Walke r (Conoco ) fo r providin g copie s o f seismi c lines DG95-8 and -10 ; P. Broad o f Fugro Geoteam fo r allowing access to, and use of data from, the WRM96 survey; and the following oil companies, who , together with th e BGS , mak e u p th e Rockal l Continenta l Margin Consortium , an d withou t whos e suppor t thi s work could no t have been undertaken : Agip , Amerad a

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Hess, BG , BP-Amoco , Conoco , Enterprise , Exxon / Mobil, Phillips, Statoil , Texaco an d TotalFinaElf. The paper i s published with the permission o f the Directo r of the British Geological Surve y (NERC).

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LATE NEOCENE , UK ATLANTIC MARGI N and contourite s i n th e norther n Rockal l Trough . Marine Geology, i n press . OKADA, H . & BUKRY , D . 1980 . Supplementar y modification an d introductio n o f cod e number s t o the low-latitude coccolith biostratigraphi c zonation (Bukry 1973 ; 1975) . Marine Micropalaeontolog\, 5,321-325. RAYMO, M.E . & RUDDIMAN , W.F . 1992 . Tectoni c forcing o f lat e Cenozoi c climate . Nature, 359 , 117-122. ROKOENGEN, K. , RlSE , L. , BRYN , P., pRENGSTAD , B. ,

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Shelves, an d Adjacent Deep-Water Areas. Britis h Geological Survey . HMSO, London . STOKER, M.S. , LESLIE, A.B., SCOTT, W.D. & 6 OTHERS 1994. A recor d o f lat e Cenozoi c stratigraphy , sedimentation an d climat e chang e fro m th e Hebrides Slope , N E Atlanti c Ocean . Journal o f the Geological Society, London, 151 , 235-249 . STOKER, M.S. , VAN WEERING, T.C.E. & SVAERDBORG , T. 2001 . A mid - t o lat e Cenozoi c tectonostrati graphic framewor k fo r th e Rockal l Trough . In : SHANNON, P.M. , HAUGHTON , P . & CORCORAN , D . (eds) Th e Petroleum Exploration o f Ireland's Offshore Basins. Geologica l Society , London , Special Publications , 188 , 411-438 . STOKER, M.S. , NIELSEN , T, VA N W E E R I N G, T.C.E. & KUIJPERS, A . 2002 . Toward s a n understandin g of the Neogene tectonostratigraphic framework of the NE Atlanti c margin between Ireland and the Faro e Islands. Marine Geology, i n press. S U M M E R F I E L D , M.A . 1991 . Global Geomorphology: an Introduction to the Study of Landforms. Longman, Harlow . THIEDE, J . & MHYRE , A.M . 1996 . Introductio n to th e North Atlanti c gateways : plat e tectonic-palaeo ceanographic history and significance . In: THIEDE , J., MHYRE , A.M. , FIRTH , J.V. , JOHNSON , G.L . & RUDDIMAN, W.F . (eds ) Proceedings o f th e Ocean Drilling Program, Scientific Results, 151. Ocea n Drilling Program , College Station , TX, 3-23 . TUCHOLKE, B.E . & MOUNTAIN , G.S . 1986 . Tertiar y palaeoceanography o f th e wester n Nort h Atlanti c Ocean. ///: VOGT, PR. & TUCHOLKE , B.E. (eds) The Geology of North America, Volume M, The Western North Atlantic Region. Geologica l Societ y o f America, Boulder , CO, 631-65C K VAIL, PR . & HARDENBOL , J . 1979 . Sea-leve l changes during th e Tertiary. Oceanus, 22, 245-276. VORREN, TO . & LABERG , J.S . 1997 . Troug h mout h fans—palaeoclimate an d ic e shee t monitors . Quaternary Science Reviews, 16, 865-881. WEAVER, P.P.E. & CLEMENT , B.M. 1986 . Synchroneity of Pliocen e planktonic foraminiferal datum s i n th e North Atlantic . Marine Micropalaeontolog\, 10 , 295-307. W I L K I N S O N , I.P . & HARLAND , R . 1999 . Appendi x 1 : Mid- to late Cenozoic biostratigraphic markers. In: STOKER, M.S . (ed. ) Stratigraphic Nomenclature o f the UK North West Margin: 3. Mid- to Late Cenozoic Stratigraphy. Britis h Geologica l Survey. Edinburgh, A1-A6. WRIGHT, J.D . & MILLER , K.G . 1996 . Contro l o f North Atlantic Deep Water circulation by the GreenlandScotland Ridge . Paleoceanography, 11 , 157-170 .

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Quantifying exhumatio n from apatite fission-track analysis and vitrinite reflectance data: precision, accuracy and latest results from the Atlantic margin of NW Europe PAUL F. GREEN, IAN R. BUDDY & KERRY A. HEGARTY Geotrack International Pty Ltd, 37 Melville Road, Brunswick West, Vic. 3055, Australia (e-mail: mail@ geotrack.com.au) Abstract: I n area s wher e significan t unconformitie s ar e present , palaeotemperature s derived fro m apatit e fission-trac k analysi s (AFTA ) an d vitrinit e reflectanc e (VR ) dat a through a vertical rock sectio n ca n be used to estimate palaeogeotherma l gradients and (by extrapolation t o a n assume d palaeo-sur f ace temperature ) amount s o f exhumation (palaeo burial). AFTA also provides a direct estimate of the timing of exhumation. These parameter s can be used to reconstruct more complet e historie s tha n those based purel y on the preserved rock record. Precision an d accurac y o f thes e estimate s ar e controlle d b y a rang e o f theoretica l an d practical factors, perhaps the most important bein g the use of appropriate kineti c models. In extracting therma l histor y informatio n fro m fissio n track s i n apatite , i t i s essentia l t o us e models tha t can describe variation i n response betwee n apatit e grain s withi n a sample. I t is also important t o recognize th e limitations o f the methods. AFT A and VR are dominated by maximum temperatures , preservin g n o informatio n o n event s prio r t o a palaeo-therma l maximum. Recognitio n o f thi s allow s definitio n of ke y aspect s o f the histor y with greate r precision. Results fro m N W Europ e defin e a serie s o f regionall y synchronou s palaeo-therma l episodes, wit h coolin g beginnin g i n Earl y Cretaceous , Earl y Tertiar y an d Lat e Tertiar y times. Latest result s sho w that Early Tertiar y palaeo-therma l effect s i n NW England can be understood a s being du e to a combination o f higher basa l hea t flow and deeper burial, and emphasize th e importanc e o f obtainin g dat a fro m a vertica l sequenc e o f samples . Comparison with similar result s fro m othe r part s o f the world suggest s that events at plate margins exer t a ke y influenc e o n th e processe s responsibl e fo r regiona l exhumation , as recognized throug h Mesozoi c an d Cenozoic time s acros s N W Europe .

Over th e las t 3 0 years, th e importanc e o f etc . (e.g . Hilli s 1992 ; Brodi e & Whit e 1994) . exhumation i n th e sedimentar y basin s an d Quantificatio n o f th e effect s o f exhumatio n i s basement terrain s o f N W Europ e ha s bee n therefor e importan t fo r man y reasons , an d a increasingly recognized . Here we us e 'exhuma - variet y o f technique s hav e bee n applie d t o tion' t o describe the process by which rock units quantif y variou s aspect s o f exhumatio n (see, that were once more deeply buried are brought to e.g . Rii s & Jense n 1992) . Her e w e focu s o n shallower depth s a s a resul t o f remova l o f palaeo-therma l methods , i.e . thos e base d o n overlying rocks . Althoug h debat e continue s increas e o f temperatur e with depth , specificall y concerning precis e definition s (e.g . Englan d & withi n th e contex t o f sedimentar y basins , Molnar 1990) , mos t worker s appea r t o b e althoug h th e principle s employe d ar e equall y comfortable wit h thi s interpretatio n o f th e applicabl e in basemen t terrains. We revie w the word, use d essentiall y to represen t th e revers e natur e o f th e informatio n obtaine d fro m thes e of 'burial' . methods , discus s th e variou s practica l factor s The proces s o f exhumatio n play s a n tha t affec t precisio n and accurac y i n estimating important rol e i n petroleu m system s o f man y bot h th e magnitud e and timin g of exhumation , regions o f N W Europ e (e.g . Gree n et al 1997 ; an d finall y revie w lates t result s fro m th e Duncan e t al . 1998 ; Dor e e t al . 1999) . Atlanti c margi n o f N W Europe , includin g Recognition o f th e proces s raise s man y compariso n with other areas and some speculation questions i n term s o f tectoni c mechanisms , o n possible mechanisms. From: DORE , A.G. , CARTWRIGHT, J.A. , STOKER , M.S. , TURNER, J.P. & WHITE , N . 2002 . Exhumation of th e North Atlantic Margin: Timing, Mechanisms and Implications for Petroleum Exploration. Geological Society, London, Special Publications, 196, 331-354 . 0305-8719/027$ 15.00 © The Geological Society of London 2002.

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Quantifying th e magnitude o f exhumatio n using palaeo-thermal methods Because o f the progressiv e increas e i n tempera ture wit h dept h withi n th e lithosphere , palaeo thermal indicator s suc h a s apatit e fission-trac k analysis (AFTA®) and vitrinite reflectance (VR), which provid e estimate s o f th e maximu m temperature attaine d b y a roc k sampl e a t som e time in the past, can also be used to assess former burial depths . Sedimentar y unit s ar e progress ively heated as they are buried, and begin to cool at th e initiatio n o f exhumation . AFT A an d V R provide quantitative estimates of the temperature of individua l sample s a t th e palaeo-therma l maximum, immediatel y befor e th e onse t o f cooling (a s explained , e.g . by Bra y e t al 1992 ; Duddy e t al 1994 ; Green e t al. 1995) . Wherea s VR value s provid e discret e estimate s o f th e maximum post-depositiona l palaeotemperature , AFTA ma y provid e eithe r lowe r o r uppe r limits or a rang e o f value s fo r th e maximu m palaeotemperature i n one , tw o o r rarel y thre e separate episodes (fo r more details, see e.g. Bray etal 1992) . No othe r palaeo-therma l indicator s ar e currently understoo d i n sufficien t detai l t o allo w quantitative estimatio n o f maximu m palaeotem peratures. Fo r som e techniques , publishe d conversions t o equivalen t V R values , e.g . Thermal Alteratio n Inde x (TAI) , Conodon t Alteration Inde x (CAI ) an d 7 max (Waple s 1985), Spor e Colou r Inde x (SCI ) (Fisher e t al . 1980), allo w estimatio n o f palaeotemperature s using th e kinetic s o f V R response , althoug h uncertainties in these calibrations may introduce additional error . As described in detail by Bray et al. (1992) , a series o f palaeotemperatur e estimate s ove r a range o f depth s i n a wel l o r borehol e (o r elevations o f outcro p sample s i n mountainou s terrains) allow s determinatio n o f th e palaeo geothermal gradient . Extrapolatio n o f th e fitte d palaeotemperature profile t o an assumed palaeo surface temperature then provides a n estimate of the amount by which the sectio n wa s once mor e deeply buried, i.e. the amount of section removed during exhumatio n (Fig . 1) . (I n th e absenc e o f direct palaeogeotherma l gradien t constraints , a range o f realistic value s may be assumed. ) A palaeotemperature profil e ca n be character ized b y a singl e valu e o f palaeogeotherma l gradient only when the profile i s linear (fro m th e surface t o th e bas e o f th e section) ; i n situations where palaeotemperatur e profile s ar e markedl y non-linear (se e below), th e analysis illustrated i n Fig. 1 is no t valid . However , i t i s importan t t o recognize tha t palaeotemperature s determine d

from AFT A and/o r V R in individual samples ar e independent o f thes e considerations , an d ca n therefore b e use d t o constrai n possibl e therma l history model s regardles s o f an y assumption s about th e for m o f the palaeotemperatur e profile .

Factors affecting th e accuracy of the magnitude of exhumatio n Basic assumptions The analysi s i n Fig. 1 depends criticall y o n th e assumptions tha t th e palaeotemperature-dept h profile is linear through the preserved sectio n and that th e profil e ca n b e linearl y extrapolate d through th e remove d sectio n t o th e assume d palaeo-surf ace temperature . Explici t estimatio n of removed section from palaeo-thermal methods requires such assumptions to reduce the proble m to a leve l wher e forma l estimatio n o f palaeo geothermal gradient s an d amount s o f exhuma tion o r remove d sectio n i s possible . Mor e importantly, thes e assumption s als o allo w determination o f th e associate d uncertaintie s (±95% confidenc e limits) , thereb y allowin g rigorous and objective assessment of the range of scenarios tha t ar e consistent with the data .

Fig. 1. Where heating is due to deeper burial , possibly combined wit h elevate d hea t flow , amount s o f exhumation (o r deepe r burial ) ca n b e estimate d b y fitting a linear palaeogeothermal gradien t to a series of down-hole palaeotemperatur e constraint s an d the n extrapolating t o a n assume d palaeo-surfac e tempera ture, a s shown . Thi s approac h depend s criticall y o n certain assumptions , as discussed i n the text.

QUANTIFYING EXHUMATIO N USIN G AFT A AN D VR

Although thi s approac h ha s been criticize d a s over-simplified (e.g . Hollida y 1993 ; Smit h e t al. 1994), result s i n well-controlle d situation s ar e usually highl y consisten t wit h estimate s o f former buria l depth s fro m othe r sources . Fo r example, i n th e Fresne- 1 wel l i n th e Taranak i Basin o f Ne w Zealan d (Kam p & Gree n 1990) , extrapolation o f a linea r profil e fitte d t o palaeotemperature constraint s fro m AFT A an d VR dat a give s a n estimat e o f sectio n remove d during Lat e Miocen e basi n inversio n tha t i s highly consisten t wit h value s derive d fro m extrapolation o f truncate d seismi c reflector s into the inversio n structur e (Gree n e t al. 1995) . In detail , o f course , th e variatio n o f tempera ture with depth is not exactly linear , but depend s on the variatio n o f thermal conductivit y through the section, which is related, in turn, to variations in lithology (see , e.g . Gretene r 1981 , Fig. 2.8-1) . Practical experienc e show s tha t i n mos t situ ations, a linea r approximatio n i s reasonabl e (e.g. Demin g 1994 , Fig . 9.4) , probabl y becaus e

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small-scale variation s i n litholog y serv e t o blu r any loca l therma l conductivit y contrast s t o produce a broadly linea r variation of temperature with depth . I n addition , th e accurac y o f typical bottom hol e temperatur e (BHT ) value s (usuall y the onl y availabl e contro l o n present-da y temperatures) i s often s o poor tha t their detailed variation wit h dept h ma y hav e mor e t o d o with recording practice , rathe r tha n rea l therma l structure withi n a sedimentar y section . Wit h typical precision s o n palaeotemperatur e esti mates o f th e orde r o f 10° C (se e discussio n below), typica l variatio n i n tru e temperature s about a linear profile usuall y means that a linear approximation introduce s negligibl e additiona l error t o the treatment. In summary , although the approach illustrated in Fig . 1 alway s represent s som e degre e o f approximation t o th e tru e situation , i t ha s th e advantage o f providing an objectiv e assessment, with result s constraine d b y th e measure d dat a and simplifyin g assumption s explicitly stated. In situations wher e th e approximatio n o f linea r gradients i s though t t o b e inappropriate , th e results of this approach ma y stil l provide a 'first pass' assessmen t o f the situation, on which more detailed treatment s base d o n hea t flo w (se e discussion below ) ca n be based . Non-linear palaeogeothermal gradients

Fig. 2 . This plot illustrates the influence o f the thermal conductivity o f th e remove d sectio n o n th e natur e o f the palaeotemperatur e profil e throug h tha t par t o f th e section. Onl y wher e th e remove d an d preserve d sections ar e identica l wil l th e therma l gradien t b e th e same throughout the entire section, but in practice the assumption of linearity appear s to give reliable results (see text) .

It situations wher e th e palaeotemperature profil e is obviousl y non-linear , th e amoun t o f sectio n removed canno t be estimated usin g the construction show n i n Fig . 1 . Non-linear palaeotempera ture profile s ma y b e expecte d i n tw o importan t types o f situation . First, large-scal e contrast s i n thermal conductivit y ca n caus e significan t non linearity i n th e temperatur e profil e withi n th e (removed and/o r preserved) section . This is most pronounced whe n the removed sectio n consiste d of ver y differen t lithologie s compare d wit h th e preserved section , a s illustrated in Fig. 2 . In this case, th e analysi s illustrate d i n Fig . 1 will giv e spurious answer s dependin g o n the nature of the eroded lithologies . The secon d exceptio n i s wher e observe d palaeo-thermal effect s ar e no t primaril y relate d to depth o f burial, but wer e cause d b y enhance d lateral hea t flow; for example, a s a result o f flow of hot fluids in an aquifer system either within the eroded sectio n o r shallo w i n th e preserve d section. Heatin g cause d by the passage of hot fluids ca n produc e a variet y o f non-linea r palaeotemperature profiles , wit h differen t form s depending on the time scale of heating (see , e.g . Ziagos & Blackwel l 1986 ; Dudd y e t a l 1994) . Most importantly , longe r tim e scale s resul t i n a

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linear temperatur e profil e belo w th e aquifer , which i s parallel t o the present-day temperatur e profile. Erosio n o f th e shallo w strat a i n suc h cases ma y leave a palaeotemperature profil e tha t mimics th e effect s o f deepe r burial , an d tha t would give erroneously hig h estimates o f former burial depths . Heating effect s owing t o mino r igneou s intrusions ca n produc e purel y loca l anomalie s as a resul t o f contac t effects , o r ma y b e mor e widespread i f the intrusion s cause circulatio n of heated fluid s o n a regional scale (e.g . Summer & Verosub 1989) . Suc h effect s ca n als o mimi c th e effects o f deeper burial, and identification of such effects require s regiona l dat a coverage . Examples o f heating tha t ar e probably du e t o the effects o f hot fluids, but that mimic the effect s of deepe r burial , hav e bee n discussed , fo r example, b y Gree n e t al. (1993a , 1997 ) i n th e Irish Se a an d b y Marshallse a e t a l (2000 ) i n the Laur a Basin o f Far North Queensland . Approaches based on heat flow Taking int o accoun t th e spatia l variatio n o f thermal conductivit y throug h th e sectio n an d temporal variatio n i n hea t flo w ca n potentiall y provide mor e accurat e prediction s o f th e variation o f temperatur e wit h dept h an d time . However, this procedure i s also subjec t to major uncertainties. For example, therma l conductivity is highl y sensitiv e t o litholog y (i n particular, t o porosity an d wate r content) , wit h apparentl y similar sample s givin g therma l conductivit y values that vary by almost an order of magnitude (e.g. Corriga n 199la) . Suc h problem s ar e particularly pronounce d i n section s tha t hav e been mor e deepl y burie d i n the past , a s therma l conductivities depen d o n the degre e o f compac tion, whic h i s no t know n (withou t direc t measurements, whic h ar e availabl e onl y rarely ) until the amount of exhumation is determined. In addition, whe n considerabl e sectio n ha s bee n removed n o informatio n i s availabl e o n litho logies (an d hence therma l conductivities ) in th e removed section. Therefore, the effects o f greater depth o f burial are difficul t t o predict accuratel y in sequence s affecte d b y exhumatio n whe n th e magnitude o f buria l i s unknow n (fo r furthe r discussion, se e Duddy et al. 1991 ; Waples e t al. 1992). A further point to note in assessing approache s based on heat flow is that all heat-flow value s are derived originall y fro m down-hol e temperatur e measurements, which are combined with thermal conductivities (usuall y assumed , rarel y measured) t o produc e th e reporte d heat-flo w value. I n th e past , ra w temperature s wer e ofte n

discarded, an d no t reporte d wit h th e resultin g heat-flow values . This allows the possibility tha t the 'measured ' hea t flow may then be combine d with a differen t se t o f therma l conductivit y values tha n thos e use d i n thei r origina l determination, whic h wil l resul t i n a therma l model tha t doe s no t honou r th e measure d temperatures fro m whic h th e hea t flo w wa s derived. I t i s no t vali d t o simpl y tak e a quote d heat-flow valu e an d combin e i t wit h a se t o f preferred o r 'off-the-shel f therma l conduc tivities t o predic t a temperatur e profile . Th e quoted heat-flow valu e is meaningful only when combined wit h th e therma l conductivitie s fro m which it was derived, s o as to accurately predict the present-day therma l structure . We should also note that treatments that use a single valu e o f therma l conductivit y an d hea t flow a s a n approximatio n t o mor e comple x situations (e.g . Gallaghe r & Brow n 1999 ) ar e equivalent t o assumin g linea r geotherma l gradients, an d ar e subjec t t o exactl y th e sam e potential limitation s a s th e approac h i n Fig . 1 (and possibly more if the combined heat-flow and conductivity value s d o no t matc h th e origina l temperature data from whic h the heat flows were derived). Palaeo-surface temperature Estimates o f remove d sectio n fro m Fig . 1 als o depend o n the assume d palaeo-surf ace temperature, independent estimates of which may or may not b e availabl e fo r a particula r region . I n th e absence o f suc h information , eithe r th e present day value can be used, or else calculations can be performed fo r a rang e o f likel y values . Uncertainties associate d wit h thi s paramete r can b e easil y assesse d fo r an y give n valu e o f palaeogeothermal gradient . Fo r instance , i f th e palaeogeothermal gradien t wa s SO^km" 1 , a 10°C rise in palaeo-surf ace temperature requires that th e estimate d erode d sectio n shoul d b e reduced by 200 m. Heating rates The therma l histor y befor e th e onse t o f coolin g from a palaeo-therma l pea k canno t b e con strained b y AFTA data (Gree n e t al. 1989a ) an d as the thermal sensitivity of VR is similar t o that of AFT A (Dudd y e t a l 1991 , 1994 , 1998 ) th e same i s tru e o f V R data . I n extractin g quantitative therma l histor y informatio n fro m AFTA an d VR , i t i s therefor e necessar y t o assume a heatin g rat e t o estimat e a specifi c temperature. Typica l value s ar e usually between 1 an d lO^Ma" 1 . Changin g th e assume d

QUANTIFYING EXHUMATIO N USIN G AFT A AN D VR

heating rat e b y a n orde r o f magnitud e i s equivalent to a change o f c. 10° C in the require d maximum palaeotemperature fo r both AFTA and VR (Green et al 1989a ) with higher heating rates requiring highe r temperature s an d vic e versa . Thus, if assumed heatin g rates are systematically high, the n amount s o f exhumatio n wil l als o b e correspondingly high . Identifying the appropriate unconformity In sedimentary sequences , accurac y in estimating amounts o f remove d sectio n a s show n i n Fig . 1 depends, i n additio n t o th e point s discusse d s o far, o n identificatio n o f th e appropriat e uncon formity fro m whic h the sectio n wa s removed, a s the extrapolatio n i n Fig . 1 i s constructe d wit h respect t o th e dept h o f tha t unconformity . I n sections containin g onl y a singl e majo r uncon formity, thi s i s straightforward , bu t i n section s containing multipl e unconformities , correc t assignment ma y be more difficult , an d erroneou s assignment produce s a systemati c erro r i n estimating th e amoun t o f remove d section . I n such cases , determinin g th e timin g o f exhuma tion independentl y usin g AFT A (se e discussio n below) ca n provid e uniqu e insigh t int o whic h unconformity represent s th e mai n phas e o f exhumation. Calibration of system response Given al l o f th e systemati c factor s discusse d t o this point , accurac y i n estimatin g amount s o f exhumation i s controlle d primaril y b y th e accuracy o f th e palaeotemperature s derive d from AFT A an d VR . This , i n turn , depend s o n using th e mos t reliabl e quantitativ e kineti c descriptions o f syste m response , an d i t i s essential t o demonstrat e tha t thes e description s are calibrated i n geological situations . Estimates of exhumation derive d usin g kinetic model s that do no t accuratel y matc h calibratio n dat a fro m well-understood situation s wil l inevitabl y introduce systemati c error s i n estimate s o f exhumation. In the studies described below, thermal history information ha s been extracte d fro m AFT A dat a using a proprietary 'multi-compositional ' kineti c model, whic h make s ful l quantitativ e allowanc e for th e effec t o f Cl content o n annealin g rates of fission track s i n apatite (Gree n e t al. 1996) . Thi s model wa s derive d directl y fro m dat a i n geological conditions , i n combinatio n wit h laboratory data . V R value s ar e converte d t o maximum palaeotemperature s usin g th e kineti c model develope d b y Burnha m & Sweene y (1989) an d Sweene y & Burnha m (1990) . Goo d

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agreement betwee n measure d an d predicted V R values (a t leas t u p t o c . 1.0% ) i n sediment s that have undergone progressive burial and are now at their maximu m post-depositiona l palaeotem peratures confirm s th e validit y o f th e kineti c model. Use of inappropriate system kinetics: apatite fission-track analysis The firs t quantitativ e kinetic mode l fo r fission track annealin g i n apatit e t o gai n widesprea d acceptance wa s that of Laslett et al. (1987). This was base d purel y o n laborator y annealin g experiments in a single, compositionally uniform apatite. Compariso n o f predictions based o n this model wit h dat a fro m controlle d geologica l conditions (Gree n e t al . 1989a ) showe d tha t the predictions wer e broadl y consisten t wit h th e measured data . But i t wa s als o recognize d a t an early stag e i n th e developmen t o f AFT A tha t apatites o f differen t compositio n annea l a t different rates , wit h C l conten t apparentl y providing th e dominan t influenc e (Gree n e t al . 1985). Thi s effec t produce s a variatio n i n th e degree o f fission-trac k annealin g i n differen t apatite grain s withi n a singl e sample , whic h cannot be described usin g a kinetic model base d on a singl e apatit e specie s suc h a s th e Laslet t et al . (1987 ) model . Th e 'multi-compositional ' model (Gree n e t al 1996 ) introduce d i n th e previous section has the same general form as the Laslett e t al . (1987 ) model , bu t use s constant s that var y systematicall y with Cl content so as to reproduce th e observed within-sampl e variation. Other worker s hav e adopte d a n alternativ e approach t o dealin g with variation i n annealing kinetics betwee n apatit e specie s (Carlso n e t al . 1999;Donelickeffl/. 1999 ; Ketcham ef a/. 1999) . However, th e performanc e o f thes e model s i n geological condition s ha s ye t t o b e rigorousl y demonstrated. A number of other kinetic models based o n laborator y annealin g of mono-compo sitional apatite s hav e als o bee n publishe d (Carlson 1990 ; Crowle y e t al . 1991 ) bu t thei r predictions ar e no t compatibl e wit h dat a fro m controlled geologica l situations , whic h ha s precluded routin e use. In addition , i t i s widel y accepte d tha t th e Laslett e t al . (1987 ) mode l under-predict s th e degree o f annealin g observe d a t lo w tempera tures (e.g . Vrolij k e t al . 1992) . Becaus e o f this , thermal histor y solution s derive d fro m apatit e fission-track dat a usin g the Laslet t e t al . (1987 ) model invariabl y suggest up t o 3 0 °C or more of Late Tertiar y cooling , whic h i s generall y regarded a s a n artefac t o f th e low-temperatur e

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behaviour o f thi s kineti c model . A n adde d contribution t o this artefactual cooling ma y well be due to compositional variation , as the majority of apatit e grain s annea l mor e rapidl y tha n th e Durango apatit e o n whic h th e Laslet t e t al (1987) model was derived. Thus, track lengths in such grains will be shortened t o a greater degre e than expecte d i n apatit e o f th e compositio n t o which th e kineti c mode l directl y relates , inevitably resultin g i n th e nee d fo r a late , anomalous cooling episode . The 'multi-compositional ' mode l (Gree n et al . 1996 ) i s base d o n bot h laborator y an d geological dat a an d provide s a muc h bette r fi t to control dat a a t lower temperatures, a s well as incorporating within-sampl e variation . Thus , thermal historie s derive d fro m thi s mode l d o not suffe r fro m th e 'Lat e Tertiar y coolin g artefact' tha t characterize s therma l histor y solutions derive d usin g th e Laslet t e t al . (1987) model . Despite th e problem s discusse d s o far, use of the Laslet t e t al . (1987 ) mode l remain s widespread i n man y studie s base d o n analysi s of fissio n track s i n apatit e (man y o f whic h

appear t o sho w Lat e Tertiar y cooling!) . Us e o f this mode l t o estimat e palaeotemperature s fro m fission-track dat a i n apatit e wil l giv e system atically hig h temperature s belo w c . 7 0 °C, an d systematically lo w temperature s a t aroun d 100-110 °C. Fo r thi s reason , us e o f thi s model wil l resul t i n anomalousl y lo w palaeo geothermal gradient s an d anomalousl y hig h amounts o f remove d section . Bu t mos t importantly, i t wil l fai l t o adequatel y describ e variation withi n a n apatit e populatio n fro m a single sample , introducin g seriou s error s t o thermal histor y solution s tha t ma y b e widel y erroneous. To summarize, accurac y of palaeotemperatur e determination from AFT A is critically dependent on us e o f kineti c model s tha t describ e variatio n between apatite grains within individual samples, for whic h incorporatio n o f compositiona l influ ences i s essential . Furthe r discussio n o f th e kinetic respons e o f the AFTA syste m is provide d below i n th e sectio n 'Quantifyin g timin g o f exhumation using AFTA, as estimates o f timing and palaeotemperature obtaine d fro m AFT A ar e inextricably linked.

Fig. 3 . This plot shows the relationship between maximum palaeotemperature s and VR values predicte d from the Burnham & Sweeney (1989) model, using variou s heating rates , and for the Barker & Pawlewicz (1986) model, which ignore s th e influenc e o f time an d relates reflectanc e directl y t o palaeotemperature. The tw o approache s show very different behaviour , particularly at reflectances abov e 1 % and below 0.5%. Fo r this reason, use of the Barker & Pawlewicz mode l wil l generall y produc e highe r palaeogeotherma l gradients , an d lowe r amount s o f removed section, compared wit h th e Burnham & Sweeney model.

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Use of inappropriate system kinetics: vitrinite reflectance

Fig. 4 . Various sets of maturity data in an offshore well are plotted against depth (RKB; below Kelly bushing). An earlie r V R datase t (grey , white and black squares) plus VR data from ne w analyses (diagonal stripes) are shown, togethe r wit h range s o f equivalen t VR level s derived from AFTA in two samples from this well, plus equivalent V R value s derive d fro m SC I an d r max values. I n th e earlie r V R analyses , reflectanc e measurements wer e assigne d t o variou s population s by th e analysts , as indicated in th e legend . In the ne w analyses, identificatio n o f indigenou s vitrinit e wa s based o n petrographic inspection of polished sections. The continuou s line show s the V R profil e expecte d if samples throughou t the sectio n ar e currentl y at thei r maximum temperatur e sinc e deposition . Ther e i s a clear mismatc h betwee n those value s from th e earlier analyses originall y interprete d a s representin g th e indigenous vitrinit e population, on on e hand , and th e new V R values plus th e equivalent V R value s defined by AFTA , SC I an d r max dat a o n th e other , whic h all define a consistently higher trend. Whereas the original data sugges t tha t al l unit s throughou t th e wel l ar e currently a t their maximum temperatures , the new VR data and the AFTA, SC I and 7 max results show that in fact maturit y levels are higher than previously thought by c . 0. 2 o r 0.3%, an d mos t unit s have been hotte r in the past . Th e highe r tren d i s sub-paralle l t o th e predicted profile , suggestin g tha t heatin g wa s du e t o deeper burial , followe d b y coolin g a s a resul t o f exhumation. It is also worth noting that the occurrence of three separat e sub-parallel trends denoting caved, in situ an d reworke d vitrinite , a s i n th e earlie r analyse s shown here , i s highl y unlikely , an d mor e probabl y arises becaus e o f measurement o f macerals othe r than true vitrinite . Th e occurrenc e o f thre e sub-paralle l trends i s usuall y a reliabl e sig n o f macera l misidentification.

In th e year s sinc e vitrinit e reflectanc e wa s adopted a s a standar d measur e o f maturit y fo r hydrocarbon exploration , a numbe r o f kineti c models have bee n suggeste d fo r th e evolutio n o f reflectance a s a function of temperature an d time . The Burnha m & Sweene y (1989 ) mode l i s undoubtedly th e mos t successful , i n term s o f accurately reproducin g observe d value s i n well controlled situation s (se e als o Morro w & Issle r 1993). However , despit e th e succes s o f thi s model, othe r treatment s remai n i n common use , and us e o f inappropriat e kineti c description s fo r VR ca n provid e seriou s systemati c error s i n estimating palaeotemperature s an d thus amounts of exhumatio n (o r forme r depth s o f burial) . At on e extreme , understandin g o f th e syste m kinetics ma y b e ignore d completely . Fo r instance, estimate s o f remove d sectio n hav e been obtaine d historicall y b y extrapolatin g V R profiles t o a valu e representin g untransforme d kerogen a s deposite d a t th e surface , typicall y around 0.2 % (e.g . Do w 1977) . However , experience show s (Cook , pers . comm. ) tha t untransformed vitrinite , a s deposite d i n a sediment, ma y hav e a reflectanc e a s hig h a s 0.32%. Thus , extrapolatio n t o 0.2 % shoul d invariably overestimat e amount s o f remove d section. Problems wit h another approac h ar e illustrated in Fig . 3 , whic h show s th e relationshi p betwee n maximum palaeotemperature s an d V R value s predicted fro m th e Burnha m & Sweene y (1989 ) model, usin g a variet y o f differen t heating rates , and from th e Barker & Pawlewicz (1986 ) model, which ignore s th e influenc e of tim e an d relate s reflectance directl y t o palaeotemperature . Th e two approache s sho w ver y differen t behaviour , particularly a t reflectance s above 1 % an d belo w 0.5%. Fo r thi s reason , us e o f th e Barke r & Pawlewicz mode l wil l generall y produc e highe r palaeogeothermal gradients , an d lowe r amount s of removed section , compare d with the Burnha m & Sweene y model .

Integration of results from different methods If th e kineti c algorith m use d t o extrac t therma l history informatio n fro m th e AFT A and/o r V R data wer e t o b e systematicall y i n error , the n palaeotemperatures from on e method or the other would b e consistentl y eithe r lo w o r high . Comparative result s fro m a wid e rang e o f situations (e.g . Dudd y e t al 1994 , 1998 ; Gree n et a l 1995 , 1997 ) sho w tha t th e tw o technique s

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give highl y consisten t palaeotemperatures . A s the calibratio n o f eac h syste m has bee n carrie d out independently, this suggests that both sets of values ca n b e regarde d a s reliable, an d that any systematic inaccurac y a s a resul t o f error s i n system response ar e not significant. Analytical problems Analytical problem s i n V R analyse s ar e wel l known, includin g problems suc h as suppression or retardatio n (e.g . Car r 2000) . Suc h problem s

usually affec t a particular horizon (e.g . suppression in hydrogen-rich source-rock facies) and can often b e recognized b y local departure s from a n overall trend. A variety of approaches have been proposed fo r dealin g wit h reflectanc e suppression, includin g utilizatio n o f fluorescenc e infor mation (e.g. Wilkins et al. 1992; Newman 1997), and development of specific kinetic models (Carr 1999, 2000). Combination of palaeotemperatures derived from differen t technique s (e.g. AFTA, as discussed above ) ca n als o allo w detectio n o f anomalous VR values.

Fig. 5 . Palaeotemperature profile s from AFTA and VR data in the Anglesea-1 well. Left: AFTA data clearly revea l two palaeo-thermal episodes, a s shown, with palaeotemperature constraint s fro m eac h method fo r the two events plotted against depth (RKB). AFTA shows that cooling began in the earlier episode between 11 0 and 95 Ma and in the later episode betwee n 60 and 10 Ma. Unconformities are present for mid- to Late Cretaceous an d Late Tertiary time, suggestin g tha t both episodes ca n be explained b y deeper burial . Palaeotemperature profile s characterizin g both episodes ar e linear, supportin g this conclusion. DB, Demons Bluff Formation ; EV , Eastern Vie w Formation . Right: statistica l analysi s o f th e palaeotemperatur e constraint s define s th e rang e o f allowe d value s o f palaeogeothermal gradien t an d remove d sectio n durin g eac h episod e o f exhumatio n withi n 95 % confidenc e limits, a s shown by the contoured regions , togethe r wit h the best-fi t values .

QUANTIFYING EXHUMATIO N USIN G AFT A AN D VR

However, a possibl y mor e commo n problem , which ma y b e mor e difficul t t o recognize , concerns th e incorrec t assignmen t o f th e i n situ vitrinite populatio n throughou t a n entir e wel l section. Thi s ca n resul t i n systematicall y lo w estimates o f palaeotemperatur e (an d therefor e palaeo-burial) an d maturit y levels . Thes e pro blems are illustrated i n Fig. 4, in which a series of datasets fro m a n offshor e wel l ar e compared . AFTA dat a i n tw o sample s fro m thi s wel l sho w consistent evidenc e o f highe r temperature s before Miocen e cooling , bu t existin g V R dat a indicated tha t al l unit s wer e no w a t thei r maximum temperatures sinc e deposition. Result s of ne w V R analyses , carrie d ou t t o resolv e this mismatch , ar e highl y consisten t wit h th e equivalent maturit y value s indicate d b y th e AFTA data , an d als o wit h equivalen t maturit y levels derive d fro m SC I an d r max value s within th e origina l dataset . Further investigatio n o f th e origina l V R dataset revealed a number o f populations withi n the measurements , a s shown i n Fig. 4 , including a populatio n o f highe r reflectances , originall y interpreted a s 'reworked ' vitrinite . Significantly, the ne w V R value s an d th e equivalen t maturit y values fro m AFTA, SC I and 7max data are highly consistent wit h these 'reworked ' vitrinit e values . All these values define a consistent trend in Fig. 4, consistently highe r tha n thos e originall y attrib uted t o th e indigenou s population . Thus , thes e original value s ar e too low and values originall y designated a s reworked, togethe r wit h the AFTA, SCI an d 7 max values , provid e th e mos t reliabl e indication o f true maturit y level s i n this well . Such experienc e i s remarkabl y common , an d illustrates a genera l tendenc y amongs t man y geologists t o underestimat e th e importanc e o f exhumation (o r deepe r burial) . Thus , th e reflectance populatio n tha t fall s closes t t o th e values expecte d o n th e basi s o f th e preserve d section i s mos t commonl y identifie d a s th e indigenous population . Thes e observations agai n highlight th e importanc e o f combinin g infor mation o n palaeo-therma l effect s fro m differen t techniques, whic h provide s a mor e objectiv e assessment. Factors affecting th e precision of the magnitude of exhumation Availability of palaeotemperature constraints over a range of depths The main facto r affectin g precision i n estimatin g amounts o f remove d sectio n i s th e rang e o f depths over which palaeotemperature constraint s

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are available, a s this controls the degree t o which the palaeogeotherma l gradien t i s constrained , which, i n turn , control s th e precisio n o f th e estimate o f removed sectio n (Fig . 1) . An extreme case o f this is where data are available only fro m outcrop samples , in which case no estimate of the palaeogeothermal gradien t i s possibl e an d a value mus t b e assumed . Uncertainties i n th e fitte d palaeogeotherma l gradient ar e magnifie d becaus e o f th e extrapol ation required to estimate the amount of removed section. Thi s also cause s a correlatio n betwee n allowed value s o f palaeo-gradien t an d remove d section, suc h tha t hig h palaeo-gradient s requir e smaller amount s o f remove d sectio n an d vic e versa. Statistica l technique s allo w definitio n o f the rang e o f allowe d value s (withi n ±95 % confidence limits ) for eac h parameter . These point s ar e illustrate d i n Fig . 5 usin g measured dat a fro m th e Anglesea-1 well locate d in th e Otwa y Basi n o f S E Australi a (Dudd y 1994). Combine d AFT A an d V R dat a fro m thi s well define two discrete palaeo-thermal episodes . The palaeotemperature profil e characterizin g the mid-Cretaceous episod e i s define d b y V R dat a over a dept h interva l o f c . 2. 5 km (AFT A dat a provide onl y minimu m limits) , whic h result s i n relatively tight constraints on the allowed ranges of palaeo-geothermal gradient s (43-57 °C km"1 ) and remove d sectio n (1750-270 0 m). I n con trast, Tertiar y palaeotemperatur e constraint s from AFT A an d V R ar e availabl e ove r a dept h of only 1. 5 km, and provide much broader ranges of allowe d value s fo r eac h paramete r (18-52 °C km"1 and 400-2300 m, respectively). These example s typif y th e level s o f precisio n available fro m thi s approach. However , it should be note d tha t althoug h th e range s o f allowe d values ar e broa d fo r bot h events , fo r an y particular valu e of gradient the range o f allowe d values of removed sectio n wil l be much smaller, typically aroun d 200-500m i n Fig. 5 . Nature of palaeotemperature constraints Other factor s affectin g the precisio n of estimat ing amount s o f exhumatio n (remove d section ) include th e qualit y o f th e V R an d AFT A dat a (poor-quality dat a wil l generall y provid e onl y broad palaeotemperatur e constraints , whic h provide littl e contro l o n th e palaeo-gradient ) and th e availabilit y o f AFT A an d V R dat a through th e section , couple d wit h th e natur e of the AFT A constraint s (whic h depend s t o som e extent o n th e natur e o f th e underlyin g therma l history). Fo r example, i f all AFTA sample s wer e totally o r nea r totall y anneale d befor e cooling , such tha t onl y a lowe r limi t t o th e maximu m

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palaeotemperature i s availabl e fro m AFTA , useful constraint s o n th e palaeogeotherma l gradient wil l b e availabl e onl y i f V R dat a ar e also availabl e throug h the section . Result s fro m the Wes t Newton- 1 wel l (N W England ) (Gree n et al. 1997 ) illustrat e thi s point . Availabilit y of VR dat a throug h th e Carboniferou s sectio n i n this well , wher e AFT A sample s wer e totall y annealed before Early Tertiary cooling, results in tight constraint s o n th e palaeo-gradien t an d th e degree o f exhumation. Quantifying timin g of exhumation using AFTA The timin g o f exhumatio n event s has tradition ally bee n inferre d fro m regiona l geologica l evidence, bu t a majo r advantag e o f AFT A i s that i t provide s a n independen t estimat e o f th e time a t whic h a sampl e bega n t o coo l fro m it s maximum palaeotemperatur e (o r a subsequen t peak value) . Thi s estimat e i s derive d fro m th e AFTA dat a alone , an d therefor e provide s a n

objective measur e o f timing . I f coolin g ca n b e attributed t o exhumation , then AFTA ca n defin e the timin g of the onse t o f exhumation. In considerin g th e accurac y an d precisio n o f timing estimate s fro m AFTA , i t i s importan t t o recognize ho w thi s informatio n i s code d i n th e AFTA data, which in turn requires explanation of the therma l respons e o f fissio n track s i n apatite . AFTA i s based o n 'fission-trac k annealing' : th e progressive reductio n i n trac k lengt h a s a function o f temperatur e an d tim e (Gree n e t al . 1986, 1989b) . Thi s reductio n i n trac k lengt h is also manifeste d a s a reductio n i n fission-trac k age (Gree n 1988 ; Gree n e t a l 1989b) . A s illustrated i n Fig . 6 , ne w track s ar e produce d throughout geologica l time , a s a resul t o f spontaneous fissio n o f uraniu m impurit y atom s within the apatite crystal lattice. In a sample that is heate d an d the n coole d (e.g . Sampl e 1 i n Fig. 6) , tracks produced u p to the time a t which cooling begin s wil l b e shortene d t o a lengt h determined by the maximum palaeotemperature, whereas track s produce d afte r th e onse t o f cooling will be longer. The time at which cooling

Fig. 6 . The thermal respons e of fission track s i n apatite t o geological therma l historie s is well understood , based on a series of observations i n laboratory an d geological conditions (Gree n e t al. 1986 , 1989a ; Laslett et al. 1987 ; Duddy e t al . 1988) . I n a therma l histor y tha t involve s heatin g an d subsequen t cooling , followe d b y mino r reheating (a) , fissio n track s i n apatite respon d i n a characteristic way (b) , whic h allow s th e main features of the history to be constrained. Detail s are discussed in the text. Understanding the way in which fission tracks in apatite respond to heating an d cooling allows us to focus on those aspects of the history tha t can be constrained (i.e . th e maximum palaeotemperature and the time of cooling), wherea s th e history before th e onset o f cooling canno t be constrained.

QUANTIFYING EXHUMATION USING AFTA AN D VR

begins, i n relatio n t o th e overal l duratio n o f th e history, determine s th e proportio n o f shorte r t o longer tracks , an d the maximum palaeotempera ture determine s th e lengt h o f th e shorte r peak . Because a n apatite may contain track s at the time of depositio n i n a sediment , thi s trac k lengt h information mus t b e combine d wit h a fission track ag e measuremen t t o establis h th e tota l duration ove r whic h track s hav e bee n retained . In sample s tha t excee d a critica l temperatur e limit, al l track s ar e 'totall y annealed ' (i.e . th e track length is reduced to zero, as for Sample 2 in Fig. 6) . Suc h sample s retai n track s onl y afte r cooling belo w thi s limi t (whic h depend s o n th e composition o f th e apatite) , an d provid e onl y a minimum estimat e o f th e maximu m palaeotem perature. However , suc h samples usually provide tight constraint s o n the tim e o f cooling, throug h the fission-trac k age .

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Determining th e time a t which cooling begin s from AFT A dat a i s thu s inextricabl y linke d t o extraction o f palaeotemperatur e information , as both aspect s o f th e overal l therma l histor y solution exer t critica l control s o n th e measure d AFTA parameters .

Factors affecting th e accuracy and precision of timing estimates from AFTA Extracting therma l histor y solution s from AFT A data involve s modellin g AFT A parameter s (Green e l al 1989b ) throug h variou s therma l history scenarios, base d on a detailed knowledge of th e kinetic s o f th e annealin g process, s o as t o define th e rang e o f maximu m palaeotempera tures and timing of cooling for which predictions are consisten t wit h the measure d data . Figur e 7

Fig. 7 . Extracting thermal history solutions from AFT A data involves modelling AFTA parameters (Green et al. 1989a) throug h various therma l histor y scenarios, usin g forma l statistica l procedure s to defin e th e rang e of maximum palaeotemperatures and timing of cooling for which predictions are consistent with the measured data. This process requires a detailed knowledge of the kinetics of the annealing process. This synthetic example, based on a notional mono-compositional example, for simplicity, illustrates the basic principles involved. Cooling from a maximu m palaeotemperature of 90 °C beginning a t 50 Ma give s the bes t fit to th e data .

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illustrates th e basic principles involved, based on a simpl e mono-compositiona l example . B y modelling th e AFT A parameter s throug h likel y thermal histor y scenarios , an d comparin g pre dictions with measured data, the range of thermal history solutions compatible with th e data withi n 95% confidenc e limit s can be defined . In practice , annealin g kinetic s depend s o n chlorine content. As illustrated in Fig. 8 , therma l history solution s ca n b e extracte d fro m dat a broken down int o discrete compositional groups , using separat e kinetic s for each group . Th e final

thermal histor y solutio n shoul d no t onl y matc h the pooled data for the whole sample but also the within-sample variation o f fission-trac k ag e an d length wit h w t % Cl, whic h provide s additiona l information b y whic h th e fina l therma l histor y solution can be constrained. The variation i n the trac k lengt h distribution s in Fig . 8 between compositiona l group s shoul d be compare d wit h tha t show n b y Gree n (1986 ) for variou s outcro p sample s fro m Norther n England. Individua l compositiona l groups withi n the singl e sampl e illustrate d i n Fig . 8 spa n a

Fig. 8 . Influenc e o f Cl conten t o n AFTA data . In detail , fission-trac k annealin g kinetic s i n apatit e depend s o n chlorine content , and thermal history solutions are extracted fro m dat a broken down into discrete compositiona l groups, using separate kinetics for each group. The final thermal history solution should match not only the pooled data characterizing th e sample as a whole but also the variation of fission-track age and length with wt % Cl. The data shown here were measured in a sample of Triassic sandston e from N W England. In the upper plot, fissiontrack age s o f individual apatite grain s are plotte d agains t Cl conten t (measure d b y electro n microprobe) . Als o shown ar e predicted pattern s o f fission-trac k ag e an d mea n trac k lengt h v. wt % Cl fo r thre e differen t therma l histories, as indicated. A maximum of 100° C at 60 Ma clearly gives the best fit to the data. The lower plot shows measured lengt h distribution s binne d int o discret e w t % C l intervals , togethe r wit h predicte d trac k lengt h distributions corresponding to the best-fit thermal history, which also give a good match to the measured data. It is essential, i n extracting thermal history information from fission-trac k dat a in apatite, to use kinetic models tha t incorporate th e effec t o f this within-sample variation.

QUANTIFYING EXHUMATION USIN G API A AND VR

similar rang e o f annealin g show n b y outcro p samples ove r a wid e area . Jus t a s i t i s no t acceptable t o extrac t a singl e therma l histor y solution fro m a collectio n o f outcro p sample s showing varyin g degree s o f annealing , i t i s equally unacceptabl e t o attemp t t o extrac t suc h information fro m dat a withi n a singl e sampl e showing a n equivalent spread . The accurac y o f timin g estimate s fro m AFTA depend s mainl y o n us e o f a n appro priate kineti c model , as above , but precisio n often depend s criticall y o n th e magnitud e o f palaeo-thermal effects . I n sample s tha t hav e been heate d onl y t o aroun d 7 0 °C o r below , th e degree of lengt h reductio n in track s forme d before th e onse t o f coolin g i s relativel y small . In suc h cases , th e shorte r componen t ca n ofte n not b e resolve d fro m th e longe r track s forme d after cooling , an d a wid e rang e o f timing s wil l be consisten t wit h th e data . Conversely , i n samples tha t hav e bee n heate d t o c. 9 0 o r 100°C, th e degre e o f lengt h reductio n i s severe and the shorte r track s can easil y be resolved, allowin g muc h tighte r constraint s o n timing. Sample s i n whic h al l track s wer e totally anneale d befor e coolin g (maximu m palaeotemperature typicall y > 11 0 -120 °C) generally giv e th e tightes t timin g constraints , although thi s depend s t o som e exten t o n th e history afte r cooling . Uncertainty in the final estimate o f timing ca n depend t o som e exten t o n th e timin g itself . I f cooling (exhumation ) bega n earl y i n the history, the effect s ma y b e difficul t t o discriminate fro m pre-depositional effect s (reflectin g th e history of sediment sourc e terrains), and a large uncertainty may result . Conversely , i f coolin g bega n relatively lat e i n th e history , th e majorit y o f tracks wil l b e affecte d an d timin g ca n b e determined wit h greater precision . In practica l terms , poo r apatit e yield s and/o r low uraniu m content s (producin g lo w trac k counts for age determination an d smal l number s of trac k lengths ) invariabl y resul t i n poo r constraints o n timing . Th e Poissonia n uncer tainty o n trac k count s (fractiona l erro r o f \IN) provides a basi c limi t t o th e fina l erro r o n th e fission-track ag e o f a n individua l apatit e grai n and/or sample , whic h i n tur n limit s th e analytical uncertaint y i n th e timin g o f cooling . Even i n th e highes t qualit y analyses , i t i s unusual fo r mor e tha n c . 100 0 track s t o b e counted, an d thu s uncertaintie s o f severa l pe r cent are to be expected in the best case. However, as stresse d elsewhere , countin g an d measurin g large number s o f track s ar e no t a s importan t a s incorporating within-sampl e variatio n int o th e analysis.

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Limitations Thermal history resolution A fundamenta l characteristic o f bot h th e AFT A and VR techniques is that they are dominated b y maximum temperatures, as discussed above . For this reason , thes e technique s ar e subjec t t o th e fundamental limitatio n tha t the y ca n provid e n o information o n th e approac h t o th e palaeo thermal maximum . (Thi s i s illustrate d b y th e track length shortenin g trajectorie s fo r Sampl e 1 in Fig . 6 , i n whic h al l track s forme d befor e th e onset o f coolin g ar e shortene d t o th e sam e degree.) AFT A ca n provid e informatio n o n th e history following the onset of cooling, because of the continuous production of tracks through time . For example, b y reference t o Sample 2 in Fig. 6, if track s forme d afte r th e initia l coolin g wer e again shortene d durin g a late r heatin g episod e and th e sampl e the n coole d an d retaine d longe r tracks t o th e presen t day , th e resultin g lengt h distribution woul d preserv e evidenc e o f thi s episode, an d dat a fro m thi s sampl e woul d thu s constrain tw o discrete events . However, because of the inherent spread i n the distribution of track lengths a s a functio n o f th e degre e o f annealing (Green e t al. 1986) , there is a limit to the amount of informatio n that ca n b e obtained . In practice, resolution of two episodes of heating and cooling is ofte n th e limi t allowe d b y eve n th e highes t quality data , although Green e t al. (200la ) have illustrated a situatio n wher e AFT A dat a fro m a single sampl e allo w resolutio n of thre e discret e episodes. I n contrast, VR data are not sensitive to the histor y afte r th e onse t o f cooling , excep t i n situations characterize d b y ver y slo w coolin g o r where a considerable tim e is spent at or near the maximum temperature. For these reasons, the approach describe d here is based on using AFTA data to rigorously defin e the maximu m o r pea k palaeotemperatur e an d timing o f coolin g i n on e o r tw o (rarel y three ) discrete episode s o f heatin g an d cooling , usin g assumed heatin g an d coolin g rate s (se e discus sion above) . AFT A data do not contain sufficien t information t o allo w definitio n o f th e entir e thermal histor y o f a sample , a s attempte d b y some worker s (e.g . Corriga n 1991b ; Lut z & Omar 1991 ; Gallaghe r 1995 ; Willet t 1997 ; Ketcham e t al . 2000) , o r eve n th e whol e o f th e history afte r th e onse t o f cooling . Suc h approaches invariabl y result in confidence limit s that are so wide that the thermal history solutions provide no useful constraints . This arises because the effect s o f variatio n withi n on e par t o f th e history ca n b e compensate d b y event s a t othe r times. Onl y b y focusin g o n thos e aspect s o f th e

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thermal histor y t o whic h th e dat a ar e sensitiv e (i.e. maximu m o r pea k palaeotemperatur e an d timing of cooling, unde r assumed heatin g rates), and thereb y reducin g th e proble m t o a manage able numbe r o f variables , ca n usefu l constraint s be obtained . Onset v. duration of cooling or exhumation Because o f th e limitation s discusse d above , AFTA usuall y defines onl y th e onse t o f coolin g (as a result , fo r example , o f exhumation ) wit h any precision , an d canno t provid e ver y precis e constraints o n the duration of episodes of cooling or rates o f cooling, althoug h broad limit s on the magnitude an d duratio n o f coolin g phase s ma y be possible . Example s ar e discussed belo w i n the contex t o f result s fro m th e N W Europea n Atlantic margin . Depth resolution v. thermal resolution Palaeotemperatures from AFT A and VR data are typically accurat e t o within 10°C , and precisio n is usuall y simila r t o o r bette r tha n this . Thi s compares favourabl y wit h present-da y tempera ture assessment , whic h is , a t best , probabl y accurate to no more than ± 10°C . However, as a consequence o f therma l gradient s typicall y between 2 0 an d 60°Ckm~ 1 , eve n unde r idea l circumstances uncertaintie s i n resultin g esti mates o f erode d sectio n ar e usuall y severa l hundreds o f metres . Fo r a palaeogeotherma l gradient o f c. 30°CkirT 1, 1 0 °C is equivalent to 300 m o f inherent uncertainty in th e estimat e of section remove d (500 m fo r a gradien t o f 20°CkirT 1 ). Thi s emphasize s tha t therma l history reconstructio n i s no t a particularl y precise metho d o f estimatin g amount s o f removed section . However , a s illustrate d i n Fig. 5 , greate r precisio n ca n b e obtaine d b y combining dat a fro m multipl e sample s ove r a range o f depths , an d allowe d range s ca n b e rigorously defined . Reburial A fundamenta l limitatio n of all palaeo-therma l techniques tha t ar e dominate d b y maximu m temperatures i s tha t th e effect s o f earl y heatin g are obscure d b y late r heatin g whe n th e magnitude o f th e mor e recen t even t exceed s that o f th e earlie r episode . Fo r example , i f a sedimentary sectio n i s reburied , followin g a n earlier episod e o f exhumation , th e palaeo thermal effect s associate d wit h coolin g durin g exhumation ar e progressivel y 'overprinted' . Therefore, i n a sectio n affecte d b y multipl e

heating episode s onl y th e maximu m palaeo temperature even t wil l b e clearl y reveale d b y most techniques , includin g VR . I n contrast , AFTA i s capabl e o f detectin g lesse r magnitud e events (fro m th e reductio n i n lengt h o f track s formed before each event), but only if they occur after the maximum palaeotemperature even t (and only if they are sufficiently separate d in time and temperature). In area s affecte d b y repeate d cycle s o f buria l and exhumation , earlie r buria l phase s involvin g lower buria l depth s ca n b e reveale d onl y i f th e palaeogeothermal gradient s a t th e tim e wer e higher tha n during late r episodes . Result s fro m Inner Mora y Firt h wel l 12/16- 1 (Gree n et al. 1995) provide a n example o f this situation.

Latest result s fro m the NW Europea n Atlantic margi n Early Tertiary palaeo-thermal effects in NW England Earliest evidenc e o f th e importanc e o f Earl y Tertiary palaeo-therma l effect s i n the UK regio n came fro m th e applicatio n o f AFT A t o out cropping Caledonia n basemen t fro m th e Southern Upland s o f Scotlan d (Hurfor d 1977 ) and th e Lak e Distric t o f N W Englan d (Gree n 1986). A s sample s wer e no t analyse d fro m vertical sequences , thes e studie s provide d n o insight int o th e magnitud e o f palaeogeotherma l gradients an d th e origi n o f th e observe d Earl y Tertiary palaeotemperatures . Similar effect s wer e subsequentl y recognize d regionally across Northern England an d the Irish Sea (Lewis e t al 1992 ; Gree n e t al 1993b) , but although sample s wer e analyse d fro m wel l sequences, tota l o r near-tota l annealin g o f al l AFTA sample s a s a resul t o f th e hig h Earl y Tertiary palaeotemperature s a t sea-be d o r out crop level again precluded an y useful constraints on palaeogeotherma l gradients . Mainl y becaus e of the regional extent of the effect , couple d with the lac k o f evidenc e (a t tha t time ) fo r elevate d basal hea t flow , a n explanatio n i n term s o f heating primaril y cause d b y deepe r buria l wa s considered mos t likely , wit h prevailin g palaeo geothermal gradients close to present-day values, and wit h subsequen t coolin g resultin g largel y from uplif t an d erosio n (exhumation) . O n thi s basis, Lewi s et a l (1992 ) calculate d tha t fo r an assumed palaeogeotherma l gradien t o f 30°Ckm~ 1 9 th e observe d palaeotemperature s required c. 3 km of section to have been removed over much of NW England.

QUANTIFYING EXHUMATIO N USIN G AFT A AND VR

Such result s wer e i n marke d contras t t o th e prevailing consensu s view , an d provoke d con siderable commen t an d criticism . I n particular , Holliday (1993 ) estimate d a probabl e rang e o f 700-1750 m for the amount of former Mesozoic cover ove r the Lake Distric t an d Pennine blocks ,

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and suggeste d tha t highe r amount s wer e no t consistent wit h geologica l informatio n fro m surrounding regions . Later studie s b y Cop e (1994 ) an d Cha d wick et al (1994 ) suggeste d a growing acceptanc e o f the genera l concep t o f kilometre-scal e Tertiar y

Fig. 9 . Summar y o f Earl y Tertiar y palaeo-therma l effect s i n N W England . AFT A an d V R dat a fro m th e Wes t Newton-1 well (upper left ) an d an elevation section from Se a Fell (uppe r right) allow reconstruction of the region immediately befor e th e onse t o f exhumation in Early Tertiary time , a s shown in the sectio n (lower). For furthe r details, se e Green e t al. (1995 , 1999 ) an d Green (2001) .

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exhumation, bu t ar e themselve s ope n t o criti cism. Fo r example , th e concep t o f a domal erosional even t centred i n the Irish Sea just nort h of Anglese y (Cop e 1994 ) i s no t consisten t wit h published palaeotemperature map s (Gree n e t al 1993b, 1997) , which show Early Tertiary palaeo thermal effect s o f maximu m magnitud e i n th e Lake District , decreasin g southward s int o Nort h Wales an d westward s t o th e Isl e o f Ma n an d Northern Ireland . Th e erosio n ma p o f Cop e (1994) also fails to reproduce th e observed Early Tertiary palaeotemperatur e maximu m i n th e Cleveland Basi n reporte d b y Gree n e t al . (1993b). Also , ove r mos t o f th e Lak e Distric t Block, the estimates of removed section reporte d by Chadwic k e t al. (1994 ) ar e outsid e the limits that Hollida y (1993 ) considere d consisten t with geological constraints , bein g > 1750m ove r al l but a small portion i n the centre of the region an d increasing t o the north, west an d south to values well outsid e th e rang e o f value s considere d acceptable. Unti l recently, therefore, the origin of the Earl y Tertiar y palaeotemperature s reveale d by AFT A i n Norther n Englan d ha s remaine d enigmatic, wit h th e natur e o f th e underlyin g processes unclear . Subsequent t o thes e studies , a n increasin g focus on hydrocarbon exploratio n i n the Irish Sea and adjacen t regions le d to major improvements in definitio n o f therma l histor y style s i n th e region, th e first signs of which were reported b y Green e t al . (1993a) . I n additio n t o recognizin g the occurrenc e o f Mesozoi c palaeo-therma l episodes, particularl y t o th e wes t o f th e regio n (Green et al. 1997) , th e increasing availability of AFTA an d V R dat a ove r a rang e o f depths , combined wit h improve d understandin g o f fission-track annealin g kinetic s (a s discusse d above), resulte d i n muc h tighte r constraint s o n the natur e o f Earl y Tertiar y palaeo-therma l effects, revealin g a majo r difference fro m sout h to north. Wells from the south of the basin define low palaeogeotherma l gradient s suggestiv e o f heating relate d t o ho t flui d circulation , wherea s the northern parts of the basin were characterized by muc h highe r palaeo-gradients , suggestin g a major contributio n o f heatin g a s a resul t o f elevated basa l hea t flow. Evidence fro m th e nort h o f th e basi n ar e typified b y AFTA an d VR data from th e onshore West Newton- 1 wel l (Gree n e t a l 1997 , 1999) . Assuming that both AFTA and VR data represen t the sam e palaeo-thermal episode , the y defin e a n Early Tertiar y palaeogeotherma l gradien t o f c. 50°Ckm~ 1 (Fig . 9) , compared wit h th e present-day gradien t o f c. 35°Ckm~ 1 , implyin g that th e Earl y Tertiar y hea t flow was u p t o 50% higher tha n the present-day value . Extrapolating

the palaeotemperature s t o a palaeo-surfac e temperature o f 2 0 °C require s aroun d 1.55k m of post-Earl y Triassi c sectio n remove d b y Tertiary exhumation . However , a s explaine d b y Green e t al . (1997 , 1999) , dat a fro m Wes t Newton-1 woul d allo w a n alternativ e interpretation whereb y th e V R dat a represen t a n earlie r episode (perhap s a t latest Carboniferous time) in which th e Carboniferou s unit s reache d thei r maximum post-depositional palaeotemperatures, whereas Early Tertiary palaeotemperatures in the Carboniferous sectio n were somewha t lower. Recently, AFT A dat a i n a serie s o f outcro p samples from variou s elevations in the vicinity of Sea Fel l hav e provided confirmation o f elevated palaeogeothermal gradient s durin g Earl y Ter tiary time (Green et al. 1999 ; Green 2001). In this region, characterized b y the highest elevations in England a t just unde r 1000m , an d locate d onl y c. 25 km to the south of the West Newton-1 well, Early Tertiary palaeotemperatures obtained fro m AFTA data define a palaeogeothermal gradient of 61 ""Ckm"1 an d require onl y c. 680m of sectio n removed sinc e Earl y Tertiar y tim e assumin g a palaeo-surface temperatur e of 20 °C (Fig. 9). The difference o f c . 870 m i n amount s o f remove d section between the location of the West Newton1 wel l an d Se a Fel l i s clos e t o th e c . 950 m difference i n elevation between the (near-coastal) location o f th e Wes t Newton- 1 wel l an d th e summit of Se a Fell (particularl y bearing in mind typical uncertaintie s of ±50-100m). Thus, thes e ne w observation s provid e a self consistent framewor k withi n whic h Earl y Ter tiary palaeo-therma l effects i n N W Englan d can be understoo d a s being du e t o a combination of higher basal heat flow and deeper burial , with the amount of sectio n remove d durin g Tertiar y exhumation generally varying between c. 0.7 km (from mountai n peaks ) an d c . 1.6k m (fro m coastal plain s and glacia l valley s near se a level) over th e regio n (Fig . 9) . Amount s o f remove d section require d to explai n the observe d Earl y Tertiary palaeotemperatures are entirely consistent wit h th e conclusion s o f Hollida y (1993 ) based o n regional geological trends . Thus, these latest result s provid e a geologicall y plausibl e mechanism fo r th e origi n o f th e observe d Earl y Tertiary palaeo-therma l effect s i n N W Englan d (although th e underlyin g cause o f th e elevate d heat flo w remain s unclear) , illustratin g th e benefits t o be gained fro m analysin g data ove r a range o f elevations. New results from Central England Green (1989 ) reporte d AFT A dat a i n a suit e of samples fro m outcrop s an d exploration wells on

QUANTIFYING EXHUMATIO N USING AFTA AND VR

the East Midlands Shel f (EMS ) and the Pennine High that revealed Earl y Tertiary palaeo-thermal effects, interprete d a s representing greater depths of burial before Tertiary exhumation. Bray et al (1992) subsequentl y incorporate d V R dat a with these AFTA data and provided a more quantitative analysis of the palaeotemperature data, which supported a n explanation of heating primarily as a resul t o f deepe r burial , wit h amount s o f removed sectio n varyin g betwee n 1 an d 2k m across th e East Midlands Shelf. Holliday (1993 ) and Smit h e t a l (1994 ) suggeste d tha t thes e estimates o f sectio n remove d fro m th e Eas t Midlands Shel f derive d fro m AFT A wer e to o high by c. 1 km. Further work has been carrie d ou t recently to investigate various aspects o f these results . Thi s work has included reanalysis o f various sample s analysed b y Gree n (1989 ) t o incorporat e subsequent advances in understanding of annealing kinetics, as described above . New AFTA and VR dat a fro m Centra l Englan d (Gree n e t al 2001/7) documen t th e transitio n fro m inverte d basinal region s i n th e Eas t Midland s Shel f t o a stable platfor m settin g i n th e sout h (Midlan d Platform). AFT A reveal s tw o discret e coolin g episodes, i n Earl y Tertiar y tim e (beginnin g between 6 5 an d 6 0 Ma) an d Lat e Tertiar y tim e (beginning between 25 and 5 Ma). Early Tertiary palaeotemperatures fro m AFT A an d V R i n samples collected from outcro p define a consistent increas e fro m <50° C i n Lowe r Cretaceou s and Uppe r Jurassi c unit s i n th e S E t o aroun d 80-90 °C in Triassic an d older unit s in the NW. Results fro m th e Rufford- 1 wel l defin e a n Early Tertiar y palaeogeotherma l gradien t o f 40.5 °C km"1 (32-5 0 °C km'1 a t ±95 % confi dence limit s (c.l.) » compare d wit h a present-day gradient o f c . 30°Ckm~ 1 ), correspondin g t o deeper buria l b y 1450 m o f additiona l section, subsequently remove d b y Tertiar y erosio n (1.1-2.2km a t ±95 % c.L) . Thus , thes e Earl y Tertiary palaeotemperatures, as in NW England, also appea r t o reflec t a combinatio n o f deepe r burial an d elevate d basa l hea t flow . Geologica l considerations sugges t a maximu m overburden of 800-900 m abov e th e bas e o f th e Lia s sequence i n th e vicinit y o f Rugb y wher e th e Early Tertiar y palaeotemperatur e a t outcro p i s similar to that near the Rufford- 1 wel l site . Th e discrepancy betwee n stratigraphi c an d palaeo thermal reconstructio n of forme r buria l depths, often note d i n earlie r studies , remain s unresolved. The Lat e Tertiar y episod e i s muc h les s wel l constrained, bu t result s fro m Rufford- 1 ma y require betwee n 91 0 an d 1650 m o f erode d section. Thu s muc h o f th e tota l amoun t o f

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removed overburde n ma y hav e bee n remove d during Late Tertiary time. Results from the Apley Barn Borehole, i n the SW of the region, revea l a significantly differen t therma l history, involving Permian cooling , whic h probabl y reflect s th e protracted effect s o f Varisca n tectonism , an d a Late Tertiar y episod e characterize d b y a highly non-linear palaeotemperatur e profile , whic h probably reflect s loca l heatin g a s a resul t o f passage of hot fluids. Results from thi s borehole show no evidence of any Early Tertiary effects .

Timing of exhumation Some worker s hav e explicitl y questione d th e Early Tertiar y timin g o f coolin g reveale d b y AFTA i n N W an d Centra l Englan d (Hollida y 1993, 1999 ; McCulloch, 1994) . A s documente d by th e mor e recen t studie s describe d above , recent analyse s usin g lates t technique s provid e very tight definition of the onset of Early Tertiary cooling i n thes e regions , wit h th e interva l 65-60 Ma being the best available estimate. Although th e overal l duratio n o f exhumatio n that began in early Tertiary time cannot be tightly constrained fro m th e data , i t i s clea r tha t i n Central England , samples underwent a considerable amoun t o f Lat e Tertiar y coolin g an d i t remains possibl e tha t muc h o f th e exhumation actually occurre d withi n th e Lat e Tertiar y episode, a s suggeste d b y Japse n (1997) . Resol ution o f th e effect s o f discret e episode s o f Tertiary exhumation remains a major objective of continuing work in this region .

Mesozoic palaeo-thermal episodes recognized in Ireland and the Central Irish Sea Basin Integration o f AFTA and V R data from onshor e Ireland reveal s a comple x therma l history , characterized b y multipl e coolin g episode s o f late Carboniferous , Jurassic , earl y Cretaceous , early Tertiary an d late Tertiary ag e (Green e t al. 2000). Peak palaeotemperatures i n each episod e fall through time to produce an overall long-term cooling tren d sinc e lat e Carboniferou s time . Thermal history styles across the region are very similar, although the magnitude of peak palaeo temperatures i n individual episodes show s som e variation. The regional nature of all these palaeothermal episodes , an d thei r correlatio n wit h regionally significan t unconformities , suggest s that heating was due primarily to greater depth of burial, with subsequent cooling representing the progressive unroofin g o f th e presen t onshor e region sinc e late Carboniferous times. In Northern Ireland, explanations of early Cretaceous and

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early Tertiar y palaeotemperature s i n term s o f greater dept h o f buria l ar e mor e difficul t t o reconcile wit h geologica l evidence , an d heatin g as a resul t o f ho t flui d movemen t appear s mor e likely. Thi s applie s particularl y t o early Tertiar y effects, fo r whic h th e Tertiar y Igneou s Provinc e provides a read y explanation . Ove r th e entir e onshore region , maximu m maturit y level s i n Carboniferous an d olde r unit s wer e reache d a t

the en d o f Carboniferous time , an d preservatio n of hydrocarbon s t o th e presen t day , throug h several tectono-therma l episodes , appear s unlikely. Results fro m well s i n th e Centra l Iris h Se a Basin (CISB ) ar e dominate d b y th e Earl y Cretaceous episode , bu t als o revea l Earl y an d Late Tertiary cooling (Dunca n et al. 1998 ; Green et al . 200la) . Palaeotemperature s i n th e CIS B

Fig. 10. Summary of the main features o f the post-Palaeozoic exhumation history of the UK and Ireland regions, based on work discussed in the text. Several palaeo-thermal episodes of regional extent and variable intensity are recognized. Understanding the timing of these episodes within the regional geological context, and the variation in magnitude of individual episodes acros s the region is a vital ste p in understanding th e processes involved in producing these palaeo-thermal effects (burial , exhumation, changes in heat flow, fluid circulation, etc.), as well as being a critical aspect of understanding the petroleum systems of the region.

QUANTIFYING EXHUMATIO N USING AFT A AND VR

appear t o be due almost solel y t o deeper burial, with up to 3 km of sectio n remove d sinc e Early Cretaceous time . However , result s fro m wel l 42/21-1, locate d i n a S W extensio n o f th e St . George's Channe l Basin , ar e dominate d b y a n Early Tertiar y onse t o f exhumation , suggesting some degre e o f loca l structura l contro l o n discrete phases of exhumation. Results from other areas Studies i n region s suc h a s th e souther n North Sea, Fort h Approache s an d th e Mora y Firt h (some result s fro m whic h wer e reporte d b y Green e t al. (1995)) , a s wel l a s continuin g research wor k in th e Clevelan d Basi n (prelimi nary results reported by Green et al. (1993b)), all show that a wide region of the U K wa s affecte d by Earl y Tertiar y cooling . Result s fro m th e Scottish Highland s (Thomso n e t al . 1999b ) ar e very simila r t o thos e fro m onshor e Ireland , showing Earl y Cretaceous , Earl y Tertiar y an d Late Tertiar y coolin g episodes , whic h presumably agai n represen t th e effect s o f progressiv e exhumation. Interestingly , towards the Atlantic

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margin an d in th e Hebridean Basins , the effect s of Earl y Tertiar y exhumatio n diminis h (Gree n et al 1999 ) and the mid- to Late Tertiary period appears t o represent th e main onse t o f exhumation, with results from the Sea of Hebrides-1 well (aka LI34/5-1 ) definin g a n onse t o f coolin g between 45 and 20 Ma (Green et al 1999) , which may represen t a thir d Tertiar y episode , distinc t from Earl y and Late Tertiary episodes discussed thus far. Figure 1 0 summarize s th e mai n conclusion s that ca n b e draw n regardin g th e histor y o f Mesozoic an d Cenozoi c exhumatio n across th e UK region, o n the basis of the results discussed here.

Comments on mechanisms o f exhumatio n Regional events in NW Europe As discusse d above , result s fro m th e N W European Atlanti c margi n sho w consisten t evidence o f a t leas t thre e pulse s o f Cenozoi c exhumation (i n Early , mid - an d Lat e Tertiar y time), varying in magnitude across the region. In

Fig. 11. Palaeogene events in the UK region (after Ziegler 1990). The region where Early Tertiary palaeo-thermal effects are recognized from AFTA appears to link the developing Atlantic margin and Tertiary igneous province in the NW with the region to the SE characterized b y the initial stages of Alpine compression an d Palaeogene basin inversion. The spatial and temporal relationships betwee n these events is suggestive of some sort of causative link.

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considering likel y mechanisms, i t seems signifi cant that majo r cooling event s that appea r to be broadly synchronou s wit h thes e episode s hav e also been identified throughout the Arctic regio n including Svalbard (Blyth e & Kleinspehn 1998) , Alaska (O'Sulliva n e t al 1993 , 1995 ) an d Eas t Greenland (Thomson e t al. 1999a) . Results fro m recent unpublished studies in the Barents Sea and the North Slop e o f Alaska have emphasized th e synchroneity of Cenozoic event s in these regions and in NW Europe. This suggests that the driving mechanisms ar e trul y regional , an d ar e mor e likely t o b e relate d t o event s a t plat e margin s rather than more local processes such as igneous underplating. In particular , a s illustrate d i n Fig . 11 , Earl y Tertiary palaeo-therma l effect s identifie d fro m AFTA in Northern an d Central Englan d an d the Irish Se a are broadly synchronou s with rifting i n the Nort h Atlantic , initia l stage s o f Alpin e compression an d basi n inversio n throughou t NW Europe . Althoug h th e exac t relationship s between thes e processes remai n t o be resolved , the spatia l an d tempora l relationship s betwee n the palaeo-therma l effect s an d thes e othe r tectonic processe s sugges t som e sor t o f geneti c relationship. At a mor e loca l level , a s detaile d b y Gree n et al (2001b) , latest results from Central England suggest that at least in the UK region major Early Tertiary exhumatio n appear s t o b e limite d t o regions underlai n b y olde r Palaeozoi c basin s whereas region s overlyin g stabl e Palaeozoi c basement remaine d dorman t unti l Lat e Tertiar y time, whe n more regiona l exhumatio n occurred . Green et al. (200 Ib) have suggested this reflect s the preferentia l reactivatio n an d progressiv e locking o f the weaker basinal regions as a result of compressiona l event s a t plat e margins , culminating i n regiona l uplif t an d erosion . Th e relationship betwee n th e regiona l Tertiar y exhumation discusse d her e an d discrete Tertiar y basin inversion events in Southern England (e.g. Bray et al, 1997 ) als o merit s consideration , but is beyond the scop e of present discussion. Accelerated burial before uplift Reconstructed therma l and/o r buria l an d exhu mation historie s fo r area s affecte d b y significan t exhumation invariabl y sho w a n acceleratio n i n the rate of burial before the onset of exhumation (see examples given by Hillis 1991 ; Green et a l 1995; Japse n 1997) . Althoug h som e hav e suggested tha t thi s results fro m usin g unreason able amount s of additiona l burial , a s discusse d above, w e believe tha t thi s i s a rea l effect , an d that th e accelerate d buria l an d subsequen t

exhumation ar e linked . I f thi s i s true , an y successful mode l fo r th e underlyin g mechanism of regiona l exhumatio n episode s suc h a s discussed her e shoul d explai n no t onl y th e amounts and regional extent of removed section, but also thi s prior burial phase. Repeated cycles of burial and exhumation Results from onshore Ireland (Green et al. 2000), the Central Iris h Se a Basin (Green e t al 200la ) and th e Scottis h Highland s (Thomso n e t a l (1999b) strongl y sugges t th e occurrenc e o f repeated cycle s o f buria l followe d b y exhumation, wit h th e overal l magnitud e o f bot h processes decreasin g ove r time . Result s fro m other area s sho w therma l historie s tha t ca n b e interpreted i n simila r term s (e.g . Eas t Iris h Se a Basin, Gree n e t a l 1997) . I n thi s respect , th e phases o f accelerate d buria l discusse d i n th e preceding sectio n becom e par t o f a progressive sequence o f events , diminishin g throug h time , giving th e appearanc e o f a dampe d oscillation . Although we have no mechanisms for explaining these observations , thei r ubiquitou s recognition suggests som e mechanisti c control , an d w e suggest tha t thi s provide s a fruitfu l avenu e fo r consideration i n developin g possibl e model s t o explain the sor t of effect s discusse d here . Comparison with results from other regions Consideration o f dat a fro m othe r region s showing simila r style s of palaeo-thermal effect s is als o usefu l i n considerin g mechanism s fo r exhumation. Well-documented examples include Cretaceous effect s i n Souther n Afric a an d S W Brazil (Gallagher & Brown 1999, and references therein), N E US A (Mille r & Dudd y 1986 ) an d SE Australia (Moore et al 1986 ; Dumitr u e t a l 1991; Dudd y & Gree n 1992 ; O'Sulliva n e t a l 1995b, 1996) . I n al l thes e area s a n associatio n between th e onse t o f exhumatio n an d con tinental riftin g o r separatio n a t adjacen t plat e margins is evident (also in NW Europe, Fig. 11) . Although detail s ar e ye t t o b e resolved , thes e empirical observations strengthen the suggestion that events at plate margins exert a key influenc e on th e processe s responsibl e fo r regiona l exhumation. Palaeo-therma l effect s i n thes e regions als o appea r t o requir e accelerate d rate s of burial before the onset of exhumation, further emphasizing previous comments on this topic. In most of these areas, particularly in Southern Africa an d i n S E Australia , palaeogeotherma l gradients at the palaeo-thermal maximum are not well constrained, because of the lack of data from vertical sections , an d therefor e amount s o f

QUANTIFYING EXHUMATION USING AFTA AND VR exhumation ar e stil l onl y poorl y understood . In this context , i t i s wort h notin g tha t th e view s expressed b y Gallaghe r & Brow n (1999 ) an d Gunnell (2000 ) tha t hea t flo w ca n generall y b e treated a s constan t throug h tim e an d therefor e that all the observed cooling can be interpreted a s being du e to denudation (exhumation) is at odds with result s fro m N W Englan d (Gree n e t al. 1999; Gree n 2001 ) an d S E Australi a (e.g . Duddy 1994 , 1997) , whic h revea l consisten t evidence fo r elevate d basa l hea t flo w i n locations wel l separate d fro m th e site s o f rifting a t th e appropriat e time . Amount s o f removed sectio n (exhumatio n o r denudation ) based o n th e assumptio n o f constan t hea t flo w should therefor e alway s be regarded a s likely t o overestimate th e tru e amount, and estimates ar e likely t o b e realisti c onl y wher e estimatio n i s based o n vertica l sequence s o f sample s tha t directly constrain t th e palaeo-hea t flo w o r palaeogeothermal gradient . AFTA® i s a registere d trademar k o f Geotrac k International Pt y Ltd.

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CARLSON, W.D . 1990. Mechanism s an d kinetic s o f apatite fission-trac k annealing . American Mineralogist, 75, 1120-1139. CARLSON, W.D. , DONELICK, R.A . & KETCHAM , R.A. 1999. Variabilit y of apatite fission-trac k annealin g kinetics: I . Experimenta l results . American Mineralogist, 84 , 1213-1223. CARR, A . 1999 . A vitrinite reflectanc e kineti c mode l incorporating overpressur e retardation. Marine and Petroleum Geology, 16, 355-377 . CARR, A . 2000 . Suppressio n an d retardatio n o f vitrinite reflectance , Par t 1 . Formatio n an d significance fo r hydrocarbo n generation . Journal of Petroleum Geology, 23, 313-343. CHADWICK, R.A. , KIRBY, G.A . & BAILY , H.E . 1994. The post-Triassic structural evolution of north-west England an d adjacen t parts o f th e Eas t Iris h Se a Basin. Proceedings o f th e Yorkshire Geological Society, 50, 91-102. COPE, J.C.W . 1994 . A lates t Cretaceou s hotspo t an d the southeasterl y til t o f Britain . Journal o f th e Geological Society, London, 151, 905-908 . CORRIGAN, J.D . 19910 . Therma l anomalie s i n th e Central Indian Ocean: evidence for de-watering of the Benga l Fan. Journal of Geophysical Research, 96, 14263-14275. CORRIGAN, J.D . I99lb. Inversio n o f apatit e fissio n track data for thermal history information. Journal of Geophysical Research, 96, 10347-10360 . CROWLEY, K.D. , CAMERON, M . & SCHAEFFER , R.L. 1991. Experimental studies of annealing of etched fission track s i n fluorapatite . Geochimica e t Cosmochimica Acta, 55, 1449-1465. DEMING, D . 1994 . Overburden rock, temperature an d heat flow. In: MAGOON , L.B . & Dow , D.G. (eds) The Petroleum System —From Source to Trap. American Associatio n o f Petroleu m Geologists , Memoir, 60 , 165-186. DONELICK, R.A. , KETCHAM, R.A . & CARLSON , W.D. 1999. Variabilit y of apatite fission-trac k annealin g kinetics: II . Crystallographi c orientatio n effects . American Mineralogist, 84, 1224-1234. DORE, A.G., LUNDIN, E.R., JENSEN, L.N., BIRKELAND, 0., ELIASSEN , P.E. & FICHLER, C . 1999 . Principal tectonic event s i n th e evolutio n o f th e northwes t European Atlanti c margin . In : FLEET , A.J . & BOLDY, S.A.R . (eds ) Petroleum Geology o f North West Europe, Proceedings of the 5th Conference. Geological Society, London , 41-61. Dow, W . 1977 . Keroge n studie s an d geologica l interpretations. Journal o f Geochemical Exploration, 7, 79-99. DUDDY, I.R . 1994 . Th e Otwa y Basin : thermal , structural, tectoni c an d hydrocarbo n generatio n histories. NGMA/PESA Otway Basin Symposium, Extended Abstracts, 14 , 35-42 . DUDDY, I.R. 1997. Focusing exploration i n the Otway Basin: understandin g timin g o f sourc e roc k maturation. APPEA Journal, 37, 178-191 . DUDDY, I.R . & GREEN , P.P . 1992 . Tectonic develop ment o f th e Gippslan d Basi n an d Environs : identification o f key episodes using apatite fission track analysi s (AFTA) . Gippsland Basin Symposium, Joint AusIMM (Melbourne Branch)-PESA

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calibration. Earth an d Planetary Science Letters, 89, 335-352. GREEN, PF . 1989 . Therma l an d tectonic histor y of the East Midland s shel f (onshor e U.K. ) and surrounding region s assesse d b y apatit e fissio n trac k analysis. Journal o f th e Geological Society, London, 146 , 755-773 . GREEN, PF . 2001 . Earl y Tertiar y palaeo-therma l effects i n Norther n England : reconcilin g result s from apatit e fission track analysis with geological evidence. Tectonophysics, 349 , 131-144 . GREEN, P.P. , DUDDY , I.R . & BRAY , RJ . 1993a . Earl y Tertiary heatin g i n Northwes t England : fluid s o r burial (o r both?) (extended abstract). In: PARNELL, J., RUFFELL , A.H . & MOLES , N.R. (eds) Geofluids '93: Contributions to an International Conference on Fluid Evolution, Migration and Interaction in Rocks, 119-123 . GREEN, PR , DUDDY , I.R . & BRAY , RJ . 1995 . Applications o f therma l histor y reconstructio n in inverted basins. In: BUCHANAN, J.G. & BUCHANAN , P.G. (eds ) Basin Inversion. Geologica l Society , London, Specia l Publications , 88 , 148-165 . GREEN, P.F., DUDDY, I.R. & BRAY, RJ. 1997 . Variation in therma l history styles around the Iris h Se a an d adjacent areas : implication s fo r hydrocarbo n occurrence an d tectoni c evolution . In: MEADOWS , N.S., TRUEBLOOD , S., HARDMAN, M. & COWAN , G. (eds) Petroleum Geology o f th e Irish Se a an d Adjacent Areas. Geologica l Society , London , Special Publications , 124 , 73-93 . GREEN, P.P., DUDDY, I.R., BRAY, R.J., DUNCAN, W.I. & CORCORAN, D . 2001« . Th e influenc e o f therma l history on hydrocarbon prospectivity in the Central Irish Sea Basin . In: SHANNON , P.M., HAUGHTON, P. & CORCORAN , D . (eds ) Petroleum Geology o f Ireland's Offshore Basins. Geologica l Society , London, Specia l Publications, 188, 171-188 . GREEN, P.F., DUDDY, I.R., BRAY, RJ. & LEWIS , C.L.E. 1993&. Elevate d palaeotemperatures prior t o early Tertiary coolin g throughou t th e U K region : implications fo r hydrocarbo n generation . In : PARKER, J.R. (ed.) Petroleum Geology of Northwest Europe: Proceedings of the 4th Conference. Geological Society , London , 1067-1074 . GREEN, P.P. , DUDDY , I.R. , GLEADOW , AJ.W . & LOVERING, J.F . 1989b . Apatit e fissio n trac k analysis a s a palaeotemperatur e indicato r fo r hydrocarbon exploration . In : NAESER , N.D . & MCCULLOH, T . (eds ) Thermal History o f Sedimentary Basins —Methods and Case Histories. Springer, Ne w York, 181-195 . GREEN, P.P. , DUDDY , I.R. , GLEADOW , A.J.W. , TINGATE, PR . & LASLETT , G.M . 1985 . Fission track annealin g i n apatite : trac k lengt h measure ments and the form of the Arrhenius plot. Nuclear Tracks, 10 , 323-328. GREEN, P.P. , DUDDY , I.R. , GLEADOW , A.J.W. , TIN GATE, PR . & LASLETT , G.M . 1986 . Therma l annealing o f fissio n track s i n apatit e 1 . A qualitative description. Chemical Geology (Isotope Geoscience Section), 59 , 237-253. GREEN, P.F. , DUDDY , I.R. , HEGARTY , K.A . & BRAY , RJ. 1999 . Earl y Tertiar y hea t flo w alon g the U K

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Sonic velocit y analysis o f the Tertiary denudatio n of the Irish Sea basin PHILIP D. WARE 1'2 & JONATHAN P . TURNER1 1 University of Birmingham, School of Earth Sciences, Birmingham B15 2TT, UK (e-mail: j.p. turner® bham. ac. uk) ^Present address: Kerr McGee Oil Crawpeel Road, Aberdeen AB12 3LG, UK Abstract: Interactio n between uplif t relate d t o the Cretaceous-Paleocene opening of the North Atlantic , Neogen e shortenin g (basi n inversion ) an d Pleistocen e glacio-isostas y i s illustrated by the complex denudation pattern o f Britain; such denudation i s greatest over the submergent Eas t Irish Sea basin, som e 500 km from the Atlantic margin . This paper reports on analysis of sedimentary porositie s using sonic velocity log s from 42 wells in the East Irish Sea basin . W e present a ne w ma p showin g th e variatio n i n exhumatio n magnitud e a t th e uppermost Mesozoi c unconformit y (i.e . thicknes s o f denude d Mesozoi c an d Cenozoi c sedimentary rocks), today burie d beneat h a thin veneer of Pleistocene sediment. It indicates that exhumatio n i s mostl y < 1500m (632-213 2 m; mea n standar d deviatio n 407m) , les s than denudation results obtained from vitrinite reflectance and apatite fission-track data. The map als o reveals substantia l variatio n i n exhumatio n ove r shor t distances , ofte n betwee n adjacent well s sited on opposing walls of individual faults . This is interpreted in terms of the influence o f Neogen e basi n inversio n o n th e exhumatio n o f th e EISB . Th e rol e o f lat e Tertiary tectonics in western U K exhumation i s therefore discussed .

Estimates o f exhumatio n (i.e . thicknes s o f denuded overburden ) var y widel y accordin g t o the method employed to evaluate it (see Dore & Jensen 1996 ; Japse n & Chalmer s 2000) . Th e main purpos e o f thi s pape r i s t o tes t existin g denudation magnitude s obtaine d fro m apatit e fission-track (AFT ) analysi s an d vitrinite reflec tance (VR ) dat a fro m th e Eas t Iris h Se a basi n (EISB) usin g soni c velocitie s (SV ) logge d i n petroleum exploratio n wells . Th e long wavelength (> 2 km) form of vertical SV profiles responds chiefl y t o compaction-drive n porosit y reduction. Consequently , i t i s a particularl y effective measur e of the former maximum burial depth o f exhume d sedimentar y succession s i n basins where transient heating episodes elevate d thermal burial proxies, such as AFT and VR data, without concomitant burial increases. The mai n processe s drivin g exhumatio n ar e thermal-isostatic effect s relate d t o continental extension, orogeni c crusta l shortenin g an d epeirogeny - uplif t o f broa d region s o f continental interior s drive n by plume s (Nadi n et al 1995) , magmati c underplating (Brodi e & White 1994) , mantl e delaminatio n (Plat t & England 1993) , post-glacia l isostas y (Lambec k 1991), intra-plat e stres s (Cloeting h 1988) , etc . Exhumation differs fro m uplift , whic h describes

vertical movement relative to the geoid (England & Molna r 1990) . Becaus e o f th e proble m i n constraining an absolute datum, very few studies are abl e t o measur e ancien t uplif t pe r s e (e.g . Abbott e t al . 1997) . Furthermore , i n man y instances, suc h a s buria l histor y modellin g o f petroleum source rocks, exhumation is actually a far more useful parameter to constrain because it leads directly to cooling an d lithostatic pressur e release, wit h attendan t implication s fo r pet roleum generatio n and retention. In th e absence of any erosion, uplift on its own results in neither cooling nor pressure decrease. The EIS B i s a n intra-cratoni c sedimentar y basin whos e Cenozoi c evolutio n record s th e influence o f Cretaceou s riftin g o f th e Nort h Atlantic an d associate d Paleocen e break-up , Oligo-Miocene shortenin g (basi n inversion ) and Pleistocen e glacio-isostasy . Fulle r discus sion o f it s geologica l an d tectoni c evolutio n is given b y Jackso n & Mulhollan d (1993) , Knip e et al . (1993) , Jackso n e t al . (1995) , Cop e (1998), Maingar m e t al . (1999) , an d references therein. Denudation analysi s ha s attracte d particula r interest i n th e gas-pron e EISB , wher e sourc e rock buria l histor y modellin g i s hindere d b y uncertainty ove r th e thicknes s o f th e erode d

From: DORE , A.G., CARTWRIGHT , J.A., STOKER , M.S. , TURNER , J.P. & WHITE , N. 2002 . Exhumation of th e North Atlantic Margin: Timing, Mechanisms and Implications for Petroleum Exploration. Geological Society , London, Special Publications , 196, 355-370 . 0305-8719/02/$15.00 © The Geological Societ y of London 2002.

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Jurassic-Cretaceous section . Give n th e pri mary relianc e o f hydrocarbo n generatio n o n a long-term ris e i n sourc e roc k temperature , cooling (and , t o a lesse r extent , pressur e release) durin g exhumatio n lead s directl y t o arrested generation . Conversely , a potentia l source roc k toda y lyin g withi n th e appropriat e depth windo w fo r hydrocarbo n generatio n may no t b e effectiv e if , befor e denudation , its generativ e capacit y wa s exceede d durin g deeper burial . Exhumatio n analysi s i n th e EISB ha s bee n a particularl y livel y topi c since th e wor k o f Lewi s e l al. (1992) , whos e results fro m AF T analysi s indicate d i n exces s of 3k m o f sedimentar y sectio n denude d fro m a broa d regio n centre d o n th e presen t Englis h Lake Distric t an d contiguou s Iris h Sea . Mor e recent AF T dat a confir m a majo r coolin g event i n th e EIS B startin g a t c . 6 0 Ma an d decelerating int o Neogen e tim e (Gree n e t al . 1997; Dunca n e t al . 1998 ) bu t the y sugges t that 3k m exhumatio n represent s a n uppe r bound. A majo r difficult y associate d wit h modellin g exhumation usin g V R an d AF T data , an d th e rationale fo r thi s work , i s thei r susceptibilit y t o transient hydrotherma l effect s suc h a s loca l igneous intrusio n an d hydrotherma l flu x (e.g . Green et al. 1993 ; Sibson 1995 ; White & Morton 1995; Hun t 1996 ; Gree n et al 2001) . Because of the increased geotherma l gradien t accompanyin g these transient heating events, they anneal fission tracks an d increas e V R withou t a concomitan t increase in buria l depth . Consequently , VR and AFT dat a tha t hav e bee n affecte d b y transien t heating events will yield exaggerated denudation magnitudes. There ar e tw o reason s t o suppos e tha t thi s study are a wa s expose d t o transien t therma l effects durin g Late Cretaceous and Tertiary time. First, th e Iceland Plume , a mantle up welling that affected a region of the North Atlantic continents of 1000k m width , generate d transien t heatin g through (1 ) extensiv e volcanis m o f th e Nort h Atlantic regio n (Whit e & McKenzi e 1989 ) an d (2) basaltic underplatin g of the continenta l crus t conjectured beneat h a wid e regio n extendin g well beyon d th e are a o f extrusiv e volcanis m (Brodie & Whit e 1994) . Second , hydrotherma l fluid flow during Oligo-Miocene basin inversio n is implicate d i n a Lat e Tertiar y (c . 2 0 Ma) cooling even t recognize d fro m combine d AF T and flui d inclusio n studie s (Atlanti c margins , Green e t al . 1999 ; Parnel l e t al . 1999 ; EISB , Hardman e t al . 1993 ; souther n England , Gree n et al . 2001) . Th e multipl e increment s o f faul t reactivation tha t typif y basi n inversio n episode s will b e characterize d b y flushin g o f an y

overpressured fluid s alon g fault s an d throug h their wal l rock s durin g repeate d cycle s o f sea l breaching an d repai r (Sibso n 1987) . Thi s faul t valving ca n transpor t extraordinar y volume s of hot flui d man y time s th e volum e o f th e por e space i n whic h the faul t valvin g is focuse d (e.g . Sibson 1995 ) an d i t ma y lea d t o substantia l elevation of wall rock temperature (e.g. Andrews etal 1996) .

Regional settin g an d stratigraph y The Irish Se a basins comprise a linked system of Mesozoic-Cenozoic depression s that , i n th e Cardigan Ba y basin, attai n a post-Carboniferous sedimentary thicknes s i n exces s o f 12km . Th e Mesozoic fil l o f th e Iris h Se a basin s display s a continental margin-typ e subsidenc e signatur e (Welch & Turner 2000) interrupte d by Paleocene epeirogenic uplift , generall y linke d t o th e formation o f th e Icelan d Plum e (Brodi e & White 1994) . However , th e Iris h Se a are a als o displays clear evidence of shortening of formerly extensional basin s (basi n inversion ) in respons e to African-Europea n plat e collisio n durin g late Oligocene an d Miocene tim e (e.g . Ziegler 1987 ; Roberts 1989) . Miocen e basi n inversio n le d t o only mino r modificatio n o f th e Mesozoi c extensional faul t geometr y (Tucke r & Arte r 1987). The principa l manifestations o f inversion in thi s stud y are a an d contiguou s basin s ar e transpressional reactivatio n o f steep , NW-trend ing fault s (Turne r 1997 ) an d thickenin g o f th e Mesozoic basi n fill by pur e shea r (cf . Eisenstadt & Withjac k 1995) . Althoug h i t i s difficul t t o isolate Miocen e basi n inversio n from th e effect s of the Paleocene Icelan d Plume, Cenozoi c uplif t of th e Celti c Se a basins , t o th e sout h o f th e present stud y area, was responsible fo r exhumation o f betwee n 600 m (Tucke r & Arte r 1987 ) and 2500 m (Menpe s & Hilli s 1995 ) o f Uppe r Mesozoic an d Cenozoic strata . The petroleu m syste m o f th e EIS B i s mainly confined t o Triassi c rock s an d fe w well s penetrate deepe r tha n this . Th e Triassi c succes sion ca n b e subdivide d int o th e sandstone-ric h Sherwood Sandston e Formatio n (Scythian : c. 1450 m thic k i n th e souther n Eas t Iris h Se a basin (Jackso n & Mulhollan d 1993) ) succeeded by the finer-grained Mercia Mudston e Formation (Anisian-Norian; > 150 0 m thick in the southern EISB (Jackso n & Mulhollan d 1993)) . Th e Sherwood Sandston e i s a distinctiv e red , fluvial-aeolian arkosi c sandston e sequenc e tha t constitutes the principal reservoir objective in the Irish Sea . Th e Merci a Mudston e consists of re d mudstone an d siltston e wit h sal t unit s attainin g up t o 450 m thicknes s (Jackso n & Mulhollan d

TERTIARY DENUDATION OF IRISH SEA BASIN

1993). Non e o f th e well s use d i n th e analysi s reported her e drille d throug h sal t o f >50 m thickness, which , wher e encountered , wa s excluded fro m th e exhumation calculations . Methods of sonic velocity analysis Theory and mechanical control over velocity Until recently , obtainin g exhumatio n fro m S V has been carrie d ou t by using the SV log to plot interval velocit y vs . midpoin t dept h o f fine grained lithofacie s units . Through it s systemati c relation wit h porosity (Wylli e et al 1956) , sonic velocity will exhibit a progressive reduction with mean effectiv e stres s (^HYDROSTATI C ~ ~ ^F ; where ^HYDROSTATI C i s hydrostati c stres s an d PP i s flui d pressure ; se e Goult y (1998) , Gile s et al (1998 ) an d reference s therein) . Conse quently, i n a uniforml y pressure d sedimentar y sequence, plot s o f S V vs . midpoin t dept h wil l highlight anomalousl y 'fast ' (low-porosity ) successions wher e rock s hav e bee n exhume d from a formerl y greate r buria l depth . Cross plotting th e exhumatio n result s derive d fro m different part s o f th e stratigraph y allow s fo r checks o n consistenc y (e.g . Menpe s & Hilli s 1995). I n attemptin g t o comput e absolut e estimates o f denudatio n magnitude , tw o

357

problems arise . First, because thi s approac h ha s no way of defining the 'normal' velocity vs. depth trend for an unexhumed succession , al l exhumation estimate s wil l onl y b e relativ e t o th e leas t exhumed log (i.e. that succession which, after the denudation episode(s) , remaine d a t or closest t o its maximum burial depth). Second, conventional SV analysis relies o n individual lithofacies units exhibiting near-identica l compactio n behaviou r across th e are a o f interes t (e.g . Issle r 1992) . In practice, i t ma y b e difficul t t o discriminat e lithological variatio n fro m short-wavelengt h changes in exhumation. In commo n wit h othe r approache s t o S V analysis, this study uses the exponential decreas e of porosit y wit h buria l dept h t o comput e exhumation magnitude: where (/> i s porosit y a t dept h jt , o i s surfac e porosity o f uncompacte d sedimen t an d b i s compaction coefficien t pe r uni t litholog y (Ath y 1930; Rube y & Hubbert 1959) . P-wave velocity , as recorde d b y soni c logs , i s a widel y use d measure of porosity : where (f> i s porosity, an d Af

Fig. 1 . Scheme of the method adopted in this study, modified fro m Magar a (1976). The pre-denudation condition is given by the combined continuous and dashed line extending, in this study, to 76 m beneath the depositional surface.

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P. D. WAR E & J. P . TURNE R

respectively, th e soni c lo g (measured) , roc k matrix an d interstitia l flui d transi t time , th e reciprocal o f velocit y (Wylli e e t al. 1956) . Assuming friction , attenuatio n an d frequenc y

dependence to be negligible, P-wave velocity can be defined in terms of the dynamic elastic moduli (Gassmann 1951) . O f these , th e shea r modulus and roc k framewor k bul k modulu s (i.e . bulk

Fig. 2. Data from soni c logs and best-fit transit time vs. depth curves (pale grey ) for the four most extensive shale sections i n the study area. O n the y-axis, positive an d negative values indicate depths above an d below the present sea floor , respectively .. Exponentia l deca y constant s an "d shif t constant" s fo r eac h wel l are : fo r 110/3-2 , " " "f " - 0.00037 ft" an d 63 JJLS ft" 1; fo r 110/3-4 , -0.002 7 ft"1 an d 59 JJLS ft" 1 ; fo r 110/6-1, -0.00035ft" 1 an d 61 ft" 1 ; fo r 110/11-2 , -0.00056ft" 1 an d 6 5 JJL S ft" 1.

TERTIARY DENUDATION OF IRISH SEA BASIN modulus o f th e minera l grains ) ar e strongl y dependent o n porosity an d ar e therefor e subjec t to potentially rapi d chang e wit h burial depth . Circumstances i n which equation (2 ) does no t work well include: (1) compaction retardation as a resul t o f intensiv e cementatio n and/o r over pressure (Erickso n & Jarrar d 1998) ; (2 ) poros ities greate r tha n c . 25 % (Falve y & Middleto n 1981; Serr a 1984) ; (3 ) non-aqueou s pore fluid , especially gaseou s hydrocarbo n (Magar a 1978) ; (4) microcrack s (Erickso n & Jarrar d 1998 ; Dewhurst e t al. 1999) . Operators ' lo g inter pretations wer e use d t o identif y zone s o f intensive cementation , overpressur e an d anom alously hig h porosit y (non e encountered) , an d

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non-aqueous por e fluid . Dat a coverin g thes e intervals, whos e porosity-dept h behaviou r i s likely t o be unpredictable, wer e removed durin g the initial phase of editing of logs before analysi s of the soni c logs . Account wa s als o take n o f evaporiti c salt , which wa s edite d ou t fro m ou r datase t befor e analysis, an d poroelasticit y durin g overburde n removal. I n situ stres s relaxatio n generate s an d opens microcrack s suc h tha t initia l microcrac k porosities o f <0.5 % ar e sufficien t t o caus e pressure-dependent velocit y variation s o f 5-50% (Erickso n & Jarrar d 1998) . Poroelasti c rebound durin g denudatio n i s a measure o f thi s microcrack porosit y chang e an d ma y caus e a s

Fig. 3. Comparison of transit time vs. depth curves for the four wells in Fig. 2, before (top left) and after (top right) their normalization . Lowe r grap h summarize s th e characteristic s o f th e resultan t supercurv e derive d fro m 1 averaging ~ the normalize d curve s show n (supercurv e exponential decay constan t —0.00034ft" ; shif t constan t

Table 1 . Summary o f denudation magnitudes an d statistical data for wells i n the East Irish Sea basin whose locations ar e given i n Figs 4 an d 6

Well number

TVDto unconformity (m)

110/2-1 110/2-2 110/2-3 110/2-4 110/2-5 110/2-6 110/2-7 110/2-8 110/3-1 110/3-2 110/3-3 110/3-4 110/6-1 110/7-1 110/7-2 110/7-3 110/7-4 110/7-5 110/7-6 110/8-2 110/9-1 110/10-1 110/11-1 110/11-2 110/11-3 110/12-1 110/12-2 110/12-3 110/12-4 110/13-1 110/13-2 110/13-5 110/13-7 110/13-8 110/13-15 110/18-1

64 66 61 62 65 113 69 72 62 80 98 22 80 66 88 136 91 130 180 99 71 111 77 75 37 65 71 71 77 61 61 61 61 73 66 59

Water depth (m)

32 30 29 32 36 34 34 36 29 22 24 22 13 32 37 28 34 42 41 28 10 20 43 44 46 34 33 38 40 27 33 28 33 29 26 19

KBE(m TVD)

32 36 35 30 30 31 35 37 33 27 37 29 34 34 36 37 32 26 31 34 37 35 34 31 37 31 38 34 37 37 35 29 29 30 34 34

Minimum ET before correction (m)

Maximum ET before correction (m)

SD (m)

(m)

95% CI(m)

Average ET (m)

Correction (m)

Corrected E, (m)

132 -85 -315 -242 -259 -229 -305 -333 -515 -396 40 -526 -65 27 201 635 -307 445 673 87 374 310 484 217 -73 602 746 430 772 237 814 157 599 260 369 270

3341 1695 4334 4047 2953 2682 5597 4324 2649 5957 3889 4372 5347 3107 2108 2565 2763 3560 2578 2751 7598 3304 2896 5698 5156 3696 3372 1380 5735 4461 5299 3008 3171 4029 5951 5941

442 307 554 454 429 315 703 468 465 425 442 380 374 333 295 247 363 351 403 418 467 335 302 382 399 341 345 204 361 373 515 326 342 413 738 671

888 444 670 799 626 522 1003 658 498 577 722 515 858 811 714 1029 944 1051 1219 711 1314 909 1037 961 797 1256 1295 713 1395 1340 1663 1161 1278 956 1519 1944

18.1 15.9 20.3 13.8 16.4 12.3 27.3 17.7 16.4 13.1 20.0 13.2 11.7 13.0 19.4 16.3 15.4 12.4 23.4 16.6 18.8 15.7 23.4 12.0 13.1 13.4 34.8 12.9 11.2 18.8 42.2 13.2 13.0 16.2 26.0 21.4

952 510 731 861 691 635 1072 730 560 657 820 537 938 877 802 1165 1035 1181 1399 810 1385 1020 1114 1036 834 1321 1366 784 1472 1401 1724 1222 1339 1029 1585 2003

140 142 140 138 141 141 146 148 138 126 137 127 123 142 149 141 143 145 148 139 123 132 153 152 159 141 147 148 153 139 144 134 139 136 136 129

1092 652 871 999 832 776 1218 878 698 783 957 664 1061 1019 951 1306 1178 1326 1547 949 1508 1152 1267 1187 993 1462 1513 932 1625 1540 1868 1356 1478 1165 1721 2132

TERTIARY DENUDATIO N O F IRISH SEA BASIN 36

1

much a s 7 % porosit y chang e followin g exhumation (Hamilto n 1976) . The chie f contro l ove r soni c velocit y i n undeformed sedimentar y succession s is therefore burial-controlled mechanica l compactio n a s a result o f porosity reduction : where At is sonic transit time at depth x, an d Ato is surface transit time of uncompacted sediment . However, althoug h Athy' s La w (1 ) correctl y predicts negligibl e porosit y a t depth , soni c velocity i n a totall y compacte d roc k wil l equa l that o f th e roc k matri x (Heasle r & Kharitonov a 1996). Consequently , th e correc t functiona l relationship betwee n soni c transi t tim e an d depth is where c i s soni c transi t tim e o f th e roc k matrix, o r shif t constant . Typica l sedimentar y rock matri x transi t time s rang e betwee n 128m s m-1 (c . 7800 m s- 1 ) fo r dolomite s an d 22 3 m s m-1 (c . 4500 m s- 1 ) fo r shale s (Schlumberge r 1989). O n a linea r plo t o f soni c transi t tim e vs . depth, the matrix transit time approximates t o the asymptote o f th e curve , which generall y flatten s out belo w depth s o f 2500m , wher e porosit y approaches zero .

Curve-fitting, geological error and derivation of exhumation To comput e magnitud e o f denudatio n fro m S V data, w e modif y th e schem e o f Magar a (1976 ) and Heasle r & Kharitonov a (1996) , i n whic h a compaction curv e i s fitte d throug h logarithmi cally transforme d transi t tim e vs . dept h dat a (Fig. 1) . B y varyin g th e shif t constan t ( c i n equation (4)) , averag e absolut e valu e an d roo t mean squar e ar e use d t o optimiz e th e fi t o f th e compaction curv e t o th e soni c data . Thi s stud y follows Heasle r & Kharitonov a (1996 ) i n assuming tha t porosit y reductio n wit h dept h i n a heterolithi c sedimentar y sequenc e ca n b e described b y a single , averag e compactio n coefficient ( b i n equation s (1) , (3 ) an d (4) ; se e also Steckler & Watts 1978 ; Tosaya & Nur 1982 ; Castagna e t a l 1985 ; Ha n e t a l 1986 ; Mario n etal 1992) . The best-fi t transi t tim e vs . dept h curv e i s extrapolated abov e th e erosio n surfac e (i.e . unconformity o r th e presen t basi n floor ) t o th e level a t whic h transi t tim e equal s tha t o f uncompacted sedimen t (Fig . 1) . Transi t time s of uncompacted sedimen t wil l vary according t o

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Fig. 4 . Variation i n denudation relativ e to the present sea floor at 42 well site s for the East Iris h Se a basin, eac h identified by quad, block and well number. Data for 36 of these wells are shown in the graphs. Bold lines give the traces of principal faults ; inset show s the location of the study area (boxed) wit h exhumed Triassic sedimentar y basins shaded grey. The diameter of each circle is proportional t o mean total exhumation, given also in bold in the corresponding graph . Soni c velocitie s ar e dar k grey ; pale gre y lin e i s th e velocity vs . dept h supercurv e in the unexhumed case. Vertical difference i n depth (kilometres o n the left-han d v-axis , thousands of feet o n the right ) between th e supercurv e an d it s correspondin g soni c datapoin t indicate s exhumatio n magnitude for tha t point. Horizontal scatte r i n soni c datapoint s (i n JJL S fiT 1) provide s a qualitativ e measure o f lithologica l variatio n an d noise in the data .

TERTIARY DENUDATION OF IRISH SEA BASIN bulk sediment porosity , water saltiness, tempera ture an d pressure . Followin g th e approac h o f Magara (1976) , a surfac e transi t tim e o f 59 1 JUL S iri (ISOfJi s ft" 1, soni c velocit y 1695 m s" 1) is adopted. The geological erro r i s define d her e a s geological erro r = poroelasticit y — temperatur e effects . Hamilton (1976 ) demonstrate d tha t poroelasti city can lead to absolute porosity change of up to 7%. Convertin g thi s int o transi t tim e give s a reduction o f 12. 6 JJL S ft" 1 fo r unburie d sediment , producing a maximum underestimate of denudation i n ou r dat a o f som e 155 m (510ft) , depending o n th e deca y constan t o f th e compaction curve . P-wav e velocitie s i n low porosity sediment s increas e b y u p to 1.7 % for a 100 °C ris e i n temperatur e (Timu r 1977) . Assuming a geotherma l gradien t o f 3 3 °C

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km \ thi s temperature-dependen t velocit y behaviour wil l lea d t o underestimatio n o f denudation b y u p t o 40 m (129ft ) i n ou r dataset, agai n dependin g o n th e deca y constant. In thi s stud y th e maximu m tota l geologica l error i s therefor e + 115 m (+37 7 ft). Give n that geologica l error s ar e positive , al l ou r exhumation value s wil l b e underestimates . Th e small magnitud e o f geologica l erro r wit h respect t o th e tota l denudatio n value s mean s that the y hav e bee n ignore d i n ou r subsequen t interpretation. Only nin e o f th e well s analyse d contai n sufficient thicknes s o f shale y sectio n and/o r depth o f wel l penetratio n t o produc e a statisti cally acceptabl e populatio n o f soni c dat a t o which a compaction curve could be fitted (Fig. 2). The solutio n wa s t o compar e thes e dat a wit h a 'supercurve', a generic compactio n curv e for the EISB derived fro m averagin g normalize d transit time vs . dept h curve s fro m th e fou r mos t extensive availabl e shal e sections . Thi s stud y

Fig. 5 . Compariso n o f denudatio n magnitude s compute d fro m wel l 110/3- 2 obtaine d b y comparin g soni c velocities in its extensive Mercia Mudstone succession with the supercurve from thi s study, and transit time vs. depth curve s derived fro m normalize d best-fi t curve s through 110/3- 2 and the souther n North Se a Rotliegende Brockelschiefer Shale curve of Marie (1975). The pale grey polka-dots indicate the transit time-depth pairs used by Colter (1978 ) to compute his exhumation magnitude for the East Irish Se a basin.

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uses a supercurv e derive d fro m soni c log s fro m wells 110/3-2 , 110/6-1 , 110/11- 2 an d 110/3- 4 (Fig. 2) , eac h comprisin g mor e tha n 8000f t (2438m) o f shale y section . Normalizatio n o f each curve , suc h tha t the y intersec t th e presen t surface wit h a transi t tim e o f 59 1 JJL S m" 1, provides u s wit h well-constraine d approxi mations o f th e for m o f th e unexhume d compac tion curve , th e averag e o f whic h yield s th e supercurve (Fig . 3) . The differenc e betwee n th e supercurv e an d each datapoin t o n a soni c lo g give s a singl e denudation magnitude , th e mea n valu e o f which yields apparen t mea n tota l exhumatio n (£ A; Table 1) . Because post-exhumationa l buria l ca n have th e effec t o f returnin g formerl y exhume d rocks t o thei r maximu m buria l depth , i t i s necessary t o ad d t o E A th e thicknes s o f sediment accumulate d sinc e exhumation , t o derive mea n tota l exhumatio n (E T; Tabl e 1) . I n this study , w e therefor e ad d th e thicknes s o f Cenozoic rock s t o £ A, t o deriv e E T relativ e t o the presen t se a floo r a t eac h wel l locatio n (Fig. 4) .

Results The graph s i n Fig . 4 sho w a larg e degre e o f scatter in almost every well, with a similar range of value s betwee n well s i n whic h th e minimu m transit tim e i s generall y ver y clos e t o th e asymptote o f th e velocity-dept h curve , an d th e maximum is close t o the supercurve. Comparing exhumation compute d fro m thes e dat a (Tabl e 1 ) with AFT- an d VR-derived exhumation, the sonic velocities yiel d consistentl y lowe r value s tha n AFT and VR. Mea n exhumatio n in the EIS B ranges betwee n 65 2 an d 2132 m (mea n 1179m , standard deviation 338m). Contrary to AFT- an d VR-based exhumatio n studies , wher e estimate s of denudation exceed 300 0 m (AFT; Lewi s el al. 1992) an d 2500m (VR; Rowley & White 1998) . we compute exhumation of > 2000 m in only one well (110/18-1 ; 2132m) , almos t al l ou r value s being < 1500m. Thes e result s accor d wit h estimates fro m subsidenc e modellin g (Rowle y & White 1998) , which, like SV data, are also less susceptible t o th e distortin g effects o f transien t heating.

Fig. 6 . Denudatio n contoure d fro m th e result s presente d i n Fig . 4 usin g a standar d convergen t least-square s gridding algorithm.

TERTIARY DENUDATIO N O F IRISH SE A BASI N

In on e o f th e fe w othe r studie s usin g S V t o compute denudatio n magnitud e i n th e EISB , Colter (1978 ) estimate d c . 2000 m exhumatio n from wel l 110/3- 2 (this study ; E T = 783m) . He compared transi t tim e vs . dept h dat a agains t a generic compactio n curv e base d o n th e Rotlie gende Brockelschiefe r Shal e o f th e souther n North Se a (Mari e 1975 ; Fig . 5) . Simila r results were obtaine d b y Jackso n e t al. (1987) . Conversely, ou r results , especiall y thos e fro m Quad 11 0 in the souther n part of this study area, are i n broa d agremen t wit h those o f Rowle y & White (1998) , wh o use d V R dat a t o constrai n forward model s of lithospheric subsidence. Discussion Exhumation result s fro m th e positiv e feedbac k between uplift , erosio n an d isostas y tha t bring s

365

formerly deepl y buried rocks to the surface. Like most majo r unconformities , th e uppermos t Mesozoic unconformit y i n th e Iris h Se a basin s is a composite surfac e recording severa l separat e uplift an d erosio n events . Thus , i t coul d b e a product o f uplif t relate d t o Nort h Atlanti c opening in Paleocene time , shortenin g in OligoMiocene time , Pleistocen e glacio-isostas y o r a combination o f these . Moreover , hydrotherma l fluid flu x associate d wit h both th e formatio n of the Atlantic Ocean margin in Paleocene time and shortening durin g Oligo-Miocen e tim e mean s that th e variou s proxie s use d t o measur e denudation magnitud e als o requir e carefu l interpretation t o identif y artefacts . A s wel l a s the susceptibilit y of therma l histor y method s t o transient heatin g episodes , w e speculate tha t the relatively large scatter in the EISB sonic velocity data ma y recor d diageneti c effect s an d differin g

Fig. 7 . Mode l o f th e uppermos t Mesozoi c unconformit y i n th e U K Iris h Sea , interpretin g i t a s a produc t o f shortening (basi n inversion ) superimpose d o n epeirogenesis (her e represente d by magmatic underplating) .

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degrees o f cementatio n relate d directl y t o Tertiary hydrotherma l fluids . Comparison o f overal l trend s i n th e S V dat a with AFT-derive d palaeotemperatur e result s reveals n o obviou s relationship . AF T data sho w both absolut e denudatio n magnitud e an d earl y Tertiary palaeotemperatur e gradient s decreasin g southward (Gree n e t al 1997 ) wherea s the SVs indicate denudatio n increasin g southwar d (Fig. 6) . Furthermore , man y of th e well s whose SV log s ar e analyse d her e hav e AFT-derive d early Tertiar y (c . 6 0 Ma) palaeotemperature s

around 110° C a t depth s o f 1 km o r les s (Green , pers. comm.). Hardman et al. (1993) reported the apparent discrepanc y betwee n V R dat a fro m a n EISB well, which indicate almost no anomalous palaeotemperatures (an d therefor e denudation) , and apatite s fro m throughou t th e sam e well , which wer e totall y annealed . Collectively , thes e observations ar e interprete d as a record o f early Tertiary hydrotherma l circulatio n withi n th e Collyhurst an d Sherwoo d Sandston e Formation aquifers drive n b y heatin g o f intra-formationa l brines during igneous intrusion and/or flushing of

Fig. 8 . Detai l o f th e variatio n i n denudatio n magnitud e acros s fault s (show n i n bold , tick s i n hangin g walls ) obtained from wells in Quad 11 0 of the southern East Irish Sea basin. Notation s on the graphs are the same as in Fig. 4.

TERTIARY DENUDATIO N O F IRISH SE A BASI N

geofluids fro m deep , overpressure d basin s (see Iliffe e t al 1999) . In th e EISB , denudatio n varie s substantiall y between adjacen t faul t block s (Fig s 4 an d 6) , thereby providin g indirec t suppor t fo r th e contribution of basin inversion to its exhumation. In Fig . 7 , thi s short-wavelengt h patter n o f denudation i s interprete d a s a recor d o f locall y variable shortenin g superimpose d o n a unifor m peneplain generate d b y therma l uplif t durin g Atlantic opening (e.g. underplating). This model is consisten t wit h th e conclusion s o f Hardma n et al . (1993) , wh o describe d larg e intra-basina l variation i n exhumatio n magnitud e withi n th e EISB, which they interpreted in terms of an early Tertiary denudation episode of c. 15 Ma duration, followed b y later denudation lasting c. 50 Ma. It is also consistent with thermal history analysis of borehole sample s fro m Oxfordshire , wher e new AFT dat a wit h lat e Tertiar y age s recor d palaeogeothermal gradient s u p t o 9 0 °C km" 1, interpreted a s a consequenc e o f flushin g o f ho t fluids durin g th e inversio n o f th e adjacen t Wessex basin (Gree n e t al. 2001) . Given the above model of superimposed basi n inversion an d therma l uplif t i n th e EISB , maximum tectonic uplift, an d therefore exhumation, woul d b e concentrate d alon g th e norther n and souther n basi n margin s whos e orientatio n was mos t favourabl e fo r reactivatio n i n th e roughly north-sout h Oligo-Miocen e compres sive stress field. Thus, denudation increases fro m the presen t centr e o f th e basi n t o th e souther n basin margin , wher e th e S V analysi s record s maximum denudatio n (see also enhance d denu dation in the hanging walls of the faults in Fig. 8). Major shortenin g a t basin margins is exemplified by the St. Tudwal's Arch, a Palaeozoic basemen t high separating th e Carnarfon an d Cardigan Bay basins, contiguous to this study area. Line-lengt h restoration o f th e uppermos t Lowe r Jurassi c marker acros s thi s hig h indicate s tha t com pression in Tertiary time cause d bul k shortenin g of some 1 5 km across a basin of c. 100 km width. We speculat e tha t th e presen t basin-and-hig h configuration o f the Irish Sea area is a product of Neogene compressiona l reactivatio n o f Palaeo zoic basemen t fault s beneat h a once-continuous basin system . Why the n i s basi n inversio n suc h a difficul t process t o recognize i n the EISB area ? First, the compressional reactivatio n o f formerl y exten sional fault s mean s tha t ne t revers e faul t displacement alon g a n 'inverted ' faul t wil l decrease downward . I n deepl y exhume d basin s such as the EISB, the preserved segments of most of th e inverte d fault s wil l ofte n retai n a ne t normal displacement, even though they may have

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undergone significan t revers e reactivation . Second, no t all , or eve n any , of th e shortenin g during basin inversion is accommodated b y faul t reactivation. I n mudd y o r overpressure d basi n fills, shortenin g wil l tak e plac e b y pur e shea r thickening of the entire basin fill, making it much harder t o measur e (se e Eisenstadt & Withjac k 1995). The significance of pure shear is demonstrated in th e inverte d St . George' s Channe l an d Cardigan Ba y basins , wher e Oligo-Miocen e shortening wa s accommodate d b y thickenin g of the c . 10k m thick , fine-graine d siliciclasti c succession, withou t muc h evidenc e o f faul t reactivation. Unfortunately , ther e ar e fe w published dat a o n th e relativ e magnitude s o f th e principal stresses during tectonic shortening and, therefore, i t i s no t possibl e t o commen t o n th e likely chang e i n th e magnitud e o f th e mea n effective stres s during the switc h from extensio n to contraction in the EISB. However, given that a condition fo r horizonta l shortenin g i s tha t th e maximum principa l stres s i s horizontal , a fraction o f the denudation magnitudes presented here ma y b e attributabl e t o shortening-induce d porosity reduction . Sonic velocit y analysis als o provide s poten tially important information for seismic time-to depth conversion . Accurat e time-to-dept h conversion i s particularl y sensitiv e t o anomalousl y slow o r fas t intervals , leadin g t o unreliabl e estimates o f dept h t o formation tops. This poin t was demonstrate d b y EIS B wel l 110/12- 3 (Fig. 8) , which drille d throug h a sonicall y slo w mudstone overburden to the reservoir, which was relatively undercompacte d compare d wit h th e same sequenc e i n surroundin g wells . Conse quently, the depth to the top of the main reservoir was significantl y shallowe r tha n ha d bee n predicted b y th e time-to-dept h conversion , wit h attendant ramifications fo r drillin g safety . Conclusions (1) Soni c velocit y analysis indicate s tha t th e EISB underwent major exhumation during which between 65 2 an d 2132 m o f overburde n wa s denuded. (2) SV-derive d estimate s o f thicknes s o f denuded overburde n ar e consistentl y lowe r than AFT and V R results. Give n the susceptibilit y of thermal histor y method s t o transien t heating , AFT an d V R ma y hav e overestimate d denuda tion, mainl y because o f the influenc e o f Tertiary hydrothermal fluid flow. (3) Contrar y t o denudatio n trend s fro m AF T studies, whic h sho w a southwar d decrease , SV s indicate a southwar d increas e i n denudation ,

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toward the faulted margins of the EISB. Together with th e distinc t short-wavelengt h variatio n i n intra-basin denudatio n betwee n adjacen t faul t compartments, thes e observation s provide indirect support fo r the importance of basin inversion in th e EISB . We ar e gratefu l to Enterpris e Oi l pi c fo r provisio n of data an d fundin g fo r P.D.W.'s MPhi l project o n which this wor k i s based . Discussion s wit h P . Green , W . Owens, M. Stephenson, R. Swarbrick, K. Thomson, G. Westbrook an d N. White were useful an d encouraging. We acknowledg e helpfu l review s b y D . Jame s an d J . Cartwright.

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TERTIARY DENUDATIO N O F IRISH SE A BASIN HARDMAN, M. , BUCHANAN , J. , HERRINGTON , P . & CARR, A . 1993. Geochemical modellin g of the East Irish Se a Basin . In : PARKER , J.R . (ed,) Petroleum Geology of Northwest Europe: Proceedings of the 4th Conference. Geologica l Society , London , 791-808. HEASLER, H . & KHARITONOVA , N.A . 1996. Analysis of soni c wel l log s applie d t o erosio n estimate s i n the Bighor n Basin , Wyoming. AAPG Bulletin, 80, 630-646. HUNT, J.M . 1996 . Petroleum Geochemistry an d Geology. 2 ; W. H. Freeman, New York. ILIFFE, J.E', ROBERTSON, A.G. , WARD , G.H.F. , WYNN , C., PEAD , S.D.M . & CAMERON , N . 1999 . Th e importance o f fluid pressures an d migration t o th e hydrocarbon prospectivity of the Faeroe-Shetland White Zone. In: FLEET, A.J. & BOLDY, S.A.R . (eds) Petroleum Geology of Northwest Europe: Proceedings of the 5th Conference. Geologica l Society , London, 601-611 . ISSLER, D.R . 1992 . A ne w approac h t o shal e compaction an d stratigraphi c restoration , Beau fort-Mackenzie basi n an d Mackenzi e corridor , northern Canada . AAPG Bulletin, 76, 1170-1189. JACKSON, D.I . & MULHOLLAND , P . 1993 . Tectoni c and stratigraphi c aspect s o f th e Eas t Iris h Se a Basin an d adjacen t areas : contrast s i n thei r post Carboniferous structura l styles . In : PARKER , J.R. (ed.) Petroleum Geology o f Northwest Europe: Proceedings o f th e 4t h Conference. Geologica l Society, London, 791-808. JACKSON, D.I. , JACKSON , A.A. , EVANS , D. , WINGFIELD, R.T.R. , BARNES , R. P & ARTHUR , M.J., 1995 . The Geology o f th e Irish Sea. United Kingdom Offshor e Regiona l Report . Britis h Geological Survey . HMSO, London. JACKSON, D.I. , MULHOLLAND , P. , JONES , S . & WARRINGTON, G. 1987 . The geological framework of th e Eas t Iris h Se a basin . In : BROOKS , J . & GLENNIE, K.W . (eds ) Petroleum Geology o f Northwest Europe. Graham an d Trotman, London, 191-203. JAPSEN, P . & CHALMERS , J.A . 2000 . Neogen e uplif t and tectonics aroun d the North Atlantic: overview. Global and Planetary Change, 24, 165-173 . KNIPE, R.J. , COWAN, G . & BALENDRAM , V.S . 1993. The tectonic history of the East Irish Sea Basin with reference t o th e Morecamb e Fields . In : PARKER , J.R. (ed. ) Petroleum Geology o f Northwest Europe: Proceedings o f th e 4t h Conference. Geologica l Society, London, 857-866. LAMBECK, K . 1991 . Glacial reboun d an d se a leve l change i n th e Britis h Isles . Terra Nova, 3 , 379-389. LEWIS, C.L.E. , GREEN , F. , CARTER , A . & HURFORD , A.J. 1992 . Elevate d K/ T palaeotemperature s throughout Northwes t England : thre e kilometre s of Tertiar y erosion ? Earth an d Planetary Science Letters, 112, 131-145 . MAGARA, K. 1976 . Thickness o f removed sedimentary rocks, paleopor e pressur e an d paleotemperature , southwestern part of Western Canada basin. AAPG Bulletin, 60, 554-566.

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The post-Variscan thermal and denudational history of Ireland PHILIP A. ALLEN 1'5, STUAR T D . BENNETT 1, MICHAEL J . M. CUNNINGHAM 1, ANDY CARTER 2, KERRY GALLAGHER 3, ERIC LAZZARETTI 1, JOSEPH GALEWSKY 1, ALEX L . DENSMORE 1' 5 , W. E. ADRIAN PHILLIPS 1, DAVID NAYLOR46 & CRISTINA SOLL A HACK 1 1

Department of Geology, Trinity College, Dublin 2, Ireland

^Research School of Geological and Geophysical Sciences, Birkbeck College and University College London, Gower Street, London WC1E 6BT, UK 3

T. H. Huxley School of Environment, Earth Science and Engineering, Imperial College of

Science, Technology and Medicine, RSM Building, Prince Consort Road, London SW7 2BP, UK 4

ERA-Maptec Ltd, 36, Dame Street, Dublin 2, Ireland

5

Present address: Department of Earth Sciences, ETH-Zentrum NO, Sonneggstrasse 5, CH-8092 Zurich, Switzerland (e-mail: philip.allen@ erdw.ethz.ch) ^Present address: Dromreagh, Durrus, Bantry, Co. Cork, Ireland Abstract: Th e thermal an d denudational history of Ireland i s evaluated using an extensive new apatit e fission-trac k (AFT ) dataset derive d fro m surfac e samples . Modelle d therma l histories ar e used to construct maps of denudation for a number of time slices from Triassi c time t o 1 0 Ma usin g a time-dependen t palaeogeotherm . Th e map s illustrat e th e spatia l variability of denudation and subsidence within each time slice. The patterns of denudation are complex , showin g considerabl e variabilit y a t th e lengt h scal e o f 10 1-102km, wit h especially hig h denudation rates found ove r known igneous centres suc h as the Mournes of County Down. Based o n the onshor e AFT data alone, ther e i s no definitive signatur e of an Irish Sea Dome extending significantly across Irelan d in Early Tertiary time. The cumulative amount o f denudation durin g Tertiary tim e varies dependin g o n the AF T annealing model used, but is generally in the region between 1 and 2 km and without clear spatial trends. High amounts o f denudatio n hav e bee n mappe d ove r th e Tertiar y intrusion s i n Count y Down , reflecting their unroofing sinc e emplacement in Paleocene time . The cumulative denudation from Triassi c tim e t o 1 0 Ma show s relatively lo w amount s of denudatio n (<2km) i n th e Irish Midland s an d th e extrem e N E o f th e island , consisten t wit h th e observatio n tha t Mesozoic-Tertiary sediment s and igneous products are preserved i n the Ulster Basin. The western flan k o f Irelan d an d th e regio n betwee n Dubli n an d Count y Dow n sho w hig h cumulative amount s o f denudatio n (<4km) , th e latte r bein g consisten t wit h th e hig h amounts of denudation interprete d fo r the Irish Sea region. This denudation pattern explain s in part th e outcro p of Precambrian an d Lower Palaeozoic rock s i n these areas . Th e spatia l integration of the denudation over the entire landmass gives the average denudation rate and the sedimen t discharg e fro m Irelan d a s a functio n o f time . Averag e denudatio n rates ar e moderately hig h in Triassic time , falling to lo w values in Cretaceou s time , an d increasing substantially i n Tertiar y time . However , th e tota l volumetri c discharg e o f sedimen t i n Tertiary tim e is an order o f magnitude smaller tha n the preserved soli d volume of Tertiary sediment i n the basins offshore western Ireland .

A comprehensiv e understanding of th e therma l Basi n in the NE of the island, and from smal l and history o f Irelan d sinc e the en d o f th e Varisca n scattere d outcrop s throughou t th e res t o f th e orogeny is hampered by the scarcity of Mesozoic landmass . In contrast, the offshore basin s (Fig. 1) and Cenozoic rocks exposed o n the landmas s of contai n a wealth of information on Mesozoic and the island. Post-Carboniferous geological history Cenozoi c sedimentation. A genera l inferenc e i s can b e directl y evaluated onl y fro m th e Ulste r tha t th e present-da y Iris h landmas s ha s suffere d From: DORE , A.G., CARTWRIGHT, J.A. , STOKER, M.S. , TURNER, J.P . & WHITE , N . 2002 . Exhumation o f th e North Atlantic Margin: Timing, Mechanisms an d Implications for Petroleum Exploration. Geologica l Society, London, Special Publications, 196, 371-399. 0305-8719/02/$15.00 © The Geological Society of London 2002.

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Fig. 1 . Ma p showin g Irelan d an d it s offshor e basins . • , exploratio n wells . Borehole s mentione d i n tex t ar e highlighted.

extensive uplif t an d erosion ove r the last 200 Ma and ha s served , a t least i n part , a s a sourc e are a for clasti c detritu s no w fillin g th e offshor e depocentres suc h a s th e Porcupin e Basi n an d Rockall margin . A majo r researc h goa l i s therefore t o investigat e th e denudationa l history o f Irelan d a s a too l fo r evaluatin g th e likely deliver y o f clasti c sedimen t t o offshor e basins i n th e geologica l past . B y evaluatin g the denudational histor y o f Ireland , w e ar e als o able t o provid e a jigsa w piec e i n th e broade r picture o f th e uplif t an d erosiona l evolutio n o f the hinterland s o f th e hydrocarbo n province s of th e N W Europea n Atlanti c continenta l

margin (Dor e 1992 ; Dor e e t al. 1999 ; Spence r et al . 1999 ) an d particularl y th e British-Iris h sector. The ai m o f thi s investigatio n is t o provid e a quantitative pictur e o f th e therma l an d denuda tional history of Ireland since Triassic tim e fro m a ne w extensiv e apatit e fission-track (AFT) database mergin g previousl y publishe d result s (McCulloch 1993 , 1994 ; Keele y e t al . 1993 ; Green e t al . 2000 ) an d unpublishe d result s (Gleadow, pers . comm. ; McCulloch , unpubl . data; Murphy , pers. comm. ) wit h data collecte d and analysed during a field campaign in 199 8 and 1999. Th e resultin g databas e contain s AF T

THERMAL AND DENUDATIONAL HISTOR Y O F IRELAND

results from localitie s throughou t Ireland, 13 9 of which hav e bee n modelle d i n thi s paper . Individual time-temperatur e trajectorie s hav e been use d t o contou r map s o f palaeotempera ture a s a functio n o f time . Thes e map s ar e converted t o estimates o f denudation (presente d here) throug h us e o f palaeogeotherma l gradi ents constraine d b y present-da y estimates , vitrinite reflectanc e (VR ) profile s an d model s of earl y Mesozoi c continenta l stretching . Thes e denudation map s provid e vita l informatio n o n the regiona l pattern s o f denudatio n acros s th e island. B y integratin g denudatio n informatio n spatially ove r th e Irish landmass , we are able t o make a n estimat e o f th e tota l discharg e o f sediment sinc e Triassi c time . Thi s tota l discharge histor y ca n b e compare d wit h th e solid volume s of sediment though t to be present in th e offshor e basins .

373

The post-Variscan record in Ireland and its offshore basin s The Irish landmass The geological evolutio n of Ireland sinc e the end of th e Varisca n orogen y (Lat e Carboniferou s time) ha s bee n summarize d b y Na y lor (1992 , 1998). I n essence , thi s lon g geologica l histor y can b e directl y evaluate d onl y fro m th e stratigraphic an d igneou s recor d o f th e Ulste r Basin an d fro m smal l and scattere d occurrences across the island (Fig. 2): (1) Permo-Triassi c an d Liassi c sedimentar y rocks o f th e Ulste r Basi n expose d alon g th e shores of Belfast Lough, the coast of Antrim and penetrated i n a numbe r o f borehole s i n th e province, and in the small outliers at Kingscourt, County Cava n (Visscher , 1971 ) an d i n Count y

Fig. 2 . Simplified ma p of Ireland showing Mesozoi c to Tertiary stratigraphy, Tertiar y igneou s product s an d othe r Mesozoic to Tertiary occurrences discussed in text.

374

P. A. ALLEN ETAL.

Wexford (Clayto n etal. 1986). In these areas, the bulk of the post-Variscan successio n is made up of Permo-Triassic red beds (1146 m an d 1203 m at boreholes near the Antrim coast a t Magilliga n and Port More, respectively, and 2880 m at Larne 2) (se e Fig. 1 fo r borehol e positions) . Liassi c fine-grained siliciclasti c deposit s occu r a t outcrops aroun d th e N E Iris h coas t an d i n th e

subsurface, reachin g 269 m i n thicknes s i n th e Port Mor e borehole . (2) A punctuate d and condense d Jurassi c t o Tertiary successio n i n the Ulster Basi n compris ing Uppe r Cretaceou s greensand s an d chalk s (91m i n th e Por t Mor e borehole) , Paleogen e basalts (up to 500 m thick) with red lateritic soils and weathere d as h falls , an d Oligocen e clays ,

Fig. 3 . Simplified map of Ireland showin g Lower Palaeozoic an d Precambrian rocks, and main igneous plutons, with location of apatite fission-track samples modelled in this study. Circle with dot, TULIP sampling campaign; square wit h dot , previously publishe d an d unpublishe d sample s (Keele y e t al. 1993 ; McCulloch 1993 , 1994, unpubl. data; Green e t al. 2000; Gleadow, unpubl. data; Murphy, unpubl. data).

THERMAL AND DENUDATIONAL HISTOR Y O F IRELAND

silicliclastic deposit s an d lignite s o f th e Lough Neagh Grou p (Parnel l e t al 1989 ) (reachin g 350m i n th e Washin g Ba y borehole , Count y Tyrone) (Wilkinso n e t al 1980) . (3) Isolate d pockets o f Mesozoic sedimen t i n fault-related depressions , karsti c solutio n hol lows o r fissure-fill s (Fig . 2) , suc h a s th e chal k breccia o f Ballydeenlea , Count y Kerr y (Wals h 1966; Evan s & Clayton 1998) , Upper Jurassic to Lower Cretaceou s clay s nea r Carrick-on-Suir , County Kilkenn y (Higg s & Jone s 1998 ) an d Middle Jurassi c reddened clay s at Cloyne, south County Cork (Higg s & Beese 1986) . (4) Sporadi c occurrences of supposed Tertiary sediments (Fig . 2 ) includin g Oligocen e non marine clay s a t Ballymacada m (Count y Tipperary), terrestria l ?Tertiar y breccia s forme d by collapse o f karstic limestone caverns at Listry (County Kerry), upper Pliocene dark, non-marine clays a t Holl y mount (Count y Carlow) , non marine upper Pliocene-Pleistocene sands within gorges an d cave s i n Carboniferou s Limeston e bedrock a t Poulnahalli a (Count y Galway) , a karstic solution pipe in Carboniferous Limestone filled wit h cla y a t Ballygadd y (Count y Offaly) , and various Late Tertiary lacustrin e records (se e summaries and discussions by Davies (1970) and Mitchell (1980)) . Dee p Tertiar y weatherin g ha s been describe d fro m Tynag h (Count y Galway) . Combined wit h th e fac t tha t th e Paleocen e basalts o f th e Antri m Lava Grou p were erupte d subaerially onto a dissected land surface, there is therefore consisten t thoug h fragmentar y evi dence tha t Irelan d ha s bee n abov e se a leve l throughout Tertiary time . The onshor e stratigraphi c record reveal s a far from passiv e post-Palaeozoi c history , wit h multiple periods of relative uplift and denudation separating period s o f relative subsidence . I n th e Ulster Basi n o f N E Irelan d (McCaffre y & McCann 1992) , majo r unconformities , locall y demonstrably angular , are present a t the base of the Cenomanian-lower Maastrichtian Hibernian Greensand an d Ulste r Whit e Limeston e Formations, as a surface of karstification beneat h the subaeriall y erupted basalt s o f th e Paleocen e Antrim Lava Group, below the Oligocene Loug h Neagh Group , an d belo w a surfac e venee r o f recent soils , alluviu m an d Pleistocen e till s (McCann 1988 , 1990 ; McCaffre y & McCan n 1992). The geological ma p of Ireland gives a simple, general impressio n o f th e time-integrate d denudation of Ireland since the end of Palaeozoic time (Fig s 2 an d 3) . Olde r rock s emerg e fro m their regional Carboniferou s cover i n the NW of Ireland fro m Donega l t o Count y Galway , exposing pre-Dalradia n gneisses , Dalradia n an d

375

Lower Palaeozoi c sedimentar y an d meta sedimentary rocks , an d Caledonia n graniti c basement (Fig . 3) . I n easter n Ireland , sout h o f Dublin, th e Leinster Massi f produces th e largest contiguous are a o f hig h topograph y i n Irelan d and is composed o f Caledonian granite s intruded into Lowe r Palaeozoi c metasedimentar y rocks . This Massif was unroofed in Late Devonian time to provide detritus for the Munster Basin (Penney 1980), again in Jurassic time to feed basins in the Celtic Se a (Petrie et al 1989) , and once again in Tertiary tim e (Davie s 1970) . Th e local presenc e of schis t roo f pendant s demonstrate s tha t th e plutons ar e currentl y expose d a t clos e t o thei r upper surface . Betwee n Dubli n an d Belfast , th e Longford-Down Massi f i s a swat h o f Lowe r Palaeozoic sedimentar y rock s wit h occasiona l plutons o f Caledonian an d Tertiary age , suc h as the Newr y an d Mourn e granites , respectively . The Tertiar y pluton s o f Countie s Dow n an d Armagh ar e especiall y informative . SHRIM P U-Pb date s yiel d 55. 5 ±1. 1 Ma an d 56.5 ± 1. 3 Ma fro m th e Mourn e an d Sliev e Gullion granites , respectivel y (Meigha n e t al . 1999). As these granites were emplaced at depths of c . 3 km in the upper crust, there has evidently been very significan t unroofin g (c . 50 m Ma"1 time-averaged) sinc e thei r crystallizatio n i n Paleocene time . The youngest rocks preserved in Ireland are in Ulster, wher e Mesozoi c strata , includin g thic k Permo-Triassic sequences, ar e preserved beneath lavas of Paleocene age . The Irish Midlands are a low-elevation, low-relief region of Carboniferous outcrop interspersed with higher-standing inliers of Lowe r Palaeozoi c rocks . A s a firs t approxi mation therefore , w e shoul d suspec t tha t th e centre of the island, and the NE in particular, has experienced lowe r tha n averag e time-integrate d erosion sinc e th e en d o f Palaeozoi c time . Model result s presente d belo w suppor t thi s approximation. The pictur e o f th e denudatio n o f Irelan d i s aided by consideration of the stratigraphic record of th e offshor e basin s (Fig. 1) . The southern offshore flank The Fastnet , Sout h Celti c Sea , St . George' s Channel an d Cardiga n Ba y basin s al l contai n a thick (<800m ) blanke t o f Tertiar y (Eocen e t o Miocene o r Pliocene) sedimentar y rock s above a regional unconformit y (Shanno n 1991) . Thi s base-Tertiary unconformit y overlie s Cenoma nian-Maastrichtian chal k (u p t o 1 km thick ; Naylor & Shannon 1982 ) throughout most of the Celtic Se a area, but cuts down more deeply int o Jurassic sediment s i n th e St . George' s Channe l

376

P. A . ALLE N ETAL.

and Cardiga n Ba y basin s o f th e souther n Iris h Sea (Blundel l 1979) . O n th e assumptio n tha t Upper Cretaceou s chal k formerl y extended over the entir e souther n an d easter n offshor e flank of Ireland, the southern Irish Sea is inferred to have undergone greater erosio n i n Early Tertiary tim e than the Bristol Channel , Celtic Se a and Fastnet basins towar d th e S W an d south . Furthermore , we can conclude from the near-complete absence of Upper Cretaceou s chalk an d Tertiary deposit s in souther n Irelan d tha t ther e wa s differentia l uplift an d erosio n betwee n th e Iris h landmas s and th e souther n offshor e flan k bot h durin g the Early Tertiar y perio d o f unconformit y an d i n the ensuin g Eocene-Miocene stage s of Tertiary time. Lower Cretaceou s sandstone s an d shale s ar e found conformabl y underlyin g th e Uppe r Cre taceous chal k i n th e Fastnet , Celti c Se a an d Bristol Channel basins. In most exploration wells these Lowe r Cretaceou s rock s ar e initiall y continental Wealden , followe d b y marin e sand stones and shales. Aroun d the basin margins, but with the exception o f the North Celtic Se a Basin, these succession s overli e a base-Cretaceou s (Upper Cimmerian ) unconformit y that cut s ou t the Middle-Upper Jurassic units. Marine Lower Jurassic shales , marl s an d limestone s occu r across th e souther n offshor e flan k o f Ireland . Lower Jurassi c sedimentar y rock s (Sinemurian) are though t t o hav e bee n derive d i n par t fro m exposed Ol d Re d Sandston e i n S W Irelan d an d from th e Leinste r Massi f i n easter n Irelan d (Petrie et al. 1989) , indicatin g uplif t an d erosio n of these area s durin g Early Jurassic time . Wher e Middle Jurassic sediments are present, a Leinster Massif sourc e i s also invoke d t o explai n uppe r Bathonian-lower Callovian red beds prograding into th e Celti c Se a area . Uppe r Jurassi c fluviolacustrine deposit s ar e foun d alon g the northern edge o f the Nort h Celtic Sea Basin. Major inversio n along WSW-ENE structures also cause d unconformitie s i n th e centra l portion o f th e Nort h Celti c Se a Basi n (Tucke r & Arter 1987 ; Menpes & Hillis 1995 ; Murdoc h et al . 1995 ) i n th e post-Maastrichtia n t o pre-mid-Eocene interval . In summary , th e souther n flan k o f Irelan d exhibits unconformitie s o f probabl e lates t Jurassic-earliest Cretaceou s (Lat e Cimmerian ) and Earl y Tertiar y ages , an d locall y withi n Mid-Jurassic tim e (Mid-Cimmerian ; Na y lor & Shannon 1982) . Th e Irish landmass was a region of predominantl y subaeria l exposur e an d con tinental depositio n durin g larg e part s o f Mid Late Jurassi c time . Thi s landmas s probably became progressivel y floode d durin g Earl y Cretaceous time. There i s stratigraphic evidence

that wherea s th e Iris h landmas s an d Iris h Se a were undergoin g profoun d erosio n i n Earl y Tertiary time , whic h remove d thei r Uppe r Cretaceous cover , th e souther n offshor e flan k suffered relativel y littl e uplif t an d denudatio n except alon g structura l line s o f inversio n i n th e North Celti c Se a Basin . Thi s allowe d th e preservation o f chal k beneat h a base-Eocen e unconformity. Durin g Eocene-Miocen e time , the southern flank acted as a depocentre wherea s the Iris h landmas s experienced a lon g perio d of subaerial landscap e development . The western offshore flank Permo-Triassic sedimentar y rocks i n th e Porcu pine, Slyne , Erri s an d Donega l basin s ar e continental (Scotchma n & Thoma s 1995 ; Chapman e t al. 2000). Marine Liassi c sedimen tary rock s ar e foun d throughou t th e region , passing u p into continental and marginal marin e Middle Jurassi c units . I n th e mai n Porcupin e Basin, sedimentatio n wa s continuou s int o th e mixed continenta l an d shallo w marin e Uppe r Jurassic units . Althoug h recognize d b y th e erosional truncatio n o f rotate d faul t block s an d onlap of Cretaceous sedimentary rocks along the Porcupine Basi n margin , the widel y recognized uppermost Jurassic-lowermos t Cretaceou s (Upper Cimmerian ) unconformit y appear s t o involve little uplift an d erosion in the central part of th e Porcupin e Basin . Lowe r Cretaceou s sedimentary rock s pas s u p int o c . 1 km o f Upper Cretaceou s chal k alon g th e basi n axi s of the Porcupine Trough. There is little stratigraphic break a t the base o f the Tertiary sequence , and a more or less continuous although highly variable succession o f Paleocen e t o Pliocen e Tertiar y sedimentary rocks (Naylor & Anstey 1987). The stratigraphy i s simila r i n th e Nort h Porcupin e Basin, but there is a stratigraphic break at the top of Uppe r Cretaceou s chalk s (Tat e 1993 ) corre sponding t o tha t see n in th e onshor e section s in the Ulster Basin. Offshore borehole s in the Donegal and Slyne Erris basins to the NW of Ireland (Fig. 1 ) provide a different pictur e from th e Porcupine Trough. In the Donega l Basi n (Texac o 13/3-1) , Miocene Pliocene sediment s res t directl y o n Carboni ferous units , indicatin g the mergin g of th e post Palaeozoic unconformitie s recognize d i n th e onshore sections into a single stratigraphic break occupying mos t o f Mesozoi c an d Tertiar y time. In contrast , i n th e N E Erri s Troug h (Amoc o 12/13-1) ther e i s a relativel y thick sectio n fro m Eocene t o Aptia n rocks , wit h thi n o r absen t Paleocene unit s an d a thic k (817m ) Lowe r Cretaceous successio n resting on Rhaetian units.

THERMAL AND DENUDATIONAL HISTORY OF IRELAND

A Lower Tertiar y unconformit y can therefore be recognized in the Erri s Trough . The olde r unconformity i s probabl y du e t o uplif t an d erosion durin g Lat e Jurassic-Earl y Cretaceou s time, perhap s relate d t o riftin g i n th e Rockal l Trough. Furthe r t o th e S W i n Amoc o 19/5-1 , Lower Cretaceou s strat a and Uppe r Cretaceou s chalks hav e bee n erode d beneat h a pre-uppe r Miocene unconformit y tha t truncate s strati graphy dow n t o Lowe r Jurassi c units . I n th e Slyne Trough , El f 27/13- 1 penetrate d a Neo gene-Quaternary cove r unconformabl y over lying a thic k (>2km ) Lower-Middl e Jurassi c succession (Truebloo d 1992 ; Dance r e t al 1999). No exploratio n well s hav e yet bee n drille d i n the Iris h secto r o f th e Rockal l margin . Th e sedimentary fil l o f th e basi n i s though t t o b e a ?Palaeozoic t o Jurassic succession that was rifted in Early Cretaceous time and overlain by post-rift Upper Cretaceou s t o Paleocen e an d Eocen e t o Recent sediment s (Shanno n e t a l 1995 , 1999) . A borehol e i n th e U K secto r (132/15-1 ) penetrated Lowe r Cretaceou s sediment s restin g on crystallin e basemen t (Musgrov e & Mitch ener 1996) , demonstratin g considerabl e erosio n of th e flan k o f th e Rockal l Basi n befor e Earl y Cretaceous onlap . I f present , th e base-Tertiar y unconformity i s poorl y develope d i n th e Rockall margin . Th e regio n ca n b e though t o f as a long-term depocentr e sinc e Late Cretaceou s time. The effects o f Early Tertiary uplift an d erosion were therefor e no t strongl y fel t i n th e mai n Porcupine Basin or Irish Rockall margin. Instead, these basins continued to subside, allowing thick Cenozoic sediments to accumulate. Likewise, the effects o f latest Jurassic-Early Cretaceous (Lat e Cimmerian) erosio n ar e mino r withi n th e mai n depocentre o f the Porcupine Basin. In the SlyneErris basins and margins of the Porcupine Basin, however, ther e i s goo d evidenc e o f a majo r pre-Early Cretaceou s o r Lat e Jurassic-Earl y Cretaceous erosiv e even t that variabl y cu t down into Middl e Jurassic , Lowe r Jurassi c o r Triassi c sedimentary rocks . Th e Lowe r Tertiar y uncon formity cause d b y relativ e uplif t an d karstifica tion o f th e Ulste r Whit e Limeston e beneat h th e Antrim lava s i n th e onshor e section s o f N E Ireland is , however, unrecognized in the western offshore flank basins with the possible exceptio n of th e norther n Erri s Trough . I t i s likel y tha t a major unconformit y beneat h th e Mio-Pliocen e sequences resulting from regional Neogene uplift of th e N W flan k o f Irelan d an d adjacen t continental shel f (Stoke r e t al . 2001 ) ha s removed evidenc e o f a n earlie r Earl y Tertiar y event.

377

The eastern and northern offshore flanks Permo-Triassic sedimentary rock s ar e thought to be extensiv e i n th e Iris h Se a region . The y ar e overlain b y thic k ( < 2700m), marin e Liassi c sediments i n th e Kis h Ban k Basi n (Brougha n et al. 1989 ) and Cardigan Ba y Basin (Woodlan d 1971). Middle-Uppe r Jurassi c sedimentar y rocks ar e generall y absent , althoug h a mor e o r less complete , entirel y marin e Jurassi c succes sion i s found i n the Cardiga n Ba y Basin . Lowe r and Upper Cretaceous rocks are also absent fro m the Iris h Sea , indicatin g considerabl e erosion , probably durin g Earl y Tertiar y time . Th e dee p erosion o f Mesozoic stratigraph y in the Irish Sea area ha s been interprete d a s due to crustal uplif t and denudatio n cause d b y igneou s underplatin g (Brodie & White 1994 ; White & Lovell 1997 ) or transient dynami c uplif t (Cop e 1994 ) relate d t o plume activity centred under the Irish Sea during Paleocene time. The Rathli n Basi n o f Ulste r extend s alon g a Caledonian tren d northeastward s toward s Scot land. I t contain s thic k continenta l Triassi c sedimentary rock s overlai n by marine transgres sive Rhaetia n units , bu t Middl e an d Uppe r Jurassic an d Cretaceou s sedimentar y rock s ar e generally absen t (Evan s e t al . 1980 ; Naylo r 1992). A thin basal laye r of sandstone and chal k deposited unconformabl y o n th e erode d Meso zoic and older rocks of the Inner Hebrides, Malin Shelf an d Firt h o f Clyd e correlate s wit h th e Hibernian Greensan d an d Ulste r Whit e Lime stone o f th e onshor e section s i n Ulste r an d therefore demonstrate s a pre-Lat e Cretaceou s erosional event . Th e subsequen t erosiona l removal o f muc h o f thi s chal k laye r record s Tertiary denudation .

Palaeothermal and thermochronological data Although palaeothermal an d palaeoburial indice s (such a s vitrinit e reflectance (VR), cla y minera l crystallinity, soni c velocities , flui d inclusions ) are availabl e t o differin g degree s o n th e Iris h landmass an d adjacen t offshore, AF T analysis is potentially extremel y valuabl e i n providin g information o n th e time-temperatur e histor y o f samples, rathe r tha n simpl y a maximu m palaeotemperature o r buria l depth . I t ha s bee n used as a thermochronological too l extensively in the Britis h Isle s regio n (e.g . Bra y e t a l 1992 ; Lewis et a l 1992 ; Gree n e t a l 1993 , 1997 ; Duncan e f a l 1998) . Maps o f VR sho w stron g heating in the sout h of Irelan d interprete d t o b e du e t o a Lat e

Table 1 . Apatite fission-trac k results from th e TULIP sampling campaign an d those o f Murphy (unpubl.) (borehole samples an d those with low track length populations were not modelled)

Sample no.

Locality (County)

Easting, Northing

BS1 Fl

The Blac k Gap (Donegal ) Barnesmore Ga p (Donegal)

204527, 36908 9 202500, 38540 0

MT1 F9 GG3 GG9 GG12 GG13 GG15 IG1 CL1 LE GL CGI OM1 OM2 LG2 LG3 LG5 LG6 LG7 LG8 LG10 LG11 LG12 LG13 LG14 LG15 LG16 LG17 LG18 LG19 LG20 LG21 LG22 LG23

Milltown (Donegal) Mourne Mts. (Down) nr. Cashel (Galway) Kilkieran (Galway) Bovroughaun (Galway) nr. Spiddle (Galway) Loughinch (Galway) Pollrevagh (Galway ) Aughrus More (Galway) Partry Mts . (Mayo) Partry Mts. (Mayo) Laghtaeighter (Mayo ) Ox Mts. (Sligo) Pontoon Bridg e (Mayo ) Carysfort Par k (Dublin) Sandycove (Dublin) Dalkey Qu. (Dublin) Dalkey Qu . (Dublin) Killiney Hil l (Dublin) Dalkey Isl . Hotel (Dublin) Church Vie w Rd. (Dublin) Ballyedmond Qu . (Dublin ) Barnaculla Qu . (Dublin ) Three Rock Mt. (Dublin ) Military Rd . (Wicklow) L. Bra y (Wicklow ) Sally Ga p (Wicklow) Carrigshouk (Wicklow ) Carrigshouk (Wicklow) Glenmacnass Riv . (Wicklow) Wicklow Ga p Rd . (Wicklow ) L. Nahanagan (Wicklow ) Turlogh Hil l (Wicklow) Tullow Lowland s (Wicklow )

184980,378454 326850, 32627 7 79610, 24240 0 82860,231453 101740, 22836 0 109360, 22206 0 122590,225470 56490, 244670 56890, 257770 104900, 267000 105900, 270800 85170,274020 164223, 32763 5 123383, 30485 2 321580, 22857 1 325700, 228000 326450, 226040 326247, 22630 4 325990, 225550 327135,226371 324100, 225459 318739,223458 317890, 22423 0 317600, 22340 0 315090, 21800 0 314290, 215140 312940, 212200 309480, 20520 0 309794, 20529 3 311310, 20399 4 310090,98140 307600, 19920 0 306500, 19850 0 288770, 18516 5

Strat. age (lithology)

Spontaneous psa (A&)

Induced pia I (M) (

dosimeter pda Nd}

Precambrian (Psammite ) Devonian (Granite ) Carboniferous (Sandstone) Silurian (Quartzite ) Devonian (Granite) Devonian (Granite) Devonian (Granite) Devonian (Granite) Devonian (Granite) Devonian (Granite ) Devonian (Granite ) Ordovician (Sandstone ) Ordovician (Sandstone ) Devonian (Granite ) Dalradian (Gneiss ) Silurian (Granite ) Devonian (Granite ) Devonian (Granite ) Devonian (Granite ) Devonian (Granite ) Devonian (Granite ) Devonian (Granite) Devonian (Granite) Devonian (Granite) Devonian (Granite) Devonian (Granite) Devonian (Granite) Devonian (Granite) Devonian (Granite) Devonian (Granite) Devonian (Granite) Devonian (Granite) Devonian (Granite) Devonian (Granite) Devonian (Granite) Devonian (Granite)

1.547(1171) 1.656(1772)

2.114(1603) 1.761 (1884 )

1.435 (1613) 1.339(592) 0.687 (902 ) 1.049(1493) 1.189(2667) 0.934(1853) 1.053 (1224 ) 0.734(1668) 0.827 (2070) 0.859 (243 ) 0.533 (153) 0.837 (847 ) 1.862(2682) 1.360(1457) 4.410(3527) 2.569(2136) 3.069 (3250) 2.117(2441) 1 .678 (2240 ) 2.426(2132) 1.856(1201) 3.883 (5264 ) 3.995 (5454 ) 3.661 (3557 ) 1.710(2495) 3.160(3083) 2.912(4295) 2.248(3081) 2.701 (3485 ) 2.885 (2911 ) 1.827(2788) 4.795 (2645 ) 4.121 (3987 ) 3.952(3111)

.886(2120) .366 (604 ) 0.679(891) .483(2111) .647 (3693 ) .429 (2837 ) .359(1580) .013(2775) .109(2775) 0.587 (166 ) 0.505 (145) 0.948 (959 ) 2.415 (3478) 1.742(1866) 5.030 (4023) 3.516(2924) 4.117(4360) 2.975 (3431 ) 2.467 (3293 ) 3.494 (3070) 2.311 (1496 ) 4.407 (5974 ) 0.439 (6059 ) 3.808 (3700 ) 1.817(2652) 3.333 (3252 ) 3.159(4659) 2.502 (3430 ) 3.339 (4309 ) 3.688(3721) 2.306(3519) 5.821 (3211 ) 4.388 (4245 ) 4.399 (3463)

^(r)h

(no. of crystals)

FT central age (Ma) (RE %) c

.235 (3423 ) .205 (6679 )

< 1 (20) 99 (22)

151 ± 8(12.5 ) 189 ± 7(0 )

.506(4159) .205 (6679 ) .168(6474) .168(6474) .168(6474) .168(6474) .168(6474) .168(6474) .168(6474) .134(3144) .134 (3144 ) .168(6474) .235 (3423 ) .235 (3423 ) .124(6234) .124(6234) .124(6234) .124(6234) .124(6234) .124(6234) .124(6234) .124(6234) .124(6234) .124(6234) .124(6234) .124(6234) .124(6234) .124(6234) .124(6234) .124(6234) .124(6234) .124(6234) . 1 24 (6234) .124(6234)

15 (29) 40(12) 60 (30) < 1 (30) 40 (30 ) 15 (30 ) 15 (25 ) 10(20) 7(30) 70(16) 90(14) 60 (30 ) < 1 (22) < 1 (25) 82.1 (36 ) 40.2 (28 ) 0(34) 14.3 (34) 28 (32) 94.7 (32 ) 69 (23) 61.3 (34) 37.2 (33) 21 (30 ) 95.9 (40 ) 70 (29) 26.6 (30) 5.4 (40) 48.1 (29 ) 31.7(31) 57.5 (40) 32.4 (24) 94.9 (30 ) 9.1 (29 )

191 ± 8(10.2 ) 197 ± 12(4 ) 197 ± 10(0.1 ) 138 ± 6(13.3 ) 141 ± 4 (2.9) 128 ± 5 (7.8) 152 ± 7 (9.2) 142 ± 6 (9.9) 146 ± 5 (10.6 ) 275 ± 2 8 (2.9) 200 ± 2 3 (0) 172 ± 8(1.4 ) 159 ± 7(15.1 ) 159 ±7 (12.3 ) 165 ± 4 ( 0) 138 ± 4(3.9 ) 142 ± 5 (14.9) 133 ± 5 (9.2) 127 ± 4(6.5 ) 131 ± 4(0 ) 151 ± 6(0.9 ) 166 ±4(0.8) 169 ± 4 ( 2 . 1) 181 ± 5 (1.3) 177 ± 5 ( 0 . 1) 178 ± 5 (0.1) 173 ± 4(2.3 ) 170 ± 5 (9.4) 152 ± 4(0.6 ) 147 ± 4 (4.9) 149 ± 4 (0.6) 156 ± 5 (5.8) 177 ± 4 (0) 169 ± 5 (7.3)

Table 1 - continued

Dosimeter pd a (Nd)

(no. of crystals)

FT central ag e (Ma) (R E %) c

2.966(1858) 2.636 (2696 ) 2.686 (1452) 0.584 (1098 ) 3.004 (2590 )

3.457 (2166 ) 2.507 (2564 ) 3.099(1675) 0.841 (1581 ) 4.375 (3772 )

1.168(6474) 1.168(6474) 1.168 (6474) 1.168 (6474 ) 1.500(4159)

< 1 (30) 40 (30 ) 30 (30 ) 5(30) 2(20)

167 ± 7(11.8 ) 205 ± 6 (0.9 ) 169 ± 7 (6.4 ) 136 ± 7 (9.2 ) 174 ± 7 (10.4 )

1.710(1989) 2.262 (2144 )

1.788(2079) 2.303(2183)

1.503 (4159 ) 1.241 (6882 )

40 (28 ) < 1 (23)

240 ± 9 (6.2 ) 204 ± 9 ( 1 1)

1.342(1406)

1.507(1579)

1.504(4159)

<1 (18 )

228 ± 1 3 (15.7)

Paleocene (Granite )

0.206 (263 )

1.365(1743)

1.461 (4628 )

98 (32 )

36 ± 3 * (0.03)

0.459 (233 ) 0.768 (316 )

0.371 (188 ) 1.105 (455 )

1.136(4723) 1.168 (6474 )

90 (22) <1 (14 )

243 ± 2 3 (0) 144 ± 1 7 (28.5)

0.283 (40 )

0.226 (32 )

1.136(4723)

80(3)

236 ± 5 7 (4.6 )

104000, 260300

Ordovician (Sandstone ) Devonian (Granite ) Caledonian (QF Porphyry) Caledonian (QF Porphyry)

0.296 (109 )

0.217 (80 )

1.124(6234)

90 (10 )

255 ± 3 8 (0)

219500, 356600 324300, 425500 330663, 32419 1

Devonian (Sandstone) Trias sic (Sandstone ) Paleocene (Granite )

4.142(213) 2.676 (623 ) 0.218 (173 )

1.175(604) 1.499 (349 ) 1.843 (1462 )

1.235 (3423 ) 1.235 (3423 ) 1.439(4605)

5(13) 10 (10 ) 89 (23 )

73 ± 6 (7.3 ) 369 ± 3 2 (14.2 ) 28 ± 2 * (0.18 )

295620, 26737 0 272642, 17379 0

Devonian (Granite ) Devonian (Granite )

2.442 (1266 ) 2.163 (1733 )

5.953 (3086 ) 5.246 (4203 )

1.241 (6882 ) 1.501 (4159 )

<1 (21) 3(23)

84 ±4(13.7) 105 ±4(11.4)

Strat. ag e (lithology)

Spontaneous psa

Sample no.

Locality (County)

LG26 LG28 LG30 LG32 LG36

Kelshabeg (Wicklow ) Mt. Leinster (Carlow ) Tomduff (Carlow ) Tullow Hi l (Carlow ) Lugnaquilla Mt. (Wicklow) Slieve Bloo m Mts . (Offaly) Slieve Bloo m Mts . (Laois ) Knockshigowna (Tipperary)

293700, 18825 0 282743, 15256 4 278140, 15311 4 286582, 17326 7 303180, 19120 0

200163, 19636 0

Devonian (Granite ) Devonian (Granite ) Devonian (Granite ) Devonian (Granite ) Devonian (Granite ) Carboniferous (Sandstone) Devonian (Sandstone ) Carboniferous (Sandstone)

MGl Unmodelled data GA GG4

Mourne Mts . (Down)

329979, 32528 5

Killary fjor d (Galway ) L. Maumwee (Galway)

86500, 262200 97320, 24900 0

DB

Lough Mas k (Galway)

103700,261100

GS

Lough Mas k (Galway) Doonan Quarr y (Fermanagh) Western Re d Bay (Antrim) Mourne Mts . (Down) Kentstown (N O 1442 ) (Meath) Quinagh (Carlow )

SB1 SB2 SB3 d

DQ1 RBI MG2d NOe LG38e

Induced (M)

Easting, Northing

223200, 20825 0 229750, 20445 0

a Track densities (p) ar e (X10 6 tracks cm 2 ); numbers of tracks counted (N) show n in parentheses. Analyses by external detector method using 0.5 for the 47r/2ir geometry correction factor . b P()f) i s probability of obtaining $value for v degrees of freedom (wher e v is number of crystals (Nc) — 1) and is the probability that the single grain ages are consistent with on e population (<5 % denote s failure a t the 95% level) . c Fission trac k (FT ) age s calculate d usin g dosimete r glas s CN- 5 (analys t Carte r £ N — 5 = 33 9 ± 5 , *analys t Murph y £C N — 5 — 330 ± 10) , calibrate d b y multipl e analyses o f IUGS apatit e an d zirco n ag e standard s (se e Hurfor d 1990) . Central ag e i s a modal age , weighte d fo r differen t precision s o f individua l crystal s (Galbrait h & Laslett 1993) . RE % is the relative error or age dispersion. Quoted age uncertainties are ± la . d Murphy (unpubl.) . e Boreholes. Sampl e depths: NO, 472-492 m; LG38, 205m.

380

P. A. ALLEN ETAL.

Carboniferous therma l even t linke d t o th e Variscan orogen y (Clayto n e t al. 1989) . A substantial thicknes s (5- 7 km) o f post-Lowe r Carboniferous rock s i s though t t o hav e bee n removed b y erosio n t o explai n th e hig h therma l maturity o f Devonian-Carboniferou s rock s i n southern Irelan d (Clayto n 1989) . Post-Carbon iferous rocks , however , have low VR levels. For example, th e Campania n chal k brecci a a t Ballydeenlea ha s V R value s betwee n 0.4 6 an d 0.53% Rm , wherea s th e underlyin g Namuria n rock ha s value s centre d o n 4 % Rm . Th e reflectance o f chalk clasts indicates that between 1 an d 1.5k m o f overburde n (probabl y lates t Cretaceous and earliest Tertiar y rocks ) has bee n removed by erosion fro m abov e the chalk breccia (Evans & Clayto n 1998) . Th e ide a o f regiona l erosion o f a significan t overburde n o f lates t Cretaceous t o earlies t Tertiar y sedimen t acros s Ireland in Early Tertiary tim e is supported by the occurrence o f karsti c cav e passage s an d pinnacles along pre-existing joints in the Ulster White Limestone of NE Ireland (Simms 1998), showing that i t wa s wel l cemente d an d fracture d befor e the eruptio n o f ashe s an d lava s o f th e Antri m Lava Group . Thi s stat e o f induratio n wa s probably achieve d b y buria l unde r a pil e o f Maastrichtian an d possibl y lowe r Paleocen e sedimentary rock , subsequentl y remove d b y dissolution an d erosion . Although a certain amoun t of AF T dat a fro m Ireland's offshore basins is available in the public domain (Dunca n e t al . 1998) , fe w dat a ar e available fro m onshor e Ireland . I n th e presen t study, previousl y published (Keele y e t al 1993 ; McCulloch 1993 , 1994 ; Gree n e t al . 2000 ) an d unpublished (Gleadow, pers. comm.; McCulloch, unpubl. data ; Murphy , pers . comm. ) AF T dat a have been ver y substantiall y augmented by dat a from a new campaig n carrie d out in 199 8 and 1999, par t o f th e continuin g TULI P Tertiar y Uplift o f Irelan d Project . Ne w sample s wer e analysed b y Carte r a t th e Universit y Colleg e London Fissio n Trac k Laborator y an d modelle d by Gallaghe r t o obtai n mos t likel y time temperature trajectorie s an d thei r uncertaintie s (see Gallaghe r 1995 ; Gallaghe r & Brown 1999) . The publishe d an d unpublishe d dat a wer e als o modelled o r remodelle d t o provid e a full y consistent se t o f results . Detail s o f procedure s involved i n th e ne w samplin g progra m an d caveats t o th e interpretatio n o f th e result s ar e given below. Sampling A wid e rang e o f lithologie s (igneous , meta morphic an d sedimentary ) acros s Irelan d wer e

sampled for apatite grains. The best apatite yields were obtaine d fro m basemen t granites . Som e lithologies, suc h a s th e Hibernia n Greensan d from Ulster , did no t yield sufficien t quantitie s of apatite. I n general , therma l aureole s wer e avoided. W e presen t i n table s th e dat a fo r al l samples that were modelled o r remodelled in this study, plu s thos e TULI P sample s tha t wer e no t modelled (se e below) . Analytica l data an d AF T age results for the TULIP and Murphy (unpubl.). samples, togethe r wit h location , litholog y an d stratigraphic interval are given in Table 1 . Table 2 contains trac k lengt h dat a binne d a t 1 [Jim intervals fo r th e sam e samples . Dat a fo r th e published an d othe r unpublishe d AF T sample s are given in Table 3. It was deemed inappropriate to mode l o r remode l som e o f th e TULI P an d other availabl e publishe d sample s becaus e o f very lo w fission-trac k counts . Fission-track analytical details (Tables 1 and 2) Spontaneous fissio n track s wer e reveale d b y etching polishe d apatit e grai n mount s wit h 5 N HNO3 a t 2 0 ± 1 °C fo r 2 0 s. Mount s wer e irradiated i n th e therma l facilit y a t th e Ris 0 Reactor a t th e Nationa l Researc h Centre , Roskilde, Denmar k (cadmiu m rati o (thermal / epithermal + fas t neutrons ) >200) , usin g th e external detector method. Fluence was monitored using Corning uranium standard glass dosimeter CN-5. Muscovit e mic a externa l detector s wer e etched afte r irradiatio n usin g 48 % H F a t 20 ± 1 °C for 55 min. Observation and measurement o f fission-trac k densit y an d lengt h wa s conducted unde r transmitte d ligh t usin g a Zeiss axioplan microscop e wit h a tota l dr y magnifi cation o f X 1250. Reflecte d light was use d a s a tool fo r discriminatin g individual tracks . Hori zontal confine d fission-trac k length s wer e measured usin g a digitizin g table t calibrate d before eac h analysis against a stage micrometer. Ages were determined using the zeta calibration method an d IUG S recommende d ag e standards (Hurford 1990) . Al l dat a ar e reporte d a s th e central ag e (moda l ag e weighte d fo r differen t precisions o f individua l crystals) togethe r wit h the percentag e ag e dispersio n o r relativ e erro r (RE % ) abou t th e centra l ag e (Galbrait h & Laslett 1993) . Thermal history and denudation chronology modelling: procedures Thermal history modelling was undertaken using a search procedure outline d by Gallagher (1995)

THERMAL AN D DENUDATIONAL HISTORY OF IRELAN D

initially usin g th e annealin g metho d o f Laslet t et al. (1987) . Thi s annealin g mode l i s derive d empirically, base d o n laborator y annealin g experiments performe d o n a mono-compo sitional Durang o apatit e wit h a n F/C 1 ratio o f c. 0. 1 (Youn g e t a l 1969) . Thi s apatit e i s moderately retentiv e compare d wit h a pur e fluorapatite. Extrapolatio n o f th e Laslet t e t a l (1987) Durango apatite model to geological tim e scales (10 6-107 a ) predict s a partia l annealin g temperature range from c. 60 °C to c. 110 °C with an uncertaint y o f c . 10°C . However , on e o f th e major limitation s of this model is that it does not appear t o predic t sufficien t annealin g a t temperatures les s tha n c . 6 0 °C. Consequently , thermal histor y simulation s ca n b e biase d towards inferrin g rapid , recen t coolin g fro m c. 60 °C to surface temperatures, wit h consequen t overestimates o f denudatio n i n th e mos t recen t time interval. Structure in this temperature range therefore need s t o b e treate d wit h caution . We also presen t result s fro m a preliminar y ne w annealing mode l tha t remove s th e effect s o f excessive denudatio n a t low temperatures (Gunnell e t al i n prep.). Thi s model i s based o n the same data as the Laslett e t al (1987 ) model, but uses the Laslett & Galbraith (1996 ) formulation. The ke y differenc e betwee n th e tw o mode l formulations i s th e assumptio n o f th e erro r distribution on the data. The Laslett et al (1987 ) model transform s th e dat a befor e fittin g them , and assume s th e varianc e i s constan t i n th e transformed data . Th e revise d mode l doe s no t make thi s assumption and fits the observe d dat a directly. Relativ e t o th e Laslet t e t a l (1987 ) model, thi s revise d mode l lead s t o mor e annealing a t a given temperature. Consequently, predictions from thi s model ar e more consistent with low-temperatur e (<60°C ) annealin g implied b y geologica l data . W e compar e th e two annealin g model s fo r coolin g an d thu s denudation during Tertiar y tim e whe n sample s were at temperatures < 60 °C as they approached the present surface . Data fro m individua l AF T sample s wer e modelled independently , and the optimal thermal history model wa s selected o n the basis of the fit to th e observe d data . Th e modellin g procedur e requires th e input of stratigraphic ag e and timetemperature window s throug h whic h th e modelled therma l histor y passe s (Gallaghe r 1995) . The window s wer e give n broa d bound s i n temperature an d allowe d t o overla p i n tim e s o as not to restrict the modelling process. The same windows were used for all samples for reasons of consistency an d ar e a s follows : 50 0 ± 10 0 Ma at 80±80°C ; 25 0 ± 5 0 Ma a t 7 0 ± 7 0 °C; 150±50Ma a t 7 0 ± 70°C ; 10 0 ± 20M a a t

381

7 0 ± 7 0 ° C; 7 0 ± 20 Ma a t 7 0 ± 7 0 °C; 5 0 ± 2 0 M a a t 7 0 ± 70°C ; 3 0 ± 20M a a t 70 ± 70°C;OMaat20°C . The choic e o f data fit statistic depend s o n the form o f the original data. For new data collecte d during the TULIP campaign , where we have the raw dat a ( a serie s o f track lengt h measurement s and the spontaneous and induced track counts for a suit e o f crystals) , w e use d a maximu m likelihood statistic . Fo r previousl y unpublishe d (Gleadow, pers . comm. ; McCulloch , unpubl . data; Murphy , pers . comm. ) an d publishe d datasets (Keele y e t a l 1993 ; McCulloch 1993, 1994; Gree n et al 2000 ) we had access to mean, pooled o r centra l age , and mea n an d standar d deviation o f th e trac k lengt h distributions , a s reported i n the original papers. A weighted leastsquares measur e o f dat a fi t wa s use d fo r thes e data. Gallaghe r (1995 ) summarize d th e variou s misfit statistic s tha t ma y b e adopte d i n thes e different situations . The for m o f a n individua l thermal histor y i s encouraged to show cooling where required to fit the data , bu t otherwis e th e modellin g approac h tries to minimize th e structure or variation i n the thermal histories . Thi s approac h ensure s tha t random variations producing unresolved heating events are minimized, whereas the cooling events required to fit the data are maintained. However, if the modelled palaeotemperatur e remain s more or les s constan t ove r a give n tim e interval , thi s does no t mea n tha t progressiv e heatin g o f th e sample t o tha t temperatur e ha s no t occurre d over tha t tim e interval . Instead , th e modelle d thermal histor y reflect s th e lac k o f resolutio n when heating from a low temperature to a higher temperature occurs . Consequently , a stati c thermal history may be inferre d durin g a perio d of burial , wit h th e maximu m temperatur e occurring a t th e tim e o f maximu m buria l an d little resolution on the earlier, lower-temperature part of the thermal history . This has the effect o f damping buria l (heating ) event s i n a therma l history. Fla t portion s o f th e time-temperatur e trajectories shoul d therefor e b e viewe d wit h some caution , but d o indicat e a n uppe r limi t t o the palaeotemperatur e durin g thi s interval . I n later section s o f thi s paper , w e wil l refe r t o this effec t i n th e therma l histor y a s deposition , although th e stric t interpretatio n i s 'n o cooling'. The outcom e o f th e modellin g proces s i s a thermal history represented as a series of discrete time-temperature points , wit h linea r interp olation between point s adjacent in time. Palaeo temperature maps for a certain time in geological history ca n the n b e constructe d a s describe d i n the following section.

Table 2 . Track length results for TULIP an d Murphy (unpubl.) samples Number o f track s i n fx m intervals' 1 Sample no .

2

3

4

5

6

7

8

9

10

11

12

13

14

15

16

17

BS1 Fl MT1 F4 F9 GG3 GG9 GG12 GG13 GG15 IG1 CL1 LE GL CGI OM1 OM2 LG2 LG3 LG5 LG6 LG7 LG8 LG10 LG11 LG12 LG13 LG14 LG15 LG16 LG17 LG18 LG19 LG20 LG21 LG22 LG23 LG26

0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0

0 0 4

0 0 0 0 0

0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0

0

0

2 3 2 2 0 4 0 0 2 2 3 0 2 1 1 1 2 6 5 4 8 3 5 0 4 6 1 1 3 5 3 3 1 1 3 0 3 4

7 3 3 3 0 15 0 3 11 4 11 6 0 3 0 5 0 13 5 8 15 9 17 2 12 7 12 6 5 20 6 6 3 9 13 7 8 16

8 13 13 4 2 24 4 10 15 6 20 21 0 3 0 5 5 26 19 16 27 16 19 8 17 24 24 20 24 33 13 17 7 17 24 22 17 22

17 23 20 5 1 48 17 30 36 22 47 30 0 4 3 23 12 25 10 29 19 23 17 11 34 52 42 28 23 47 18 38 17 38 39 42 36 31

16 38 38 4 5 66 15 56 29 21 42 37 4 11 7 41 38 32 26 30 31 22 32 13 46 37 49 41 40 44 31 36 23 47 57 55 49 28

36 30 32 6 8 26 33 46 53 27 26 47 7 11 16 46 53 44 40 59 47 38 45 15 47 45 32 50 45 39 26 55 24 43 42 39 51 48

14 26 17 4 8 12 34 40 37 24 38 41 5 5 21 24 29 38 33 37 36 33 37 11 33 19 28 35 27 5 10 34 19 37 17 27 28 29

8 13 6 2 8 3 13 12 13 17 9 13 1 2 13 4 8 11 16 12 13 13 18 4 9 9 6 11 13 2 2 10 6 8 4 7 2 14

0 4 2

0 0 0 0 0 0 3 0 0 1 0 0 1 0 2 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0

1

0 0 0 2 0 0 0 0 0 0

1

0

1

0 0 0 0

1

0 0 0 0 0 0 0 0 2 0 0

1

0 0 0 0 0 0 0 0 0 0 0

1

1

0

0 0 0 0

0 0 0

0 0 0 0 0

0 0

0 0

0 0 0 0 0 0 0 0

0 0 2 0 0 0 0 0 0 0 2

1 1 1

1

0 0

1

0 0 0 0 0 0 0

1 1 1

1 1

2

1

1

1

2 0 0 0 0 0 0

1

0 0

1

0 0 0 0 0

1

0

1

0 0 0

1

0 0 0 0

1 1

0 0

1

0 3 2 0

1 1

1

0

1 1 0 1 2 1 0 1 1 0

1

0 0 2 2 2 1 2 1 2 1 1 0 0 3 0 1 0 1 0 0

1

0

2 0 11 3 5 5 3 5 0 4 0 3 1 2 3 2 0 0 2 0 0 0

2 5

1

Mean trac k length (|JLm) b

SDC

No. o f tracks' 1

12.75 ± 0.1 6 12.84 ± 0.1 4 12.35 ± 0.2 0 11.93 ±0.4 5 12.87 ± 0.2 5 12.98 ± 0.1 1 13.80 ±0.1 5 13.11 ± 0.1 0 12.82 ± 0.1 3 13.09 ± 0.1 7 12.50 ± 0.1 2 12.98 ±0.1 1 12.79 ± 0.5 8 12.60 ± 0.3 0 13.23 ±0.1 9 12.79 ± 0.1 1 13.27 ± 0.1 1 12.43 ±0.1 5 12.91 ±0.1 6 12.85 ± 0.1 2 12.42 ±0.1 5 12.85 ± 0.1 5 12.60 ± 0.1 5 12.53 ± 0.2 7 12.55 ± 0.1 2 12.24 ± 0.1 2 12.28 ±0.1 3 12.78 ± 0.1 2 12.72 ±0.1 2 12.50 ±0.1 4 12.22 ± 0.1 7 12.71 ± 0.1 1 12.93 ± 0.1 5 12.66 ± 0.1 1 12.28 ±0.1 0 12.45 ± 0.1 2 12.38 ±0.1 3 12.47 ± 0.1 5

1.70 1.77 2.33 2.57 1.47 1.49 1.71 1.37 1.85 2.00 1.75 1.57 2.64 1.91 1.51 1.40 1.31 2.11 1.97 1.72 2.09 1.87 2.14 2.20 1.73 1.71 1.85 1.64 1.60 2.02 1.78 1.61 1.54 1.62 1.47 1.68 1.88 2.15

108 155 140 33 34 200 131 200 203 132 200 200 22 42 67 150 150 200 159 200 201 160 195 68 205 201 200 193 180 204 111 202 101 203 200 204 201 200

Table 2 - continued Number o f tracks in |ji m intervals 3 Sample no. LG28 LG30 LG32 LG36 SB1 SB2 SB3 MGl d Unmodelled samples GA GG4 DQ1 RBI MG2d N0e LG38e

1 1 0 0 0 0 0 0

0

0 0 0 0 0 0 0

0 0 0 0 0 0 0

1 0 1 1 0 0 0

1

2 0 0 0 0 0 0

0 0 0 0 0 0

1

0 0 0 0 0 0

0 3

1

1

1

0

0 0 0 0 0

0 0 0 2 0 0

0 0 0 0 0 0 0

0 0 0 0 0 0 0

0 0 0 0 0

1

1

0

10

11

12 1

3

14

15 1

61

Mean trac k 7 lengt h (jjLm) b SD

C

No

. oftracks d

5 6 6 0 2 3 0 0

4 14 11 3 6 4 2 0

10 40 16 12 12 9 9 1

44 43 23 39 23 16 8 2

47 48 55 45 50 53 28 1

49 27 33 34 34 39 34 3

24 12 39 18 19 28 15 6

11 2 14 7 5 5 3 9

3 0 3 1 0 0 1 6

0 0 0 0 0 0 0 3

12.60 ± 12.56 ± 12.74 ± 12.57 ± 12.52 ± 12.75 ± 12.84 ± 13.98 ±

0.1 3 0.1 4 0.1 3 0.1 2 0.1 3 0.1 2 0.1 5 0.3 1

1.83 1.92 1.83 1.50 1.62 1.45 1.47 1.70

200 200 201 160 152 159 101 31

0 2 0 0 0 2 3

0 1 0

0 2 0 0 0 2 10

0 1 0 2 0 7 6

1 4 0 6 1 16 10

1 5 1 10 2 40 32

5 2 2 7 3 21 21

5 1 2 4 3 8 3

2 0 2 2 0 1 0

0 0

14.94 ± 12.38 ± 14.54 ± 12.72 ± 13.90 ± 13.23 ± 12.82 ±

0.3 3 0.4 7 0.4 1 0.2 7 0.3 8 0.1 6 0.2 1

1.18 1.95 1.07 1.51 1.31 1.57 1.94

14 18 8 32 10 100 90

1

0 2 4

1 1

0

0 0

a Each track length interval covers the range n — 1 to n, where n is the interval heading. N o tracks counted <2 |xm and > 1 7 fjim i n length. b Quoted uncertaintie s o f mean an d standard deviatio n o f track lengt h distribution s ar e ± 1 cr. c No track length data are available for samples D B an d GS owing to low apatite yields and anomalousl y low uranium concentrations . d Murphy (unpubl.) . e Borehole samples .

Table 3 . Published an d unpublished apatite fission-trac k results modelled fo r this study (McCulloch 1993, 1994, unpubl. data; Keeley e t at., 1993 ; Gree n e t al. 2000 ; Gleadow (unpubl . data); Murph y (unpubl. data)) Sample no. a Gleadow (unpubl.) Gil G12 G13 G14 G15 G16 G17 G18 G19 G110 Gill G112 G113 G114 McCulloch (unpubl.) Mel Mc2 Mc3 Keeley et al. (1993) 659 660 664 665 666 667 668 671 672 McCulloch (1993) 1 4 5 7 8 9 11 12

Locality

Easting, Northin g

Strat. ag e (lithology)

Fission trac k ag e (Ma) b

Mean trac k length (|jLm)c

SDC

No. o f tracks

The Rosse s (Donegal ) The Rosse s (Donegal ) The Rosse s (Donegal ) The Rosse s (Donegal ) SE of Dunglo w (Donegal ) SE of Dunglo w (Donegal ) N o f Ardar a (Donegal ) E of Cregganbau n (Mayo ) Doonloughan (Galway ) nr. Roundstone (Galway ) nr. Roundstone (Galway ) Carna (Galway ) Ballyknockan (Wicklow ) Wicklow Ga p (Wicklow )

182500,418000 183000,418000 180000, 416000 182500, 415800 179000,409900 180000, 407300 173000,393000 86000, 27400 0 57000, 24550 0 71000,239000 73000, 24200 0 79000, 23200 0 301000, 207500 307000, 20100 0

Devonian (Granite ) Devonian (Granite ) Devonian (Granite ) Devonian (Granite ) Devonian (Granite ) Devonian (Granite ) Devonian (Granite ) Devonian (Granite ) Devonian (Granite ) Devonian (Granite ) Devonian (Granite ) Devonian (Granite ) Devonian (Granite ) Devonian (Granite )

188.1 ± 3. 6 195.0 ± 7. 0 197.1 ± 6. 8 200.9 ± 7 . 1 194.1 ± 8. 0 179.9 ± 8. 8 175.5 ± 9. 3 124.1 ± 5. 1 119.8 ± 6. 1 149.7 ± 6. 4 163.0 ± 7. 5 182.7 ± 5. 3 200.8 ± 4.6 208.2 ± 6. 4

12.25 ±0.2 2 12.05 ±0.1 9 11.88 ± 0.2 2 12.23 ± 0.2 0 12.22 ± 0.2 3 12.34 ±0.23 12.39 ±0.1 8 12.26 ± 0.2 4 12.19 ± 0.1 8 12.44 ± 0.2 1 12.39 ± 0.2 0 11.93 ±0.2 2 12.54 ± 0.1 7 12.20 ± 0.2 3

2.15 1.93 2.15 2.04 2.42 2.27 1.88 2.36 1.77 2.10 1.94 2.17 1.81 2.30

96 103 96 104 11 1 97 109 97 97 100 94 97 113 100

S o f Creeslough (Donegal ) Cushendall (Antrim)

206200, 427800 324300, 426100

159.0 ± 11. 0 394.0 ± 24. 0

12.41 ± 0.2 1 12.37 ± 0.2 0

1.65 1.59

62 63

62000, 24500 0

Devonian (Granite ) Devonian (Sandstone ) Ordovician? (Metagabbro?)

154.0 ± 9. 0

12.11 ± 0.3 4

2.18

41

275300, 10142 0

Devonian (Sandstone )

226.1 ± 11 . Ic

12.38 ±0.1 5

1.58

113

271240, 13942 0 221100,97380 199820, 6388 0

Devonian (Sandstone ) Devonian (Sandstone ) Devonian (Sandstone )

130.3 ±7.3 c 154.5 ± 13.9 c 211.8 ± 12.2 c

12.38 ± 0.1 9 12.56 ± 0.3 1 12.10 ±0.2 9

1.67 1.55 1.98

74 25 47

219760, 76520 230060, 12130 0 162420,40690 184900, 10170 0 23 1 240, 8922 0

Devonian (Sandstone ) Devonian (Sandstone ) Devonian (Sandstone ) Devonian (Sandstone ) Devonian (Sandstone )

185.7 ± 9.2 c 165.8 ± 7.9 c 154.2 ±5.6 c 191.6 ± 8.3 c 192.1 ± 10.3 c

12.06 ± 12.37 ± 12.65 ± 12.68 ± 12.63 ±

0.3 3 0.1 6 0.1 5 0.1 4 0.1 6

1.92 1.58 1.44 1.41 1.51

34 101 92 99 90

81000, 27430 0 89800, 73000 31600, 96700 93000, 58300 73000, 26290 0 1 1 3200, 22300 0 145000, 32370 0 140200, 31450 0

Devonian (Granite) Devonian (Sandstone ) Dev )nian (Sandstone) Dev )nian (Sandstone) Silu ia n (Sandstone ) Devonian (Granite ) Dev >nia n (Granite ) Devonian (Granite )

152.0 ± 156.3 ± 167.6 ± 164.6 ± 170.8 ± 115.5 ± 168.7 ± 154.8 ±

12.89 ± 0.1 9 12.59 ± 0.1 9 11.80 ± 0.2 4 12.61 ± 0.2 6 13.27 ±0.2 3 12.24 ± 0.2 3 12.49 ± 0 . 1 6 12.55 ±0.1 6

2.00 1.45 2.07 1.88 2.15 2.13 1.54 1.53

114 62 78 53 92 90 96 94

Mannin Ba y (Galway ) nr. Hoo k Hea d (Wexford ) Graiguenamanagh (Kilkenny) Col ligan wood (Waterford ) Ballycotton (Cork ) Ardmore Hea d (Waterford) Comeragh Mts . (Waterford ) Old Hea d o f Kinsal e (Cork) Ballyderown (Cork ) Creadan Hea d (Waterford ) Cregganbaun (Mayo ) Gortamullin (Kerry ) Slea Hea d (Kerry ) Caha Mts . (Cork) nr. Tull y Cros s (Galway ) Spiddle (Galway) Ox Mts . (Sligo ) Ox Mts . (Sligo)

14.8 p 26.8 m 15.0 p 22.8 p 25 . 4p 13.0 p 15.6 p 17.6 p

Table 3 - continued

Sample no. a

Locality

Easting, Northing

Strat. ag e (lithology)

Fission trac k ag e (Ma)b

Mean trac k lengt h (|jLm)c

SDC

No. o f tracks

14 16 17 18 19 20 21 22 23 24 30 31 32 39 43 45 46 47 48 51 54 55 56 57 60 84 64 68 75 76 77 80 82 83 86 88 90 McCulloch (1994) 2 3 4 6 7

Cronagort (Clare ) Diomond Hil l (Galway) E of Ome y Islan d (Galway) Doonloughan (Galway ) Roundstone (Galway) Spiddle (Galway) Barna (Galway) Aughrim (Wicklow) Ballinacarrig (Wicklow ) Kilmanoge (Wicklow) N of Ardara (Donegal ) Doonbeg Ba y (Clare ) Mullaghmore (Sligo ) Crocknaconspody (Fermanagh ) Cratloe (Clare ) Ballymastocker Bay (Donegal ) N of Port Laois e (Laois ) Cushendall (Antrim ) near Enniskillen (Fermanagh ) W of Carlow (Laois ) L. Beltra (Mayo) L. Corrib (Galway) Kilkee (Clare ) NE of Leenane (Mayo ) Oughterard (Galway) Clonloskan (Cavan ) Belfast Loug h (Down) Beaghbeg (Tyrone) Ballyvoyle (Waterford ) Slieveardagh Hill s (Tipperary) Mouth o f Shannon (Kerry) Graiguenamanagh (Kilkenny) Shannon estuar y (Clare ) Kerry Head (Kerry) Tagoat (Wexford) nr. Foynes (Limerick ) Slieve Aught y Mts. (Galway )

108500, 19480 0 72800, 25600 0 57400, 25600 0 56900, 245300 72300, 240300 113200,222200 123400, 22250 0 313100, 17970 0 318300, 18550 0 325100, 18890 0 172500, 39330 0 95700, 16750 0 171000,357900 239700, 35470 0 148100, 16160 0 225100, 438000 239500, 20900 0 324400, 427600 230200,341100 266200, 17670 0 107500, 29730 0 98100, 24980 0 88400, 15980 0 93500, 26570 0 111700, 24260 0 235200, 30070 0 340900, 380700 268300, 38200 0 233600, 94800 228700, 15420 0 87600, 14490 0 270800, 14340 0 84800, 15240 0 69000, 13210 0 310300, 11140 0 124700, 15210 0 154700, 201300

Carboniferous (Sandstone ) Precambrian (Granite ) Devonian (Granite ) Devonian (Granite ) Devonian (Granite ) Devonian (Granite ) Devonian (Granite ) Devonian (Granite ) Devonian (Granite ) Devonian (Granite ) Devonian (Granite ) Carboniferous (Sandstone ) Carboniferous (Sandstone ) Carboniferous (Sandstone ) Devonian (Sandstone ) Devonian (Sandstone ) Devonian (Sandstone ) Devonian (Andesite ) Silurian (Sandstone ) Carboniferous (Sandstone ) Devonian (Sandstone ) Precambrian (Schist ) Carboniferous (Sandstone ) Ordovician (Sandstone ) Devonian (Granite ) Devonian (Granite ) Carboniferous (Sandstone ) Ordovician (Tuff ) Carboniferous (Sandstone ) Carboniferous (Sandstone ) Carboniferous (Sandstone ) Devonian (Granite ) Carboniferous (Sandstone ) Devonian (Sandstone ) Ordovician (Tuff ) Carboniferous (Sandstone ) Devonian (Sandstone )

150.0 ± 47.8 m 156.4 ± 17.2 p 135.2 ± 15.0 p 129.1 ± 11 . 8p 127.3 ± 14.8 p 115.9± 10.2 p 119.1 ± 11 . Op 181.9 ± 16.6 p 184.2 ± 24.2 m 217.2 ± 19.8 m 180.7 ± 11 . 6p 112.7 ± 17.6 p 165.7 ± 15.0 p 177.4 ± 17.8 p 205.2 ± 27. 4 m 300.8 ± 25.4 m 179.4 ± 16.4 p 267.2 ± 26.2 p 176.2 ± 15.0 p 153.7 ± 18.0 p 186.5 ± 15.2 p 150.3 ± 11 . 4p 138.4 ± 22.4 m 152.2 ± 13.0 p 129.6 ± 14.2 p 200.7 ± 14.6 p 273.8 ± 26.8 p 225.2 ± 21.4 p 194.9 ± 12.2 p 147.0 ± 14.4 p 133.9 ± 28.4 m 149.0 ± ll.S p 132.9 ± 30.0 m 154.3 ± 26.0 m 219.8 ± 19.4 p 143.5 ± 29.6 m 192.3 ± 16.6 p

12.15 ± 0.4 3 13.22 ± 0.1 3 13.01 ± 0.2 1 12.66 ± 0.1 6 12.77 ± 0.2 2 12.54 ± 0.2 0 12.12 ± 0.1 7 11.81 ±0.4 5 12.03 ± 0.2 1 13.02 ± 0.1 6 12.50 ± 0.1 6 13.20 ± 0.2 6 12.45 ± 0.2 1 11.48 ± 0.2 0 12.52 ± 0.1 8 12.85 ± 0.2 8 11. 19 ±0.22 12.71 ± 0.1 6 12.28 ± 0.2 0 12.14 ± 0.3 4 12.05 ± 0.2 4 12.66 ± 0.2 0 12.54 ± 0.2 3 12.62 ± 0.1 8 12.37 ± 0.2 5 12.13 ±0.1 8 12.13 ±0.1 5 12.40 ±0.1 7 12.94 ± 0.1 7 12.03 ± 0.2 6 12.28 ±0.1 9 12.35 ±0.1 3 11.77 ±0.2 6 11.56 ± 0.2 4 12.28 ± 0.1 9 12.40 ± 0.2 1 12.88 ±0.1 8

2.30 1.27 2.07 1.61 2.15 1.95 1.66 2.37 1.91 1.61 1.56 1.45 1.92 1.67 1.70 2.37 2.15 1.61 2.07 2.18 2.42 1.94 1.36 1.81 1.91 1.76 1.54 1.74 1.67 1.92 1.85 1.76 2.58 2.43 1.85 2.24 1.71

30 103 100 101 99 101 105 29 83 100 101 32 86 69 89 74 101 101 100 43 104 99 37 95 60 100 104 103 100 54 95 172 99 96 95 113 92

NW of Roundwood (Wicklow) NW of Enniskerry (Wicklow ) Three Roc k Mt. (Dublin ) N of Slan e (Meath) E o f Banbridge (Down)

316000, 20860 0 319200, 21930 0 317800, 22320 0 295500, 279300 327800, 34450 0

Devonian (Granite) Devonian (Granite ) Devonian (Granite ) Ordovician (Tuff ) Devonian (Granite )

136.7 ± 126.5 ± 161.9 ± 212.1 ± 134.4 ±

12.27 ± 12.27 ± 11.73 ± 12.78 ± 11.74 ±

0.1 7 0.1 6 0.2 1 0.2 4 0.1 5

1.71 1.62 1.96 2.42 1.46

101 100 86 101 101

13.0 p 6.4 p 33.4 m 31.6 p 12.2 p

Table 3 - continued

Sample no. a

Locality

Easting, Northin g

Strat. ag e (lithology)

Fission trac k ag e (Ma) h

Mean trac k lengt h (fjim) c

SDC

No. o f tracks

8 9 11 12 13 14 16 Green et al. (2000) GC345-2 GC345-3 GC345-7 GC458-1 GC458-2 GC458-5 GC543-63 GC543-64 GC543-65 GC543-72 GC543-73

Loughshinny (Dublin) N o f Skerrie s (Dublin ) E of Banbridge (Down ) Dundalk Ba y (Louth) S Dundal k Bay (Louth) Clogher Hea d (Louth ) Balbriggan (Dublin )

327000, 25770 0 325300, 26060 0 321000, 334800 307200, 30260 0 311500,293500 317100, 28460 0 320300, 264400

Carboniferous (Sandstone ) Silurian (Sandstone ) Devonian (Granite ) Silurian (Sandstone ) Silurian (Sandstone ) Silurian (Sandstone ) Ordovician (Tuff )

70.2 ± 30.4 p 73.4 ± lO.O p 58.3 ± 7.4 p 64.5 ± 14.2 m 46.0 ± 15.6 m 50.8 ± 6.4 m 128.4± 22.8 p

12.56 ±0.5 1 12.20 ± 0.5 0 13.94 ±0.1 9 13.42 ± 0.2 0 13.72 ± 0.1 6 13.07 ±0.2 5 13.54 ±0.3 8

2.52 2.45 1.20 1.30 1.17 1.95 2.41

25 25 42 43 54 60 40

Moll's Ga p (Kerry ) Lough Muckros s (Kerry ) Kingscourt (Cavan ) Mullaghmore (Sligo ) Ballysadare (Sligo ) Slisgarrow (Fermanagh ) Strangford Loug h (Down ) Craigavad (Down ) Scrabo North Quarr y (Down ) Waterfoot (Antrim ) Runabay Hea d (Antrim)

86000, 7750 0 93400, 8550 0 285155, 28158 2 171000, 358000 165100,329600 202525, 35177 3 340477, 36957 5 342500, 38140 0 347500, 37310 0 324700, 425000 325600, 437000

Devonian (Sandstone ) Carboniferous (Sandstone ) Carboniferous (Sandstone ) Carboniferous (Sandstone ) Carboniferous (Sandstone ) Carboniferous (Sandstone ) Ordovician? (Sandstone ) Carboniferous (Sandstone ) Triassic (Sandstone ) Devonian? (Alluvia l sands ) Precambrian ('Dalradian' )

177.0 ± 150.8 ± 78.5 ± 164.0 ± 164.2 ± 148.1 ± 318.9 ± 321.7 ± 72.9 ± 312.9 ± 291.9 ±

12.57 ±0.1 9 12.87 ± 0.1 8 12.12 ±0.2 7 11. 76 ±0.24 13.31 ± 0.3 0 11.95 ±0.3 0 12.33 ± 0.1 7 12.62 ± 1.1 8 13.21 ±0.2 3 11.89 ± 0.1 9 12.51 ± 0.2 8

1.90 1.82 2.72 2.38 1.49 2.40 1.77 1.76 1.51 1.96 1.50

103 100 105 100 25 65 108 100 44 106 28

7. 1 p 11. 1 c 6. 9 c lO.O p 21. 6 c 13. 6 p 25. 3 p 18. 7 p 5. 6 p 29. 2 p 43. 5 p

a Sample number s given a s in the literatur e wher e applicable . b Fission track age types given by: c, central age; p , pooled age ; m , mean age. McCulloc h (unpubl. ) and Gleadow (unpubl. ) age types unknown. Quoted ag e uncertainties are ± \cr, except thos e o f McCulloch (1993 , 1994 ) whic h are ±2o \ McCulloc h (unpubl. ) ag e uncertaintie s unknown. c Quoted uncertaintie s of mea n an d standar d deviatio n of trac k lengt h distributions are ± 1 a.

THERMAL AND DENUDATIONAL HISTOR Y OF IRELAND

Palaeotemperatures ca n b e converte d t o estimates o f denudatio n b y choosin g a certai n time slice of known duration an d dividing by an estimate o f th e palaeogeotherma l gradien t a t each o f th e sampl e sites . A physicall y mor e appropriate approach i s the use of heat flow and thermal conductivity , the latte r bein g dependen t on th e litholog y o f th e roc k sectio n bein g considered. Here , w e avoi d th e uncertaintie s inherent i n estimatin g th e therma l conductivit y of the missin g section an d assum e it is constant. In the absence of information to the contrary, it is common to assume that heat flow or temperatur e gradients hav e bee n constan t ove r tim e (e.g . Gallagher & Brown 1999) . I n this situation, it is an assumption tha t the modelled thermal histor y reflects onl y coolin g cause d b y denudation . However, i n regions o f active tectonics, particu larly i n region s o f lithospheri c stretching , hea t flow and geothermal gradients ar e likely t o vary strongly a s a function o f time (McKenzie 1978) . Higher palaeogeotherma l gradient s impl y smal ler amounts of denudation to explain the cooling . We have therefore mapped denudatio n with both constant an d variabl e geotherma l gradient s (se e section o n 'Evaluatio n o f th e geotherma l gradient'), bu t fo r reason s o f spac e presen t results only for the variable geother m case . Contouring method Palaeotemperatures a s a function o f time derived from AF T analysis were contoured . The invers e distance weighte d (IDW ) interpolato r wa s use d to produc e a temperatur e gri d fo r eac h chose n time slice , a t a cel l siz e o f 5 km X 5 km. Thi s method determine s outpu t cel l value s usin g a linearly weighted combination of a set of sample points. Th e weigh t i s a functio n o f invers e distance. The input number of sample points was limited t o 12 , usin g th e neares t neighbou r method. Th e powe r paramete r i n th e ID W interpolation control s th e significanc e o f th e surrounding points upon the interpolated value , a higher powe r resultin g i n a smalle r influenc e from distan t points . Th e powe r use d i n th e interpolation o f th e AF T grid s wa s 2 , i.e . low . The outpu t value for a cell using IDW is limited to the range of values used to interpolate, that is, the maximu m an d minimu m values constraine d by th e origina l palaeotemperatur e range . A s the IDW is a weighte d average , the interpolate d value canno t be greate r tha n the highes t o r les s than th e lowes t inpu t (Watso n & Phili p 1985) . The resulting surface is the best estimate of what quantity i s o n th e actua l surfac e fo r eac h location. Th e interpolation preserve s th e existing data poin t value s o n th e temperatur e grid .

387

Although there may be significan t gradient s and discontinuities i n th e AF T palaeotemperatur e data, w e hav e constructe d map s o f palaeotem perature (and thereby denudation presented here) purely usin g a n ID W interpolato r s o a s t o b e transparent an d unbiased . However , th e reade r needs t o bea r i n min d tha t thi s metho d necessarily smoothe s potentia l discontinuitie s in th e origina l data . Whe n presentin g th e denudation map s we found tha t the conventiona l 'equal interval ' metho d o f classificatio n smoothed ove r trend s i n th e dat a i n tim e slice s when muc h o f Irelan d underwen t relativel y similar amount s o f denudation . A 'natura l breaks' classificatio n schem e wa s use d instea d to identif y breakpoint s betwee n classe s usin g a statistical formula, allowin g pattern s o f denuda tion t o b e mor e readil y picke d out . N o user defined discontinuities , suc h a s faul t bloc k margins, hav e bee n incorporate d int o th e contouring procedure . In contourin g palaeotemperatur e maps , dat a points wer e remove d fro m th e interpolatio n exercise whe n thei r palaeotemperature s exceeded 120°C , as higher temperature s are not resolvable wit h th e AF T technique . Conse quently, th e interpolatio n i s carrie d ou t wit h a variable number of data points, particularly at old time slices. Although this results in a poorer data density, i t remove s th e possibilit y o f bias to th e results cause d b y holdin g palaeotemperature s at thei r maximu m resolvabl e leve l (120°C ) arbitrarily. Evaluation of the geothermal gradient A knowledg e o f palaeogeotherma l gradient s i s important fo r th e correc t conversio n o f palaeotemperature data to estimates of denudation. The available present-day heat-flow data from Irelan d are relativel y limite d (se e Broc k (1989 ) an d Brock et al (1991)) . Th e typica l rang e i s 52-87 mW m , wit h a mea n valu e o f 6 7 mW m"2 clos e t o th e Europea n averag e o f 6 4 mW m"2 (Cerma k 1979) . Thi s somewha t limite d spatial variatio n o f hea t flow , an d a lac k o f evidence t o the contrary, justifies our assumption of a constant geothermal gradien t spatially . A numbe r o f estimate d palaeogeotherma l gradients ar e availabl e fro m V R profile s i n boreholes. Thi s techniqu e relie s o n th e extra polation o f best-fi t V R profile s upward s t o th e appropriate surfac e temperature . Ther e ma y b e problems i n th e us e o f a linea r extrapolatio n where therma l conductivit y contrasts character ize the upper part of the basin fill (Holliday 1993; Allen e t al. 1998) , i n the parameterization o f the temperature-reflectance relatio n (Burnha m &

388

P. A. ALLEN ETAL

Sweeney 1989 ; Rowle y & Whit e 1998) , an d i n the recognition an d removal of the effects o f fluid flow on reflectanc e value s (Gree n e t al 1997) . Nevertheless, V R profiles , augmente d wher e possible b y dat a fro m AF T analysis , hav e bee n shown t o b e usefu l i n th e estimatio n o f palaeogeothermal gradient s (Corcora n & Clayton 1999) . Palaeogeothermal gradient s estimate d fro m VR an d AF T profile s var y considerabl y (e.g . Bray et al, 1992; Gree n 1986 , 1989 ; Lewi s et al 1992; Murdoc h e t a l 1995 ; Scotchma n & Thomas 1995 ; Gree n e t al 1997 ; Dunca n e t a l 1998; Corcora n & Clayto n 1999 ) eve n withi n relatively small sub-basins . For example, palaeo geothermal gradient s o f betwee n 1 0 an d SC^Ckm"1 hav e bee n estimate d i n th e Eas t Irish Se a and Solway basin s (Gree n e t al 1997) . There i s therefor e som e meri t i n estimatin g palaeogeothermal gradient s fo r th e Iris h land mass throug h a therma l mode l derive d fro m subsidence history . A large numbe r of borehole s show a period o f Permo-Triassic to Early Jurassic stretching followe d b y therma l relaxatio n (Rowley & White 1998) , wit h stretc h factor s of less than 1.4. For example, the subsidence history and V R profil e o f Gul f 42/21- 1 (locatio n i n Fig. 1 ) can be satisfactoril y explaine d usin g two periods o f stretchin g (i n Triassi c Sherwoo d Sandstone an d Rhaetia n Penart h Grou p times ) with stretc h factor s (fi) o f betwee n 1. 2 an d 1. 3 and a radiogenic heat contribution from the upper

crust (Alle n e t al 1998) . Th e palaeogeotherma l gradients in this model range from c . 40 °C km~] in Triassi c tim e t o 22°Ckm ~ l i n Lat e Tertiar y time (se e caption s t o Fig s 4-8) . Simila r geotherms ar e implie d (bu t no t directl y given) in th e muc h more extensiv e study by Rowle y & White (1998) . In th e calculatio n o f denudatio n i n Ireland , therefore, w e hav e use d a variabl e geother m t o capture th e earl y stretchin g and post-rif t history of th e region . Th e geother m ha s not , however , been varie d t o accoun t fo r possibl e therma l pulses associate d wit h th e onse t o f widesprea d igneous activit y in th e Britis h Tertiar y igneou s province. Green e t al (1999 ) foun d n o evidence for elevate d Earl y Tertiar y hea t flow s alon g th e UK Atlanti c margin . Fo r simplicity , w e hav e applied th e time-dependen t geother m uniforml y to al l o f th e therma l historie s i n thi s study , although acknowledgin g that the palaeogeother mal gradien t is likely to have also been spatially variable i n respons e t o differen t amount s o f lithospheric stretching , t o th e distributio n o f variably radiogeni c crus t acros s Irelan d an d t o the differin g therma l conductivitie s o f th e sedimentary cover . Maps of denudation since Triassic time We presen t map s correspondin g t o period s o f geological tim e tha t ar e bounde d b y know n

Fig. 4 . (a) Denudation in the Triassic tim e slice (250-210 Ma) using the annealing model of Laslett et al. (1987) and a geothermal gradient of 40 °C km"1, (b) Denudation in the Early Jurassic time slice (210-179 Ma) using the annealing model o f Laslett e t al. (1987 ) an d a geothermal gradien t of 36 °C km"1.

THERMAL AND DENUDATIONAL HISTOR Y O F IRELAN D

stratigraphic 'events ' suc h a s unconformities, o r corresponding t o establishe d biostratigraphi c boundaries. The map s show n i n Fig s 4- 8 illustrat e th e total amount of denudation during that time slice. Denudation rate s ca n b e easil y obtaine d b y dividing the denudation by the time period . Before discussin g th e results, i t is appropriat e to conside r briefl y th e precisio n o f th e denuda tion estimates . Th e precisio n an d resolutio n o f the palaeotemperature var y in space and time, as well a s dependin g o n factor s suc h a s th e uncertainties in the annealin g model s and the estimation o f th e geotherma l gradien t (o r hea t flow and thermal conductivity). Furthermore, th e inferred timin g o f coolin g episode s ha s a n inherent uncertainty . Consequently , i t i s a nontrivial tas k t o combin e thes e multi-dimensiona l uncertainties wit h th e result s int o a visuall y accessible form . As a guide, for a given sample location, we can consider an uncertainty of 10°C on the temperature differenc e betwee n tw o time points in the thermal history. This, coupled with an uncertainty o f 20 % o n a geothermal gradien t of 30 °C, implies a combine d uncertaint y o f abou t 400-500 m o n th e denudatio n estimate . Thus , spatial variations on the maps should be considered with thi s leve l o f uncertaint y i n mind . Th e uncertainties i n timin g o f coolin g ar e les s problematical whe n th e result s ar e considere d over th e relativel y lon g time interval s we adop t here.

389

Triassic time (Fig. 4a) The total amoun t of denudation in Triassic tim e is low, with most of Ireland experiencing < 1 km over the 40 Ma interval. Some areas experience d heating as a result of burial, especially i n the NE of th e island , corresponding t o deposition in the Ulster Basin . The results ar e therefore consisten t with th e know n Triassi c basi n developmen t i n Ulster. Th e result s als o suppor t th e view tha t Permo-Triassic sedimentatio n di d no t cove r Ireland completely , and was restricte d to faul t grabens o r half-graben s (Naylo r 1992) . Maxi mum denudatio n rate s ar e c . 30 m Ma"1, wit h most o f Irelan d experiencin g rate s o f 0-20m Ma"1 ove r th e duratio n o f th e Triassi c time slice . Jurassic time (Figs 4b and 5a and b) The amoun t o f denudatio n continue d t o b e relatively low during Jurassic time. The map for Early Jurassi c tim e show s depositio n i n th e N E and east , whic h i s consisten t wit h know n occurrences o f Liassi c sedimentar y rock s i n Ulster an d th e Kis h Ban k Basin . However , th e general patter n acros s Irelan d i s on e o f wea k denudation a t < 20m Ma"1 . Thi s averag e denudation rat e i s difficul t t o reconcile wit h th e notion tha t th e entir e are a wa s undergoin g subsidence unde r a n extensiv e Earl y Jurassi c sea. Th e rate s o f denudatio n remai n lo w i n

Fig. 5. (a) Denudation i n the Mid-Jurassic time slic e (179-152 Ma) using the annealing mode l of Laslett e t al. (1987) an d a geothermal gradient o f 33 °C km"1, (b ) Denudatio n i n the Late Jurassi c tim e slic e (152-131 Ma) using the annealin g mode l of Laslett e t al. (1987 ) an d a geothermal gradient o f 30 °C km-1.

390

P. A. ALLE N ETAL.

Mid-Jurassic time , wit h area s o f heatin g (sub sidence) i n the north, NE, east and , interestingly, south (Countie s Cor k an d Waterford) o f Ireland . The souther n par t o f th e Leinste r Massi f show s denudation rate s o f 30 m Ma"1, supportin g th e view tha t th e Massi f provide d sedimen t fo r basins locate d i n th e Celti c Se a area . Th e presence o f Middle Jurassi c reddened, continen tal clay s a t Cloyne , Count y Cor k (Fig . 2) , supports th e AFT results. The Late Jurassic ma p shows ver y lo w amount s o f denudation , wit h almost al l th e islan d experiencin g rate s o f < 20 m Ma~1 . The AFT results therefore sugges t limited realm s o f subsidenc e an d wea k denuda tion during Jurassic time, with localized area s of uplift suc h as parts o f the Leinste r Massif. Ther e is no obviou s signatur e of a n early Mid-Jurassi c (mid-Cimmerian) erosiona l event . However, thi s may b e becaus e o f th e poo r resolutio n i n th e thermal mode l o f reheatin g event s withi n a n overall coolin g trajectory , as described above . Cretaceous time (Fig. 6a and b) The ma p fo r Earl y Cretaceou s tim e show s tw o large zone s o f n o coolin g (equivalen t t o deposition) i n th e central-sout h an d S E o f Ireland an d i n th e north . Th e extrem e wes t (County Galway ) an d eas t o f Irelan d sho w denudation rates locally reaching 40 m Ma~l an d 30m Ma"1, respectively . Thi s suggest s tha t basement massif s suc h a s thos e o f Connemar a and Leinste r wer e reactivate d durin g Earl y Cretaceous tectonism , wherea s larg e part s o f

Ireland wer e encroache d b y Earl y Cretaceou s basins. Th e larg e swat h o f Earl y Cretaceou s heating o r subsidenc e i n th e sout h o f Irelan d links wel l wit h th e presenc e o f thic k Lowe r Cretaceous deposit s i n th e Celti c Sea . Th e presence o f a zone of deposition i n the north and NE (alon g th e tren d o f th e O x Mountains Fintona block) suggests that the offshore area s of the Ulster Basin and northern Irish Sea were once covered wit h Lowe r Cretaceou s sediment s tha t have sinc e bee n remove d b y erosio n befor e th e end o f Cretaceou s tim e o r durin g Tertiary time . During Lat e Cretaceou s time , mos t o f Irelan d underwent subsidence , wit h th e notabl e exception o f th e are a o f central-eas t Irelan d i n Counties Lout h an d Meath . Thi s are a experi enced denudatio n rates o f c . 50 m Ma"1 , an d as high as 80 m Ma ' around the shores of Dundalk Bay. Th e prediction s o f denudatio n aroun d th e Antrim coas t ar e no t consisten t wit h th e know n occurrences o f Campania n t o lowe r Maastrich tian Ulste r Whit e Limeston e i n thi s region . W e suspect that some Early Tertiary cooling has been mapped int o th e Lat e Cretaceou s tim e slic e during th e therma l modelling . Nevertheless , th e broad pictur e is supportiv e o f th e ide a tha t with the exceptio n o f east-centra l Ireland , th e are a underwent flooding under Late Cretaceous chalk seas. Thi s is fully consisten t with th e occurrence of th e chal k brecci a a t Ballydeenlea , Count y Kerry (Fig . 2) , an d wit h th e widesprea d distribution of flint pebbles i n tills across Ireland. The hig h amoun t o f Cretaceou s denudatio n i n east-central Irelan d pose s th e interestin g

Fig. 6 . (a) Denudation i n the Early Cretaceous time slice (131-96 Ma) using the annealing model o f Laslett et al. (1987) and a geothermal gradien t o f 27 °C km"1 , (b) Denudatio n i n the Late Cretaceous tim e slic e (96-66 Ma) using th e annealin g mode l o f Laslett e t al. (1987 ) an d a geothermal gradien t o f 24 °C km"1 .

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Fig. 7 . Denudation in the Paleocene-Eocene time slice (66-36 Ma) with a geothermal gradient of 22 °C km \ using (a ) th e annealin g mode l o f Laslett e t al. (1987 ) and (b ) th e annealin g mode l o f Gunnell e t al. (i n prep.).

possibility tha t the absenc e o f Cretaceou s rock s in th e centra l Iris h Se a ma y b e partl y du e t o Cretaceous erosio n an d non-depositio n rathe r than solel y to erosion durin g a Paleocene crusta l uplift event . Paleocene-Eocene time (Fig. la and b) The ma p fo r Paleocene-Eocen e tim e base d o n the Laslett et al. (1987) annealin g model show s a significant chang e i n denudatio n patterns . I t shows a strong spatia l variabilit y (even allowing for uncertaintie s in the estimates), indicative of a patchwork o f localize d area s o f stron g denuda tion (up to 80 m Ma"1) adjacent t o areas of weak denudation o r subsidence . Thes e variation s appear t o b e maintaine d b y th e modifie d annealing mode l (Gunnel l e t al . i n prep.) . Individual faul t block s o r geologica l compart ments presumably acted like piano keys, moving relative to adjacent blocks. In some regions, suc h as the Shanno n Estuar y an d nort h Count y Clar e and Cork harbour, there appear s t o be no readily available explanatio n fo r th e locall y hig h denudation rate s inferre d fro m th e Laslet t e t al . (1987) model. The fact that these regions become areas o f anomalousl y lo w denudatio n i n Oligo cene-mid-Miocene tim e suggest s tha t th e thermal modelling has generated a strong change across th e 3 6 Ma boundary , whic h woul d disappear o r b e dampene d wit h th e choic e o f a different tim e slicing . Also , thes e spatia l variations ar e reduce d whe n th e result s o f th e modified mode l ar e considered and s o should be

treated with caution. The centre of Ireland i s the only broa d zon e o f n o coolin g (deposition ) during Paleocene-Eocene time. There is no clear indication of an increase i n denudation towards a putative Iris h Se a Dom e (se e Cop e 1994 , 1998 ) in th e Paleocene-Eocen e map . Th e averag e amount o f denudatio n i s 500 m give n b y th e Laslett e t al. (1987) annealin g mode l an d 600m for th e modifie d model . Th e latte r therefor e predicts slightl y more denudatio n (althoug h this is probably a t the leve l o f uncertainty) implying samples wer e no t a t lo w enoug h temperature s during this time slice for the differences between the two models t o become apparent . Oligocene-mid-Miocene time (Fig. 8a and b) Almost al l o f Irelan d underwen t denudatio n during thi s tim e interval , wit h rate s varyin g between 2 0 an d 50 m Ma"1 sprea d fairl y uniformly acros s the island. A notable localized region o f ver y hig h denudatio n i s th e Mourn e region o f Count y Down , wher e abou t 3k m o f denudation wa s experience d i n thi s interval , representing a denudatio n rat e o f 120 m Ma"1. This is supported by independent evidence of the unroofing o f th e 5 6 Ma Sliev e Gullion-Mourn e Granites t o thei r present-da y level . Th e pre dominance of denudation over subsidence during Oligocene t o mid-Miocen e tim e i s consisten t with th e widesprea d occurrenc e o f continenta l and lacustrine Tertiary deposits across the island. There i s a suggestio n tha t east-centra l Irelan d (Louth, Meath , Dublin , Wicklow ) experience d

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Fig. 8. Denudation in the Oligocene-mid-Miocene time slic e (36-10 Ma) wit h a geothermal gradien t of 22 °C km"1, using (a) the annealing model of Laslett etal. (1987) and (b) the annealing model of Gunnell etal. (in prep.).

weak denudatio n an d som e subsidence , whic h may lin k wit h th e preservatio n o f middl e t o Upper Tertiar y sedimentar y rock s i n Iris h Se a basins. It should be borne i n mind, however, that the resolutio n o f th e therma l mode l a t tempera tures belo w th e to p o f th e apatit e partia l annealing zon e i s limited , particularl y fo r th e Laslett et al (1987 ) model. This is highlighted by the differenc e i n mea n denudation s betwee n th e Laslett e t al (1987 ) an d modified models, 970 m and 560 m, respectively, the former being a likely overestimate. No map is shown of the 10-0 Ma time interval because o f uncertainties i n the annealing models at suc h low temperatures . The cumulativ e amoun t o f denudatio n fro m Triassic t o mid-Miocen e tim e (1 0 Ma) i s shown in Fig . 9 , demonstratin g th e broa d patter n o f Mesozoic-Tertiary denudatio n (mea n valu e o f 2.7km). Th e area s o f leas t denudatio n ar e N E Ulster, th e Iris h Midland s an d th e extrem e S E (Counties Wexfor d an d Waterford) . Th e lo w values o f cumulativ e denudatio n (<2km ) ove r the sit e o f th e Ulste r Basi n i s a strong , independent validation of the thermochronology, as thi s are a o f Irelan d contain s preserve d Mesozoic-Cenozoic stratigraphy . Th e area s o f maximum denudatio n fro m Triassi c t o mid Miocene tim e ar e easter n Irelan d fro m Dubli n Bay t o County Down, an d to a lesser exten t in a tract o f wester n Irelan d fro m Donega l t o Kerry , and i n th e Leinste r Massi f an d it s southwester n subsurface extension .

The cumulativ e amoun t o f denudatio n i n Tertiary tim e (66-1 0 Ma) show s les s variatio n than th e Triassic t o mid-Miocene ma p (compar e Figs 9 an d 10) . Th e averag e denudatio n fo r Ireland i s c . 1.5k m an d c . 1.2k m usin g th e Laslett e t al . (1987 ) an d modifie d annealin g models, respectively . Bot h model s satisfactorily account fo r th e estimat e o f < 1.5 km Cenozoi c denudation a t Ballydeenlea , Count y Kerry , bu t the match with the new annealing model is better. There i s littl e structur e i n th e ma p o f Tertiar y denudation (Fig . lO a an d b) . Ther e i s n o convincing regional tren d of increasing amounts of denudatio n t o th e eas t tha t migh t reflec t th e effects o f a long-wavelengt h Earl y Tertiar y domal uplif t centre d o n th e Iris h Sea . However , the high amount of denudation recognized alon g the coasta l trac t betwee n Dubli n Ba y an d Dundalk Ba y i s compatibl e wit h th e existenc e of a n Irish Se a high to th e east . The averag e denudatio n i n th e las t 1 0 Ma i s 650m a s give n b y th e Laslet t e t al . (1987 ) annealing mode l an d 230 m a s give n b y th e modified annealin g model. Total denudationa l efflu x o f the Iris h landmass The maps shown in Figs 4-8 portra y amounts of denudation durin g chose n tim e slices . W e ar e also abl e t o monito r th e denudatio n rat e integrated across the entire present-day landmass

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Fig. 9 . Cumulativ e amoun t of denudatio n fro m Triassi c tim e (25 0 Ma) t o 1 0 Ma usin g the annealin g mode l o f Laslett e t al. (1987). Inset show s histogram o f denudation (in m).

of Ireland as an essentially continuou s (time step 1 Ma) functio n o f time . Th e tota l ne t sedimen t discharge o f Ireland a s a function o f time can be simply calculate d b y multiplyin g th e averag e denudation rate per square kilometre b y the total land surfac e area o f Ireland . Two assumptions need to be considered. First , we assume a zero porosity of bedrock undergoing denudation. Second , w e assum e tha t al l o f th e denudation o f th e Iris h landmas s resultin g i n delivery o f particulat e sedimen t loa d t o neigh bouring depocentre s i s b y physica l processes .

Whereas th e first assumption is trivial bearing in mind th e orde r o f magnitud e estimate s bein g made, th e secon d assumptio n i s potentiall y important. Th e rati o o f solut e t o particulat e loads i n present-da y Iris h river s i n S E Irelan d varies in relation to the bedrock lithologies of the individual catchments . However , th e genera l pattern i s tha t solut e loa d dominate s ove r particulate loa d (Malon e 2001 , unpubl . data) . Consequently, an y estimat e o f th e particulat e sediment discharg e fro m Irelan d t o offshor e basins as a function of geological tim e is likely to

Fig. 10. Cumulative amount of denudation from 66 to 10 Ma, using (a) the annealing mode l of Laslett et al. (1987) and (b ) th e annealin g model o f Gunnell et al. (i n prep.). Insets sho w histograms o f denudation (i n m).

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be a significan t overestimate. Th e exten t o f th e overestimation wil l b e a comple x functio n o f largely unconstraine d climatic , topographi c an d bedrock variations throug h geological time in the area no w occupied b y th e Irish landmass . The sedimen t yield s an d discharge s o f individual ancien t sedimen t routin g system s can be calculate d wit h knowledg e o f th e forme r spatial exten t o f palaeocatchments . Thi s pro cedure form s par t o f continuin g research an d i s not reporte d here . The long-ter m patter n (Fig s 1 1 an d 12 ) i s o f moderate rate s o f denudatio n an d discharg e during Earl y t o Mid-Jurassi c tim e decreasin g through Late Jurassic time to low values through much o f Earl y an d Lat e Cretaceou s time . Denudation rate s an d sedimen t discharge s increase throug h Lat e Cretaceou s time , wit h peaks i n Maastrichtian-Paleocene , Eocen e an d Oligocene time . Thi s genera l patter n i s retained with a constan t geother m wit h tim e an d wit h a geotherm involvin g post-stretchin g therma l relaxation usin g th e annealin g mode l o f Laslet t et al. (1987 ) (Fig . 11) . A secon d se t of curves is shown fo r th e modifie d annealin g mode l o f Gunnell et al. (i n prep.), which reduces th e high cooling rate s a t lo w temperature s predicte d b y the Laslett et al (1987) model (Fig. 12) . With the

Laslett et al. (1987) annealing model, denudation rates an d discharge s increas e furthe r throug h Neogene time , wherea s wit h th e ne w annealing model, rate s fal l slightl y afte r a pea k i n Eocene-Oligocene time . A compariso n o f these curve s als o implie s tha t peak s i n th e denudation chronolog y inferre d fro m th e modified annealin g model ten d t o occu r slightly earlier tha n i n th e Laslet t e t al . (1987 ) model . However, thes e timin g difference s ar e poten tially withi n th e uncertaintie s inheren t i n th e modelling. The result s in Figs 1 1 and 1 2 can be compared with estimate s o f soli d sedimen t volume s in th e offshore basins . Th e Porcupin e Basi n ha s a surface are a o f c. 20 000 knr. The tota l volume of Cenozoi c soli d sedimen t i n th e Porcupin e Basin (Jone s 2000 , unpubl . data ) i s 54 900 ± 6300km 3. The combined Erri s Trough and eastern flank of the Irish Rockall Basin has a surface are a of 20 000 km2 (Spencer et al 1999) . Approximating th e averag e Cenozoi c soli d sediment thicknes s a s 3- 4 km, th e volum e o f Cenozoic sedimen t i s o f th e orde r o f 60 000-80 000 km3. Th e smal l surfac e are a (c. 5000km 2) an d thi n cove r o f Cenozoi c sediment i n the Slyn e Trough imply that relative to th e large r Porcupin e an d Iris h Rockal l

Fig. 11 . Average denudatio n rat e an d maximum volumetri c discharg e fo r the Irish landmas s since Triassic tim e for a constant geotherma l gradien t (blu e lines), and a time-dependent geotherm (re d lines).

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Fig. 12 . Average denudation rat e an d maximum volumetri c discharg e fo r the Irish landmass sinc e Triassi c tim e with a ne w annealin g mode l (gree n lines ) tha t remove s th e artefac t o f hig h denudatio n a t near-surfac e temperatures (Gunnel l e t a l i n prep.) . Result s fro m th e Laslet t e t al . (1987 ) annealin g mode l show n fo r comparison (re d lines). The maximu m volume of sedimen t exporte d fro m the are a now occupie d by the Iris h landmass i s a n orde r o f magnitud e smalle r tha n th e preserve d soli d volum e o f Cenozoi c sedimen t i n th e Porcupine-Rockall basin system.

depocentres i t i s unimportan t i n calculatin g sediment volumes . Th e tota l Cenozoi c soli d sediment volum e i n th e offshor e basin s t o th e west o f Irelan d i s therefor e wel l i n exces s o f 100000km3. The average denudatio n rat e of Ireland durin g Cenozoic tim e derive d fro m th e ne w annealin g model is 20-25m Ma"1, andc. 40m Ma"1 using the Laslet t et al, 198 7 annealing model. Ove r a period o f 6 5 Ma, wit h a present-da y are a o f Ireland of 8344km 2, this gives a maximum total sediment volum e o f between c . 1 0 000km3 and 20 000km 3, about a factor of 5-10 smaller than the soli d volum e occupyin g th e Porcupin e an d Irish Rockal l basins . I f w e wer e t o deriv e thi s volume o f sedimen t solel y fro m th e Iris h mainland, w e woul d nee d denudatio n rate s o f about 180-23 0 m Ma"1, an d thi s woul d impl y cooling rate s o f 5-8 °C Ma"1 fo r a 30°Ckm" 1 geotherm. W e would then predict al l sample s t o be at temperatures i n excess o f 120°C , that is, at total annealing, around 15-25 Ma. We would not expect t o measur e an y fission-trac k age s olde r than this . A t suc h hig h denudatio n rates , th e assumption o f a linea r geother m woul d b e

invalid, but inferre d fission-trac k age s woul d be even younge r tha n 2 5 Ma. Clearly , i t i s no t possible t o invok e the curren t Iris h mainlan d as the dominan t sourc e regio n fo r th e offshor e basins on the Atlantic margin. In searching for a source o f sedimen t fo r th e wester n offshor e basins, i t i s necessar y t o loo k fa r beyon d th e present-day landmass of Ireland, such as the Irish Sea an d th e continenta l shel f borderin g th e Rockall margin . A ke y proble m fo r futur e researc h i s th e delineation o f th e catchment s feedin g th e offshore basin s i n Tertiar y time . Continuin g studies are attempting to map out palaeochanne l systems an d mode l th e palaeotopographi c evolution o f Irelan d s o a s t o mak e prediction s about Tertiar y sedimen t routin g patterns i n thi s sector o f th e N W Europea n Atlanti c margin. In addition, i t i s necessar y t o conside r th e tim e correlation betwee n denudatio n an d deposition, so as to refine the spatial links between the source areas an d sedimen t delivery into specific basins . Finally, it will be necessary t o undertake detailed provenance studie s on the sediment s themselves and th e potentia l sedimen t sourc e region s t o

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fingerprint an d correlat e thei r geochemica l o r geochronological signatures .

Conclusions Although ver y littl e i s know n abou t th e post Variscan histor y o f Irelan d base d o n it s fragmentary stratigraphi c record , w e ar e abl e t o make som e progres s i n assessing its thermal and denudational history through apatite fission-track analysis o f sample s collecte d fro m surfac e exposures. Reconstructe d therma l historie s ar e used to create map s of denudation for time slice s since th e beginning of Triassic time. The denudatio n map s i n genera l sho w a patchwork o f relativel y small-scal e spatia l variations i n denudation . I n Triassi c time , mos t of Irelan d underwen t denudatio n a t rate s o f 0-30m Ma"1, wit h som e areas , suc h a s N E Ireland, undergoin g subsidenc e during this time. This patter n i s supported b y th e preservatio n o f Permo-Triassic stratigraph y i n th e Ulste r Basin . Denudation continue d t o b e relativel y lo w i n Jurassic time , wit h zones o f subsidence , suc h as southern Irelan d adjacen t t o know n Jurassi c depocentres i n th e Nort h Celti c Sea , an d loca l regions o f crusta l uplift , suc h a s th e Leinste r Massif. Durin g Earl y Cretaceou s tim e tw o larg e zones o f subsidenc e wer e establishe d i n th e central-SE an d nort h o f Ireland , wherea s base ment massif s suc h a s thos e o f Connemar a an d Leinster underwen t denudation at < 40m Ma"1. In Late Cretaceous time most of Ireland subsided with th e exceptio n o f a regio n centre d o n Dundalk Ba y i n N E Leinste r an d S E Ulster , which experience d denudatio n rate s a s hig h a s 80m Ma"1. This wid e distributio n o f heating o r subsidence support s the idea that most of Ireland was covere d b y a venee r o f Uppe r Cretaceou s chalk. Th e denudatio n ma p fo r th e Paleocene Eocene interva l i s spatiall y variable , suggestin g that faul t block s o r geologica l compartment s reacted lik e pian o key s t o a regiona l tectoni c event. There is no strong indication of a regional increase i n denudatio n acros s Irelan d toward s a putative Iris h Se a dome , bu t th e relativel y local high denudation rates in central-east Ireland are compatible wit h th e existenc e o f a n Earl y Tertiary focu s o f denudatio n i n th e Iris h Se a that impinged o n the region now occupied b y the Irish landmass . Almost al l o f Irelan d underwent moderate rate s o f denudation in the Oligocenemid-Miocene time, with the notable exception of the Mourn e regio n o f Count y Down , whic h suffered c . 3k m o f denudatio n i n thi s tim e interval (120 m Ma"1 ), consisten t wit h th e required rapi d unroofin g o f th e 5 6 Ma Mourn e and Sliev e Gullio n granites .

The tota l discharg e o f sedimen t fro m Irelan d as a function of time, and the average denudation rate, show s moderat e value s i n Triassi c time , falling throug h Jurassic time t o ver y lo w values in Cretaceou s time . Discharge s increas e signifi cantly into Tertiary time. However, the maximum volumetric discharg e o f sedimen t i n Cenozoi c time is considerably smaller than that implied by the preserved solid volume of Cenozoic sediment in the Porcupine-Irish Rockall basi n system , by a facto r o f abou t 5-10 . Consequently , i n searching for a source for the Cenozoic sediment of the offshore basins, one mus t look fa r beyond the present-da y landmass of Ireland. We ar e gratefu l t o th e Irelan d Frontie r Exploratio n Licence 02/9 7 (ARCO , BG , Anadarko ) fo r fundin g during 1998-1999 , an d S . Bergma n (ARCO ) i n particular fo r assistanc e an d advice . W e receive d assistance i n samplin g fro m G . Sevastopul o an d J. Graha m (Trinit y Colleg e Dublin) , an d I . Meigha n (Queen's Universit y Belfast), and thank A. Gleadow, P. O'Sullivan an d J . Murph y fo r providin g u s wit h thei r unpublished fission-trac k dat a fro m Donega l an d th e Mournes. W e als o gratefull y acknowledg e th e collaboration an d review s o f N . Whit e an d S . Jone s at Cambridge University.

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Prediction of the hydrocarbon system in exhumed basins, and application to the NW European margin A. G. DORE1, D. V. CORCORAN2 & I. C. SCOTCHMAN1 l

Statoil (U.K.) Ltd, lla Regent Street, London SW1Y 4ST, UK (e-mail: agdo@ statoil.com)

2

Statoil Exploration (Ireland) Ltd, Statoil House, 6, St George's Dock, IFSC, Dublin 1, Ireland Abstract: Uplift , erosio n and removal of overburden have profound effects o n sedimentary basins an d the hydrocarbon system s they contain. Thes e effects ar e predictable from theor y and from observation of explored exhumed basins. Exhumed basins are frequently evaluate d in the same way as 'normal' subsiding basins, leading to errors and unrealistic expectations . In this paper we discuss the consequences o f exhumation in terms of prospect risk analysis, resource estimation , an d overall basin characteristics . Exhumation shoul d be taken into account whe n assigning risk factors used to estimate th e probability o f discover y fo r a prospect. I n general, exhumatio n reduces th e probabilit y o f trapping o r sealin g hydrocarbons, except wher e highly ductile seals suc h as evaporites are present. Exhumatio n modifie s th e probabilit y o f reservoi r i n extrem e cases ; fo r example , where a unit may hav e been burie d s o deeply befor e uplift tha t it is no longe r a n effectiv e reservoir, o r wher e fracturin g o n uplif t ma y hav e create d a n entirel y ne w reservoir . Th e probability o f sourcin g o r chargin g i s affecte d b y multipl e factors , bu t primaril y b y th e magnitude o f th e post-exhumatio n hydrocarbon budge t an d th e efficienc y o f remigration . Generally ga s will predominate as a result of methane liberatio n fro m oil , formation water and coal , an d becaus e o f expansio n o f ga s trappe d befor e uplift . Thes e factor s i n combination ten d to result i n gas flushing of exhumed hydrocarbon basins. Compared wit h a similar prospect in a non-exhumed basin, resource level s of a prospect in an exhume d basin ar e generall y lower . Higher level s o f reservoir diagenesi s influenc e th e standard parameters used to calculate prospect resources. Porosity, water saturation and netto-gross rati o ar e adversel y affected , an d (a s a consequenc e o f al l three ) lowe r recover y factors ar e likely . Hydrostati c or near-hydrostati c fluid pressure gradient s (a s observe d i n exhumed NE Atlantic margin basins) will also reduce the recovery factor and, in the case of gas, wil l adversel y affec t th e formation volum e factor . Hydrocarbon system s in exhume d setting s show a common se t of characteristics . The y can include : (1 ) large , basin-centre d ga s fields ; (2 ) smaller , peripheral , remigrate d oi l accumulations; (3 ) two-phase accumulations; (4 ) residual oil columns; (5) biodegraded oils; (6) underfilled traps . Many basins on the N E Atlantic seaboard underwen t kilometre-scal e uplift durin g Cenozoi c tim e an d contai n hydrocarbo n system s showin g th e effect s o f exhumation. Thi s knowledg e ca n constrai n ris k an d resourc e expectatio n i n furthe r evaluation of these basins, and i n unexplored exhumed basins.

Global oi l an d ga s resource s ar e finit e an d availabilit y an d thu s t o location . I n bot h cases , depleting rapidly. Estimates as to when world oil th e result is a drive to explore th e more difficult , production wil l begi n it s termina l declin e vary , highe r ris k basins . but al l authorities agree that this must take place Althoug h n o hydrocarbo n basi n i s withou t within a fe w decade s (e.g . Campbel l 1996) . exploratio n problems, a n 'ideal' basin is perhaps Natural gas , th e logica l short-ter m replacemen t on e containin g abundan t reservoir s an d ric h for oil , is more abundant and may provide global sourc e rocks , whic h i s continuousl y subsidin g supply fo r abou t a centur y at projecte d rate s o f an d wher e hydrocarbon s are bein g generate d a t consumption (Lerch e 1996) . Therefore , th e th e presen t day , replenishin g thos e tha t lea k t o search fo r oi l ma y b e sai d t o b e enterin g it s th e surface . Example s includ e th e Norther n 'end game' , characterize d b y increasin g diffi - Nort h Sea , th e Gul f o f Mexic o an d th e Sout h culty i n locating majo r ne w reserves. I n the case Caspia n Basin. Most such basins are now known of gas , economic attractivenes s i s tied to market an d unde r production . Exhume d basins , o n th e From: DORE , A.G., CARTWRIGHT, J.A. , STOKER, M.S. , TURNER, J.P. & WHITE , N . 2002 . Exhumation o f th e North Atlantic Margin: Timing, Mechanisms an d Implications for Petroleum Exploration. Geologica l Society, London, Special Publications, 196, 401^29. 0305-8719/027$ 15.00 © The Geological Society of London 2002.

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Fig. 1 . Location ma p o f the NW Europea n Atlanti c margin , showin g basin s affecte d b y Cenozoic exhumation .

PREDICTION OF HYDROCARBON SYSTE M IN EXHUMED BASIN S

other hand , belon g i n th e highe r ris k category , and will become increasingly important as global resources diminish . For th e purpose s o f thi s paper , a n exhume d basin is loosely defined as one that has undergone uplift an d erosion , suc h tha t th e sedimentar y rocks tha t constitut e th e petroleu m syste m (source, reservoi r an d seal ) ar e significantl y shallower no w than in the past (see mor e forma l definitions give n b y Rii s & Jense n (1992 ) an d Dore & Jensen (1996) and Dore et al (2002)) . In such basins, the rock properties an d hydrocarbon systems will be radically different fro m those at a similar depth i n a continuously subsiding basin. These properties ca n b e studie d b y observatio n of drille d exhume d basins , o r predicte d b y modelling an d experimentatio n (e.g . Nylan d et a l 1992 ; Dor e & Jense n 1996) . Althoug h most effect s ar e individuall y understood , the y are rarel y studie d systematicall y i n th e initia l exploration o f a n exhume d basin . Together , however, the y constitut e a powerfu l predictiv e tool. In th e past , inappropriat e compariso n o f th e exploration potentia l o f exhume d basin s wit h 'classic' subsidin g basin s ha s resulte d i n unfulfilled expectations . Realizatio n early i n the exploration proces s tha t a basi n ha s bee n exhumed give s ris e t o a differen t approac h t o hydrocarbon exploration , an d ca n hel p t o constrain resourc e prediction . I n thi s pape r w e systematically describe the effects o f exhumation with reference t o two of the standard procedure s of petroleu m exploration : (1 ) estimatio n o f th e probability o f findin g hydrocarbon s i n a pro spect; (2 ) calculatio n o f its volumetri c resourc e potential. Then , base d o n thes e discussions , w e derive a generalized set of key characteristics for exhumed petroleum basins. It i s no w generall y understoo d tha t majo r uplift an d erosion took place in the circum-North Atlantic borderland s durin g Cenozoi c time , transforming a regio n dominate d b y lo w relie f and shelf seas (Late Cretaceous) to one bordered by highlands such as Norway, Scotland and East Greenland (e.g . Rii s & Fjeldskaa r 1992 ; Rii s 1996; Dor e e t a l 1999 ; Japse n & Chalmer s 2000). I t i s als o apparen t tha t man y o f th e offshore basin s margina l t o th e landmasse s underwent Cenozoi c uplif t an d erosion . Thes e include basin s wher e hydrocarbo n system s ar e proven (th e wester n Barent s Se a an d Hord a Platform (Norway) , Inne r Mora y Firth , Wes t Shetland Basi n an d Eas t Iris h Se a Basi n (UK ) and the North Celtic Se a and Slyne-Erris Basin (Ireland)) in addition t o many unexplored basin s (Fig. 1) . Exhumatio n i n thes e area s ha s bee n quantified usin g numerou s methods , includin g

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seismic velocities , shal e velocities , vitrinit e reflectance, apatit e fissio n track , mas s balanc e and basin restoratio n (e.g . Rii s & Jensen 1992) . These measurement s sho w tha t Cenozoi c uplif t around th e Nort h Atlanti c wa s geographicall y variable an d took plac e i n several phases . Thre e of thes e phase s hav e particularl y widesprea d significance: a n uplif t o f Paleocen e age , gener ally though t to be associate d wit h effect s o f the Iceland plume and incipient opening of the North Atlantic (e.g . Whit e 1988 ; Whit e & McKenzi e 1989; Milto n e t a l 1990) , a n Oligo-Miocen e episode usuall y associate d wit h inversio n (e.g . Underhill 1991 ; Murdoc h e t a l 1995 ; Parnel l et a l 1999 ) an d a Neogen e (primaril y Plio Pleistocene) event of more enigmatic origin (e.g. Solheim et al 1996 ; Dor e et al 1999 ; Japse n & Chalmers 2000) . A discussio n o f alternativ e uplift mechanism s is not within the scope of this paper, except to note that explanations essentially fall int o three groupings : (1) isostati c (respons e to erosiona l unloading) ; (2) therma l (associate d with a mantl e plume) ; (3 ) compressiona l (intraplate stres s an d inversion) . Categorie s (1 ) and (2 ) ar e broa d regiona l effects , wherea s category (3 ) may be very loca l in nature. In most cases, exploratio n o f these basins ha s occurred withou t a ful l understandin g o f thei r exhumed nature . A particularl y instructiv e example i s th e wester n Barent s Se a (offshor e Norway), wher e licencin g i n th e earl y 1980 s carried th e hop e o f a majo r Nort h Se a rift-type hydrocarbon province , bu t wher e expectation s were radicall y revise d a s th e effect s o f uplif t became apparen t durin g earl y exploratio n (Nyland et al 1992 ; Dore 1995) . This experienc e alerted researchers, particularl y in Norway, to the widespread natur e o f th e exhumatio n and , importantly, t o it s commercia l consequences . We therefore believe that a template for prospect evaluation o f th e typ e presente d here , althoug h largely qualitative , ca n b e beneficia l i n futur e exploration an d resourc e evaluatio n o f suc h provinces. Although specific reference is made to the N W Europea n basins , thi s appproac h is applicable to any exhumed basin. Prospect risk analysis in exhumed basins Introduction Explorationists addres s prospec t ris k a t tw o levels: (1) risk with respect t o the validity o f the prospect, i.e . the chance of success; (2 ) the range of possibl e reserves. I n thi s account , w e simpl y address the chance before drilling of discoverin g any hydrocarbon accumulatio n withi n a mapped prospect. Method s use d t o calculat e thi s figur e

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Fig. 2 . Factor s t o b e considere d whe n carryin g ou t ris k analysi s o f a prospec t i n a n exhume d basin . Arrow s indicate increased or decreased probabilit y in exhumed settings.

PREDICTION OF HYDROCARBON SYSTEM IN EXHUMED BASIN S

vary widel y withi n th e oi l industry , but ca n b e generalized i n an equation of the type where P hc i s th e probabilit y o f findin g hydro carbons, oi l o r ga s (thi s valu e i s sometime s regarded a s th e chanc e o f findin g an y hydro carbons at all, some specified minimum quantity, or hydrocarbons capable of flowing to the surfac e (see, e.g. Snow et al. 1996)) , P T is the probability that a reservoi r roc k i s present , capabl e o f holding oil or gas, Ps is the probability tha t there is a source rock that has charged the prospect and Pt i s th e probabilit y tha t a seale d tra p exist s capable o f holdin g oi l o r gas . Th e chanc e o f a particular hydrocarbo n phas e bein g presen t (assuming a single-phas e accumulation) is then given by where P 0 an d P g ar e th e probabilitie s o f discovering oi l an d gas , respectively , a s th e main phase. All o f these factors are subjectively estimated by th e petroleu m geoscientis t befor e drilling . Where basins ar e mature in terms of exploration they ca n b e calibrate d agains t know n succes s rates. Th e influenc e o f exhumatio n o n thes e factors i s summarize d i n Fig . 2 an d discusse d below. Probability of reservoir Reservoir rock s ca n b e bot h enhance d o r degraded i n a n exhume d terrane , dependin g o n the typ e o f reservoi r an d th e natur e o f th e process. Degradatio n compare d wit h a reservoir at a similar depth in a subsiding basin can occur as a result o f inheritance o f a compactional an d diagenetic stat e reflectin g a previousl y greate r burial (e.g . Walderhau g 1992) . Improvemen t may occu r becaus e o f fractur e enhancemen t of porosity an d permeabilit y durin g uplif t (e.g . Aguilera 1980) . However , risk analysis does not take into account the quality of a reservoir, only whether a suitable reservoi r exist s or not. Thus, exhumation i s probabl y neutra l fo r th e prob ability o f reservoi r excep t i n extrem e cases ; fo r example, where a unit has been buried so deeply before uplif t tha t i t i s n o longe r a n effectiv e reservoir, o r wher e fracturin g o n extrem e uplif t has created an entirely ne w reservoir. Th e much more importan t effec t o n reservoi r volumetri c parameters (porosity , net/gros s rati o an d hydrocarbon saturation ) an d recover y facto r i s examined i n th e sectio n o n prospec t resourc e estimation.

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Probability of source and charge The probabilit y tha t a sourc e roc k i s presen t i s independent of whether or not the basin has been exhumed. However , th e probabilit y tha t suc h a source roc k ha s bee n effectiv e i n chargin g a reservoir is strongly influenced by exhumation. A complex interpla y o f positiv e an d negativ e factors mus t be considered . In exhume d basins , an y sourc e roc k wil l b e more matur e than expected for its present depth . Therefore, ther e will be an increased probabilit y that source rocks now lying shallower than the oil generation window will have generated oil in the past. Whethe r suc h oi l ha s survive d uplift i s a separate question . Thi s reasonin g ha s bee n applied t o evaluat e ris k i n margina l basin s around Norway, where the Upper Jurassic sourc e rock i s shallo w (Ghazi 1992 : Jensen & Schmidt 1993). Conversely , a more deepl y burie d sourc e rock may previously have been below oil window depths, increasin g th e chanc e o f ga s o r tha t th e source rock is 'burn t out' (postmature) . This risk has been evaluated for uplifted Upper Palaeozoic source rock s i n th e easter n Norwegia n Barents Sea (Thei s e t al . 1993) . I n mos t case s som e indication o f th e degre e o f exhumatio n ca n b e obtained even before drilling (e.g . from regiona l setting, seismi c velocitie s an d structura l model ling) an d can therefore b e used to derive a firstorder estimat e o f sourc e rock maturit y attained before uplift . A critica l observatio n i s that , whateve r th e maturity state, generation is curtailed once uplif t commences. Usin g th e widel y accepte d kineti c model fo r generatio n o f hydrocarbon s fro m kerogen (Tissot & Espitalie 1975), which stresses the importance of temperature rather than time, it follows tha t any sourc e rock that is significantl y uplifted throug h a therma l fram e o f referenc e will ceas e hydrocarbo n generation . N o gener ation can occur until the basin subsides again and the previou s maximu m temperature i s reached . In exhumed basins, no new hydrocarbons will be available to charge traps vacated during uplift by spillage an d sea l failure , or newl y created trap s (for example , fold s forme d durin g inversion related uplift) . Thus , an y remainin g origina l charge must have survived both the uplift and any subsequent leakage (e.g . by diffusio n i n the case of gas ; Krooss et al. 1992) , thereby significantl y increasing risk . An exceptio n occur s wher e th e uplifted basin is on the migration rout e for oil or gas generate d in an adjacent, continuously subsiding basin. On the N W Europea n margi n suc h area s includ e parts o f th e Hord a Platform , Inne r Mora y Firt h Basin an d Wes t Shetlan d Basin . I n suc h area s

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hydrocarbons los t durin g th e uplif t proces s ca n be replenished , an d th e probabilit y o f sourcin g must therefor e tak e int o accoun t migratio n timing, rout e an d efficienc y fro m th e adjacen t kitchen (e.g . Skjerv0 y & Sylt a 1993 ; Goodchil d et al 1999 ; Parnel l et al 1999) . Hydrocarbon chargin g followin g an episode of uplift, i n th e absenc e o f newl y generate d hydrocarbons, ca n tak e plac e onl y wit h hydro carbons alread y presen t i n th e basin . Th e simplest o f thes e processe s i s b y remigration . In uplifte d basin s numerou s processe s ca n displace an d redistribut e hydrocarbon s fro m pre-existing accumulations . Vertical remigratio n is a well-know n phenomeno n i n subsidin g basins, wher e reservoi r overpressur e build s until i t exceed s th e sealin g capacit y o f th e caprock, causin g hydrocarbon s t o escap e t o shallower level s (e.g . i n th e Gul f o f Mexic o (Lopez 1990) , i n th e Nort h Se a Centra l Grabe n (Taylor et al 1999 ) an d in the Faeroe-Shetland Basin (Illiff e e t a l 1999)) . A s show n late r (se e probability o f trap an d seal ) an d by Corcora n & Dore (2002) , sea l failur e can als o b e anticipated during uplift . This means tha t hydrocarbons ma y subsequently accumulat e i n previousl y uncharged shallowe r levels . A notabl e exampl e of such charging i s in the Zagros fold belt in Iran and Iraq , wher e sea l failur e i n Cretaceou s reservoirs durin g th e Lat e Miocene-Recen t Zagros uplif t expelle d oi l upward s int o th e highly prolifi c Oligocene-Miocen e Asmar i reservoir (Al a 1982 ; Bordenav e & Burwoo d 1989). However , i n man y case s hydrocarbon s must b e los t t o th e surfac e durin g uplift . Assessment o f thi s ris k mus t tak e int o accoun t the positio n an d geometr y o f th e shallowe r reservoirs, an d th e amoun t o f erosio n o f shallower level s tha t ha s take n plac e durin g uplift. Assuming tha t uplift i s not perfectly uniform, lateral remigratio n wil l occur through tilting and spilling, resultin g i n los s o f som e hydrocarbon s from pre-existin g trap s an d migratio n updip . Although som e hydrocarbon s wil l accumulate in updip traps , ther e wil l b e a ne t decreas e i n hydrocarbon budge t a s a resul t o f migratio n losses an d escap e t o th e surface . I n pre-existin g two-phase accumulation s th e oi l le g wil l preferentially spill . I n th e Hammerfes t Basi n area of the Barents Sea, residual oi l legs in field s such a s Sn0hvi t an d Askeladde n sho w tha t significant oi l spillag e too k plac e durin g Plio-Pleistocene uplift (Kjemperu d & Fjeldskaar 1992; Nylan d e t al 1992 ; Dor e & Jensen 1996) . It has bee n assume d b y som e worker s tha t most of the oi l from th e Middle Jurassi c reservoi r ha s been los t t o th e surface , bu t recen t carefu l

modelling o f the direction o f remigration ha s led to the discovery of a new oilfield (Goliath) on the periphery o f th e basi n (B . Wandaas , pers . comm.). Al l o f th e basin s immediatel y adjacen t to the Norwegian landmass (e.g. Horda Platform, Egersund Basin) will have been tilted westwards during th e Cenozoi c uplif t o f Scandinavia , an d hydrocarbon redistributio n b y spillag e ca n confidently b e inferre d ther e (Dor e & Jense n 1996). Risk associate d wit h charging i n a n exhume d basin can be mitigated in the following ways: (1) by assessin g th e residua l hydrocarbo n budge t after 'switchin g of f o f sourc e roc k maturation; (2) by recognizing the hydrocarbon displacement drivers i n th e basin ; (3 ) by identifyin g the post exhumation regiona l spil l direction ; (4 ) b y determining remigration pathways and bypassed areas ('shadows') . Timin g o f tra p formatio n (pre-, syn - o r post-uplift ) i s als o critica l wit h respect t o charg e adequacy . Probability of oil versus gas Perhaps th e mos t radica l effec t o f uplif t o n hydrocarbon basin s i s th e shif t toward s gas dominated systems. Several phenomena conspire to produce thi s effect, a s follows . (1) Ga s exsolutio n fro m oil . Assumin g a reservoired flui d i s a t o r belo w it s bubble-point pressure at maximum burial, a gaseous phase will be exsolve d on uplift an d consequent pressure temperature reduction . Unles s a tra p i s initially underfilled, oi l will therefore b e driven fro m th e trap (Nylan d et al 1992 ; Dore & Jensen 1996) . (2) Ga s expansio n a s a resul t o f pressure temperature reduction . This will result in net loss from gas-fille d trap s and , again , preferentia l displacement o f th e oi l le g i n two-phas e accumulations. Nylan d e t a l (1992 ) estimate d that over two billion barrels of oil were lost fro m the Sn0hvi t Fiel d i n th e Barent s Se a becaus e of the combine d effec t o f ga s exsolutio n an d expansion durin g regional uplift . (3) Methan e liberatio n fro m formatio n brine . At constant pressure, th e solubility of methane in water attain s a minimu m between abou t 6 0 an d 90 °C, the n increase s wit h temperature . A t constant temperature , methan e solubilit y increases wit h pressure . Therefore , i n general , methane solubilit y i n formatio n brine s wil l increase wit h buria l and , conversely , fre e methane wil l b e liberate d durin g uplif t a s pressure an d temperatur e decreas e (e.g . Culberson & McKetta 1951 ; Price 1979 ; Crame r et a l 1999 ; Crame r & Poelcha u 2002) . Thi s mechanism i s perhaps th e mos t poten t source of gas durin g exhumatio n a s ga s wil l b e liberate d

PREDICTION OF HYDROCARBON SYSTE M IN EXHUMED BASIN S

over the whole basin, tapping methane originally generated by dispersed organi c matter as well as from ric h sourc e rocks . Th e potentia l fo r ga s liberation is vast. All of the giant dry gas fields in the Wes t Siberia n Basi n probabl y deriv e fro m this process (Cramer et al 1999) , a s do the major basin-centred ga s fields in uplifted basins i n the western US A suc h a s th e Alberta, Denve r an d San Jua n Basin s (Dor e & Jense n 1996 : Pric e 2002). Ga s field s i n uplifte d basin s i n Centra l Europe (Pannonian, Vienna and Po Basins) have been tie d directl y t o a groundwate r origin vi a noble ga s marker s (Ballentin e e t al . 1991) . Exsolution from water probably also accounts for a significan t part o f th e hydrocarbo n budge t i n gas-dominated uplifte d basin s o n th e N W European margi n suc h a s the Barent s Sea , Eas t Irish Se a Basin , Slyne-Erri s Basin s an d Nort h Celtic Se a Basin . T o date , however , th e onl y quantitative wor k i n suc h area s know n t o th e authors i s tha t reporte d fo r th e Barent s Se a b y Dore & Jensen (1996) . (4) Methan e expulsio n fro m coal . Coa l bed s are widesprea d i n man y petroleu m basin s an d expel ga s a s th e coa l i s progressivel y burie d through maturatio n thresholds . However , coal bed methan e studie s sho w tha t significan t quantities o f ga s ar e retaine d i n th e coa l o n internal surface s (adsorption ) o r withi n th e molecular framewor k o f th e organi c matte r (absorption). Thi s methan e wil l migrat e ou t o f the coa l by desorptio n and diffusio n durin g reduction i n pressur e (Littk e & Leythaeuse r 1993, Fig. 4; Rice 1993) . Expulsion during uplift is als o aide d b y increas e o f macroporosit y an d permeability i n the coal s compare d wit h deeply buried coals of similar rank, presumably partly as a result of fracture dilation (Littke & Leythaeuser 1993, Fig . 5) . (5) Hydrodynami c flow . Durin g subsidence , compaction-driven wate r flo w outward s t o th e flanks of the basin is normal, whereas the outcrop of aquifer s an d developmen t o f topograph y during uplif t an d erosio n ma y revers e thi s situation an d creat e gravity-drive n wate r flo w towards th e basi n centr e (se e furthe r discussio n by Corcora n & Dor e (2002)) . A s show n b y Cramer e t al . (1999) , thi s flo w provide s a recharge mechanis m whereb y methan e ma y b e brought in from outside the normal drainage area of a gas field and liberated a s a result of drop in reservoir pressure . Crame r e t al . estimate d tha t about 12 % o f th e ga s reserve s o f th e gian t Urengoy Field (Western Siberia ) wer e emplace d in this way. Additional biogenic methane may be introduced during or after uplif t b y groundwater flow throug h coa l beds , a proces s tha t ca n stimulate bacteria l acivit y an d ga s productio n

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from coa l o f an y ran k (Ric e 1993) . Finally , introduction of fresh groundwater into the system is also likely to promote bacterial biodegradation of any shallow oils within the aquifer, leading t o the formatio n o f heav y oi l residue s an d agai n changing the oil-gas balance . A changed oil-gas balance ca n also occur via retrograde condensation . I n thi s cas e oi l i s favoured a s a result of the dropping out of liquids from a wet gas during a reduction in pressure and temperature induce d b y uplif t (e.g . Piggot t & Lines 1991 ; Dunca n et al. 1998) . Otherwis e the overwhelming tendency is for an increase in gas exsolution from oils , brines and coals. It may be predicted tha t significan t exhumatio n o f a petroliferous basi n wil l produc e a massiv e ga s bloom i n th e basi n centre , drivin g oi l t o peripheral locations , t o mor e shallo w depth s where it will be biodegraded, o r to the surface . Assessment o f th e ris k o f ga s flushin g durin g uplift an d pressure-temperature decrease can be carried ou t b y geochemica l modellin g o f th e original oil-gas balance i n a prospect. Inpu t of gas b y exsolutio n fro m formatio n brine ca n b e assessed fro m volumetri c calculation s on the aquifer draining into the prospect (Dore & Jensen 1996; Crame r e t a l 1999) . Knowledg e o f formation wate r salinit y wil l improv e suc h calculations, becaus e mor e salin e brine s ca n dissolve les s methan e (Maximo v e t al . 1984) . Contribution o f exsolutio n ga s fro m hydrodyn amic flo w ca n b e estimate d b y mappin g hydraulic gradients , a s show n by Crame r e t al . (1999). In all cases control points such as nearby wells will , o f course , increas e th e accurac y of such estimates . Probability of trap and seal Traps ca n b e eliminated , o r thei r volum e decreased, b y th e effect s o f regiona l tiltin g during exhumation . Similarly , ne w traps ca n b e created b y tiltin g o f three-wa y di p closure s ('noses'). Extrem e exhumatio n may , o f course , breach pre-existing accumulations at the surface. In area s where uplift i s associated wit h faulting , tectonic breachin g o f trap s ca n occu r throug h fault displacemen t o f th e sea l (caprock) , faul t juxtaposition o f hydrocarbo n reservoirs agains t aquifers or thief zones, or the formation of crestal extension fracture s ove r dome s an d anticlines . Where inversion (i.e. compressive reactivatio n of an extensiona l basin ) i s involved , trap s suc h a s horsts o r tilte d faul t block s ma y b e destroyed , whereas ne w trap s ca n b e create d b y (fo r example) reverse rejuvenatio n of half-grabe n o r bulge of the basin centre. MacGregor (1995) has shown fro m a globa l databas e tha t exploratio n

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success rate s in strongl y inverte d rift s ar e lowe r than thos e i n locall y inverte d rift s an d muc h lower than those in uninverted rifts. In all cases, a key consideratio n i s th e timin g o f generatio n compared wit h th e timin g o f uplif t an d restructuring. A s alread y demonstrated , durin g uplift n o ne w hydrocarbon s ca n b e generate d from sourc e rock s t o compensat e fo r redistri bution losses . Newl y create d structure s mus t therefore b e fille d b y remigratio n o r b y hydrocarbons (principall y gas ) from exsolution . A secon d an d equally important consideration is th e performanc e o f sealin g lithologie s during uplift. In general, as a claystone sea l is buried it becomes mor e compacte d an d stronger . Whe n exhumed i t shoul d retai n th e tensil e strengt h o f its maximu m buria l depth , an d thu s wil l b e stronger tha n a claystone a t the sam e dept h in a continuously subsiding basin. This observation is supported b y LOT s (leak-of f tests ) an d FIT s (formation integrit y tests ) o n seal s i n uplifte d Atlantic margin basins (Corcoran & Dore 2002) . Additionally, greate r compactio n tha n 'normal ' (and corresponding decreas e i n pore throa t size ) should increas e th e capillar y retentio n capacit y of a n exhume d claystone . Therefore , unde r certain conditions, the seal capacity of a prospect in an exhumed basin may be superior to that in a subsiding basin , a t th e equivalen t depth . However, severa l importan t factor s combin e t o diminish this capacity, a s follows. (1) Brittlenes s of the seal . A ductile rock ca n accommodate mor e strai n (u p t o 10% ) before fracturing tha n a brittle rock (<3%) . Changes in ductility wit h increasin g buria l dept h ar e complex an d depen d o n composition , tempera ture, confining pressure and fluid pressure (Davis & Reynold s 1996) . I n general , however , brittle ness i n claystone s ca n b e sai d t o increas e wit h density, wit h the transition from ductile to brittle behaviour takin g plac e ove r th e rang e 2.2-2.5 gcirT3 (Hoshin o e t al 1972) . Conse quently, exhume d clayston e seal s ma y b e mor e brittle tha n norma l seal s a t th e sam e depth . Brittle seal s wil l b e mor e likel y t o ruptur e an d leak i n respons e t o stres s (e.g . fro m tectoni c bending o r hydrofracturing ) tha n ductil e ones , which wil l defor m elasticall y an d plasticall y before fracturing . A crucia l question , therefore , is whether embrittlement of a potential claystone seal ha s take n place befor e uplift . Evaporite s o r mudstones containing evaporitic minerals, which deform plasticall y unde r a ver y wid e rang e o f pressure-temperature conditions, for m th e mos t efficient seal s i n uplifte d basin s (see , fo r example, wor k on the Eas t Iris h Sea Basi n by Seedhouse & Race y (1997 ) an d Cowa n e t a l (1999)).

(2) Hydrauli c fracturing . Hydrauli c leakag e may occu r whe n rapi d exhumation , unde r conditions o f lo w differentia l stres s an d dis equilibrium flui d pressures , result s in failur e o f brittle seals . Thi s failur e ma y b e manifeste d as extensional shea r fractures , dilatio n o f faul t planes, o r hydrofracturing . Shear fracture s wil l be formed in conditions of high differential stres s in th e caproc k an d wil l b e promote d b y disequilibrium pore pressures durin g rapid uplift . Pre-existing fracture s an d fault s ma y als o b e induced t o fai l i n thes e circumstances . Th e orientation of the new fractures tha t form, and of the pre-existin g fracture s tha t reactivate , wil l depend o n th e directio n o f th e principa l compressive stres s (crj) . Unde r condition s o f low differential stres s and similarly high retained disequilibrium fluid pressures, hydrofractures are likely t o for m b y tensil e failur e o f th e caproc k (Corcoran & Dore 2002 ; see also Sibso n 1995) . (3) Diffusion . Leakag e o f hydrocarbon s b y means of molecular transport through caprocks is thought t o b e usua l i n petroleu m basin s (e.g. Krooss e t al. 1992) . Wherea s diffusio n rate s for oil ar e probabl y negligibl e becaus e o f th e larg e size o f th e oi l molecules , gas wil l diffus e mor e readily throug h water-saturate d clayston e caprocks. Ther e i s considerabl e debat e i n th e literature on diffusion rates , but there seems little doubt that over a moderate geologica l tim e scal e (say, th e lengt h of th e Cenozoi c period , 6 5 Ma) diffusion losse s from a shale-sealed gas field can be considerabl e (e.g . Leythaeuser e t al . 1982; Krooss et al. 1992) . In a subsiding basin the fill of a ga s fiel d wil l b e determine d b y th e rati o o f diffusion losse s throug h th e sea l t o newl y generated ga s enterin g th e trap . Afte r exhuma tion, however , th e suppl y o f ne w hydrocarbon s will be arrested, allowin g the trap to be gradually depleted vi a diffusion . Th e diffusio n rat e o f methane throug h evaporites i s s o lo w a s t o b e negligible, agai n showin g tha t evaporite s ar e highly efficien t seal s tha t ca n preserv e hydro carbons ove r significan t geologica l tim e i n exhumed basi n setting s (see , fo r example , Kontorovitch e t al . (1990) , o n Proterozoi c ga s reservoirs in the Lena-Tunguska Basin , Russia). In summary , i n exhume d basin s th e ris k associated wit h tra p an d sea l i s significantl y increased. Underfille d trap s and near-hydrostatic reservoir pressures are commonly encountered in uplifted Atlanti c margi n basin s (Corcora n & Dore 2002) , presumabl y reflectin g pressur e depletion throug h th e sea l durin g exhumation , lack o f ne w hydrocarbon s fro m sourc e rock s once uplif t commenced , lac k o f ne w exsolution products once uplif t stopped , subsequen t escape of gas by diffusion an d contraction of gas during

PREDICTION OF HYDROCARBON SYSTE M IN EXHUMED BASINS

post-exhumation reburial . Uplift-relate d depres suring in low-permeability rock s ca n also resul t in transien t underpressuring , i.e . pressur e gra dients belo w hydrostati c (Lu o & Vasseur 1995) . Underpressuring i s als o a characteristi c o f th e basin-centred ga s fields in uplifted basin s i n the western USA . Som e o f thes e fields , suc h a s Elmworth i n th e Albert a Basin , ar e actuall y synclines (Master s 1984 ) i n whic h th e ga s accumulation probabl y represent s a disequili brium conditio n an d wher e th e underpressurin g may b e attributabl e t o therma l contractio n o f formation fluid s (Pric e 2002) . Apar t fro m a single wel l i n th e Barent s Se a (Dor e & Jense n 1996) underpressurin g has not yet been reporte d in the Atlantic margin basins. The probabilit y o f trappin g i n a n exhume d basin settin g i s bes t assesse d b y structura l modelling, whereb y the timing of trap formation and modification is compared wit h the timing of charging. Knowledg e o f maximum burial depth s (and henc e maximu m pressure s an d tempera tures) can indicate whethe r a shale sea l i s likel y to hav e becom e embrittle d befor e exhumation . Evidence o f fractur e trend s an d present-da y stress systems , combine d wit h modellin g o f pressure evolution, can help in assessing whether hydraulic failur e o f seal s i s likel y t o hav e occurred. Publishe d dat a o n methan e diffusio n rates, in combination wit h estimation o f the tim e elapsed sinc e uplift of a trap, can quantify likel y diffusion losse s throug h a seal . Evidenc e o f evaporites i n a n uplifte d basin , eve n a t a ver y preliminary stag e o f evaluation , significantl y

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enhances th e probabilit y tha t som e trappin g capability wil l hav e bee n retaine d durin g exhumation. Prospect resource estimation in exhumed basins Introduction Potential hydrocarbo n reserve s i n a prospect ar e estimated fro m a n equation of the typ e RR = G R V x N / G X 0 X S h c X F V F x R F (3 ) where R R ar e th e recoverabl e hydrocarbo n reserves, expresse d a s a volume ; GR V i s th e gross roc k volume , i.e . th e volum e o f th e reservoir withi n th e potentia l trap ; N/ G i s th e net-to-gross ratio , i.e. the fraction of the reservoi r that i s capabl e o f containin g movabl e hydro carbons; 4> is th e interconnecte d porosit y i n th e reservoir, expresse d a s a fraction ; S hc i s th e hydrocarbon saturation , i.e . th e fractio n o f por e space take n u p b y hydrocarbons ; FV F i s th e formation volume factor, a multiplier taking into account th e expansio n o f ga s o r th e contraction of oil (owing to liberation o f dissolved gas ) as the hydrocarbons are brought to surface pressure and temperature condition s durin g production; RF is the recover y factor , i.e . th e proportio n o f th e hydrocarbons i n th e prospec t tha t ca n b e recovered t o surfac e give n a n assume d pro duction method . Thes e factor s ar e estimated based o n th e predicte d dept h o f th e prospect , local an d regiona l data , an d experience .

Fig. 3 . Factors to be considered when carryin g out recoverable reserves calculation for a prospect in an exhumed basin. Arrows indicate improvement or deterioration in reservoir parameters.

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Uncertainty i n reserves estimation is traditionall y addressed b y assuming a range of values for each prospect paramete r (GRV , N/G , porosity , S hc, FVF, RF ) and by applyin g stochasti c simulation procedures t o generat e a rang e o f reserves . Predictive capabilit y is , o f course , mainl y a function o f local dat a density and quality. Where evaluation is mainly dependent o n regional data , it i s especiall y critica l t o kno w whethe r a basi n has been exhumed. The effects o f exhumation on the volumetri c parameter s ar e show n i n Fig . 3 and discusse d below . Gross rock volume Because the measuremen t of the tota l potentia l volume withi n a tra p i s mad e o n a prospectspecific basis , thi s valu e i s independen t o f whether th e tra p ha s bee n exhumed . However , the actual volume also depends o n the degree-of fill factor , i.e . th e proportio n o f th e availabl e vertical closur e take n u p b y th e hydrocarbo n column. For continuously subsiding basins, Sales (1993) an d other s hav e show n tha t a rang e o f trap-fill value s i s possible , base d o n th e relationship betwee n th e vertica l closur e (an d hence th e potential upwar d buoyancy pressure of any hydrocarbon fill ) an d the sealin g capacity of the caprock. Becaus e ther e is often a continuous supply o f newly generate d hydrocarbons i n such basins, th e sealin g capacit y become s th e mai n limiting factor . I f th e sea l i s efficient , th e structure has a good chanc e of being completel y hydrocarbon-filled: fo r example , Sale s (1993 ) asserted tha t mos t ga s field s i n th e Norwegia n North Se a ar e ful l t o spill . I n contrast , i n a n exhumed basin , trap-fil l i s limite d no t onl y b y sealing capacit y (which , a s shown i n the sectio n on trap and seal, ca n be catastrophically reduced during exhumation ) but also b y the hydrocarbo n budget generate d o r liberate d a t th e tim e o f th e last exhumation . Processe s subsequen t t o exhumation (e.g . diffusio n an d reburial contraction o f gas) wil l serv e t o diminis h th e trap-fill . Thi s observation i s strongly supported b y observatio n of underfilled gas fields in NW European uplifte d basins suc h a s th e Barent s Se a (Spence r e t al 1987), West Shetland Basi n (e.g. Goodchild e t al. 1999), East Irish Sea (e.g. Stuart & Cowan 1991) , Southern Ga s Basi n (e.g . Hillie r & William s 1991) an d Slyne-Erris Basi n (in-hous e dat a o n the Corrib Field) . Net-to-gross ratio, porosity and hydrocarbon saturation A sedimentar y roc k tha t ha s bee n uplifte d with resulting removal of overburden will preserve the

compactional an d diageneti c stat e associate d with it s maximu m buria l depth . I n mos t case s these highe r levels of compaction an d diagenesi s (stylolitization, quart z precipitatio n an d authi genic cla y minera l formation ) wil l involv e porosity loss . Consequently , ther e wil l b e a decrease i n hydrocarbon saturation as a result of the increased proportion of a given pore occupied by th e wettin g wate r phase . Net-to-gros s rati o will als o b e impaire d a s a resul t decreas e i n porosity o r permeability o f some rock belo w th e threshold considere d t o defin e a n effectiv e reservoir. Thus , overal l reservoi r qualit y wil l usually b e impaire d compare d wit h tha t a t a similar dept h i n a subsidin g basin. This general principle i s illustrated in the Barent s Sea, wher e Middle Jurassi c sandstone s i n th e Hammerfes t Basin ar e petrographicall y simila r t o thos e tha t form major reservoirs off Mid-Norway and in the North Sea . Th e Barents Sea reservoirs, however, consistently sho w higher levels of stylolitization (Walderhaug 1992 ) an d quart z precipitatio n (Berglund e t al. 1986) , wit h consequently lower porosities (Olausse n e t al . 1984) , becaus e o f a maximum burial depth some 150 0 m greater than at present. As indicate d b y Pric e (2002) , coolin g o f formation water s during uplift wil l also result in the precipitation of solutes (e.g. silica) and hence the occludin g o f porosity . However , a s Parnel l (2002) points out, this effect ma y be minor and is not widel y observed; furthermore , some mineral species (notabl y carbonates ) becom e increas ingly solubl e a t lowe r temperatures , thereb y introducing the possibility of secondary porosity development. As discusse d i n th e sectio n o n prospec t ris k analysis, exhumatio n create s a n increase d probability o f hydrodynami c flo w an d th e introduction o f meteori c wate r int o th e basi n aquifers. Suc h groundwate r wil l usuall y contain dissolved oxyge n and wil l b e acidi c (principally as a result of dissolved carbo n dioxide) , leadin g to th e possibilit y o f oxidatio n an d aci d dissolution i n reservoirs. The effec t o n reservoi r quality wil l b e comple x an d depen d o n th e chemistry o f th e reservoir , th e formatio n wate r and th e introduce d water . Oxidatio n o f ferro magnesian mineral s ma y for m pore-cloggin g iron oxides, whereas acid dissolution of feldspars and carbonate s ca n creat e substantia l secondar y porosity (se e th e muc h fulle r discussio n b y Parnell (2002)) . Thes e effect s ar e likel y t o b e most prevalen t clos e t o th e surfac e wher e meteoric wate r flo w i s strongest . Nevertheless , they ca n for m a n importan t modifie r t o th e overall negativ e implication s of exhumatio n on reservoir quality ; in basin s that have undergone

PREDICTION OF HYDROCARBON SYSTE M I N EXHUMED BASIN S

repeated exhumatio n an d reburia l episodes , improved reservoi r qualit y a s a resul t o f dissolution may be forecast below unconformity surfaces (e.g . Shanmuga m 1988) . Knowledge tha t a basi n ha s bee n exhume d allows th e interprete r t o plac e constraint s o n predicted reservoir quality . The most useful dat a for thi s proces s are , o f course, loca l wel l descriptions, whic h wil l giv e direct evidenc e a s to th e detrita l an d authigeni c mineralogy o f th e reservoir. However , eve n withou t suc h data , reconstruction o f th e maximu m buria l dept h (maximum temperatur e exposure ) o f th e reser voir is important. It allows a first-pass prediction as t o whethe r th e reservoi r ha s exceede d temperature thresholds for kinetically controlle d poroperm-reducing minera l transformations ; for example, quart z cementatio n an d th e develop ment of authigenic illite (e.g. Nadeau et al 1985 ; Bj0rkum e t al 1993) . Formation volume factor Uplift o f a hydrocarbon accumulation will result in lowe r temperature s an d pressures , wit h exsolution o f dissolve d ga s fro m oi l an d expansion o f reservoire d gas . Therefore , fo r a given hydrocarbo n por e volume th e expectatio n will be for more oi l (because of lower shrinkag e on production ) an d les s ga s (becaus e o f les s expansion o n production). I t ca n be argue d tha t these factor s ar e independen t o f whethe r th e reservoir ha s bee n uplifted , an d ar e simpl y a function o f pressure-temperature-volum e relationships a t a give n depth . Althoug h this i s undoubtedly true , th e occurrenc e o f near hydrostatic pressur e gradient s resultin g fro m pressure dissipation on uplift (a s observed on the NE Atlanti c margin ) wil l creat e a tendenc y towards lowe r shrinkag e oil s an d lowe r expan sion gase s compare d wit h a continuousl y subsiding basin. Recovery factor Recovery factor can be influenced positively and negatively b y uplif t an d exhumation . Generally, the lowe r porosit y an d permeabilit y o f a n uplifted reservoir for a given depth should impair recovery factor for both oil and gas. Additionally, the dissipation of overpressure durin g uplift wil l limit th e amoun t o f hydrocarbo n tha t ca n b e produced b y simpl e pressure depletion , wherea s a reservoir in a subsiding basin at the same depth may retain overpressure and hence have a higher initial reservoir pressure. Running counter to this argument, th e developmen t o f ope n pressur e systems an d hydrodynami c flo w a s a resul t o f exhumation ma y provid e pressur e maintenanc e

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(water drives ) durin g commercial depletio n o f a field. Fractures ar e a n extremel y importan t com ponent o f productiv e reservoir s worldwide , an d can b e attribute d to diastrophis m (fo r example , over fold axes , e.g. Aguilera 1980 ) or to removal of overburde n stres s b y exhumatio n (e.g . Aguilera 1980 : Sibso n 1995 : Corcora n & Dor e 2002). A s shown in the section o n probability of trap an d seal , shale s tha t hav e bee n embrittle d during buria l ma y fractur e durin g uplift . Reservoir lithologie s ar e generall y mor e brittl e than sealin g lithologie s suc h a s shale s an d evaporites. Tigh t sandstones , quartzite s an d dolomites ar e th e mos t fracture-pron e an d limestones ar e th e mos t ductil e o f th e potentia l reservoir rock s (Handi n e t al . 1963 : Stearn s & Freidman 1972: Dore & Jensen 1996). Therefore, reservoirs ma y fractur e withou t correspondin g rupture of the caprock, a situation that creates the basis fo r globall y importan t hydrocarbo n resources suc h as the Asmari fractured carbonat e fields of the Zagros fol d bel t (e.g . Daniel 1954) . Fracturing durin g uplif t ca n creat e reservoir s from non-reservoi r lithologie s (e.g . basement o r siliceous shales) , contribute to both porosity and permeability i n low-poroper m reservoirs , an d enhance por e connectivit y (an d henc e per meability) i n highe r poroper m reservoirs . Frac tures therefore have the potential to significantl y boost recover y i n uplifte d terrane s wher e poroperms woul d otherwis e b e unacceptabl y low. A critica l issu e t o wel l location , wel l completion and recover y in fracture d reservoir s is identificatio n o f th e fractur e set s tha t ar e dilatational (an d henc e contribut e th e mos t t o fluid flow) . Dilatio n wil l occu r mos t readil y i n fractures orthogona l t o th e leas t compressiv e stress directio n (cr 3), whic h wil l b e approxi mately horizonta l i n a n extensiona l regim e an d approximately vertica l i n a compressiona l regime. Fo r a simplifie d cas e o f su b vertical fractures, natura l fractures ar e more likel y to be open and support fluid flow if they strike close to the maximum horizontal stress (S hmax) direction, an observatio n supporte d b y globa l waterfloo d studies on producing fields (Heffer & Dowokpor 1990). Th e N E Atlanti c margi n a t presen t i s under a mil d NW-S E compressiv e regime , probably attributabl e t o ridge-pus h fro m th e Atlantic spreading centre (Dore & Lundin 1996). Where evidenc e exists , i t appear s tha t ope n fractures have a strike close to the NW-SE S hmSLX direction. Th e Clai r Fiel d i n th e uplifte d Wes t Shetland Basi n ha s a reservoi r consistin g o f fractured Uppe r Palaeozoi c sandstone . Detaile d studies sho w tha t specifi c fractur e set s aligne d close t o Shma x ar e dilatational , allowin g a

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recovery progra m t o b e devise d base d o n exploiting thes e fracture s with directional well s (Coney e t al 1993) . Borehole data , regiona l geology , seismi c reconstruction an d basi n modellin g ca n hel p t o assess whethe r fractur e enhancemen t o f a n uplifted reservoi r i s likel y t o hav e occurred . Key criteri a are : (1 ) prognose d reservoi r lithology; (2 ) maximum burial dept h an d degre e of uplif t o f th e reservoir ; (3 ) probabilit y o f overpressure developmen t an d dissipation ; (4 ) direction o f seismicall y mappabl e faults ; (5 ) direction o f fractures from borehole s o r regional outcrop data ; (6 ) present-da y stress-fiel d orientation.

Synopsis: key characteristics of hydrocarbon systems in exhumed basins On th e basi s o f th e foregoin g discussions , i t i s possible t o describ e a suit e o f phenomen a prevalent i n exhume d hydrocarbo n system s (Fig. 4) . Althoug h thes e characteristic s ca n als o occur i n subsidin g basins , th e occurrenc e o f many o r al l factor s togethe r wil l b e a stron g signature o f exhumation . Conversely , prio r knowledge tha t a basi n ha s bee n exhume d (from, fo r example , it s truncate d stratigraphi c record o r characteristi c structura l style ) allow s such factor s t o b e anticipated . The y ar e a s follows: (1) near-hydrostaticall y pressure d o r under pressured reservoir s a s a resul t o f catastrophi c pressure releas e durin g exhumation . Under pressure ma y deriv e fro m depressurin g and/o r thermal contraction in low-permeability aquifers. (2) Underfille d traps resulting from reductio n in sealin g capacity , spillag e losses , an d remigration inefficiency durin g exhumation, followed by cessation of the hydrocarbon supply , diffusion o f gas an d reburia l shrinkag e o f ga s afte r exhumation. (3) Large , basin-centre d ga s deposit s liber ated during uplift fro m oils , formation brines and coals, an d furthe r displacin g pre-existin g oi l b y gas expansion. Thes e accumulation s often over lie th e deepes t par t o f th e basi n becaus e o f th e thicker sedimentar y successio n availabl e t o generate exsolve d gas , inversio n o f th e basi n centre t o creat e ne w structura l traps , slo w dissipation o f the gas bloom i n low-permeability lithologies, and hydrodynamic focusing . (4) Two-phas e accumulation s a s a resul t o f gas exsolutio n fro m oi l an d retrograd e conden sation o f liquids fro m we t ga s during uplift . (5) Residua l oi l column s lef t behin d b y sea l failure o r i n technicall y breache d traps , b y

spillage, an d b y risin g o f th e gas-oi l an d oilwater contacts during post-exhumation diffusio n of th e ga s cap . (6) Small , remigrated periphera l oi l deposits : oil that is not driven completely from th e system by seal failure, tilting and gas flooding is likely to accumulate i n trap s o n th e basi n margin , fo r example, hanging-wall traps. (7) Heav y oil deposits formed by the remigration of oils t o shallo w levels o f the basin , where washing b y meteori c wate r inflo w an d bacteria l biodegradation ca n occur. The occurrenc e o f these characteristics on the NW Europea n margi n i s discusse d belo w wit h reference t o selecte d basin s (Fig s 5-10 ) an d summarized i n Fig. 11 .

Examples: exhumed provinces on the NW European margin In th e cas e historie s an d i n Fig s 5-1 0 th e following term s o f referenc e ar e used : (1 ) exhumation i s uplif t o f ke y referenc e horizon s above maximu m buria l depth ; (2 ) two-phas e accumulations ar e counte d a s bot h oi l an d gas ; (3) succes s rat e indicate s th e numbe r o f discovered pool s o f testabl e hydrocarbon s divided b y th e numbe r o f exploratio n wells ; i t does no t represen t th e rat e o f commercia l success. Western Barents Sea (Fig. 5) The Barents Sea consists of a complex mosaic of basins an d platforms , whic h i n th e wester n (Norwegian) sector become younge r towards the North Atlanti c Ocean. I n the east, the Nordkapp Basin i s a NE-SW-trending graben , initiate d in Late Palaeozoi c tim e an d dominate d toda y b y near-surface sal t dome s an d wall s formed fro m Upper Carboniferous-Lowe r Permia n halite . Farther west , th e Hammerfes t Basin , i s a n e n echelon continuation of this trend, but in contrast the last significant rif t episod e wa s late r (in Late Jurassic-Early Cretaceou s time) . The Hammer fest Basi n i s cross-cu t t o th e wes t b y a north south lin e o f dee p Cretaceou s depocentre s (Bj0rnoya an d Troms 0 Basins) , whic h ar e i n turn supercede d t o th e wes t b y Tertiar y depocentres (e.g . S0rvestnage t Basin ) clos e t o the continent-ocean boundary (Gabrielsen et al. 1990). Major exhumation s too k plac e durin g Cenozoic time , roughl y synchronous with uplif t of th e Fennoscandia n mainland . Thes e include d an episod e o f Paleogen e uplif t probabl y associ ated wit h incipient opening of th e N E Atlantic,

PREDICTION O F HYDROCARBON SYSTEM I N EXHUMED BASINS

and a particularl y sever e Plio-Pleistocen e epi sode emphasized by repeated glacial erosion and isostatic re-equilibration. The Nordkapp Basin i s deeply exhumed, wit h a thin laye r of Quaternar y sediments overlying truncate d Cretaceou s rocks. Some Tertiar y sediment s ar e preserve d i n th e

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Hammerfest Basin , bu t ther e i s a majo r unconformity betwee n th e Paleocen e an d th e Pliocene sequences (e.g . Westr e 1983) , and well data indicat e remova l o f abou t 1500 m o f overburden (e.g . Nylan d e t al 1992 ; Walderhau g 1992).

Fig. 4. Highly schematic and vertically exaggerated basin cross-section illustrating some effects o f exhumation on the hydrocarbon system, (a) A simple rift geometry containing an oil-dominated hydrocarbo n system , used as the starting point, (b) Effects takin g place during exhumation. Regional exhumation, in this example, is accompanied by inversio n of the basin centre, (c ) Processes afte r exhumatio n has ceased an d minor reburial has taken place .

Fig. 5. Structural features, hydrocarbon discoveries and data template for the western Barents Sea. In the inset table: (1) exhumation is uplift o f key reservoir horizons above maximum burial depth; (2) two-phase accumulations are counted as both oil and gas finds; (3) success rate does not represent commercial success , it indicates the number of discovered pool s of testabl e hydrocarbons divided by th e numbe r of exploratio n wells.

Fig. 6. Structural features, hydrocarbon discoveries an d data template fo r the Wes t Shetland Basin . (Fo r qualifiers to inset table , se e Fig. 3.)

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Exploration drillin g o f the Norwegian Barent s Sea commence d i n 198 0 an d a t tim e o f writin g about 6 0 well s hav e bee n drilled . Althoug h success rates have been fairl y hig h (about one in three) th e result s hav e bee n commerciall y disappointing, largel y becaus e o f the dominanc e of ga s an d th e remotenes s o f th e are a fro m potential ga s markets . Discovere d resource s ar e currently about 10. 5 TCP (trillion cubic feet) gas with abou t 30 0 MMB (millio n barrels) oil . Mos t of th e ga s discovere d t o dat e i s concentrate d i n three fields (Sn0hvit , Askeladden an d Albatross) in th e axia l par t o f th e Hammerfes t Basin , comprising faul t blocks and horsts with a Middle Jurassic reservoi r (St 0 Formation ) seale d b y Upper Jurassi c shales . Sourcin g i s fro m th e Upper Jurassi c Hekkinge n Formation , whic h attained maturit y befor e uplif t i n norther n an d western part s o f the Hammerfes t Basin . The hydrocarbo n syste m show s man y classi c attributes o f exhumation . Th e Middl e Jurassi c reservoirs hav e anomalousl y hig h level s o f diagenesis (e.g . Walderhau g 1992) . Th e centra l gas accumulations are underlain by thin oil discs or residua l oi l legs , a resul t o f spillin g o f pre existing oil deposits . Modellin g of the larges t field, Sn0hvit, suggests that the oil was evacuated by gas expansion an d that current underfilling of the tra p i s consisten t wit h diffusio n losse s an d gas contractio n durin g reburia l (Nylan d e t al. 1992). Methane exsolution from oi l and brine has probably als o contribute d t o th e dominanc e o f gas (Dor e & Jense n 1996) . Discret e oi l accumulations ar e small , an d includ e th e Myrsilde discovery , i n Lowe r Cretaceou s sand s in a hanging-wal l tra p agains t th e norther n margin o f th e Hammerfes t Basin , an d th e recent Goliat h discover y o f 70-10 0 MMB remigrated oi l i n a Middl e Jurassi c reservoir . Current exploratio n effort s focu s on : (1 ) area s that hav e no t bee n recentl y exhume d (e.g . th e western margi n o f th e Barent s Sea , whic h received erosio n product s fro m th e Cenozoi c denudations); (2 ) area s wit h potentia l evaporit e seals suc h a s th e Nordkap p Basi n (wher e wel l 7228/7-1 i n Licenc e P20 2 recentl y foun d oi l and ga s i n Triassic rocks, partly seale d by a salt overhang); (3 ) prospect s favourabl y locate d t o receive oi l spille d fro m th e mai n Hammerfes t Basin traps . West Shetland Basin (Fig. 6) The Wes t Shetlan d Basi n (WSB ) an d adjacen t Faeroe-Shetland Basi n (FSB ) hav e a NE-S W grain, cross-cu t b y NW-trendin g transfe r zone s (e.g. Dea n e t al 1999) . Th e structura l histor y of th e are a i s complex , wit h multipl e riftin g

events i n Permo-Triassic , lat e Jurassic-earl y Cretaceous, mid-lat e Cretaceou s an d Paleo cene time s (Dor e e t a l 1999) . I n th e WS B th e dominant riftin g even t wa s i n Permo-Triassi c time, afte r whic h th e basi n underwen t repeate d exhumations whil e th e FS B continue d t o subside. Majo r Cretaceou s unconformitie s and / or non-sequence s i n th e WS B presumabl y represent basin-flan k uplif t associate d wit h rifting i n th e FSB . A t th e en d o f Cretaceou s time th e WS B an d it s souther n continuation , the Sola n Basin , wer e uplifte d a s par t o f a major emergenc e o f th e Scottis h massif . Further uplif t o f th e WS B i n Oligo-Miocen e time wa s probabl y connecte d t o a n episod e o f inversion tha t cause d broa d doma l structurin g in th e FS B (e.g . Turne r & Scrutto n 1993 : Herries e t al . 1999 ; Parnel l e t a l 1999) . A s a result o f repeate d uplif t an d erosion , Cenozoi c deposits ar e thi n o r absen t ove r th e WS B an d Solan Basin . About 6 0 wells have been drille d i n the WSB and Sola n Basi n wit h a succes s rat e o f on e i n seven. Prove n recoverabl e resource s ar e i n th e order o f 28 0 MMB oi l an d 0.3 5 TCP gas . Onl y one accumulation , the Clai r oilfield , i s currently considered commercia l (Cone y e t a l 1993) . Hydrocarbons i n th e WS B wer e source d b y Upper Jurassi c marin e shale s wit h a probabl e Middle Jurassi c lacustrin e componen t (Baile y et a l 1987 ; Scotchma n e t al 1998) . Reservoir s range i n ag e fro m Lewisia n basemen t t o Earl y Cretaceous. Chargin g occurre d durin g lates t Cretaceous and Cenozoic time as pulsed episodes of migration from th e adjacent FSB (Parnell et al 1999). Earl y oi l charge s hav e frequentl y bee n lost as a result of biodegradation o r breaching of traps durin g uplift , a s show n by fluid inclusions and residua l shows (e.g . Goodchild e t al 1999) . Replenishment fro m a continuousl y subsidin g kitchen i n th e oi l windo w probably explain s the dominance of oil over gas in the basin, despite its exhumed nature. In th e Clair Fiel d 3- 5 billion barrels o f heavy oil ar e hel d i n fracture d basement an d Devono Carboniferous re d beds , althoug h onl y abou t 200 MMB are thought to be recoverable becaus e of th e lo w porosit y an d permeabilit y o f th e reservoir. Ope n fracture s ar e importan t fo r optimizing deliverabilit y i n th e reservoi r (Coney et al 1993 : se e also sectio n o n recover y factor). Th e oi l i s a mixture of biodegraded an d fresh oil , a functio n o f th e multipl e charging . Two smal l non-associate d ga s cap s o n Clai r represent a late gas charge, or possibly exsolution gas. I n the nearb y Victor y ga s field the earl y oil charge t o th e Lowe r Cretaceou s reservoi r ha s been los t a s a resul t o f breachin g o f th e tra p

PREDICTION O F HYDROCARBON SYSTE M IN EXHUMED BASINS

during early to mid-Cenozoic uplift . This has lef t residual biodegrade d oi l withi n an d belo w th e current gas column, which occupies less than half the vertica l closur e o f th e structur e (e.g . Good child e t al 1999) . Farthe r sout h i n th e Sola n Basin, smal l oi l accumulation s (Sola n an d Strathmore) totallin g 60MM B occu r i n Uppe r Jurassic an d Triassic truncatio n traps . Th e non biodegraded oi l wa s source d fro m a limite d kichen area , th e Eas t Sola n Basin , wher e th e Upper Jurassi c sourc e rock s ar e a t earl y oi l maturity (Herrie s e t al . 1999) . Herrie s e t al . considered tha t th e field s wer e charge d i n tw o Cenozoic pulses , separate d b y th e Oligo Miocene inversio n episode . W e note , however , that becaus e of the curren t ver y thi n post Paleocene cover , suc h generatio n mus t impl y episodes o f reburia l an d re-exhumatio n durin g Cenozoic time . Inner Moray Firth Basin (Fig. 7) The Inner Moray Firth Basin (IMFB) lies off the NE coast of Scotland, between the Grampian and Northern Highlands . The basin forms a westerly extension o f the trilete Mesozoic graben syste m and is separated fro m th e eastern part, the Outer Moray Firt h Basin , b y th e Halibu t Horst . Th e structural histor y o f th e IMF B i s dominate d b y the effects o f Permo-Triassic and Jurassic rifting, subsidence i n Late Jurassic to Late Cretaceousearliest Tertiar y tim e (Andrew s et al. 1990 ) an d subsequent uplift . Exhumation of the IMFB has been assigned to Paleocene tim e b y Hillis e t al. (1994) , based o n the assumption of synchronicity with the onset of denudation o f th e Scottis h Highlands . Younger uplift episodes are not, however, precluded by the data. For example, Underhill (1991) identified an Oligo-Miocene inversio n phas e o f possibl e far field 'Alpine' origin, which reactivated Mesozoic faults an d gav e ris e t o differentia l relie f withi n the basin. Uplift followe d by subsidence towards the Nort h Se a grabe n syste m ha s imparte d a n eastwards tilt to the IMFB, such that Jurassic and Lower Cretaceous rocks subcrop at the sea bed in the westernmos t par t o f th e basi n an d ar e succeeded eastwar d b y Uppe r Cretaceou s an d Tertiary subcrop s (Fig . 7) . Hilli s e t al . (1994 ) used soni c velocit y dat a t o sugges t tha t abou t 1 km of erosion took place over most of the basin during early Cenozoic time. However, because of later Cenozoi c sedimentatio n th e apparen t erosion (ne t uplif t sensu Rii s & Jense n 1992 ) decreases eastward s t o zero a t about 1°W . Approximately 80 exploration wells have been drilled i n the IMFB wit h a success rate of one in eight, much lower than in the adjacent North Sea.

417

Underhill (1991 ) attribute d poor succes s rates in this are a t o breachin g o f trap s b y reactivate d faults, som e of which extend to the surface. Total proven resources are of the order of 680 MMB oil and 0.5 1 TCP gas. I n the western, mos t uplifte d part o f th e basi n a singl e commercia l oi l discovery ha s bee n made , th e Beatric e Fiel d (155 MMB recoverable ) alon g wit h som e much smaller uncommercia l oi l an d ga s pools . Th e Beatrice hydrocarbo n syste m consist s o f a Middle Jurassi c reservoi r i n a tilte d fault-bloc k trap sealed b y Oxfordian-Kimmeridgian shales . The oil was co-sourced by Devonian and Middle Jurassic mudrock s (Peter s et al. 1989) . Chargin g occurred durin g Lat e Cretaceou s time , afte r which generation must have ceased a s a result of uplift. Remarkably , the oil accumulation appears to hav e remaine d intac t fo r th e duratio n o f Cenozoic time , preservin g a 335 m oi l colum n that fill s th e structur e t o spil l (Steven s 1991) . Retention may be partly due to the efficienc y of the shale seal (which may have retained ductility before uplift : se e Corcora n & Dor e 2002) , an d partly du e t o th e waxy , viscou s natur e o f th e crude. Th e ver y lo w energ y o f th e oi l (gas-oi l ratio of 126 SCF per barrel, bubble point pressur e 635psig) ma y testif y t o ga s depletio n b y diffusion o r lack o f an original gas charge . A cluster of fields of mixed phase in the east of the IM F (Captain , Blake , Ross , Cromarty , Phoenix) li e withi n th e uplifte d area , althoug h some o f thes e field s ma y b e receivin g charg e from currentl y generatin g kitchens . I n th e Captain Field , overlyin g the wester n en d o f th e Halibut Horst , shallowl y burie d Lowe r Cretac eous sandstones contain recoverable oi l reserves of 350 MMB sourced from Uppe r Jurassic rocks. Most o f th e Cenozoi c successio n i s missin g above th e field , an d th e hydrocarbo n accumu lation carrie s a stron g signatur e o f exhumation. The oi l i s heavil y biodegrade d an d include s a residual oi l colum n i n th e eas t o f th e fiel d attributed b y Pinnoc k & Clithero e (1997 ) t o easterly tiltin g durin g Cenozoi c time . Th e field has a small cap of thermogenic gas introduced as a lat e charg e (Pinnoc k & Clithero e 1997 ) an d probably representin g exsolutio n gas . Notably , however, the field is full t o its spil l point. East Irish Sea Basin (Fig. 8) The Eas t Iris h Se a Basi n (EISB ) i s a preserve d remnant o f a late Palaeozoi c t o earl y Mesozoi c extensional basi n syste m (Knip e e t al . 1993) . Subsequent uplif t an d denudatio n ha s remove d most o f th e post-Triassi c cove r fro m th e basin . Post-Triassic burial-uplif t histor y i s therefor e difficult t o reconstruct , and relies o n technique s

Fig. 7 . Structura l features, hydrocarbon discoverie s an d dat a template fo r th e Inne r Mora y Firt h Basin . (Fo r qualifiers t o inse t table , se e Fig . 5.)

PREDICTION O F HYDROCARBON SYSTE M IN EXHUMED BASINS

Fig. 8. Structural features, hydrocarbon discoveries and data template for the East Irish Sea Basin. (For qualifiers to inset table, see Fig. 5.)

such a s apatit e fissio n track , shal e velocit y an d other s hav e assigne d earlies t uplif t t o lates t vitrinite reflectance . Som e worker s hav e Cretaceous-Paleocen e time , concident wit h modelled a majo r uplif t phas e i n Earl y Cretac - Nort h Atlantic opening and facilitated by thermal eous time (e.g . Dunca n e t al. 1998) , wherea s uplif t o r underplating (Cop e 1994 : Cowa n e t al.

420

A. G . DORE ETAL.

Fig. 9 . Structura l features , hydrocarbo n discoveries and dat a templat e fo r th e Slyn e an d Erri s Basins . (For qualifiers t o inset table , se e Fig. 5.)

Fig. 10 . Structural features , hydrocarbo n discoverie s an d dat a template for the Nort h Celti c Se a Basin. (Fo r qualifiers t o inse t table, se e Fig. 5.)

422

A. G. DORE ETAL.

1999: War e & Turner 2002). Later exhumation is even mor e difficul t t o constrain , bu t War e & Turner (2002 ) hav e proposed a short-wavelength contribution fro m Eocene—Miocen e inversion . The EIS B wa s ice-covere d durin g Pleistocen e time, an d a glacio-isostati c componen t canno t therefore b e rule d out . Estimate s o f maximu m exhumation vary between 1 and 3 km, with recent work suggestin g 2 km a s an upper limi t (Cowa n et al 1999 ; War e & Turner 2002) . The EIS B i s a prolifi c hydrocarbo n provinc e containing 1 0 gas fields , tw o oilfield s an d nin e undeveloped hydrocarbo n discoveries . Abou t 6 0 wells hav e bee n drilled , wit h a on e i n thre e success record . Ga s reserve s o f 8. 2 TCP and oil reserves o f abou t 230MMB O hav e bee n ident ified i n an area o f 3500km 2 (Quirk e t al 1999) . The hydrocarbon syste m in the EISB consists of a Triassic aeolian-fl u vial reservoi r i n structura l traps, charged fro m a Namurian sourc e roc k and sealed b y Upper Triassi c evaporite s and shales . The hydrocarbo n accumulation s hav e a com plex evolutio n intimately relate d t o the exhumation history . Th e basi n i s characterize d b y a distinct northern ga s province an d a southern oil and ga s province , wit h approximatel y 70 % o f proven reserve s reservoire d i n th e tw o More cambe ga s fields . Earlies t ga s an d oi l emplace ment i s believe d t o hav e occurre d durin g Earl y Jurassic time. Breaching o f seals during exhumation, fo r exampl e i n Sout h Morecambe , resulte d in los s o f th e initia l oil-ric h charge followe d b y later stag e (?Earl y Tertiary ) rechargin g wit h thermogenic ga s an d present-da y underfillin g (Stuart & Cowan 1991 ; Stuar t 1993) . Breachin g or spillag e o f traps lef t behin d palaeo-oil-water contacts, indicated o n seismic data (Francis et al 1997) an d b y illit e cementatio n i n Morecamb e South. Residual columns of biodegraded oi l also testify t o th e origina l charg e (Bushel l 1986 ; Woodward & Curti s 1987) . Similarly , th e Douglas an d Lenno x oilfield s sho w evidenc e o f multiple charging , wit h th e earlies t oi l charg e being totall y degrade d t o bitumen, a subsequent higher maturit y charg e bein g partiall y biode graded, followe d b y a fina l condensat e charg e (Haig e t a l 1997 ; Yali z 1997) . Th e Formb y oilfield, a pool o f biodegraded oi l trapped i n th e Sherwood Sandston e b y Pleistocen e till , i s evidence o f ver y recen t remigratio n acros s th e Formby Poin t Fault (Fig. 8 ) from a breached tra p (Francis e t al 1997) .

Recent modelling of the EISB has stressed th e importance o f remigration , drive n b y ga s exsolution fro m oi l an d ga s expansio n durin g Cenozoic uplif t (Cowa n e t a l 1999) . Ga s exsolution fro m formatio n wate r ha s no t bee n incorporated int o thes e models , bu t w e sugges t that it provides a powerful additional mechanism for th e lat e ga s charge . Oi l wa s drive n updi p to the peripher y o f th e basi n b y th e lat e ga s flu x (Duncan e t a l 1998) . Th e complexit y o f th e remigration process is , however, indicated by the juxtaposition o f undersaturate d oil s (e.g . Dou glas) with dry gas accumulations (e.g. Hamilton). Seedhouse & Race y (1997 ) an d Cowa n e t a l (1999) hav e show n tha t th e presenc e o f halit e beds i n th e basa l par t o f th e sea l i s a critica l success facto r fo r hydrocarbo n entrapmen t an d oil-gas balance. Gas will escape through the seal in shallo w structure s excep t wher e th e basa l evaporite is present. Conversely, this discharge of the gas leg allows oil to be preferentially trapped in th e souther n part of the basi n where the basal evaporite i s absent (Quirk et al 1999) .

Slyne-Erris Basin (Fig. 9) The Slyne-Erri s Basi n (SEB ) i s a narrow , elongate, NE-SW-trendin g basi n syste m 60k m off northwester n Ireland. It consists of a series o f asymmetric half-graben s separate d b y cross cutting transfe r zones . I t experience d a multi phase riftin g an d inversio n history, althoug h the preserved basi n morpholog y i s primaril y th e result o f Mid-Lat e Jurassi c riftin g (Chapma n et a l 1999 ; Dance r e t a l 1999) . A strikin g characteristic o f th e southerl y Slyn e Troug h i s the truncated stratigraphic record wit h an almost complete absenc e o f post-rif t sediments . A thi n cover of Miocene sediments rests unconformably on synrif t sediment s o f Lat e Bajocia n t o Bath onian age . However , i n th e northerl y Em s Trough, mor e tha n 1 km o f Cretaceou s strat a are locally preserved. Multiple phases of regional exhumation and local inversio n affected th e area. These include d rift-related footwall uplif t event s in Lat e Jurassic-Earl y Cretaceou s an d Aptia n time, regional uplif t i n Paleocen e tim e probably associated wit h Atlantic opening, and inversionrelated uplif t i n Oligo-Miocene time. Maximu m exhumation is of the order o f 2000 m, although it is difficul t t o establis h wha t proportio n too k

Fig. 11 . Summary ma p showin g oil-ga s balanc e an d exhumation-relate d phenomen a i n exhume d basin s wit h proven hydrocarbo n system s o n the NW European margin . Example s o f fields or wells are given for exhumationrelated characteristics i n each area. Oil and gas quantities are related usin g oil industry standards, which are based on calorific value : 1 barrel o f oil approximatel y equal s 600 0 standard cubi c fee t of gas .

PREDICTION OF HYDROCARBON SYSTE M IN EXHUMED BASIN S

place i n Cenozoi c tim e (Scotchma n & Thoma s 1995). Sporadic exploratio n in this area over the past 25 years ha s resulte d i n th e drillin g o f si x exploration wells , whic h hav e yielde d a singl e gas discover y (th e Corri b Field ) i n th e Slyn e Trough. Tota l discovere d resource s t o dat e ar e approximately 1 TCF gas, all in the Corrib Field, an underfille d faulte d anticlina l structur e (Cor coran & Dore 2002) . Th e mai n ga s exploratio n play consist s o f a Lowe r Triassi c sandston e reservoir i n structura l traps , charge d fro m th e underlying Namurian-Westphalia n cla y stones and coal s an d seale d b y Uppe r Triassi c evaporites an d shale s (Scotchma n & Thoma s 1995; Dance r et al. 1999) . A Jurassic petroleu m system is considered proven in the Slyne Trough by Spencer et al. (1999), based on the presence of palaeo-oil accumulations . Biodegraded , residua l oil show s fro m a Lowe r Jurassi c sourc e hav e been encountered in Middle Jurassic reservoirs in wells 27/13-1 , 27/5- 1 an d 18/20-1 . Thes e residual column s ar e consisten t wit h breachin g of trap s and/o r freshwate r flushing durin g uplif t to shallo w levels . I t ha s ye t t o be demonstrate d that an y producibl e accumulation s fro m th e Jurassic hydrocarbo n syste m hav e survive d th e Cenozoic exhumatio n o f the basin. North Celtic Sea Basin (Fig. 10) The North Celtic Sea Basin (NCSB) is a NE-SW trending Mesozoi c extensiona l basi n locate d t o the sout h of Ireland. I t is bounded by a series of Palaeozoic ridge s an d platform s an d contain s a thick Triassi c t o Cretaceou s sedimentar y fill . Major riftin g episode s occurre d durin g Lat e Jurassic an d Earl y Cretaceou s tim e (Rowel l 1995), but post-rift subsidenc e was terminated by regional uplif t an d inversio n durin g Cenozoi c time. Th e exhumatio n resulte d i n subcro p o f Cretaceous Chalk at the sea floor in the centre of the basin, and the complete remova l o f Cretaceou s sediments i n th e NE of the basin. Two Cenozoic erosional events have been documented: regional uplift durin g Paleocen e tim e an d inversio n characterized b y basin doming and fault reversa l during Oligo-Miocen e time . Ne t exhumatio n i n excess o f 1100 m i s interpreted i n the N E of th e basin (Murdoc h et al. 1995) . The NCS B ha s prove d t o b e a somewha t enigmatic petroleu m province . Abou t 7 0 exploration and appraisal wells have been drille d to date , wit h a succes s rat e o f on e i n six . Onl y two accumulation s are producing, Kinsal e Hea d and Ball y cotton, containin g prove n reserve s o f 1.6 TCF (Taber et al. 1995) . A further seve n subcommercial oi l an d ga s discoverie s (e.g . Seve n

423

Heads, Helvick , Ardmore ) have been identified . The main reservoir i n the producing gas fields is the shallo w marin e Albia n Greensand . Second ary productio n occur s fro m th e fluvia l Wealde n reservoirs. Elsewher e i n th e basi n oi l ha s bee n tested fro m thes e stratigraphi c level s an d fro m Oxfordian fluvial sandstones and Middle Jurassic shelf limestone s (Casto n 1995) . Seal s ar e provided b y th e Albian-Cenomania n Gaul t Clay an d intraformationa l claystone s withi n th e Bathonian to Aptian succession . Typically fo r a n exhumed basin , th e NCSB is dominated b y centra l ga s deposit s an d i n thi s case b y a singl e accumulation , Kinsal e Head . This structure, a basin-centre anticline , may have had som e pre-Cenozoi c expressio n bu t wa s greatly emphasize d durin g Tertiar y inversio n (Taber e t al . 1995) . Atypicall y fo r field s i n N E Atlantic exhume d basins, th e tra p is ful l t o spil l (Taber e t al . 1995) . Thi s suggest s tha t th e maximum dept h o f buria l o f th e Gaul t Cla y caprock (1700-1800m ) ma y no t hav e bee n enough fo r th e clayston e t o achiev e embrittle ment, allowin g i t t o defor m plasticall y durin g exhumation an d compressiv e overprin t (Cor coran & Dor e 2002) . Publishe d model s fo r maturation an d expulsio n fro m th e Lowe r Jurassic sourc e rock s indicat e tha t pea k ga s generation woul d have occurred towards the end of Cretaceou s tim e (Murph y e t al . 1995) . Because the structure probably developed durin g Tertiary time , afte r generatio n fro m th e sourc e rocks would hav e bee n curtaile d becaus e o f exhumation, i t seem s unlikel y tha t th e presen t gas represent s th e origina l thermogeni c charge . Even give n a ductil e clayston e caprock , ga s losses a s a result o f diffusion an d underfillin g o f the tra p woul d b e expected . Therefore , activ e charging o f th e Kinsal e Hea d Fiel d ma y hav e continued unti l recen t geologica l time . A possible mechanis m fo r additiona l ga s charging is exsolution of methane from groundwater in the Greensand, Wealden and older aquifers, possibly focused b y groundwate r flow during uplift . Other signature s o f a n exhume d hydrocarbon system i n th e NCS B includ e a residua l oi l column i n Kinsal e Head , probabl y indicativ e of an earlier oil charge later displaced b y gas (Taber et al . 1995) , mino r periphera l hanging-wal l oil accumulations suc h a s Helvic k (Casto n 1995 ) and two-phas e accumulation s wit h biodegrade d oil (Seve n Heads) . Discussion As summarize d i n Fig . 11, man y o f th e Nort h Atlantic basin s contai n hydrocarbo n system s

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A. G. DORE ETAL

PREDICTION OF HYDROCARBON SYSTE M I N EXHUMED BASIN S

showing indicator s o f exhumation . Acknowl edgement o f th e importanc e o f exhumatio n ca n constrain futur e exploratio n strateg y i n thes e basins, and a similar approac h ca n be applied t o any exhume d basin. Exploration ris k analysi s i n exhumed terrane s should take int o account a decreased probabilit y for sea l o r trap , an d shoul d addres s a comple x interplay o f positiv e an d negativ e factor s whe n assessing probabilit y o f sourc e o r charge. Ther e is a n increase d chanc e tha t th e dominan t hydrocarbon phas e wil l b e gas , bu t oi l ca n stil l be predicted b y taking into account factors such as sea l integrity , displacemen t fro m trap s an d remigration pathways . Although we have given a qualitative guide to such a risk analysis, it is not possible t o provide numerical values. These will vary accordin g t o the unique geological charac teristics o f a n area ; fo r example , th e qualit y of regional seal s withi n th e basin . I t i s frequentl y said that exploration ris k analysi s i s a subjective procedure, of use as a comparative rather than an absolute measure . However , a mor e objectiv e view of risk (in any basin, exhumed or otherwise ) can b e gaine d b y carryin g ou t a n audi t afte r a period o f drilling, whereby the actua l discovery rates are compared wit h the predicted ones . Thus, a ris k analysi s constraine d b y knowledg e o f exhumation ca n be checke d an d modifie d base d on exploration history . The strateg y fo r targetin g resource s shoul d also b e contraine d by knowledg e of exhumation levels; fo r example, i n the identification of area s of porosit y preservatio n and potentia l fracture prone lithologies . Th e recover y strateg y fo r oi l and gas in uplifted fields should be influenced by knowledge o f present-da y stres s an d fractur e directions. As in the cas e of ris k analysis , we provide n o numerica l value s fo r resourc e assessment i n thi s paper , bu t agai n poin t ou t that a n audi t o f drillin g result s ca n provid e a n objective compariso n o f predicte d an d actua l volumes a t a n intermediat e stag e o f exploration of a basin. Timing o f exhumatio n i s a ke y elemen t i n prediction o f the hydrocarbon system. As shown in the case studies on the North Atlantic margin, most suc h area s hav e undergon e multipl e exhumations durin g Cenozoic tim e an d in som e cases exhumatio n bega n eve n earlier . I t i s difficult, an d require s patien t analysis , t o disentangle th e effect s o f th e various events and to assign relative importance to them. In general, however, i t ma y b e predicte d tha t th e mor e extreme effect s o n th e hydrocarbo n syste m (for example, th e flushin g effec t o f a centra l ga s bloom) ar e mor e likel y t o b e observe d wher e exhumation ha s bee n ver y recent ; a s i s seen ,

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for example , i n th e Barent s Sea , wher e th e Plio-Pleistocene regiona l uplif t episod e wa s particularly importan t (Nylan d e t al. 1992) . The effects o f olde r exhumation s ma y b e mute d b y the slo w dissipatio n o f ga s b y diffusion , an d overprinted b y reburia l an d th e introductio n o f new hydrocarbons . Th e variabilit y i n degre e o f exhumation withi n a particula r provinc e i s als o important, and relates to the uplift mechanism. I n an are a o f broa d regiona l ('epeirogenic' ) uplif t the effect s o n the hydrocarbon system should be similar ove r a wid e area , wherea s i n basin s inverted b y compressio n thes e effect s ma y b e more local in nature as a result of selective uplif t of intrabasina l structures . Th e fac t tha t bot h forms of exhumation are superimposed in several of th e Nort h Atlanti c basin s (e.g . Inne r Mora y Firth, Easter n Iris h Sea ) provide s a n additiona l challenge. The author s thank J. Parnell and M. Tate fo r thoroug h and constructive review s of the manuscript, M . Stoke r for editoria l handlin g an d J . Kipp s fo r graphics . Thi s paper wa s published by permission of Statoil.

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Geological and geochemical consequences of basin exhumation, and commercial implications LEIGH C. PRICE

In th e past , petroleu m basin s wer e traditionall y largely viewe d a s eithe r stati c o r uniforml y evolving entities . I t i s no w mor e widel y recognized tha t petroleu m basin s ar e subjec t to various intens e geologica l processes , an d henc e substantial changes , durin g thei r evolutionar y histories. On e suc h geologica l proces s i s exhumation, whic h i s usuall y accompanie d b y significant erosion , decreas e i n bot h buria l temperatures an d flui d pressures , an d ofte n decrease in geothermal gradients. These changes can hav e profoun d consequence s regardin g hydrocarbon (HC ) deposits . Beside s al l th e obvious consequences (halting source-rock generation, damaging or destroying seals, expansion of gas with trap flushing, etc.), les s obvious , but none the less equall y meaningful , consequence s also result . For example , i n basin s wit h hig h geother mal gradient s (e.g . Centra l Sumatra , Lo s Angeles), migration-accumulatio n processe s result i n mos t o f th e basin' s reservoire d oi l being emplace d withi n th e firs t 1-2 km o f th e surface. Stron g erosio n ca n largel y destro y almost th e entir e oi l resourc e o f th e basin , leaving a gas-onl y provinc e (e.g . Sa n Jua n Basin, USA) . Por e water s a t hig h tempera tures an d pressure s carr y larg e amount s o f both dissolve d H C gase s an d inorgani c mineral specie s (ions) . Significan t fall s i n burial temperature s fro m exhumatio n thu s cause tw o result s i n th e deepe r region s o f petroleum basins : (1 ) basin-centre d ga s deposits; (2 ) widesprea d destructio n o f deep basin porosity , bu t especiall y permeability , from wholesal e precipitatio n o f dissolve d mineral specie s a s diageneti c minerals . I n the cas e o f basin-centre d ga s deposits , i n going updi p fro m th e basi n depocentre , eventually a locatio n i s reache d wher e insufficient H C ga s wa s dissolve d i n th e pore wate r t o allo w ga s t o exsolv e an d for m a free-gas phas e tha t ca n excee d it s critical flui d saturation level . Thus , th e free-ga s bubble s remain immobil e an d caus e a two-phas e permeability bloc k (th e Jami n effect) , whic h becomes th e leadin g updi p sea l fo r a downdi p

basin-centred ga s deposit . Typ e example s o f the result s o f thi s proces s ar e presen t i n th e Denver an d Sa n Jua n Basin s (USA ) an d th e Alberta Basi n (Canada) . Thi s proces s als o results i n tight-ga s deposit s i n th e basi n deep s (e.g. Gree n Rive r an d Piceanc e Basins , USA) , which ar e a varian t o f basin-centre d ga s deposits. Wit h al l basin-centre d ga s deposits , commercial productio n i s dependen t o n finding 'swee t spots ' wher e reservoi r destruc tion fro m authigeni c mineralizatio n wa s impeded b y a ga s phas e predatin g th e exhumation. Rejuvenation o f basi n formatio n wit h conse quent basi n downwar p an d sedimentatio n ca n result i n th e mos t prospectiv e o f al l oil exploration targets : burie d faul t zones . Wit h resurgent sedimentation , thic k sequence s o f unfaulted sediment s ca n b e deposite d ove r th e previously exhume d sedimentar y section , wit h two important results: (1) marine source sections that eithe r wer e no t burie d deepl y enoug h t o generate an d expel HCs , or had been generatin g but were interrupted by the exhumation, will now be burie d deepl y enoug h t o achiev e th e hig h ranks necessar y t o commenc e H C generation in hydrogen-rich marin e organi c matter ; (2 ) th e thick sectio n o f overlyin g unfaulte d shale s capping th e burie d faul t zone s (a ) focuse s vertically migratin g oi l int o th e firs t reservoi r of the trap, (b) prevents the HCs from substantial migration (excep t limited updi p migration), and (c) serve s a s a n excellen t seal , allowin g microseepage bu t no t macroseepage . Th e ke y to suc h petroleu m system s i s to identif y deepe r reservoirs, whic h are connected t o deeply burie d known sourc e rocks, while placing less exploration emphasi s o n th e overlyin g unfaulte d sediments. Lee Price of the US Geological Survey, Denver, died in August 2000, shortly afte r th e conference fro m whic h these papers are taken. We include Leigh's abstract to reflect hi s livel y participatio n i n th e conference , an d his significan t contributio n t o th e understandin g o f uplifted basin s of the USA .

A. G. Dore for the editor s

From: DORE , A.G., CARTWRIGHT , J.A., STOKER , M.S. , TURNER , J.P . & WHITE , N. 2002 . Exhumation o f th e North Atlantic Margin: Timing, Mechanisms an d Implications for Petroleum Exploration. Geologica l Society, London, Specia l Publications, 196 , 431 . 0305-8719/027 $ 15.00 © The Geological Society o f London 2002 .

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Diagenesis and fluid flow in response to uplift an d exhumation JOHN PARNEL L Department of Geology and Petroleum Geology, King's College, University of Aberdeen, Aberdeen AB24 SUE, UK (e-mail: j.parnell@ abdn.ac.uk) Abstract: Uplif t o f sedimentar y rock s i s accompanie d b y a wid e rang e o f physica l an d chemical changes that contribut e t o diagenesis an d modify fluid flow regimes. Topograph y becomes a major drivin g force behin d fluid flow patterns, and meteoric water may penetrate to several kilometre s below th e surface. Typica l diageneti c processes include alteratio n an d leaching o f feldspar s an d other unstabl e minerals , precipitatio n o f iro n oxide s an d kaolin , and leachin g o f carbonat e an d sulphat e cements . Reservoire d oi l ma y b e degrade d b y near-surface waters , but reservoir rocks may become more oil-wet. Brittle fracturin g is enhanced nea r the surface, an d fluid flow may become predominantl y fracture-bound as fractures dilate . Uplift also cause s tiltin g of fluid contacts and remigratio n of hydrocarbons . Exsolutio n an d expansio n o f ga s similarl y cause s remigratio n o f oi l t o peripheral traps. Although basi n uplif t i s generall y regarde d a s bein g detrimenta l t o hydrocarbo n prospectivity, especiall y as a result o f breaching of traps, ther e is also an enhanced potential for hydrocarbo n play s based o n reserves o f exsolve d gas , condensat e dropout , periphera l traps an d fracture d reservoirs .

Although muc h researc h wor k o n exhume d basins focuse s upo n th e origin s o f uplif t an d exhumation, the resultant changes in geothermal patterns and the consequences of exhumation for regional sedimentatio n patterns , one of the most important economic aspect s of exhumation i s the effect upo n basin diagenesi s an d fluid flow. This is particularl y significan t t o hydrocarbo n pro spectivity wher e i t influence s reservoi r quality. Although treatments of diagenesis rarely address exhumation i n a specifi c manner , ther e ar e numerous diageneti c processe s tha t ar e com monly associate d wit h uplift an d ma y therefor e be predictabl e consequence s o f exhumation . As many cas e studie s o f diagenesis, particularl y those undertake n befor e read y availabilit y o f offshore wel l cores , hav e bee n carrie d ou t i n onshore basins, there are many data for rocks that have experienced exhumation. All rocks exposed at outcro p hav e bee n uplifte d sinc e maximu m burial an d therefore ca n be sai d t o be exhumed . Exhumation is taken as the upward displacement of rock s with respect t o th e surfac e (Englan d & Molnar 1990) , th e drivin g forc e fo r whic h i n many case s i s basi n inversion . Thi s revie w outlines th e rang e o f processe s involve d i n diagenetic modificatio n of exhume d basins an d the consequence s fo r reservoi r quality , and th e controls o n changin g flui d flo w patterns . Th e

review include s aspect s o f diagenesi s an d flui d flow relevant to uplift in general, includin g case s where rock s hav e subsequentl y bee n reburied . Given th e importanc e o f exhumatio n i n N W Europe (e.g . wes t of Shetland, nort h Celtic Sea , east Iris h Sea ; Hilli s 1995) , i t i s valuabl e t o b e aware o f thes e processe s an d thei r potentia l consequences.

Processes during exhumation The effect s o f exhumatio n includ e a rang e o f physical, chemical and biological processes . The physical processe s includ e rheologica l effects , particularly fracturing , bu t als o geometrica l effects a s a result o f uplift . Chemica l processe s involve change s i n por e flui d chemistry , an d hence mineral precipitation or dissolution, owing to change s i n pressure-temperatur e an d flui d flow regimes . Biologica l processe s involv e degradational modification of hydrocarbon fluids as a resul t o f th e ingres s o f surficia l water s t o uplifted reservoirs . Thes e variou s type s o f process ar e clearl y interrelated . Fo r example , new fractur e system s pla y a n importan t rol e i n allowing circulation of fluids that cause chemical and biological changes .

From: DORE , A.G., CARTWRIGHT , J.A., STOKER , M.S., TURNER , J.P . & WHITE , N. 2002 . Exhumation of th e North Atlantic Margin: Timing, Mechanisms and Implications for Petroleum Exploration. Geological Society, London, Special Publications, 196, 433-446. 0305-8719/02/$15.00 © The Geological Society of London 2002.

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Physical processes Changes in pressure and temperature, which bot h generally decreas e durin g exhumation , hav e diverse consequence s fo r th e propertie s o f rocks, mineral s an d entraine d fluids , a s follows . (1) Rheological behaviour. At shallow depths , rocks tha t ma y hav e deforme d i n a ductil e manner before inversio n may behave instead in a brittle manne r (Downe y 1994 ; Ingram & Ura i 1999). Th e importan t aspect s o f fracturin g ar e discussed below . (2) Mineral solubility . Change s i n temperatur e have a stron g influenc e o n th e solubilit y o f th e main mineral s tha t cemen t sedimentar y rocks , i.e. silic a an d carbonates . Silic a solubilit y decreases markedl y wit h fallin g temperatur e (Fournier 1983) . Carbonat e solubility , including calcite, b y contras t increase s wit h fallin g temperature, bu t i s subjec t t o a greate r influence fro m PCO 2, whic h fall s wit h pressur e drop durin g uplift , an d s o ca n allo w calcit e precipitation (Woo d 1986) . (3) Mineral transformation. Th e conversion o f one minera l t o anothe r o f a differen t chemistr y necessitates interactio n wit h ion s i n th e por e fluids, but may ultimately b e temperature driven , such as the conversion from smectite to illite with rising temperature . However , transformation s of this type may not be readily reversible , henc e the widespread occurrenc e o f authigeni c illit e i n exhumed rocks . (4) Minera l phas e changes . Temperatur e ma y also contro l transformatio n o f on e minera l polytype t o anothe r (e.g . kaolinite t o dickite ; chemically identical ) or fro m on e mineral phas e to another withi n a single chemica l syste m (e.g. gypsum t o anhydrit e plu s water ; effectivel y dehydration). Thes e change s ma y als o b e les s readily reversed , a s show n b y th e metastabilit y of dickit e an d anhydrit e a t th e surface . Thus , exhumation generall y ha s littl e consequenc e

for change s i n mineralogy . Nevertheless , prolonged interactio n o f anhydrit e wit h wate r should conver t i t bac k t o gypsum , wit h a n accompanying volum e increase . (5) Flui d phas e changes . Unlik e minera l phase changes , flui d phas e change s occu r readily wit h fallin g temperatur e an d pressure . In term s o f hydrocarbon s i n th e shallo w crust , this involve s exsolutio n o f ga s fro m oi l an d drop-out o f som e liquid s fro m ga s an d ga s condensates (e.g . Piggott & Lines 1991) . Ga s i s also release d fro m dissolutio n i n brine s a s th e pressure falls , althoug h ga s solubilit y ca n increase i n solutio n wit h fallin g temperatur e (Tiab & Donaldso n 1996) . Existin g gase s experience expansio n wit h fallin g pressure . Clearly, thi s releas e o f ga s cause s a n increas e in th e volum e occupie d b y hydrocarbo n fluids , and consequen t shiftin g o f flui d contact s withi n reservoirs, an d ma y includ e displacemen t o f hydrocarbons beyon d th e spil l poin t int o ne w reservoir compartments . Thi s ma y resul t i n successive tra p structure s updi p holdin g mor e oil a t shallowe r level s (e.g . Gusso w 1954 ; Fig. 1) , or los s fro m th e tra p structur e (Nyland et al 1992) . Onc e a tra p i s fille d wit h gas , no further oi l wil l enter it , i.e. a tra p fille d wit h oil is stil l a potential gas trap , but a tra p filled with gas i s no t a n effectiv e oi l tra p (Gusso w 1954) . The exsolutio n o f ga s fro m oi l o r wate r ca n also hel p t o develo p lo w permeabilitie s tha t function a s seals. Wher e pore s contai n tw o fluid phase s (e.g . oil an d gas , water an d gas) that hav e no t becom e separated , th e effectiv e permeability t o th e individua l phases i s muc h lower tha n t o a singl e pore-fillin g flui d phas e (Chapman 1983 ; Osborne & Swarbric k 1997) . (6) Flui d viscosity . Th e viscosit y o f oi l decreases wit h fallin g temperature , henc e migration o f oi l i s slowe r a t shallowe r buria l depths an d ma y b e effectivel y stoppe d i n low-permeability pathways.

Fig. 1 . Example o f varyin g flui d contact s i n successiv e tra p structure s i n a tilted system , showin g mor e oi l at shallower levels, Leduc Reefs, Albert a (modified fro m Gusso w 1954) .

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Chemical processes Starting fro m first principles, th e two media that are introduced from the surface upon exhumation are ai r an d surface-derive d water . I n term s o f reactive species , thi s involve s oxygen , carbo n dioxide an d water . Th e processe s tha t ma y b e induced b y thes e specie s ar e oxidation , aci d dissolution (solutio n o f carbo n dioxid e i n water is acidic ; dissolutio n i n neutra l wate r i s negligible) an d hydration . Proton s fo r acidit y are als o contribute d b y th e oxidatio n o f organi c matter an d carbonat e bufferin g reactions . A s meteoric wate r penetrate s t o severa l kilometre s depth (se e below) , thes e processe s ma y affec t significant proportion s of a sedimentary basin. An additional component introduced from th e surface i s bacteria , whic h mediat e man y chemical reactions, including sulphate reduction, and caus e th e degradatio n o f oil . I t i s wort h emphasizing th e importanc e of the water, as the carrier o f dissolve d gases , salt s an d bacteria . Although rock s tha t ar e bein g exhume d wil l i n most cases alread y be saturate d with water, new water enhance s th e likelihoo d o f diageneti c alteration. Diagenesis involving near-surface fluids The processes of oxidation, aci d dissolutio n an d hydration i n uplifte d rock s hav e importan t effects, particularl y i n sandstones, a s follows. (1) Oxidation cause s alteration o f any unstable detrital minerals (ferromagnesians , etc. ) tha t had survived previou s diagenesis , an d alteratio n o f authigenic minerals that had grown during burial diagenesis. This is most obvious in the formation of secondaril y reddene d sandstones , distinc t from sand s wit h re d grain-coating s forme d during deposition . Th e distinctio n i s important, as iro n oxide s forme d lat e i n diagenesi s ca n b e sufficiently abundan t to be pore-clogging , as opposed to the ver y fine grain-coating s tha t develop a t th e surfac e i n ari d t o semi-ari d environments (Walke r e t al. 1978) . Secondar y reddening ca n be widespread , as evinced by th e distribution o f reddene d Carboniferou s sand stones belo w th e sub-Permia n unconformit y in the norther n Britis h Isle s (Wan g 1992) ; i n tha t case, the reddened sandstone s are porous because of associate d carbonat e cemen t dissolution . I t is not simply a near-surf ace effect: th e reddening is recorded a t ove r 0.5k m belo w th e palaeo surface, fo r example , i n th e easter n Iris h Se a and Nort h Se a (Jackso n e t al . 1987 ; Cowa n 1989). (2) One of the most widely quoted examples of alteration o r dissolution o f feldspars by meteori c

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waters is in the Jurassic Brent Group sandstones of th e Nort h Sea , wher e dissolutio n t o creat e secondary porosit y ha s bee n partl y relate d t o Cimmerian uplif t (e.g . Sommer 1978 ; Glasman n et al 1989 ; Bj0rlykke et al. 1992). The degree to which thi s occurre d durin g uplif t a s opposed t o burial ha s bee n debate d (Burle y e t al . 1985 ; Nedkvitne & Bj0rlykke 1992), but the abundance of kaolin in the secondary porosity does suggest a meteoric influenc e (se e below) . (3) Dissolutio n o f carbonat e an d sulphat e cements by surface-derived waters may give rise to substantial porosity. This is widely observed in the Permo-Triassic sediment s of western Britain, which contai n calcite , dolomite , gypsu m an d anhydrite cements, but which have been leached both b y meteori c water s followin g Tertiar y inversion an d b y moder n groundwater s (e.g . Walton 1981 ; Burle y 1984 ; Stron g & Milo dowski 1987) . O f course, carbonate and evaporite rock s ma y b e dissolve d o n a larger scale , t o leave large cavities or collapse features. (4) Anothe r produc t o f th e acidit y o f near surface waters is the precipitation o f kaolin. Th e widespread occurrenc e of kaolin at unconformity surfaces (e.g . Esteoule-Chou x 1983 ) i s wel l known, bu t i t i s als o commo n t o fin d kaoli n associated with secondary porosit y afte r carbon ate cemen t dissolution , a s bot h ar e associate d with low-p H fluids . Example s hav e bee n give n by Curtis (1983 ) an d Parnell (1987) . (5) As natural fluids tend to be silica saturated and reservoir fluid s ar e ofte n oversaturate d (e.g . Bazin et al 1997) , this suggests that uplift should be a majo r caus e o f silic a cementatio n a s th e solubility fall s wit h temperature . However , except in cases wher e hot fluids are cooled very quickly (i.e. a more rapid process than uplift; see , e.g. Rossi et al 2000) , diagenetic studie s do not record thi s a s a process . Th e fac t i s tha t th e volume o f por e wate r i s inadequat e t o caus e much silic a precipitatio n an d a t th e lo w temperatures tha t accompan y shallo w depth s the precipitatio n rat e fo r silic a is extremely low (Bj0rlykke & Egeber g 1993) . Thi s exemplifies the general principle that chemical reaction rates increase wit h ris e i n temperature , i.e . low temperature processe s ar e no t kineticall y favoured. Smal l quart z outgrowth s linin g pore s could reflect silica deposition a s a result of uplift, but thi s proces s coul d neve r hav e a significant effect o n reservoir quality. (6) Given the widespread occurrence of basalts on the European Atlanti c margin, particularly o f Tertiary age , i t i s wort h notin g tha t basal t mineralogy i s especially susceptibl e t o alteration as i t di d no t for m i n equilibriu m wit h surfac e conditions. Alteratio n ma y als o occu r durin g

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burial an d hydrotherma l processes , whic h ma y be difficul t t o distinguis h fro m surficia l altera tion. A n exampl e o f basal t alteratio n (t o clays , zeolites an d silica ) tha t ca n b e attribute d t o surficial processes ha s been described b y Benson & league (1982) . The distinctio n betwee n ingres s b y meteori c water an d b y seawate r i s importan t becaus e meteoric wate r contain s muc h lowe r concen trations o f dissolve d ions . Th e lo w sulphat e content o f meteori c wate r enhance s th e like lihood o f siderite an d dolomit e precipitatio n (Kastner 1984) . Where meteoric water s mix with other water s o f differen t compositio n (e.g . connate waters) , ther e i s potentia l fo r minera l precipitation, althoug h densit y contrasts betwee n the fluids may limit th e degree to which mixin g occurs (Bj0rlykk e 1994) .

Porosity trend s Table 1 summarizes control s o n porosity declin e or enhancemen t durin g buria l an d uplift . Porosities may remain higher tha n normal durin g burial fo r variou s reasons , particularl y i f hydrocarbon emplacemen t (Worde n e t al. 1998 ) or clay coatin g (Ehrenber g 1993 ; Sulliva n et al . 1999) inhibit s cementation, overpressurin g inhi bits compaction (Osborn e & Swarbrick 1997 ) or secondary porosit y i s generate d (Schmid t & McDonald 1979a) . Thu s pre-exhumatio n poros ities ca n be high. O n the other hand , deep burial may hav e largel y eliminate d porosity , suc h tha t 'over-compaction' i s a typica l sig n tha t a basin has bee n uplifte d sinc e maximu m burial. Overcompaction i s widel y observe d i n basin s exhumed durin g Tertiar y tim e i n N W Europe , including those o f the UK (Hillis 1995) . Inverted basins inevitabl y exhibi t lowe r porositie s tha n expected fo r their current burial depth, as a result of th e effect s o f additiona l compactio n o r cementation. Dronkers & Mrozek (1991) showed a marked difference i n Triassic sandstones in the Broad Fourteen s Basin , betwee n tigh t strat a i n

the inverted basin area and porous strata in a noninverted platfor m sit e a t the sam e curren t buria l depth. Upon uplift , differen t processe s hav e ver y different consequence s for increasing or decreasing porosity . Fracturing , an d dissolutio n o f carbonate cement s o r unstabl e grains , ar e bot h widespread processe s tha t increas e porosity , whereas an y typ e o f cementatio n (silica , kaolinite, iro n oxide ) wil l decreas e it . A s explained above , kaolinit e precipitatio n ma y closely follow the creation of secondary porosity. The importanc e o f thes e trend s fo r hydro carbon systems depends on the relative timing of hydrocarbon migration . Exhumatio n inevitably turns of f th e hydrocarbo n generatio n process , because o f a dro p i n th e temperature , a s hea t i s needed t o drive the chemical reaction s involved . Thus post-exhumatio n porosit y i s relevan t t o remigrated existin g accumulation s or t o hydro carbons generated later after renewe d burial and a suitabl e therma l history . Hillie r & Marshal l (1992) describe d an exampl e in whic h hydro carbon generatio n i n th e Devonia n Orcadia n Basin wa s arreste d durin g Varisca n inversion , then recommence d durin g subsequent Mesozoic burial. An example of gas generation interrupted by inversion before regeneration was recorded by van Wijh e et al. (1980 ) i n th e Sol e Pi t Basin. The erosio n of rock, associate d wit h exhumation, ha s a n effec t o n pore flui d pressure , whic h has been investigated b y Luo & Vasseur (1995). The precis e manne r o f flui d pressur e declin e depends upo n permeabilit y an d reboun d beha viour ( a limited amount of compaction is elastic and reversible rather than plastic), but in general terms there is a time lag between commencement of erosio n an d reductio n i n effectiv e stres s (Fig. 2) . Thi s i s becaus e effectiv e stres s i s no t reduced unti l th e por e pressur e fall s belo w th e hydrostatic pressure , whic h doe s no t occu r instantaneously, especiall y i f th e syste m i s initially overpressure d (Lu o & Vasseu r 1995) . The erosio n rat e control s th e degre e t o whic h

Table 1 . Causes o f change i n porosity during burial an d uplift Porosity los s Burial Compaction by sediment loadin g Cementation (especiall y carbonates , quartz ) Uplift Cementation (especiall y kaolinite , iron oxides )

Porosity preservation or gain Grain o r cement dissolutio n a t dept h Preservation b y oi l or gas emplacement Preservation b y overpressuring Limited elasti c reboun d Dilation o f fracture s Grain o r cement dissolution near surface

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mechanical compaction , withou t th e compli cation o f cementation . I n reality , cementatio n (chemical compaction ) dominate s ove r mech anical compactio n abov e c . 9 0 °C (Bj0rku m et al 1998) . No t onl y i s overpressur e usuall y dissipated durin g uplift , bu t ther e ma y b e a slight developmen t o f underpressur e relate d t o decline i n por e flui d pressur e durin g sedimen t unloading (Fig . 4) . Secondary porosity and sandstone framework stability Fig. 2 . Evolutio n pathwa y fo r effectiv e stres s durin g burial, erosio n an d reburia l (i.e . developmen t o f unconformity), showin g tim e la g betwee n erosio n and declin e i n effectiv e stres s (afte r Lu o & Vasseur 1995).

porosity continue s t o declin e afte r commence ment o f erosio n (Fig . 3) , whic h i s anothe r consequence o f th e dela y i n reductio n o f effective stress . A t slowe r erosio n rates , porosity i s mor e likel y t o sho w a n eventua l reversal i n th e porosity-dept h tren d a s a result of elasti c rebound , althoug h o f ver y mino r magnitude. Thes e plot s ar e base d upo n

The identificatio n o f secondar y porosit y rathe r than primary porosity is important because it has very different geometr y an d other properties that are o f interes t t o petroleu m engineers . I t i s inhomogeneously distributed , an d the associate d permeability an d th e por e surfac e are a ar e different fro m thos e o f primar y porosit y (Schmidt & McDonald 1979b) . Secondary porosit y can of course be generate d at considerabl e depth s withi n a basin , particu larly followin g decarboxylatio n o f organi c matter i n sourc e rock s (Schmid t & McDonal d 1979a). Near-surfac e secondar y porosit y gener ation is also particularly a consequence of carbon dioxide availability , bu t fro m meteori c sources .

Fig. 3 . Influence of erosion rat e o n evolution o f porosity durin g erosion (afte r Lu o & Vasseur 1995) .

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Fig. 4 . Modelled effect o f erosion on pore pressure evolutio n fo r a particular burial history (detail s give n by Luo & Vasseur 1995) : (a ) por e pressur e evolutio n wit h depth; (b ) overpressur e evolutio n with depth ; (c ) overpressur e evolution wit h time . Arrow s indicat e directio n o f pathway . Continue d porosit y los s afte r erosio n reflect s continued increas e in effective stress.

The significan t differenc e i n the respons e of th e rock i s tha t sandston e framewor k stabilit y (Nagtegaal 1978 ) i s critica l a t depth, bu t les s s o in th e shallo w subsurface , i.e . th e surroundin g fabric i s mor e likel y t o collaps e afte r cemen t dissolution unde r th e hig h pressur e o f dee p burial. However, i f a sandston e ha s bee n deepl y buried the n exhume d i t i s likel y t o hav e developed a very rigi d framework , conferred b y tight grain packing, welded grai n boundaries and quartz overgrowths. I f secondary porosity ca n be generated i n exhume d rock s i t shoul d b e mechanically ver y stabl e unles s i t occupie s s o much volume that a 3D grain framework does not remain. Secondary porosit y ma y als o b e generate d from th e dissolutio n o f unstabl e grain s a s exemplified i n th e Bren t Grou p sandstones . Grain dissolution is more likely to happen during deep burial, fo r example, b y late diagenetic illit e formation fro m K-feldspa r an d kaolinit e (e.g . Ehrenberg & Nadea u 1989) , a s b y th e tim e sandstones are exhumed most unstable grains are already lost . Thi s i s wort h emphasizing : exhumed sandstone s are , i n general, mineralogi cally mor e mature tha n thos e a t the sam e buria l depth that have not been deeply buried. However, feldspars can survive deep burial (e.g. sandstones lacking kaolinite) , an d i f abundan t enoug h ca n lead to framework collapse upon dissolution afte r uplift. A dataset of uplifted Cambria n sandstones in Shropshir e (Parnell 1987 ) show s low porosity where ther e ar e fe w feldspar s t o b e leached , high porositie s wher e mor e feldspar s coul d b e leached, bu t lo w porosities wher e s o man y

feldspars wer e presen t tha t th e sandston e framework collapse d afte r dissolutio n (Fig . 5) . Meteoric wate r flow can feasibl y penetrate t o several kilometre s dept h (Bethk e e l al. 1988) , although the greatest flow rates will occur closest to the surface. Flow rate s at shallow depth ar e up to severa l order s o f magnitud e greate r tha n compaction-driven flo w rate s (Gile s 1987 ; Harrison & Summ a 1991) , an d s o hav e th e potential t o accoun t fo r relativel y rapid minera l dissolution an d precipitation . Thi s i s eviden t beneath present-da y lan d surfaces , where , fo r example, Longstaff e (1984 ) reporte d kaolinit e growth down to severa l hundred metres dept h in the Albert a Basin , an d Bat h e l a l (1987 ) recorded feldspa r dissolutio n a t simila r depth s in Britis h Triassi c reservoirs , bot h cause d b y penetration b y meteori c fluids . Similarly , leaching by meteoric waters is evident beneath palaeoexposure surfaces , i.e . belo w unconformitie s (Shanmugam 1988) . Figur e 6 show s ho w kaolinite dominate s th e cla y mineralog y o f th e upper part s o f Lowe r Carboniferou s sections in Northern Ireland , wherea s illit e occur s i n th e lower parts . Wher e the y occu r together , illit e occurs earlie r i n th e diageneti c sequenc e tha n kaolinite, s o thi s distributio n i s no t simpl y a consequence o f kaolinit e reactin g wit h feld spars t o produc e illit e (se e Ehrenber g & Nadeau 1989) . Rather , th e kaolinite , an d associated hig h porosities , reflec t diagenesi s below th e sub-Permia n unconformit y (Parnel l 1991). Sequences wit h complex buria l historie s ma y include multipl e episode s o f uplift-relate d

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Fig. 5 . Cross-plo t o f percentag e oi l residu e (representin g porosity ) agains t quart z grai n conten t i n Cambria n sandstone, Shropshir e (after Parnell 1987) , showing importanc e o f feldspar leachin g t o form secondar y porosity , but framewor k collapse wher e feldspar content i s too high. diagenesis, includin g Atlanti c margi n basin s e t al 1998 ) show s th e predominanc e o f kaoli n where Cretaceous-earl y Tertiar y uplif t i s wide - an d iro n oxid e cementatio n an d cemen t dissol spread in addition to more recent exhumation. An utio n durin g bot h uplif t phases , an d als o example i n Fig . 7 fro m N E Brazi l (afte r Garci a degradatio n of oil during the palaeo-uplift event.

Fig. 6. Section s throug h the Lowe r Carboniferou s succession acros s Norther n Ireland , showin g distribution of predominant clay mineral cement (K, kaolinite; I , illite), an d most porous zones (starred). Kaolinite affect s uppe r parts of sections a s a result of ingress of oxidizing fluids during Permian uplif t (see Parnell 1991). Co, Courceyan ; Ch, Chadian; Ar, Arundian; Ho, Holkerian; As, Asbian; Br, Brigantian.

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Fig. 7 . Diageneti c sequenc e fo r Lowe r Cretaceou s Serrari a sandstones , central Sergipe-Alagoa s Basin , N E Brazil, showin g consequences of both palaeo-uplift and recent exhumation events (modified fro m Garci a et al. 1998).

Fracturing Fracturing i s widel y develope d i n inverte d basins, a s tectonic activit y i s fundamental to the inversion process . O n th e Atlanti c margin , fracture set s includ e thos e associate d wit h extensional reactivatio n o f syn-inversio n com pressional structure s (Dor e et al. 1999) . Fractur ing ma y als o b e a resul t o f th e reductio n i n effective stres s owin g t o uplift , regardles s o f th e tectonic settin g (Sibso n 1995) . Several aspect s o f th e fracturin g relate d t o inversion ar e important , a s follows. (1) Th e adjustmen t fro m ductil e t o brittl e deformation affect s rocks that function as seals to trapped hydrocarbons , particularl y i n mudrocks , which ma y consequentl y allo w leakag e throug h new fractures . Ductilit y decrease s wit h density , so mudrocks tha t have been buried then exhume d are particularl y likel y t o exhibi t fracturin g (Ingram & Urai 1999 ; Corcora n & Dore 2002) . Potentially, brittl e deformatio n coul d als o develop i n evaporite s (Downe y 1994) , althoug h this is less likel y than i n mudrocks, a s shown by hydrocarbon distributio n i n th e eas t Iris h Se a (Seedhouse & Race y 1997) . Th e coherenc y o f ductile seal s i s therefor e critica l t o th e prospec tivity of exhumed basin s (Gabrielse n & K10vja n 1997; Spai n & Conra d 1997 ; Corcora n & Dor e 2002), althoug h thi s als o depend s upo n othe r factors, particularl y th e styl e o f th e inversio n process. (2) Fracturing , an d increasin g dilatanc y o f existing fractures , ma y caus e flui d flo w t o

become predominantl y fracture-boun d rathe r than through matri x permeability . Wher e permeabilities wer e alread y ver y lo w a s a resul t o f cementation o r compactio n durin g burial , frac tures ma y accoun t fo r almos t al l effectiv e permeability. Figur e 8 (after Dronkers & Mrozek 1991) show s a n example fro m the margi n o f th e inverted Broa d Fourteen s Basin , i n whic h th e effective permeabilit y i s i n fracture s relate d t o reverse faulting an d tight folding. Other planes of weakness, suc h a s beddin g plane s o r stylolites , may als o sho w increase d dilatancy . Bolto n e t al. (2000) suggeste d tha t horizontal effectiv e stres s decreases a t a slowe r rat e tha n vertical effectiv e stress durin g unloadin g o f clay-ric h sediments , so that horizontal fractures develop, leadin g to a substantial increas e i n anisotropi c permeability . (3) Compressiona l deformatio n durin g inver sion is likely to produce fractur e networks with a marked anisotropy , i.e. fractures show a preferred alignment, o r a t leas t th e degre e o f fractur e dilatancy i s greate r i n a preferre d orientation . The resul t is fluid flow patterns tha t are focuse d along preferre d fractur e orientations . Thi s i s i n contrast t o th e pattern s o f flo w throug h matri x permeability, whic h tend to be controlle d by th e dip o f th e aquifers . A n exampl e o f directiona l permeability i n fracture s relate d t o inversio n i n the Bristo l Channe l basi n ha s bee n recorded b y Nemcok e t al . (1995).Tw o othe r structura l aspects o f uplif t have a significan t influenc e on fluid flow, as follows. (4) Uplif t almos t certainl y wil l caus e relativ e tilting o f bot h lithologica l boundarie s an d flui d

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Fig. 8 . Cross-plot o f porosity an d permeability fo r an oil wel l a t th e margi n o f th e Broa d Fourteen s Basin , showing enhance d permeabilit y i n fracture d sand stones compare d wit h rock s i n whic h permeabilit y i s through matrix (afte r Dronker s & Mrozek 1991) .

contacts. Bot h ma y caus e remigration o f hydrocarbons (Gusso w 1954) . In som e case s thi s wil l involve only minor movements within an aquifer, but i n other s i t wil l pus h hydrocarbon s beyon d their spil l point , whic h coul d involv e migratio n into previousl y unuse d reservoi r compartment s (e.g. Cornfor d 1990) . I n reservoir s wher e cementation ha d bee n arreste d o r inhibite d b y hydrocarbon emplacement , bu t continue d i n the water-filled zone , tiltin g migh t mov e hydrocar bons int o rocks o f poorer reservoi r quality , with the geometrica l consequenc e tha t th e hydro carbons take up a greater volum e of rock. (5) Uplif t increase s th e likelihoo d tha t topography i s a drivin g forc e behin d fluid flow (Bredehoeft e t al 1992 ; Demin g 1994) . Thi s i s partly becaus e th e aquife r i s probabl y close r t o the surfac e an d mor e susceptibl e t o downward flowing meteoric waters , but more importantl y i f the exhumed basin is topographically higher than surrounding area s ther e ma y b e a significan t hydraulic hea d developed . I n th e Broa d Four teens Basin , offshor e Netherlands , whic h ha s been inverte d b y a t leas t 2km , the n reburie d (Hooper e t al 1995) , modellin g (Verwei j 1999 ; Verweij e t al. 2000) emphasizes th e importanc e of topography-driven flow syn- or post-inversion, where i t ha s contribute d t o alteratio n o f oil s

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where they have been expose d t o surface-derive d waters (De Jager et al. 1996).I n addition to these structural aspect s o f flui d flow , on e o f th e mos t fundamental consequence s o f exhumatio n i s o f course tha t seal s o r aquitard s ar e breache d b y erosion an d allow fluid leakage (an d also ingress of surficia l fluids ; se e above) . Leakag e include s drainage o f hydrocarbons , whic h i s believe d t o have occurre d i n th e Bristo l Channe l Basi n vi a inversion-related fracture s (Nemco k e t al . 1995). Fluid flow pathways and diagenetic processe s are closely linked , a s diagenesis require s impor t and export of ions. Once fracturing becomes th e major flow pathway, this is where muc h minera l deposition occurs, i.e. fractures become sink s for ion precipitation, and can be infilled t o the point of closure . Th e fracturin g o r faultin g associate d with inversio n allow s cross-formationa l flo w o f fluids tha t contribute s t o diagenesis . Thi s includes accessin g o f water s fro m aquifer s which ma y b e eithe r abov e and/o r belo w th e horizon o f interest . Fo r example , studie s o f Lower Permia n Rotliegen d sandstone s i n bot h the North Sea (Sullivan et al. 1994) and Germany (Platt 1994 ) sho w tha t cementatio n involve d a contribution o f fluid s fro m th e overlyin g Zechstein units , whic h coul d onl y ente r th e Rotliegend sandstone s after fracturin g relate d t o Early Cretaceou s inversion . I n thi s example , rapid movemen t o f larg e volume s o f flui d occurred becaus e overpressure d Zechstei n por e fluids were suddenly able to drain into the lower pressure Rotliegend sandstone .

Consequences for oil-bearing reservoirs The influenc e o f near-sur f ace water s o n th e chemistry o f oi l ha s bee n widel y describe d elsewhere (e.g . Tisso t & Welte 1984) , an d wil l not b e reviewe d her e excep t t o emphasiz e th e general effec t o f degradin g oi l qualit y (i.e . increasing it s viscosity) , particularl y throug h water washin g an d biodegradation . I n Europe , for example , th e distribution of biodegraded oil s offshore Netherland s reflect s th e effect s o f surface-derived water s followin g inversio n (D e Jager et al. 1996). There is debate over the degree to whic h biodegradatio n i n reservoir s occur s before an d after entrapment ; fo r example, i n th e case o f th e hug e deposit s i n th e Lowe r Cretaceous sandstone s o f Albert a (Crean y & Allan 1990) . I n inverte d basins , biodegradatio n specifically linke d t o exposur e t o aquifer s mus t be mor e likely , an d a wid e spectru m o f oi l preservation i s foun d i n this settin g (Macgrego r

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1993, 1996) . A n interestin g aspec t o f uplifte d basins is that prior deep burial may 'sterilize ' the oil so that biodegradation i s prevented following uplift (Wilhelm s et al 2001) . The mineralogica l change s tha t occu r follow ing uplif t affec t th e wettabilit y o f sandston e reservoirs. Th e precipitation o f grain-coatings of iron oxid e (Wan g & Guidr y 1994) , clay s including kaolinit e (Robi n e t a l 1995 ) an d biodegraded oi l residue s (Gonzale z & Midde a 1987) al l ten d t o mak e th e sandstone s oil-we t rather tha n water-wet. The chemistry o f the pore waters also influence s wetting characteristics, in a comple x wa y (Anderso n 1986) . Althoug h oxidizing water s generall y enhanc e oil-wetting, low-salinity water s enhanc e water-wetting . A s near-surface water s ma y b e bot h oxidizin g and low salinity , their effec t o n wettin g properties i s difficult t o predict . However , th e grain-coatin g phases ar e likel y t o b e predominant , an d th e

oil-wet sandstone s tha t they engende r ma y bot h influence hydrocarbo n flo w behaviou r an d restrict subsequen t cementation. The influ x o f relativel y fres h water s ma y reduce th e resistivit y contrast betwee n oi l an d water, makin g difficul t th e detectio n o f hydro carbons base d o n resistivity measurements. This is comparable wit h the problems arisin g fro m th e use o f freshwater-base d drillin g mud s (Ride r 1996). As explaine d above, ga s expansion , reservoi r tilting and leakage alon g fractures may al l cause displacement o f liqui d hydrocarbon s fro m reservoir compartments , s o tha t underfille d traps, palaeo-fluid contacts , residual oil columns and periphera l oi l accumulation s are al l typical features o f exhume d basins , includin g those o n the Atlanti c margin (Dor e & Jensen 1996 ; Dor e et al. 2000). Basin-centre d gas accumulations are also a potential result (Cramer e t al . 1999) . Ga s

Fig. 9 . Hydrocarbon fluid-fill history of Permo-Triassic reservoir compartments, eastern Irish Sea (modified from Duncan e t al . 1998) . Ke y episodes: (a ) lat e Jurassic-earl y Cretaceou s oi l accumulation ; (b) earl y Cretaceou s inversion, breaching o f traps, lateral migration to peripheral traps; (c) early Tertiary renewe d burial, new oil and gas charge ; (d ) mid-Tertiar y uplift , furthe r breaching, gas expansion and condensate dropout.

EFFECTS OF UPLIFT AN D EXHUMATION migration ma y als o b e directl y manifes t a s ga s chimneys an d sea-floor seepages . A n example of the displacemen t processe s i n th e easter n Iris h Sea has been describe d b y Duncan et al (1998) , and i s summarize d i n Fig . 9 . Permo-Triassi c reservoirs, fille d fro m Carboniferou s sourc e rocks, revea l a comple x fluid-fil l stor y relate d to their burial history. Oi l generated befor e earl y Cretaceous (Cimmerian ) inversio n wa s los t o r displaced by inversion, leaving breached traps of good reservoi r quality , an d ne w basin-fring e oi l accumulations. Renewe d buria l durin g earl y Tertiary time cause d refillin g o f trap s wit h oil , and some charging by gas. Tertiary uplif t cause d renewed breaching , ga s expansio n an d condensate dropou t fro m gas . I n thi s example , futur e exploration wil l require identification o f areas to which remigratio n ha s occurre d (Dunca n e t al . 1998).

Conclusions In summary , man y aspect s o f uplif t an d exhumation ar e detrimenta l t o hydrocarbo n prospectivity, including : (1 ) structura l an d erosional breachin g o f traps ; (2 ) degradin g o f oil quality by influence of surface-derived water; (3) minera l depositio n (excep t carbonates ) a s a result o f decreasing solubilit y a t lower tempera tures. However , give n th e constraint s o f sea l integrity, ther e i s potentia l fo r distinc t style s of hydrocarbon play : (1 ) large reserves of exsolve d gas, especiall y a t basi n centres ; (2 ) reserve s o f condensate dropou t fro m gas ; (3 ) remigrated oi l at shallo w field-periphera l sites ; (4 ) fracture d reservoirs, wit h enhance d porosit y an d permeability. I a m mos t gratefu l to A . G . Dore an d a n anonymous reviewer for critical commen t on an earlier versio n of the manuscript , an d t o J . B . Fulto n fo r essentia l cartographic assistance .

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development o f it s petroleu m systems : a n inte grated 2 D basi n modellin g approach . Journal o f Geochemical Exploration, 69-70 , 635-639. WALKER, T.R. , WAUGH , B . & CRONE , A.J . 1978 . Diagenesis i n first-cycl e deser t alluviu m o f Cenozoic age , southwester n Unite d State s an d northwestern Mexico . Geological Society o f America Bulletin, 89, 19-32. WALTON, N.R.G . 1981 . A Detailed Hydrogeochemical Study of Ground\vaters from the Triassic Sandstone Aquifer o f South West England. Repor t o f th e Institute of Geological Science s 81/5 . WANG, F.H.L . & GUIDRY , L.J . 1994 . Effec t o f oxidation-reduction condition s o n wettabilit y alteration. Society o f Petroleum Engineers Formation Evaluation, 9, 140-148 . WANG, W.H . 1992 . Origin of reddening and secondary porosity i n Carboniferou s sandstones , Norther n Ireland. In: PARNELL, J. (ed.) Basins on the Atlantic Seaboard: Petroleum Geology, Sedimentology and Basin Evolution. Geologica l Society , London , Special Publications , 62, 243-254. WILHELMS, A. , LARTER , S.R. , HEAD , I. , FARRIMOND, P., DIPRIMIO , R . & ZWACH , C . 2001 . Biodegra dation o f oi l i n uplifte d basin s prevente d b y deep-burial sterilization . Nature, 411 , 1034-1037 . WOOD, J.R . 1986 . Therma l mas s transfe r i n systems containing quart z an d calcite . In : GAUTIER , D.L . (ed.) Roles o f Organic Matter i n Sediment Diagenesis. Societ y o f Economi c Paleontologist s and Mineralogists , Specia l Publications , 38 , 169-180. WORDEN, R.H. , SMALLEY, P.C. & OXTOBY , N.H. 1998. Can oi l emplacemen t sto p quart z cementatio n i n sandstones? Petroleum Geoscience, 4 , 129-138 .

Uplift-related hydrocarbo n accumulations: the release o f natural ga s from groundwate r BERNHARD CRAMER 1, STEFA N SCHLOMER 2 & HARALD S. POELCHAU3'4 1 Federal Institute fo r Geosciences and Natural Resources (BGR), Stilleweg 2 , 30655 Hannover, Germany (e-mail: bernhard.cramer@ bgr.de) 2 EniTecnologie SpA, Via F. Maritano 26, 20097 San Donato Milanese, Italy 3 Institute of Petroleum and Organic Geochemistry, Forschungszentrum Julich GmbH, D-52425 Julich, Germany 4 Present address: Kansas Geological Survey, University of Kansas, 1930 Constant Avenue, Lawrence, KS 66047, USA Abstract: Vertica l tectoni c movement s ofte n chang e th e structura l styl e an d physico chemical habitat of sedimentary basins. Changes in pressure, temperature and salinity of the groundwater cause d by tectonic uplift may result in the release of previously dissolved gas . This process of gas exsolution from groundwater is shown t o be an important mechanism in the formation o f gas accumulations in uplifted basins . Two principal types of gas release are discussed. A hydrodynami c typ e i s activ e whe n groundwate r flow s int o area s o f lowe r pressure or mixes with water of different temperatur e or salinity. It is anticipated that this effect i s more of local importance, but over long periods of groundwater flo w larg e volumes of ga s ma y b e exsolved. The hydrostati c typ e o f gas release can occu r i n any sequenc e of sedimentary rocks where uplift causes a drop in pressure and temperature. This phenomenon may ac t basin-wide. Mass balance calculations sho w tha t th e largest gas accumulations on Earth, suc h a s the Urengo y fiel d i n West Siberia , could have been forme d b y thi s process .

During uplift of sedimentary rocks, maturation of sedimentary organi c matte r an d associate d hydrocarbon generatio n ceas e a s a resul t o f th e drop i n temperature . I n addition , variou s uplift related processes ar e known to be responsible for the destructio n o f petroleu m reservoirs . Th e understanding o f thes e mechanism s le d t o th e evaluation o f man y inverte d sedimentar y basin s as non-prospective fo r commercia l hydrocarbo n accumulations. Th e influenc e o f uplif t o n hydrocarbon system s o f sedimentar y basin s i s much mor e comple x an d ma y eithe r caus e destruction o f hydrocarbo n accumulation s o r induce redistributio n o f hydrocarbon s int o new , uplift-related type s o f accumulation s (Dor e & Jensen 1996) . Processe s tha t influenc e th e distribution o f hydrocarbon s i n exhume d basi n settings include : (1 ) th e dismigratio n o f hydro carbons a s a resul t o f structura l tiltin g an d fracturing o f cap rocks; (2 ) the diffusional losse s of ligh t hydrocarbon s fro m th e reservoir , whic h are no t replenishe d a s hydrocarbo n generatio n ceases durin g uplift; (3 ) anomalous rock properties such as mature or cemented rocks at shallow depth; (4) the presence of fluids in disequilibrium

leading t o ga s exsolutio n fro m por e wate r o r liquid hydrocarbons , expansio n o f fluids , especially gas , an d retrograd e condensation. O f these phenomena w e address here the exsolutio n of ga s fro m formatio n wate r a s a n uplift-related process resulting in new accumulations of natural gas. Afte r a shor t revie w o f th e solubilit y an d occurrence o f ga s i n th e dee p hydrosphere , th e mechanisms o f uplift-relate d ga s releas e fro m groundwater are discussed. Solubility of gas in water The solubilit y o f ga s component s i n wate r depends o n a variet y o f factors . Pressur e (P) , temperature (7) , concentration an d compositio n of inorganic components i n the water, as well as the contributio n o f othe r ga s component s i n solution are the most important known factors. A large numbe r of measurement s of ga s solubilit y in wate r ove r wid e ranges o f pressure, tempera ture an d salinit y ar e available , wit h mos t published experimenta l dat a focuse d o n th e solubility of methane in pure water or water with a singl e electrolyt e suc h a s sodiu m chlorid e o r

From: DORE , A.G., CARTWRIGHT , J.A. , STOKER , M.S., TURNER , J.P . & WHITE , N . 2002 . Exhumation o f th e North Atlantic Margin: Timing, Mechanisms an d Implications for Petroleum Exploration. Geologica l Society , London, Special Publications, 196, 447-455. 0305-8719/02/$ 15.00 © The Geological Societ y of London 2002 .

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B. CRAME R ETAL.

calcium chlorid e (Culberso n & McKett a 1951 ; O'Sullivan & Smith 1970 ; Sultanov et al 1972; Bonham 1978 ; Price 1979 ; Cramer 1980 ; Pric e et al . 1981 ; Rettic h e t al . 1981) . Som e experimental dat a ar e availabl e als o fo r ethan e (Culberson & McKetta 1950 ; Rettich et al. 1981 ; Crovetto e t al . 1984) . Investigation s o f th e solubility o f binar y o r ternar y hydrocarbo n ga s mixtures i n wate r ar e limite d t o moderat e P- T conditions (Amirijafari & Campbell 1972) . Only one serie s o f measurement s ha s bee n publishe d on th e solubilit y o f a real natura l ga s i n oilfiel d brines unde r a limite d rang e o f pressur e an d temperature condition s (Dodso n & Standin g 1945). Calculated wit h a model based on experimen tal dat a (Haa s 1978) , Fig . 1 summarize s th e solubility o f methan e fo r temperature s betwee n 20 an d 20 0 °C an d pressure s rangin g fro m 1 to lOOMPa i n brine s wit h thre e NaC l concen trations. Some general conclusion s can be drawn regarding th e solubilit y o f methan e unde r pressure, temperatur e an d salinit y condition s typical fo r groundwate r withi n th e uppermost 6-8 km o f th e Earth' s crust : (1 ) A t constan t temperature, th e solubilit y of methan e increase s with increasin g pressure . (2 ) Betwee n abou t 6 0 and 9 0 °C th e solubilit y o f methan e i n wate r has a minimu m fo r a constan t pressure . A t temperatures abov e thi s minimum, the solubility

increases wit h rise in temperature. The influenc e of temperatur e exceeds th e effec t o f pressur e o n methane solubility . (3) Increasing salinit y of th e brine suppresse s th e solubilit y o f methan e (salting out) . Difference s betwee n electrolyte s in th e salting-ou t effec t appea r t o b e smal l compared wit h the overall effect . (4 ) Solubilities of hydrocarbo n mixture s ar e greate r tha n th e solubilities o f th e pur e component s a t th e sam e pressure an d temperature . (5 ) Fo r th e norma l covariant ris e i n pressur e an d temperatur e wit h depth, methan e solubilit y increase s steadily , although no t a t constant rates . To predic t methan e solubilit y unde r geologi cally relevan t condition s a variet y o f mathe matical model s i s availabl e fro m th e literature . The concept s fo r thes e model s var y fro m semi empirical equation s mainl y base d o n curve fitting procedures (Haa s 1978 ; Coco & Johnson 1981; Pric e e t a l 1981 ; Battin o 1984 ) to mor e theoretically base d model s applyin g th e Pitze r phenomenology fo r th e liqui d phas e (Bart a & Bradley 1985 ) and a n equatio n o f stat e fo r th e vapour phas e (Dua n e t al . 1992) . Althoug h differences ar e obviou s i n th e precisio n o f these model s an d i n th e physicochemica l conditions the y cover , i t i s believe d tha t the y all ar e suite d t o predic t methan e solubilit y i n water unde r P- T condition s relevan t t o sedimentary basins .

Fig. 1 . Solubility of methane i n water a s a function o f pressure, temperatur e and NaCl concentration, calculated with the mode l o f Haas (1978) .

GAS RELEASE FROM GROUNDWATE R

Dissolved ga s in the Earth's crust The principa l ga s component s dissolve d i n th e crust down to the Mohorovicic discontinuity ar e methane, carbo n dioxide , hydroge n an d wate r vapour (Kortsenshtejn 1979). The composition of gas dissolve d i n near-surfac e groundwate r i s governed b y th e mai n atmospheri c ga s components nitroge n an d oxygen . Wit h increasin g depth, the influence of the atmosphere decrease s and ga s fro m bacteria l an d therma l degradatio n of sedimentar y organi c matter (e.g . CH4, CO 2), gas fro m recrystallizatio n processe s o f minerals (e.g. N 2, Ar) , a s wel l a s ga s fro m mantl e degassing (e.g . He , Ar ) contribut e t o th e dissolved gas phase of the groundwater. Secondary processe s suc h a s th e decompositio n o r generation o f individua l ga s components , a s a

449

result o f bacteria l activit y o r th e therma l degradation o f hydrocarbons , ma y significantl y change th e compositio n o f th e ga s dissolve d i n deep groundwater . A t leas t withi n th e dept h range o f hydrocarbon generatio n fro m sedimen tary organi c matter , methan e i s by fa r th e mos t important ga s component dissolve d i n ground water (Barka n & Yakutseni 1981). The rol e o f groundwate r a s a vas t storag e medium fo r ga s i n th e subsurfac e wa s empha sized by Kortsenshtejn (1979). H e estimated tha t at leas t 1 X 1019m3 ga s (al l volumes o f ga s ar e given i n m 3 ST P (standar d temperatur e an d pressure); 15. 6 °C, 1.01 3 kPa) ar e dissolve d i n water of the subsurface hydrosphere. This equals about twic e th e volum e o f th e Earth' s atmos pheric gas. The volum e o f ga s dissolve d i n fre e groundwater o f th e uppe r 5 km o f sedimentar y

Fig. 2 . Solubilit y regime s o f methan e i n subsurfac e groundwate r fo r hydrostati c an d lithostati c pressur e conditions an d salinity rang e o f 5-100g I" 1 NaCl . A geothermal gradien t o f 30 °C km^1 wa s assumed. I n the example, 2 km uplift results i n a drop i n solubility o f methane fro m 3.7 to 1. 9 m3 m , with 1.8m ' m release d from th e groundwater .

450

B. CRAMER ETAL.

rocks is estimate d to be at leas t 1.5 X 10 15m3, with methan e a s th e mai n constituen t (Kortsenshtejn 1979) . Thi s i s 1 0 time s th e estimated volum e o f th e globa l conventiona l gas reserve s (c . 0.15x1 0 m 3; Barthe l et al. 1999). In th e geologica l environmen t th e factor s influencing ga s solubilit y i n wate r describe d above lead to a general increas e o f gas solubility with increasin g dept h (Fig . 2) . The temperatur e minimum belo w 9 0 °C show n i n Fig . 1 i s compensated i n th e subsurfac e b y th e effec t o f increasing pressure. Nevertheless, becaus e of the temperature minimum in solubility, the temperature range u p to 90 °C, corresponding to a dept h down t o 2500-3500m , i s characterize d b y a lower rate of increase i n methane solubility than at greate r dept h (Fig . 2) . Unde r hydrostati c pressure within the uppermost 5 km of the Earth's crust, th e solubilit y o f methan e ca n excee d 5m 3 m~ 3 (Fig . 2) . Unde r lithostati c pressur e > 10m 3 m~3 of methane can be dissolved i n the pore water a t about 5 km depth . In the past, th e interest i n dissolved ga s in the subsurface wa s mainly focused on the economi c potential of gas dissolved i n brines (Kuuskraa & Meyers 1983 ; Marsde n 1993 ) a s wel l a s o n th e role of dissolved ga s for the deep gas potential of sedimentary basins . Becaus e wate r i s abl e t o store ga s effectivel y ove r a long time , dissolve d gas i n por e wate r i s believe d t o suppor t hydrocarbon potentia l eve n belo w th e dept h o f main hydrocarbo n generatio n (Barka n e t al . 1984; Pric e 1997) .

Uplift-related ga s release fro m groundwater Until recently , th e processes responsible for gas release fro m groundwate r i n th e dee p hydro sphere have not been investigated in much detail. However, i t i s clea r no w tha t th e principa l reasons fo r ga s exsolutio n ar e change s i n th e physicochemical habita t o f th e groundwater ; uplift-related change s in pore pressure , tempera ture an d salinit y ma y caus e th e releas e o f previously dissolve d gas . Processe s o f ga s release fro m groundwate r ca n b e classifie d as : (1) hydrostati c effect , no t relate d t o hydro dynamic activity ; (2 ) hydrodynami c effect , related t o th e flo w o f groundwater . Hydrostati c gas releas e ma y occu r i n al l sedimentar y sequences wher e regiona l uplif t cause s a dro p in subsurface pressure and temperature. Depend ing o n th e initia l ga s conten t o f th e por e wate r and o n th e amoun t of uplift , a critical point will be reache d a t whic h ga s solubilit y drop s

sufficiently t o initiat e ga s releas e fro m th e water. I n contrast , ga s releas e relate d t o hydrodynamic activit y may occu r i f a change i n pressure, temperatur e o r salinit y o f th e wate r is induced b y groundwate r flow , eithe r b y wate r flowing t o a regio n o f lowe r pressur e o r b y mixing wit h cooler or more salin e groundwater . During th e exhumatio n o f a sedimentar y basi n with a n activ e hydrodynami c system , bot h effects, hydrostati c a s wel l a s hydrodynamic , will simultaneousl y contribut e t o th e releas e o f gas. Hydrostatic effect To illustrat e th e potentia l o f thi s proces s t o release gas , methane-saturate d groundwate r i s assumed to have a salinity equivalent to 10 0 g 1~ l NaCl i n a sedimentar y formatio n a t 5 km dept h with a geotherma l gradien t o f 3 0 °C k m ] an d hydrostatic pressur e (Fig . 2) . Th e basi n is lifte d 2000 m and the solubility of methane drops fro m c. 3. 7 t o 1.9m 3 m~ 3 . Abou t 1.8m 3 m~ 3 o f methane coul d potentiall y b e release d fro m th e water. I n a water-saturate d sedimentar y layer of 10m thicknes s with a n averag e porosit y o f 8 % and a lateral exten t o f 1 km2 thi s uplif t woul d release abou t 1. 4 X 106m3 o f methan e fro m th e water. Th e duratio n of th e uplif t an d th e rat e of upward movemen t ar e no t considere d i n thi s calculation, becaus e it is not believe d tha t tim e plays a significan t rol e i n thi s typ e o f ga s exsolution. However, time becomes crucial when looking at the processes of gas migration to a trap and diffusiona l losse s fro m an accumulation. Hydrostatic ga s releas e ofte n occur s a t a regional t o basin-wide scale. The amoun t of gas released ca n b e enormous . T o evaluat e th e economic potential of this process, i t is important to conside r th e geologica l factor s promotin g effective ga s releas e fro m groundwate r an d accumulation o f th e ga s i n accessibl e hydro carbon traps. Favourable conditions are: (1 ) thick aquifer system s wit h hig h porosity ; (2 ) hig h contents of methane dissolved in the water (full y saturated) at maximum burial depth; (3) effectiv e buoyancy-driven migratio n o f th e release d ga s into (4 ) hydrocarbo n trap s wit h effectiv e ca p rocks under prevailing conditions. Factors (1 ) an d (2) , whic h determin e th e volumetrics o f th e ga s release , see m t o b e mutually exclusive. In general, good aquifers are found a t comparatively shallow depth, where the solubility o f methan e i n wate r i s stil l low . I n contrast, hig h ga s contents i n the pore water are expected t o b e foun d a t greate r dept h (Fig . 2) , where th e porosit y o f sediment s i s reduced . Exsolution o f large volumes of ga s require s that

GAS RELEASE FROM GROUNDWATE R

451

Table 1 . Published mass balance calculations o n the hydrostatic ga s release during uplift

Region General model (1 ) General model (2) Barents Sea (3) West Siberia (4)

Salinity o f Chang e in groundwater solubilit y 1 3 (g I" ) (m m~ )

Methane release below drainag e area (m 3)

Gas field

Initial dept h of aquifer (m )

Assumed amount o f uplift (m )

_

6000

3000

350 4.

1

_

-

6000

3000

100 8.

3

-

Sn0hvit

3400

1500

100-165 2.

0

(50-200) X 109

Urengoy

1800-4500

7

<9X10 1 2

600-1000

10-25 <1.

1, Barkan & Yakutseni (1981); 2, Maximov e t al. (1984); 3 , Oygard an d Eliassen, cited by Dore & Jensen (1996) ; 4, Cramer et al. (1999) .

at leas t on e o f th e tw o factor s i s favourable . Therefore, effectiv e ga s release fro m por e wate r can b e expecte d a t al l dept h ranges , bu t i n th e case o f shallo w aquifer s dow n t o c . 2500 m depth, sediment s wit h hig h porosit y ar e a prerequisite. Although th e formatio n o f ver y larg e ga s an d gas-condensate accumulation s a s a result o f ga s release fro m groundwate r i s believe d t o b e o f global importance (Maximo v et al. 1984), reports on actua l case s ar e sparse . Tabl e 1 summarizes published mass balance calculations of the effec t of uplift-related gas exsolution. Whereas the first two estimate s ar e mor e genera l i n natur e (Table 1) , the calculation s fo r th e Sn0hvi t fiel d in the Barents Sea and the Urengoy field in West Siberia ar e actua l cas e studies . Dor e & Jense n (1996) postulated that over the entire Barents Sea area vast amounts of gas may have been release d during Plio-Pleistocen e uplif t an d tha t th e formation o f majo r ga s accumulation s suc h a s the Shtokmanovsko e fiel d ca n b e attribute d t o this process. Fro m th e overall model o f the West Siberian Basi n a s a syste m o f hug e aquifer s discharging t o th e north , Crame r e t al . (1999 ) deduced tha t al l dr y ga s field s i n th e nort h o f West Siberi a wer e source d b y thi s process , an d that th e regio n o f a ga s releas e fro m wate r extends into the Kara Sea. In summary, the entire region fro m th e Middle O b in West Siberia ove r the Kar a Se a int o th e Barent s Se a experience d uplift durin g Cenozoic tim e an d wa s potentially subjected to release of gas from groundwate r and charging o f th e gian t ga s accumulation s identified i n this area . Basin-centred ga s field s (suc h a s th e Albert a Deep Basin or the San Juan Basin) are also likely to hav e bee n source d b y a n exsolutio n mechanism (Dor e & Jense n 1996) . Thes e ga s

accumulations li e i n dee p part s o f inverte d basins, downdi p fro m wate r wit h n o apparen t intervening permeabilit y barrier . Accordin g t o Price (cite d b y Dor e & Jense n 1996 ) th e underlying mechanis m is , mos t probably , tha t gas i s exsolved fro m wate r an d form s stati c ga s bubbles tha t bloc k the por e throat s and preven t further migration . Th e blocking o f pore spac e by gas exsolve d fro m groundwate r ha s als o bee n emphasized b y Kuo (1997). Hydrodynamic effect In all cases wher e Darcy flow of water through a permeable roc k a s a result o f latera l difference s in pressur e i s active, the water ma y pass a point where th e dro p i n hydrostati c pressur e i s sufficient t o initiat e a releas e o f dissolve d gas. From thi s point , th e wate r wil l continuousl y release ga s durin g it s passage throug h th e rock . This hydrodynamic gas release is not restricted to uplifted basi n settings . However , th e tectoni c tilting o f a sedimentar y basi n ca n activat e hydrodynamic systems , becaus e o f a n increas e in pressur e i n the continuousl y subsiding region and a drop i n pressure in the uplifted parts . This mechanism wa s show n t o b e activ e i n th e Wes t Siberian Basi n (Crame r e t al . 1999) , wher e a recent hydrauli c gradien t o f 6 x l O ~ 5 m m" 1 within th e artesia n Cretaceou s aquife r cause s a groundwater flow with an average linear velocity of about 20km Ma"1. The difference i n pressure between th e souther n an d th e norther n edg e o f the Urengo y anticline , whic h i s i n th e flo w direction of the groundwater, causes a drop in the solubility of methane of up to 0.017 m3 m~3 over this distance. This amoun t of gas was potentially released b y eac h cubi c metr e o f groundwate r passing th e Urengo y recharg e area , probabl y

452

B. CRAMER ETAL.

over a lon g tim e span . Crame r e t al (1999 ) calculated tha t u p t o I x l 0 1 2 m 3 o f methan e were release d durin g Cenozoic uplift , a s a result of th e hydrodynami c effec t withi n th e recharg e area o f th e Urengo y field . Therefore , thi s mechanism account s fo r a considerabl e portion , up to 12% , of the gas in place within the Urengoy field. Other possibl e mechanism s fo r activatin g groundwater flo w ar e upward-directe d wate r flow along fault s tha t were opene d a s a result of the uplift, o r an uplift-related depressurization of a seale d compartment , cause d b y ca p roc k o r fault sea l failure , in whic h larg e volume s o f ga s were dissolve d i n th e por e water . I n bot h cases , the subsequen t mixin g o f groundwater s wit h different salinitie s ma y accelerat e ga s release . Methane exsolutio n fro m wate r ma y als o b e enhanced whe n wate r flows past sal t dome s an d become mor e salin e durin g migratio n (Ku o 1997). Hydrodynamic gas release can be a regional or local phenomenon . I n compariso n wit h hydro static gas release, thi s process has a much smaller potential t o generat e economic ga s accumu lations. However , th e tim e facto r play s a n important role ; ga s exsolutio n relate d t o a longlasting hydrodynami c syste m ca n potentiall y release larg e volume s o f gas .

Efficiency o f gas release from groundwate r Accumulations o f natura l ga s ar e dynami c systems wit h continuou s diffusiona l losse s o f gas throug h th e ca p rock . A process suc h a s the release o f ga s fro m groundwate r ca n generat e significant accumulation s onl y i f th e rat e o f charging exceeds the rate of loss. To illustrate the efficiency o f gas release from groundwater, Fig. 3 displays a compariso n o f rate s o f therma l methane generation , ga s releas e fro m ground water an d diffusiona l losses , calculate d fo r th e Urengoy ga s fiel d o f Wes t Siberia . Fo r th e calculations, the underlying model of the Urengoy field (includin g th e thicknes s o f th e Poku r Formation an d pressur e an d temperatur e data ) was taken from Cramer et al. (1999). It wa s assume d tha t methan e thermall y generated fro m th e terrestria l organi c matte r of the Poku r Formatio n contribute d t o th e Wes t Siberian ga s field s (Galimo v 1988) . T o compar e this ga s generatio n wit h ga s releas e fro m groundwater, th e rat e o f therma l methan e generation (Fig . 3 ) wa s calculate d applyin g a specific se t of reaction kinetic data for th e Poku r Formation (Crame r et al. 1998) . Th e continuous line in Fig. 3 indicates the generation rates (up to 3.4m 3 m 2 Ma ' ) fo r the thermal histor y of the Pokur Formation beneath the Urengoy field given

Fig. 3 . Rates of thermal methan e generation , ga s release from groundwater an d diffusional losse s for the Urengo y gas field , Wes t Siberia . Methan e rate s ar e normalise d t o cubi c metr e o f methan e (STP ) pe r squar e metr e an d million years . (Fo r furthe r explanatio n se e text.)

GAS RELEASE FROM GROUNDWATE R

by Littk e e t al (1999) . Th e dotte d lin e displays the maximu m generatio n rate s (u p t o 33.5m 3 m~ 2 Ma" 1 ) for a 2KMa~ 1 heatin g rate. The rate s fo r ga s releas e fro m groundwate r are taken fro m the detaile d calculation s of Crame r et al . (1999) , wit h u p t o 3.6m 3 irT 2 Ma~ 1 methane released as a result of the hydrodynamic effect an d up to 33.3 m3 m~2 Ma-1 a s a result of the hydrostatic effect . Diffusional losse s o f ga s fro m th e Urengo y field (Fig . 3 ) wer e calculate d base d o n experi mental measurement s o f methan e diffusio n under i n situ conditions through a sample o f the cap rock, the Kuznetsov Formation. Th e methane diffusion throug h a water-saturated rock plug of about 1 cm thicknes s and abou t 2.8 cm diameter was measured in a triaxial flow cell at 35 °C and lOMPa pore fluid pressure. Experimenta l detail s have bee n give n b y Schlome r (1998) . Th e effective diffusio n coefficien t fo r th e sampl e under investigatio n wa s determine d t o b e 2.5 X 10~10 m s" 1. Th e cumulativ e amoun t o f methane los s fro m th e reservoi r an d th e steady state diffusion rate s were calculated according to the relationship for diffusiv e transpor t through a plane shee t (Kroos s e t al . 1992a , 1992b) . Considering th e effectiv e diffusio n coefficien t and th e bulk-roc k methan e concentratio n under the relevant subsurface conditions (0.036 kg CH4 m~ 3 rock) th e highest steady-stat e diffusiv e los s rate through the 8 0 m rock sequenc e amount s to c. 5.0 m3 m~2 Ma~! . Taking into account a larger thickness of the overlying cap rock (including the overlying 40 0 m thick, fine-graine d rock s o f th e Berezov Formation ) th e compute d rate s o f diffusive losse s fro m th e ga s reservoi r decreas e to 0.8m 3 m" 2 Ma~ 1 . It shoul d b e mentione d tha t al l methane rate s presented i n Fig . 3 ar e relate d t o 1 m2 o f th e reservoir area . Althoug h diffusional losse s fro m the accumulatio n ar e restricte d t o th e reservoi r area, th e processe s o f charging , i.e . ga s releas e and thermal generation, have to be related to the effective recharg e area . I n th e cas e o f th e Urengoy anticline, this recharge area is up to 4.5 times large r tha n th e recen t are a o f th e ga s reservoir (Crame r et al. 1999) . This enlarges the effective rates of the charging processes also by a factor o f 4.5. From th e compariso n o f calculate d methan e rates i t i s obviou s tha t ga s releas e fro m groundwater wa s rapi d enoug h t o charg e th e Urengoy field. The rate of diffusional losse s is at least one order of magnitude smaller than the gas release caused by the hydrostatic effect o f uplift . In contrast , th e hydrodynami c effec t o f ga s release a t th e Urengo y fiel d alon e ha s simila r rates t o th e diffusio n loss . Also , methan e

453

generation rate s fro m th e Poku r Formatio n (Fig. 3 , continuou s line ) i n thi s cas e wer e no t sufficient t o kee p u p wit h diffusio n los s an d t o charge th e Urengo y field . Eve n i f diffusiona l losses ar e neglected , th e effectiv e methan e generation rat e relate d t o th e recharg e are a cannot accoun t fo r th e hug e ga s accumulations (Schaefer e t al 1999) . These findings also argue for ga s releas e a s th e effectiv e proces s o f ga s accumulation in the Urengoy field. Conclusions During subsidenc e o f sedimentar y basin s ga s solubility i n wate r increase s wit h increasin g burial an d thermall y generate d methan e i s continuously dissolve d in the por e water . The estimated amoun t o f methan e dissolve d i n th e Earth's groundwate r by fa r exceed s th e amount of conventional gas reserves. Subsurface water is a vas t tra p fo r long-ter m storag e o f ga s i n solution, because dissolved ga s is excluded fro m rapid, buoyancy-drive n migratio n processes . Uplift movement s o f basi n setting s ca n initiat e the releas e o f dissolve d ga s a s a resul t o f th e associated dro p i n pressur e an d temperatur e (hydrostatic effect ) an d change s i n groundwater flow (hydrodynami c effect) . Bot h processe s ar e shown t o be appropriat e to release sufficien t ga s to charge ga s fields. For the huge Urengo y field of Wes t Siberia , i t i s show n that hydrostatic gas release i s th e dominan t ga s chargin g process , overwhelming diffusiona l los s throug h th e ca p rock b y a t least on e order of magnitude. In general , th e proces s o f ga s releas e fro m groundwater in exhumed basins is believed to be of globa l importance . Ga s field s i n basin s tha t experienced uplif t i n th e recen t geologica l pas t should b e re-evaluate d with regard t o th e effec t of ga s releas e fro m groundwater. References AMIRIJAFARI, B . & CAMPBELL , J.M . 1972 . Solubilit y of gaseou s hydrocarbo n mixture s in water. Society of Petroleum Engineers Journal, (Feb.) , 21-27. BARKAN, E.S. & YAKUTSENI, V.R 1981 . Perspective of the ga s potentia l i n grea t dept h (i n Russian) . Sovjetskaya Geologiya, 4, 6-15. BARKAN, E.S. , TIKHOMIROV , V.V. , LEBEDEV , B.A . & ASTAF'EV, V. R 1984. New data on the prospectivit y of natural gas dissolved i n brines a t great dept h (i n Russian). Sovetskaya Geologiya, 2, 11-20 . BARTA, L . & BRADLEY , DJ . 1985 . Extensio n o f th e specific interactio n mode l t o includ e ga s solubi lities i n hig h temperatur e brines . Geochimica e t Cosmochimica Acta, 49, 195-203 . BARTHEL, F. , REMPEL , H. , HILLER , K . & 1 2 OTHERS 1999 . Reserven, Ressourcen und

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Verfugbarkeit vo n Energierohstoffen. Bundesmi nisterium fii r Wirtschaf t un d Technologie , BMWi-Dokumentation 465 , 62 . BATTING, R . 1984 . Th e solubilit y of methan e i n water between 29 8 and 627 K at a total pressure betwee n 0.5 and 200 MPa. In: CLEVER , H.L . & YOUNG, C.L . (eds) Methane. Pergamon , Oxford , 24-44. BONHAM, L.C . 1978 . Solubilit y of methane i n water at elevated temperature s an d pressures . AAPG Bulletin, 62 (12), 2478-2488. Coco, L.T . & JOHNSON , A.E . 1981 . A correlatio n o f published dat a o n th e solubilit y o f methan e i n H 2 O-NaCl solution . In : BEBOUT , D.G . & BACHMAN, A.L . (eds ) Geopressured-Geothermal Energy Conference. Proceedings of the 5th US Gulf Coast Conference, Baton Rouge. Energy Program s Office, Bato n Rouge , LA, 215-220. CRAMER, B. , KROOSS , B.M . & LITTKE , R . 1998 . Modelling isotop e fractionatio n durin g primar y cracking of natural gas: a reaction kinetic approach . Chemical Geology, 149 , 235-250 . CRAMER, B. , POELCHAU , H.S., GERLING , P. , LOPATIN , N.V. & LITTKE , R . 1999 . Methan e releas e fro m groundwater—the sourc e o f natura l ga s accumu lations i n norther n Wes t Siberia . Marine an d Petroleum Geology, 16 , 225-244. CRAMER, S.D . 1980 . Th e solubility o f methane, carbon dioxide, an d oxygen i n brines from 0 ° to 300°C. US Bureau o f Mines Repor t o f Investigation 8706. CROVETTO, R. , FERNANDEZ-PRIMI , R . & JAPAS , M.L . 1984. The solubility of ethane in water up to 473 K. Berichte der Bunsengesellschaft fiir Physikalische Chemie, 88, 484-488. CULBERSON, O.L . & MCKETTA , J.JJ . 1950 . Phas e equilibria i n hydrocarbon-water systems—II—the solubility o f ethan e i n wate r a t pressure s t o 100.000 psi. Petroleum Transactions, American Institute of Mining, Metallurgical and Petroleum Engineers, 189, 319-322 . CULBERSON, O.L . & MCKETTA , J.JJ . 1951 . Phas e equilibria i n hydrocarbon-wate r systems—III — the solubilit y o f methan e i n wate r a t pressure s t o 10,000 PSIA. Petroleum Transactions, American Institute of Mining, Metallurgical and Petroleum Engineers, 192, 223-226 . DODSON, C.R . & STANDING , M.B . 1945 . Pressure volume-temperature an d solubilit y relation s fo r natural-gas-water mixtures . Drilling an d Production Practice, 1944, 173-179 . DORE, A.G . & JENSEN , L.N . 1996 . Th e impac t o f lat e Cenozoic uplif t an d erosio n o n hydrocarbo n exploration: offshor e Norwa y an d som e othe r uplifted basins . Global an d Planetary Change, 12 , 415-436. DUAN, Z. , MOLLER , N. , GREENBERG , J . & WEARE ,

J.H. 1992 . Th e predictio n o f methan e solubilit y in natural water s t o hig h ioni c strengt h fro m 0 t o 250 °C an d fro m 0 t o 160 0 bar. Geochimica e t Cosmochimica Acta, 56, 1451-1460 . GALIMOV, E.M . 1988 . Source s an d mechanism s o f formation o f gaseous hydrocarbon s i n sedimentary rocks. Chemical Geology, 71 , 77-95. HAAS, J.L., JR 1978. An empirical equation with tables of smoothed solubilities of methane in water and

aqueous sodium chloride solutions up to 25 weight percent, 36 0 °C, an d 138 MPa. U S Geologica l Survey Ope n Fil e Repor t 78-1004 . KORTSENSHTEJN, V.N . 1979 . A n estimat e o f globa l reserves o f ga s i n th e subsurfac e hydrosphere . Doklady Akademii Nauk SSSR, 235 , 223-224 . KROOSS, B.M. , LEYTHAEUSER , D . & SCHAFER , R.G . \992a. The quantification o f diffusive hydrocarbo n losses through cap rocks of natural gas reservoirs— a re-evaluation. AAPG Bulletin, 76, 403-406. KROOSS, B.M. , LEYTHAEUSER , D. & SCHAFER , R.G . 19926. The quantificatio n o f diffusive hydrocarbo n losses through cap rocks of natural gas reservoirs— a re-evaluation : reply . AAPG Bulletin, 76 , 1842-1846. Kuo, L.-C . 1997 . Ga s exsolutio n durin g flui d migration an d it s relatio n t o overpressur e an d petroleum accumulation . Marine an d Petroleum Geology, 14(3) , 221-229. KUUSKRAA, V.A . & MEYERS , R.F . 1983 . Revie w o f world resource s o f unconventiona l gas. Th e Fifth IIASA Conference o n Energy Resources. Inter national Institut e fo r Applie d Syste m Analysis , Laxenberg, Austria , 409-458. LITTKE, R. , CRAMER , B. , GERLING , P. , LOPATIN, N.V., POELCHAU, H.S., SCHAEFER , R.G . & WELTE , D.H. 1999. Gas generation and accumulation in the West Siberian Basin . AAPG Bulletin, 8 3 (10) , 1642-1665. MARSDEN, S . 1993 . A surve y of natura l gas dissolve d in brine . In : HOWELL , D.G . (ed. ) Th e Future o f Energy Gases. US Geological Survey , Professional Papers, 1570,471-492 . MAXIMOV, S.P. , ZOLOTOV , A.N . & LODZHEVS KAYA, M.I . 1984 . Tectoni c condition s fo r oi l and ga s generatio n an d distributio n o n ancien t platforms. Journal o f Petroleum Geology, 1 (3), 329-340 . O'SULLIVAN, T.D. & SMITH , N.O. 1970 . The solubility and partial molar volume of nitrogen and methane in wate r an d aqueou s sodiu m chlorid e fro m 5 0 t o 125° an d 10 0 t o 60 0 Atm. Journal o f Physical Chemistry, 7 4 (7) , 1460-1466 . PRICE, L.C . 1979 . Aqueou s solubilit y of methan e a t elevated pressure s an d temperatures . AAPG Bulletin, 63, 1527-1533 . PRICE, L.C . 1997 . Origins , characteristics , evidenc e for, an d economi c viabilitie s of conventiona l and unconventional ga s resourc e bases . Geologic Controls of Deep Natural Gas Resources in the United States. U S Geologica l Surve y Bulletin , 2146, 181-207 . PRICE, L.C. , BLOUNT , C.W. , MACGOWAN , D . & WENGER, L . 1981 . Methan e solubilit y i n brine s with applicatio n t o th e geopressur e resource . In : BEBOUT, D.G . & BACHMAN , A.L . (eds ) Geopressured-Geothermal Energy Conference. Proceedings of the 5th US Gulf Coast Conference, Baton Rouge. Energ y Program s Office , Bato n Rouge, LA, 205-214. RETTICH, T.R. , HANDA , Y.P. , BATTING , R . & EMMERICH, W. 1981 . Solubility of gases in liquids. 13. High-precisio n determinatio n o f Henry' s constant fo r methan e an d ethan e i n liqui d water

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at 275 to 328 K. Journal of Physical Chemistry, 85, SCHLOMER , S . 1998 . Sealing Efficiency o f Pelitic 3230-3237. Rocks —Experimental Characterisation an d GeoSCHAEFER, R.G. , GALUSHKIN , Y. , KOLLOFF , A . & logical Relevance (i n German) . Bericht e de s LITTKE, R . 1999 . Reactio n kinetic s o f ga s Forschungszentrum s Jiilich GmbH 3596, 204. generation i n selecte d sourc e rock s o f th e Wes t SULTANOV , R.G. , SKRIPKA , V.G . & NAMIOT , A.Y . Siberian Basin : implication s fo r th e mas s balanc e 1972 . Solubility o f methan e i n wate r a t elevate d of early-thermogenic methane. Chemical Geology, temperature s and pressures (in Russian). Gazovaya 156, 41-65. Promyshlennost', 17 , 6-7 .

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Depressurization of hydrocarbon-bearing reservoirs in exhumed basin settings: evidence from Atlantic margin and borderland basins D. V. CORCORAN1 & A. G. DORE 2 l

Statoil Exploration (Ireland) Ltd, Statoil House, 6, George's Dock, IFSC, Dublin, Ireland (e-mail: DVC@ statoil.com) 2

Statoil (UK) Ltd, lla Regent Street, London SW1Y 4ST, UK

Abstract: Depressurizatio n o f reservoirs i n petroliferous basins commonly occurs through cap-rocks a t structural crests where pore pressures are locally elevated because o f either the presence of a hydrocarbon column or the redistribution of overpressures by water flow along laterally extensiv e incline d permeabl e aquifers . I n exhume d petroliferou s basin s thi s deflation o f excess por e pressure s i s enhanced b y th e denudatio n process, whic h results in the large-scal e remova l o f overburden during regional uplift . Evidenc e fro m th e exhume d basins of the Atlantic margin indicates that hydrocarbon accumulations in these basins are commonly characterize d b y underfille d trap s an d hydrostaticall y pressure d o r modestl y overpressured reservoirs . Thes e observation s ar e reviewe d i n th e contex t o f th e generi c mechanisms by which top-seals leak, the properties o f cap-rocks and the physical processes that occu r during exhumation. Water-wet shal y cap-rock s ca n for m a capillar y sea l t o a hydrocarbo n colum n whil e simultaneously accommodatin g brin e flo w an d equilibratio n o f pressure s betwee n th e reservoir an d the top-seal. I n contrast, thick, low-permeabilit y shal e o r evaporite sequence s may for m pressur e seal s tha t restrict vertica l brin e an d hydrocarbo n flow and preven t the equilibration o f aquifer pressures abov e an d below th e seal . In any sedimentary basin , the presence o f regiona l pressur e seal s ca n resul t i n a layere d hydrogeologica l regim e wit h hydrostatically pressure d strat a decouple d fro m over - o r underpressure d cells . Recentl y exhumed basin s typicall y sho w limite d overpressurin g an d i n a numbe r o f thes e basin s underpressured reservoirs have been described. Post-exhumation overpressure generation is primarily drive n by tectonic compression, aquathermal pressuring an d hydraulic head . The flui d retentio n capacit y o f an y cap-roc k litholog y durin g exhumatio n is dependen t upon the physical and mechanical characteristics o f the cap-rock a t the time of exhumation and th e timin g an d condition s o f th e associate d deformatio n relativ e t o th e timin g o f hydrocarbon emplacement . Th e permeabilit y an d deformationa l characteristic s o f halit e render i t a n excellen t cap-roc k wit h a hig h retentio n capacity , eve n unde r condition s o f exhumation. However , mudrock s ma y als o for m effectiv e cap-rock s i n exhume d basin s when th e deformatio n associate d wit h exhumatio n occur s befor e embrittlemen t an d th e shale cap-roc k exhibit s ductile behaviour. Shale and evaporite cap-rocks form the main regional seals to hydrocarbon accumulations in exhume d basin s o f th e Atlanti c margi n an d borderlands . Syn-exhumation top-sea l efficiency (flui d retention capacity) is a major exploration risk in these basins, although postexhumation top-seal integrity in these basins may be relatively high under certain conditions . Consequently, a majo r exploratio n ris k facto r i n exhume d basi n setting s pertain s t o th e limited hydrocarbo n budge t availabl e post-regiona l uplif t an d th e efficienc y o f th e remigration process .

Abnormal por e pressure s ar e commonl y generatio n o f th e abnorma l por e pressur e observed i n sedimentary basins . These abnorma l condition s an d th e hydrauli c evolutio n o f th e pore pressure s ar e denote d a s overpressures basi n durin g an d afte r th e creatio n o f thes e where flui d pressure s ar e i n exces s o f th e abnorma l pore pressure conditions . I n petroliferhydrostatic gradien t a t a specifi c dept h an d ar e ou s basins , por e pressur e evolutio n ha s a n denoted a s underpressures wher e th e flui d importan t bearin g o n th e generation , migratio n pressure i s les s tha n hydrostati c (Fig . 1 ) an d retention of petroleum in hydrocarbon traps. (Martinsen 1994) . Present-da y por e pressur e Fo r example , differentia l pressure s betwee n th e distributions ar e a function o f the mechanis m o f hydrocarbo n sourc e roc k an d carrie r be d ar e From: DORE , A.G., CARTWRIGHT, J.A. , STOKER, M.S. , TURNER, J.P. & WHITE , N . 2002 . Exhumation o f th e North Atlantic Margin: Timing, Mechanisms an d Implications for Petroleum Exploration. Geologica l Society, London, Special Publications, 196, 457^83. 0305-8719/02/$15.00 © The Geological Society of London 2002.

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Fig. 1. Abnormal formation pressures; with pore pressure plotted v. depth. For any depth, formation pressures that are less than the hydrostatic pressure are termed underpressures and formation pressure s in excess of hydrostatic are termed overpressures . recognized a s th e drivin g forc e o f primar y migration (Magar a 1968 ; Palciauska s & Dome nico 1980) . Th e rol e o f overpressures i n th e secondary migratio n an d remigratio n o f hydro carbons ha s bee n highlighte d b y Gile s e t al . (1999), Iliff e e t al (1999 ) an d others. Tempora l changes i n hydrauli c gradients , carrie r be d interconnectivity, cap-rock an d fault sea l properties, an d basi n structur e influenc e migratio n o f oil and gas and the rates o f ingress an d egress of petroleum fro m th e hydrocarbon trap . Pore pressur e increase s abov e th e hydrostatic level ca n b e maintaine d onl y whe n th e por e fluids, both hydrocarbons an d brine, ar e retained in th e reservoi r b y cap-rock s wit h sufficientl y low permeability . Bradle y & Powle y (1994) , i n their revie w o f pressur e compartment s i n sedimentary basins , hav e differentiate d open hydraulic systems, whic h ar e characterize d b y fluid (brine , oi l an d gas ) continuit y throughout , hydrostatic pressure s i n the aquife r and possibl e brine flo w acros s a capillar y seal , fro m closed hydraulic systems, whic h manifes t n o flui d continuity acros s th e boundin g seals , ma y b e under-, over- o r normally pressured an d have no fluid flow across the pressure seal. Changes in the pore volum e an d flui d volum e o f an y hydrauli c system resul t fro m changes i n temperature , effective stres s an d por e pressure , whic h ar e

induced b y uplif t an d denudatio n o r subsidenc e and deposition . A close d hydrauli c syste m wil l respond t o these 'external ' force s by a change in fluid pressure wherea s an open hydrauli c syste m will respond by fluid flow in or out of the system. In the context of petroliferous basins, top-seals can b e define d a s rocks that prevent th e vertica l migration o f hydrocarbon s ou t o f traps . Unde r conditions of a closed hydraulic system top-seals may also act as a barrier t o brine flow (a pressure seal) and prevent the dissipation of elevated pore pressures in the reservoir. During the evolution of a petroliferous basin any lithology can serv e as a top-seal fo r a hydrocarbo n accumulatio n pro vided that its capillary entry pressure exceeds the upwards buoyanc y pressur e exerte d b y th e hydrocarbon colum n i n th e underlyin g accumulation. In practice, however , the vas t majority of effective sea l rock s ar e evaporites , fine-graine d clastic rock s an d organic-ric h mudrock s (Downey 1994) . The basi c physica l principle s governin g th e effectiveness o f petroleu m cap-rock s an d press ure seal s ar e wel l establishe d (Hubber t 1953 ; Berg 1975 ; Schowalte r 1979 ; Watt s 1987 ; Bradley & Powley 1994) . Lithology , uniformity of stratigraph y an d thicknes s ar e factor s tha t influence sea l capacit y (Downe y 1984) . How ever, the fundamental roc k properties that control

DEPRESSURIZATION OF HYDROCARBON-BEARING RESERVOIR S

seal performance ar e the capillary entry pressure of the seal (dominantl y controlled by pore-throat diameter) an d th e ductilit y o f th e sea l rock , which is a function o f pressure, temperature an d lithology. Althoug h top-sea l integrit y i s recog nized a s a majo r exploratio n ris k facto r i n exhumed basins , fe w systemati c studie s o f top seal performanc e i n thes e setting s hav e bee n published t o dat e (Gabrielse n & K10vja n 1997 ; Seedhouse & Racey 1997; Spain & Conrad 1997 ; Cowan e t al 1999) . In all petroliferous basins the adequac y of the hydrocarbon charg e together wit h the timing and rate of fill, spill and vertical or lateral leakage are key determinant s of the present-day in-plac e oil and gas volumes preserved in hydrocarbon traps. However, i n exhume d basins , th e interpla y between top-sea l performanc e an d hydrocarbon fill, spil l an d leakag e i s mor e critica l a s th e 'switching off' o f hydrocarbon generation during regional uplift ma y result i n a lower probabilit y of tra p replenishment post-exhumatio n (Dore & Jensen 1996) . Physical processe s that may affec t cap-rocks durin g exhumatio n includ e erosion , tectonic deformation, shea r failure, hydrofracturing a s a resul t o f disequilibriu m por e pressur e conditions and a changing hydrodynamic regime. Furthermore, ther e i s a n increase d ris k o f ne t hydrocarbon losse s becaus e o f diffusio n wher e gas accumulation s ar e dependen t upo n porou s and permeable shale seals . The primar y focu s o f thi s pape r i s t o offe r some insight s wit h respec t t o th e physica l properties an d processe s tha t affec t top-sea l performance i n exhume d basi n settings . Th e implications of regional uplif t for abnormal por e pressures (overpressure s an d underpressures) ar e also reviewed . Empirica l evidenc e fro m exhumed basin s o f th e Atlanti c margi n an d borderland basins is then discussed in the context of thes e observations .

Top-seal leakage mechanisms: a summary There ar e fou r generi c mechanism s b y whic h top-seals leak : tectoni c breaching , capillar y leakage, hydrauli c leakage an d molecular transport (diffusion) . Figur e 2 i s a summar y o f th e physical principles governing these mechanisms, which hav e bee n articulate d b y numerou s workers (e.g . Schowalte r 1979 ; Gretene r 1981 ; Krooss e t a l 1992 ; Davi s & Reynold s 1996) . These fou r generic leakage mechanisms ar e here reviewed i n th e contex t o f exhume d basi n settings of the Atlanti c margin.

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Tectonic breaching Where deformatio n o f a cap-roc k occur s post emplacement o f hydrocarbon s ther e i s a n increased ris k o f tectoni c breachin g an d cap rock leakage . Top-sea l failur e vi a tectoni c breaching i s th e mos t readily recognize d for m of sea l failure at the scal e o f seismically define d hydrocarbon traps . I n addition , tectoni c breach ing ma y be facilitated by sub-seismi c resolution faulting wher e th e top-sea l consist s o f mudrock layers interbedded wit h permeable siltstone s and sandstones. Tectonic framework, basin evolution and fault reactivatio n are also important controls on fluid flow and distribution of overpressures in sedimentary basins . For example , in the Nort h Sea Centra l Graben , pre-Cretaceou s structur e controls the flow of pore fluids below the regional Jurassic-Cretaceous pressure seal , wit h vertica l fluid escape, throug h this pressure seal , focuse d on axial faul t bloc k crests (Darby et al. 1998) . The styl e an d magnitud e o f tectoni c defor mation in any sedimentary basin is influenced by a numbe r o f factors , includin g plat e tectoni c setting, pre-existin g structura l grai n an d th e presence o r absenc e o f detachmen t layers . During basi n inversion , compressional , trans pressional or reactivated extensional deformation may resul t i n leakag e throug h cross-faul t juxtaposition o f reservoirs fro m differen t strati graphic level s (Fig . 2a(i) ) o r throug h th e development o f a connected networ k o f juxtaposed leak y bed s withi n th e cap-roc k interva l (Fig. 2a(ii)) . I n addition , radia l extensio n fractures wil l develo p abov e th e neutra l surfac e of inversion folds an d may result in leakage into the overlying sediment s (Fig. 2a(iii)) . Hall e t al. (1997) have shown from cas e studie s in the deep Central Grabe n o f th e Nort h Se a tha t reservoi r objectives lyin g abov e o r clos e t o th e neutra l surface o f a n inversio n fol d hav e a highe r probability o f bein g breache d tha n reservoir s lying below. Tectonic breachin g is an importan t leakag e mechanism in the exhumed basin settings of the Atlantic margi n wher e syn-exhumatio n exten sional fault reactivations are probable in addition to th e overprin t o f compressiona l deformatio n resulting fro m th e far-fiel d signatur e o f Alpin e orogenesis an d ridge-pus h phenomen a (Under bill 1991; Murdochs al. 1995; Dore et al. 1999) . Capillary leakage The driving force for petroleum movement in the subsurface i s buoyanc y influence d b y over pressure and hydrodynamics. Th e force opposin g the movement o f petroleu m i s th e capillar y

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CORCORAN & A.G. DOR E

Fig. 2 . The fou r generic mechanism s b y whic h top-seal s leak , (a ) Tectoni c breaching ; ther e ar e three commo n modes: (i ) fault offse t lea k path; (ii ) fault-linked leak path; (iii ) dilatant fracture lea k path, (b ) Capillar y leakag e (for capillary leakage t o occur th e buoyancy force generated b y the hydrocarbon column (Pbuoy). plus any excess overpressure (AC/ ) in the reservoir relative to the seal, must exceed the capillary resistanc e o f the porous, water wet, cap-rock s (P cap)). (c ) Hydrauli c leakage ; thi s ca n resul t fro m th e developmen t o f tensio n fracture s (hydrofractures), whic h arise fro m changin g effective stres s conditions. Tension fractures occu r under conditions

DEPRESSURIZATION OF HYDROCARBON-BEARING RESERVOIR S

of low differential stress (small cr\ — (J3), when pore-fluid pressure in the cap-rock reduces the minimum effectiv e horizontal stres s belo w zer o t o th e tensil e strengt h o f th e rock , (d ) Molecula r transport ; thi s i s primaril y th e diffusion o f methan e throug h water-saturate d shal y cap-rocks . (Adapte d fro m Kroos s e t al. 1992 ; Davis & Reynolds 1996; Hall etal 1997 ; Ingram & Urai 1999; capillary leakage equations from Schowalter 1979; Clayton & Hay 1994. )

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resistance o f porou s rock s (Fig . 2b) . Standar d equations have been develope d t o describe thes e opposing force s Pb uoy and Pcap a t the interface of a hydrocarbo n reservoi r an d cap-roc k (Hubber t 1953; Ber g 1975 ; Schowalte r 1979 ; Watt s 1987 ; Clayton & Ha y 1994) . Theoretically , fo r capillary leakag e t o occur the upwards buoyanc y pressure o f a hydrocarbo n colum n plu s an y excess overpressur e o r hydrauli c hea d mus t exceed th e P cap o f the top-seal . However, recen t discussions i n th e literatur e sugges t that , i n th e case o f a continuou s wate r phas e betwee n a water-wet petroleu m reservoi r an d a water-we t top-seal, overpressur e i n th e aquife r wil l no t contribute t o capillar y leakag e whe n ther e i s capillary pressure-gravit y equilibriu m throughout th e hydrocarbo n colum n (Bj0rku m e l al 1998). Wher e a finit e permeabilit y t o wate r i s present a t th e junction o f a n oi l colum n wit h a top-seal the n ther e wil l b e a minut e bu t finit e water flow from th e reservoi r int o th e cap-rock , which induce s a dynami c pressur e dro p i n th e water phas e o f th e reservoir , thereb y increasin g capillary pressur e (Rodger s 1999) . I n contrast , Clayton (1999 ) argued tha t there i s no continuity in the water phase, whic h is immobile, even in a water-wet reservoir , and that overpressure i n the aquifer wil l indee d contribut e t o pushin g hydrocarbons throug h a water-wet cap-rock . Clayton & Ha y (1994 ) hav e modelle d th e capillary sea l capacit y o f a mudston e sea l i n a continuously subsidin g basin , base d o n appropriate figure s fo r interfacia l tension , contac t angle, larges t interconnecte d pore-throa t radi i and subsurfac e densit y differenc e o f ga s an d water (Fig. 3). The computed sea l capacity curve for methan e indicate s tha t th e modelle d mud stone would retain a gas column of 500-1000 m depending upo n depth . In a n exhume d basi n th e predicte d capillar y retention capacit y o f a n averag e mudston e i s likely t o b e highe r a t an y present-da y depth , a s the highe r compactio n stat e o f th e exhume d mudstone wil l resul t i n smalle r interconnecte d pore throats. Wit h respect to the exhumed basins of th e Atlanti c margin , th e magnitud e o f al l known ga s column s discovere d t o dat e i s les s than 500m , considerabl y les s tha n th e capillar y seal retentio n capacit y modelle d fo r a n averag e mudstone (Fig . 3) . Hydraulic leakage Where th e capillary entry pressures to a cap-rock (evaporite o r super-tigh t shale ) ar e s o hig h tha t capillary failur e i s implausible , hydrauli c leak age may occur through brittle top-seals as a result of th e generatio n o f ne w tensio n fracture s

(hydrofractures), shea r fracture s o r th e dilatio n of pre-existin g faul t plane s (Fig . 2c) . Hydraulic fracturing ca n occu r independen t o f tectoni c breaching an d result s fro m change s i n effectiv e stress condition s i n the cap-rock. Thes e change s may b e induce d b y th e developmen t o f disequilibrium por e pressur e condition s o r b y changes i n th e tectoni c loa d (Fig . 4) . Fo r example, a reductio n i n th e minimu m effectiv e compressive stres s (0-3) , induce d b y extensio n during regiona l uplift , ma y pe r s e resul t i n th e formation o f dilatan t shea r fracture s o f certai n orientation withi n th e cap-roc k (Fig . 4a) . Shea r fractures wil l also be formed when the prevailing stress fiel d result s i n condition s o f hig h differential stres s i n th e cap-roc k (larg e °"i ~ ~ °"3) - I* 1 tms scenari o th e gradua l elevatio n of pore-fluid pressure s before exhumation, or the removal o f overburde n withou t th e re-equili bration o f elevate d pore-flui d pressures , wil l result in Coulomb failure along planes in the rock that mak e appropriat e angle s wit h cr l (Fig . 4b) . Pre-existing fractures , joint s o r fault s (i n fact , any plane s of reduced cohesion ) have important implications fo r th e mechanica l behaviou r o f cap-rocks durin g exhumation . Whe n fracture s are present their physical characteristics an d their orientation mus t b e known , t o evaluat e thei r structural significanc e (Gretener 1981) . Hydrofractures occu r unde r conditions of low differential stres s whe n pore-fluid pressure a t the cap-rock-reservoir interfac e reduce s th e mini mum effective horizontal stres s below zer o to the tensile strengt h o f th e roc k (Fig . 2c) . I n extensional basins , wher e th e minimu m com pressive stres s (cr 3) i s significantl y les s tha n the maximum compressiv e stres s (crO , thes e hydro fractures ar e invariabl y vertical t o semi-vertical in orientatio n an d ar e perpendicula r t o th e minimum horizonta l stres s (cr 3). Fo r hydrofractures t o develo p i n preferenc e t o shea r fracture s the condition s P? = 0-3 + T an d <j\ — cr3 < 4 T must be satisfie d (P f i s the pore-fluid pressure, cr 3 is th e minimu m horizonta l stres s an d T i s th e tensile strengt h o f th e cap-rock ) (Hubber t & Rubey 1959 ; Seco r 1965 ; Sibso n 1995) . Thes e conditions ca n occu r i n highl y overpressure d systems undergoin g continuou s subsidenc e o r during exhumation , whe n rapi d denudation , without re-equilibration of overpressures, results in tensil e failure . I n eithe r case , whe n thi s stat e prevails, pervasive tension fractures ma y develop in th e cap-rock , whic h wil l resul t i n a catastrophic loss of any pre-existing hydrocarbon fill and the potential re-equilibratio n or reduction of overpressures. I t has been argue d by Bj0rkum et al. (1998) that, in water-wet reservoir an d caprock systems , the buoyancy force (overpressure)

DEPRESSURIZATION OF HYDROCARBON-BEARING RESERVOIR S

Fig. 3 . Modelle d capillar y sea l capacit y wit h dept h fo r a mudston e i n a continuousl y subsidin g basi n (fro m Clayton & Hay 1994). The inferred equivalent curve for an exhumed basin settin g is schematically show n (dashe d line) togethe r wit h th e magnitud e o f th e maximu m hydrocarbo n column s i n ga s accumulation s fro m som e exhumed Atlanti c margin an d borderland basins.

exerted b y th e hydrocarbo n colum n relativ e t o the wate r phas e is balance d by the downwarddirected interfacia l tensio n forc e a t th e flui d interface an d that the presence o f a hydrocarbon column wil l no t increas e th e ris k o f hydro fracturing th e top-sea l t o th e overpressure d reservoir. Hubbert & Rube y (1959 ) demonstrate d tha t when th e pore-flui d pressur e i n a sedimentar y basin approaches the lithostatic pressure the fluid pressure i s released by rock failure . Palciauska s & Domenic o (1980 ) supporte d thes e obser vations by showin g theoretically tha t microfractures ca n develo p i n overpressure d sedimentar y

beds whil e undergoin g burial . Capuan o (1993 ) provided direc t petrographi c evidenc e o f th e occurrence o f microfractures in situ at depths of 3-5 km, in geopressured Oligocene shale s of the Gulf Coas t Basin . Furthermore , th e compute d fracture permeabilities (o f the order of 1 0 m 2) in thes e shale s combine d wit h parageneti c relationships indicat e tha t flui d flo w occurre d preferentially throug h thes e microfracture s rather tha n throug h th e matri x o f thes e shales . The development o f these hydrofracture s during burial i s facilitated by a mechanism of episodi c tensile failur e i n a lo w differentia l stres s environment (Sibso n 1995) . Unde r thes e

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Fig. 4 . Mohr circle representatio n o f the development of dilatant shear fractures durin g exhumation: (a) induced by extension (reduction of minimum compressive stress , 03) during regional uplift, Coulom b failure in shear will occur along planes that make a n angle B with the orientation of o~i ; (b) under conditions of high differential stres s (large (TI — cr3) th e remova l o f overburde n durin g exhumation , withou t th e re-equilibratio n o f pore-flui d pressures, wil l resul t i n th e residua l overburde n loa d bein g disproportionall y carrie d b y th e pore-flui d pressur e CPf), a lowering of the effective stres s levels to cr/ and cr 3', and Coulomb failure in shear along planes that make an angle 9 with the o- j axis .

DEPRESSURIZATION O F HYDROCARBON-BEARING RESERVOIR S

conditions, th e raising o f the pore-fluid pressur e will result i n the crackin g o f the rock i n tension and th e releas e o f th e flui d pressure , followe d by a raising o f fluid pressure an d a repetition o f this cycl e (Fig . 2c) . In contrast t o this pressure cyclicity , which is manifest i n a continuousl y subsidin g basin , depressurization o f a reservoi r durin g exhuma tion coul d potentiall y occu r a s a 'singular ' catastrophic event. This may arise because of the reduced potentia l fo r build-u p o f overpressures , as som e o f th e processe s tha t produc e over pressure (Osborn e & Swarbric k 1997) , suc h a s disequilibrium compaction , dehydratio n reac tions and kerogen transformation, will have been arrested onc e exhumatio n begins . Onc e cata strophic failur e vi a fracturin g ha s occurre d during exhumation , th e cap-roc k uni t ca n onl y regain 'sea l status' whe n th e high-permeabilit y open fracture s ar e heale d o r annealed . Fractur e closure, durin g o r post-exhumation , ca n occu r through a rang e o f mechanisms , includin g cementation an d increase d horizonta l compres sive stress . Cementatio n i n th e fractur e ma y b e caused b y th e coolin g o f upward-flowin g fluid s with the resultin g redistributio n of silic a and other minera l phase s or , i n th e absenc e o f fluid flow , th e chemica l diffusio n o f solid s into th e fractur e drive n b y loca l thermo dynamic potentials (Pedersen & Bj0rlykke 1994). Molecular transport (diffusion) Diffusion i s a continuous and ubiquitous process in sedimentary basin s and its role in hydrocarbon migration ha s been analyse d by severa l worker s (Fig. 2d) (Krooss et al 1992 ; Montel et al 1993 ; Schlomer & Kroos s 1997) . Migratio n o f hydrocarbons vi a diffusio n transpor t can be important under geologica l condition s tha t prohibi t separ ate phas e flo w o f hydrocarbon s (Man n e t al . 1997). Additiona l prerequisite s includ e th e presence o f hydrocarbo n component s tha t hav e a high wate r solubility and a high concentration of thes e component s i n a specifi c area . Ga s exsolution fro m groun d water i s potentiall y a n important mechanis m fo r th e formatio n o f ga s accumulations i n uplifte d basin s (Crame r & Poelchau 2002) . The diffusiv e transpor t mechanis m primaril y pertains t o th e dismigratio n o f natura l ga s accumulations i n certai n circumstance s an d ha s little relevanc e fo r oi l dismigratio n becaus e o f the increase d siz e o f oi l molecule s relativ e t o shale pore-throa t dimensions . Molecula r diffu sion ca n occu r independen t o f th e pressur e conditions i n th e reservoi r an d cap-rock . However, pressur e i s a n importan t contro l o n

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the solubilit y o f methan e i n formatio n water s (Cramer & Poelcha u 2002) . Althoug h th e modelling studie s o f Kette l (1997 ) indicate d that methane diffusion constant s for rock salt are non-zero ove r geologica l time scales , diffusio n losses fro m ga s accumulation s cappe d b y thic k evaporitic seal s ar e considered minimal . Empirical observatio n o f long-live d ga s accumulations in Uppe r Proterozoi c reservoir s seale d beneat h Lower Cambrian sal t in the extensively exhumed Lena-Tunguska provinc e o f th e forme r Sovie t Union support s thi s conclusio n (Kontorovitc h et al. 1990) . In contrast , Letythaeuse r e t al . (1982 ) hav e demonstrated tha t ga s ma y diffus e throug h water-saturated cap-rock s ove r geologica l tim e scales. Thi s diffusio n mode l suggest s tha t th e evolution an d preservatio n o f natura l ga s accumulations i s dependen t upo n th e rati o o f gas supply to the trap and gas losses through the cap-rock. Fo r example , a cas e stud y o f th e Harlingen ga s field , offshor e Netherlands , indicated tha t hal f th e 68BC F (billio n cubi c feet) contained in the Lower Cretaceous reservoi r would be lost by diffusion throug h the shales and marls o f the Hauterivia n cap-rock (39 0 m thick) in 4.5 Ma (Letythaeuser et al. 1982) . Subsequen t re-evaluation of these estimate s b y Kroos s e t al. (1992) suggeste d tha t th e rat e o f diffusiv e hydrocarbon losse s throug h th e cap-roc k a t Harlingen ar e a n orde r o f magnitud e lowe r (c. 7 0 Ma t o dissipat e hal f th e in-plac e ga s vi a diffusion throug h the cap-rock) . In a petroliferous basin that is characterized by continuous subsidence , hydrocarbo n escap e b y diffusion an d othe r processe s ca n b e wholl y o r partly offse t b y a n activ e generatio n an d migration system . However , i n a n exhume d basin setting , where the hydrocarbon generation and migratio n syste m i s 'switche d off ' durin g regional uplift , diffusiv e losse s throug h water saturated shaly cap-rocks will increase with time since uplif t an d may be significant . Cap-rocks: some physical properties The examinatio n o f cap-rock s t o hydrocarbo n accumulations i s primaril y concerne d wit h th e properties o f th e weakes t poin t o f th e reservoir-top-seal interface . A s highlighte d b y Downey (1994) , measure d propertie s o f a random cor e sampl e ma y no t be relevan t t o th e physical propertie s o f th e cap-roc k a t th e lea k point. Furthermore , geohistor y i s a n importan t control o n th e sealin g propertie s o f top-seal s (Knipe e t al . 2000 ) an d th e location o f th e potential lea k poin t throughou t the evolutio n of the hydrocarbon trap. Water-wet shaly cap-rocks

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can for m a capillar y sea l t o a hydrocarbo n column whil e simultaneousl y accommodatin g brine flow and equilibration of pressures between the reservoi r an d th e top-seal . I n contrast , thick, low-permeabilit y shal e o r evaporit e sequences ma y for m pressur e seal s tha t restric t vertical brin e an d hydrocarbo n flo w an d prevent th e equilibratio n o f aquife r pressure s above an d belo w th e seal . Som e o f th e petrophysical an d mechanica l propertie s tha t most influenc e top-sea l performanc e ar e summarized below .

Lithology, porosity and permeability Evaporites and mudrocks are commonly found as effective top-seals to hydrocarbon accumulation s because the y typically possess ver y low porosity and permeability, hig h capillar y entr y pressures , are relativel y ductil e an d ar e ofte n laterall y continuous a t th e basi n scale . However , othe r lithologies, suc h a s siltstone s an d sandstones , have been identified as having capillary retention capacity an d ca n for m th e top-sea l t o a hydrocarbon colum n (Spai n & Conrad 1997) . In mos t sedimentar y basins , mudston e poro sities rang e fro m 5 t o 80% , dependin g upo n compaction stat e (Sclate r & Christi e 1980) . Mudstone permeabilitie s var y b y te n order s o f magnitude (10~ 4 -10~ 1 5 mD) an d b y thre e orders o f magnitud e a t a singl e porosity , primarily a s a resul t o f grain-siz e variation s (Dewhurst e t al. 1999) . However , th e larges t interconnected pore-throa t diamete r i s th e critical facto r with respect t o the capillar y entr y pressure o f th e mudstone . I n tigh t mudrock s (permeability 10~ 9 D range) the risk of capillary failure an d Darcy flow through the matrix is low, as th e capillar y entr y pressur e commonl y exceeds th e buoyanc y forc e o f an y potentia l hydrocarbon colum n (Fig . 3) . I n thi s case , th e top-seal retentio n capacit y o f th e mudroc k i s a function o f th e ductilit y o f th e mudston e an d the potentia l fo r th e formatio n o f dilatan t fractures unde r tectonic deformation or changing pore-fluid pressur e conditions . Halite forms an excellent top-seal (capillary and pressure seal ) as a result of two characteristics: i t has a practically infinit e capillar y entr y pressur e and it flows plastically under deformation. When it form s a continuou s laye r ove r th e potentia l hydrocarbon tra p and is immediately juxtaposed above the hydrocarbon-bearing reservoir, the seal risk fo r tha t tra p i s considerabl y reduced , even where th e tra p ha s experience d post-emplace ment tectoni c deformatio n an d exhumation . Mudrocks tha t ar e exceptionall y ric h i n organi c

matter als o ca n for m excellen t capillar y an d pressure seals . Fo r example , i n th e exhume d Williston Basi n th e thi n bu t prolifi c Uppe r Devonian-Lower Mississippian Bakken Shale is an extremel y ric h sourc e roc k (wit h sourc e potential Rock Eval pyrolysis parameters of total organic carbo n (TOC ) 10% , S 2 30-7 0 mgg'1 rock an d hydroge n inde x (HI ) >600mgg~ 1 TOC), has very low porosity (<3%) and vertical permeability (10~ 2 -10~ 3 nD ) an d ha s retained residual overpressure s (>12MP a presen t day) , which were developed via oil generation a t peak maturity between 75 and 50 Ma, before exhumation (Burru s 1998) . At bot h the hydrocarbon tra p and basin scale, cap-rock unit s can manifes t vertical an d lateral heterogeneities. Interna l lithostratigraph y of th e unit ca n vary , with mud rock s ofte n interbedde d with larg e amount s o f leak y strat a suc h a s siltstones or sandstones. These lithologie s will be more prone to leakage than the mudrocks and can result i n multimoda l pore-throat diamete r distri butions withi n th e cap-roc k interva l an d th e development o f 'wast e zones' , i f th e siltstone s are located immediately above the reservoir unit. In a stud y o f hydrocarbo n seal s i n th e exhumed East Irish Sea Basin, Seedhouse & Racey (1997) utilized mercur y injectio n porosimetr y t o ident ify pore-throa t distribution s i n th e Merci a Mudstone Grou p (MMG ) sea l an d t o describ e some o f th e heterogeneitie s observe d i n thi s cap-rock interval . Thes e worker s identifie d th e presence o f halit e immediatel y abov e th e reservoir a s a ke y facto r i n th e retentio n o f hydrocarbon column s i n th e exhume d Triassi c Sherwood Sandston e Grou p reservoir . I n addition, thei r stud y foun d tha t Sherwoo d Sandstone accumulation s wit h significan t hydrocarbon columns , bu t tha t ar e directl y capped b y mudrocks , invariabl y manifeste d hydrocarbon show s withi n th e cap-roc k inter val u p t o th e leve l o f th e firs t halit e be d encountered abov e th e reservoir . Thi s suggested tha t th e buoyanc y force s exerte d by th e individua l hydrocarbo n column s wer e sufficient t o overcom e th e capillar y entr y pressure o f th e heterogeneou s MM G a t thes e locations, bu t no t th e capillar y resistanc e o f the halit e units . Strength, ductility and brittleness The mechanica l respons e o f rock s t o a n applie d stress varies under different condition s so a valid comparison o f the strengt h and ductility of rocks can b e mad e onl y i f th e condition s o f deformation ar e als o known . Ductility is a roc k property that pertains to the amount of strain that

DEPRESSURIZATION OF HYDROCARBON-BEARING RESERVOIRS

a material ca n withstand before brittle failure if it undergoes brittl e failur e a t all . Ductil e rock s respond t o a n applie d stres s b y a n initial , although limited , elasti c deformation , followe d by sustaine d plasti c deformatio n befor e failure . Brittle rocks respond to an applied stress by first shortening elasticall y an d the n failin g b y th e formation of discrete fracture s an d faults. A rock is considere d ductil e whe n it ca n accommodat e strains o f 8-10 % withou t fracturin g and brittl e when strai n i s <3 % befor e fracturin g (Fig . 5) . Rock ductilit y i s a functio n o f lithology , confining pressure, pore-flui d pressure, tempera ture, differentia l stres s an d strai n rat e (Davi s & Reynolds 1996) . Because the matrix permeability o f buried and compacted mudrock s i s extremel y low , i t i s th e

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fracture permeabilit y tha t primarily controls sea l capacity of these rocks. More relevant definitions have been offered b y Ingram & Urai (1999), who described a ductil e mudroc k a s on e tha t ca n deform withou t dilatanc y an d th e creatio n o f fracture permeability , an d a brittl e mudroc k a s one tha t dilate s an d develop s fractur e permeability. Experimental studie s hav e show n that , fo r most lithologies, bot h rock strength and ductility increase wit h risin g confinin g pressur e (Handi n et al. 1963 ; Gretene r 1981) . Thi s suggest s tha t sedimentary rocks , includin g mudstones , increase i n ductilit y durin g buria l becaus e confining pressur e (lithostatic pressure I total overburden stress) increase s wit h depth . How ever, thi s inferenc e result s fro m treatin g th e

Fig. 5 . Mechanica l propertie s o f mudrocks . Relationshi p betwee n ductilit y ( % strai n a t fracturing ) an d bul k density (g cm~3), at a constant confining pressur e of 1000kg cm"2, for a population of Neogene mudrocks from Japan. Mudrock s with densities in excess of c. 2.5 g cm~3 exhibit brittle behaviour an d fracture a t strains o f < 3% (data from Hoshin o e t al 1972) .

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Fig. 6 . Porosity-dept h behaviou r fo r a suit e o f Neogen e mudrock s fro m Japa n (afte r Niagar a 1968) . Th e organically rich , overpressured , mudrock s o f th e Teradomar i an d Nanatan i formation s for m th e cap-roc k t o hydrocarbon accumulation s in the underlying fractured volcanic reservoirs of the Miocene Shiunji Tuff Formatio n and equivalents. A general increas e i n shal e densit y with depth (normal compactio n trend) is observed dow n to

DEPRESSURIZATION OF HYDROCARBON-BEARING RESERVOIRS

compaction o f mudrock s a s onl y a linea r mechanical proces s an d ignore s th e effect s o f overpressures and of chemical compaction i n the deeper par t o f sedimentar y basin s (>2-3km , 70-100°C) (Bj0rlykke 1999) . Bulk densit y (pb ) i s on e measure d paramete r that can be used to indicate th e compaction stat e of mudrocks. However, mudrock density is also a function o f matrix mineralogy , porosity , applie d load, temperatur e an d pore-flui d pressure . Experimental result s (usin g a populatio n o f Neogene mudstones from Japan) of Hoshino et al (1972) have indicated that, at constant confinin g pressures, ductilit y actuall y decrease s wit h increasing density . Thi s support s th e vie w tha t chemical compactio n and diagenetic change s can alter th e picture , wit h respec t t o ductility , a t increased buria l depth s i n th e sedimentar y column (Fig . 5) . Shale s wit h densities les s tha n c. 2.2gcm~ 3 exhibi t ductil e behaviour ; shale s with densitie s o f c . 2.2-2.5gcm~ 3 ar e transi tional an d probabl y exhibi t a wid e rang e o f mechanical behaviour ; shale s wit h densitie s >2.5gcm~ 3 manifes t brittl e behaviou r b y fracturing a t strains of < 3%. If mudrock densit y can be used as a proxy for ductility then the onset of shale embrittlement can be estimated when the density-depth o r porosity-dept h behaviou r o f the claystone is known (Fig. 6). The collection of Neogene mudrock s presented b y Magara (1968) also indicate d th e potential rol e o f overpressur e in defining th e shale embrittlemen t threshol d fo r any give n basin . Th e presenc e o f overpressur e retards compactio n b y supportin g mor e o f th e lithostatic load (confining pressure) and reducing effective stres s within the mudrock . A s a result, the shal e embrittlemen t threshol d (p b c. 2.5 g cm~3 ) will be reached a t a deeper buria l depth than i n th e cas e o f a hydrostaticall y pressured sedimentar y column . Geotherma l gradient an d palaeotemperatur e histor y als o exert a critica l contro l o n diagenesi s an d henc e the rheological evolutio n of mudrocks. With respect to cap-rock integrity during basin inversion th e ke y factor s ar e th e timin g an d magnitude o f th e deformatio n an d th e mechan ical behaviou r (brittl e v . ductile) o f the cap-roc k at th e tim e o f deformation . Brittl e shale s ar e more likel y t o ruptur e an d lea k tha n ductil e shales and evaporites, which may exhibit plastic

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flow unde r th e applie d deformation . However , even evaporitic rocks, whic h serve a s extremel y efficient ductil e seal s whe n overburde n buria l exceeds 100 0 m, can manifest brittle behaviour at shallow depths (Downey 1994) . Bolton e t al . (1998 ) hav e show n experimen tally tha t althoug h elevate d pore-flui d pressure s reduce effectiv e stres s an d enhanc e shea r deformation, i t i s th e consolidatio n stat e a t th e onset o f shea r tha t i s th e crucia l facto r wit h respect t o th e deformatio n styl e an d resultin g permeability. Timing o f overpressure i s particularly relevant in this regard, as it can change th e consolidation stat e of a mudstone cap-rock wit h respect t o effective stress. Ingra m & Urai (1999) have indicated tha t claystone cap-rock s tha t have undergone substantia l uplif t ar e pron e t o th e formation o f dilatant fractures, as they are likely to b e over-consolidate d an d hav e anomalou s strength. However , i n th e contex t o f exhume d basins, the mechanical behaviou r of a mudstone cap-rock will be dependent upo n whether or not embrittlement ha s bee n achieve d befor e exhumation. Physical processes that occur during exhumation A number o f physical processe s may affec t cap rocks durin g exhumation , includin g erosion , tectonic deformation, shear failure, hydrofracturing a s a resul t o f disequilibriu m por e pressur e conditions and a changing hydrodynamic regime. Each of these processes mus t be examined in the context of the evolutionary change s that may be occurring i n th e cap-rock , i n th e petroleu m system and at the basin scale . For example, bot h the shea r an d tensil e strengt h o f mudrock s increase throug h burial and compaction (Fig . 7), hydrocarbons may migrate int o the trap thereb y locally increasin g overpressur e i n th e trap , an d hydraulic head ma y b e develope d a s a result of the uplift o f the basin margins.

Changing hydrodynamic regime Fluids i n th e subsurfac e ma y manifes t stati c behaviour (hydrostati c condition ) o r dynami c

2200m. Belo w thi s dept h mudston e porosit y deviate s fro m th e 'norma l compaction ' trend , a s a resul t o f overpressure. (Th e relativ e enrichmen t o f th e Teradomar i Formatio n i n low-densit y organi c matte r ma y als o contribute to the deviatio n fro m th e 'normal ' trend. ) Unde r condition s o f 'norma l compaction', embrittlemen t (Pb = 2.5gcm~ 3 ) o f thes e Neogen e claystone s woul d occu r a t 2700m ; however , actua l shal e embrittlemen t probably occurs below a burial dept h o f 3000 m, because of the presence of overpressures.

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Fig. 7 . Mohr-Coulomb failure envelopes for claystones at different levels of compaction. Bot h shea r strengt h and tensile strengt h increas e wit h mechanical an d chemical compaction . (Afte r Cartwright & Lonergan 1996. )

behaviour (hydrodynami c condition) . Hubber t (1953) demonstrate d tha t unde r hydrodynami c conditions accumulation s o f oi l o r ga s wil l invariably exhibi t incline d oil - o r gas-wate r interfaces an d in suc h case s th e computatio n o f hydrocarbon columns , base d o n th e assumptio n of hydrostatic conditions, will be spurious. Water flow i n sedimentar y basin s i s drive n b y geographical variation s i n wate r potential , which ca n chang e considerabl y i n natur e an d distribution durin g th e evolutio n o f a basi n (Wells 1988) . The early compaction histor y of an extensional basin , characterize d b y continuou s subsidence, will result in up-dip water flow to the basin margins. The driving force of this system is the relativel y hig h wate r potentia l i n th e basi n centre generate d b y wate r release d throug h th e processes o f mechanica l compactio n an d cla y mineral transformations . Thi s water-potentia l system change s i f significan t topographi c relie f forms adjacen t t o th e basin , owin g t o isostati c effects o r som e othe r mechanism . Th e earlie r hydraulic syste m o f th e basi n i s the n reversed . Elevated wate r table s alon g th e basi n margi n create a hydraulic head, whic h drives water flow inward toward s th e basi n centre , provide d upward discharge i s possible there . A significant consequence of basin inversion is a change in hydrodynamic conditions within the basin. The pattern of exhumation convolved with

the pre-existing basi n morpholog y ma y result in the outcro p o f ke y aquifers , a redistributio n o f recharge an d discharge areas and a change in the direction o f gravity-drive n fluid flow within the basin. The existence o f regional, topographicall y driven groundwate r flow-system s ha s bee n documented fo r severa l exhume d sedimentar y basins (Well s 1988 ; Demin g e t al 1992 ; Bredehoeft e t a l 1994 ; Demin g 1994 ; Crame r etal 1999) . Hydrodynamic effect s o n sea l capacit y may, for al l practica l purposes , b e ignore d excep t i n those basin s tha t manifes t clea r evidenc e o f hydraulic gradients . I n thes e basin s hydrodyn amic flow may modify sea l retention capacity by either increasin g o r decreasin g th e drivin g pressure agains t th e sea l (Alle n & Allen 1990) . When th e hydrodynami c forc e ha s a n upwar d vector, i t add s t o th e buoyanc y force , thu s reducing th e hydrocarbo n colum n height s th e seal ca n support . I n th e cas e o f a downwar d vector i t reduces the buoyancy force on the sea l and permit s th e retentio n o f a n increase d hydrocarbon column. Abnormal pore pressures in exhumed basins Abnormal formatio n pressures , eithe r abov e o r below the hydrostatic condition, occur in a wide

Fig. 8. Schematic illustration of burial history and forces of overpressure geoeration and deflatio for a basin characterized characterized by (a) continuous subidence and (b) exhumation. (Adapted form Gaarenstroom et al.; Luo & Vasseur 1995

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range o f geographica l an d geologica l setting s (Swarbrick & Osborn e 1998) . Th e globa l compilation o f La w & Spence r (1998 ) suggest s that abnorma l formatio n pressure s ar e observe d in some 15 0 basins with overpressure recorded in 148 o f thes e region s an d underpressure s manifest i n onl y abou t 1 2 areas. Thi s apparen t bias ma y i n par t b e du e t o th e fac t tha t underpressure i s mor e difficul t t o identif y during conventiona l drillin g operation s tha n overpressure, a s mos t well s ar e drille d wit h mud weigh t slightl y 'over-balance' . At an y tim e durin g th e evolutio n o f a sedimentary basin , th e distributio n o f abnorma l pore pressure s i s a functio n o f th e balanc e between force s tha t generat e overpressur e an d forces tha t dissipat e overpressur e (Lu o & Vasseur 1995) . Fo r example , i n a basi n characterized b y continuou s subsidence , suc h as the Nort h Se a Centra l Graben , th e curren t 'snapshot' o f por e pressur e evolutio n suggest s that forces of overpressure generatio n exceed th e forces o f dissipatio n (Fig . 8a ) (Gaarenstroo m et al. 1993) . I n contrast , i n th e exhume d basin , such a s th e Ordo s Basi n (China) , a clos e t o hydrostatic equilibriu m ha s bee n achieve d b y pervasive depressurizatio n durin g uplif t an d denudation betwee n 105-6 5 Ma (Fig . 8b ) (Luo & Vasseur 1995). Mechanisms fo r th e generatio n o f abnorma l pore pressures have been extensively reviewe d i n the literatur e (Gretene r 1981 ; Neuzi l & Pollock 1983; Martinse n 1994 ; Osborn e & Swarbric k 1997; La w & Spence r 1998 ; Swarbric k & Osborne 1998) . I n summary , thre e generi c mechanisms hav e bee n identifie d b y thes e workers a s primary potentia l contributor s t o the generation o f overpressure s i n subsidin g petro liferous basins : stress-relate d mechanism s (vertical loading , horizonta l tectoni c com pression); pore-flui d volum e expansio n (increased temperature , dehydratio n reactions , hydrocarbon generatio n an d oi l t o ga s crack ing); flui d movemen t (osmosis , hydrocarbo n buoyancy an d hydrodynami c flow) . Disequilibrium compactio n i s th e mos t com monly cite d mechanis m fo r overpressur e gener ation i n young , rapidl y subsidin g basin s wit h overpressures primaril y controlled by the vertical permeability o f th e shal y facie s containin g th e overpressures (Burru s 1998) . Th e ephemera l nature o f thes e overpressures ha s bee n empha sized b y Demin g (1994) , wh o offere d th e view that rock s ar e no t capabl e o f sustainin g zer o effective permeabilit y t o wate r ove r extende d periods o f geologica l time . I n olde r (pre Cenozoic) rock s keroge n transformatio n i s th e most commonl y cite d mechanis m fo r over -

pressure generatio n in subsidin g basin s (La w & Spencer 1998) . Whe n a subsidin g basi n i s exhumed, mechanica l compaction , minera l transformations involvin g dehydration reactions, and hydrocarbo n generatio n an d maturatio n processes wil l b e arrested , an d thes e force s will b e unabl e t o contribut e t o overpressur e generation i n an exhumed basi n settin g (Fig . 8) . Underpressured regime s ar e commonl y characterized b y shallowl y buried (0.6-3.Okm ) permeable rocks , which are, at least temporarily, hydraulically isolate d withi n low-permeabilit y mudrocks (Swarbric k & Osborne 1998) . Under pressured rocks typically occur in exhumed basin settings, whic h suggest s som e causa l linkag e with the change in matrix and pore-fluid volumes resulting fro m decrease s i n pressur e an d temperature, and hydraulic realignment, induced by th e exhumatio n process . Propose d mechan isms for underpressurin g cite d by Swarbric k & Osborne (1998 ) include : differentia l recharg e and discharg e rate s i n a topographicall y driven flow system ; rapi d migratio n o f exsolve d gas , from a low-permeabilit y reservoir durin g uplift , relative t o th e rat e o f ingres s o f ga s t o th e reservoir; dilatio n o f pore s i n shallowl y burie d mudrocks a s vertica l loa d i s remove d vi a denudation; aquatherma l contraction . Under pressured reservoirs are common in the exhumed Laramide basin s of the USA and Canada, such as the Denver , Piceance , Sa n Jua n an d Albert a basins (Belit z & Bredehoef t 1988 ; Bach u & Underschultz 1995) . However , wit h th e excep tion o f on e wel l fro m th e Barent s Se a Basin , reported b y Dor e & Jense n (1996) , under pressuring has not been documented i n exhumed Atlantic margin basins, although it is anticipated. In any sedimentary basin the present-day por e pressure distribution s ar e a functio n o f th e mechanism o f generatio n o f th e abnorma l por e pressure condition s an d th e hydrauli c evolution of the basin during and after th e creation o f these abnormal por e pressur e conditions . Bot h over pressuring an d underpressurin g ar e a disequilibrium condition , whic h wil l chang e wit h th e evolution o f th e basin . For example , i n a basin where thic k low-permeabilit y shal y successions are presen t a s a pressur e seal , th e hydrauli c evolution may involve pulses of hydrofracturing , which generate s a transien t permeability i n th e pressure sea l an d permit s th e escap e o f fluids . The timin g o f thi s hydrauli c fracturin g i s controlled primaril y b y th e prevailin g stres s conditions i n th e basi n an d th e properties o f th e pressure seal , s o that it can occur both during the burial an d subsequen t exhumation phase o f th e basin (Cosgrov e 2001) . A n assessmen t of when and wher e fracturin g ca n occu r i n th e evolution

Fig. 9 . Mohr diagrams with failure envelopes illustratin g computed stres s condition s i n a sandstone, unde r assume d conditions , a t 0.5 km intervals throughout buria l an d exhumation. Downward pointing arrows indicate the evolution of effective vertical stress (greatest effective principal stress tri) and effective horizontal stres s (least effective principal stress 03) during burial. For example, a dramatic increase i n pore-fluid pressure between 3 and 3.5 km results in a decrease in the magnitude of a-\ and cr 3 throug h this interval, fro m 4 8 an d 9 MPa t o 2 2 and — 17 MPa, respectively . I n this scenario , th e failur e envelop e i s breached an d fracturing wil l occu r t o reliev e the pore-flui d pressure build-up . Upward pointin g arrow s indicat e th e evolution o f vertical an d horizontal compressiv e stresse s during uplift . If the rock is permitted to extend durin g regional uplift significan t fracturin g will develop. Conditions favourable to hydraulic fracturing occur during burial at 0-1.5 km and 3-5 km , and during uplift at 3-Okm (from Davi s & Reynolds 1996) .

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of a n exhumed basin ca n be made b y analysin g the stres s condition s o f a rock , unde r assume d conditions, a t intervals o f 0.5 km through buria l and uplif t (Fig . 9 ) (Davi s & Reynold s 1996) . Although th e formatio n o f hydrauli c fracture s during th e earl y stage s o f burial an d diagenesi s (< 1.5k m burial ) i s an important mechanis m for fluid escape from low-permeability, semilithified sediments, thes e fractures are rarely preserved in the roc k record . However , Cosgrove (2001 ) ha s recently demonstrate d th e serendipitiou s preser vation o f thes e earl y hydrauli c fractures i n th e low-permeability Merci a Mudstone s o f th e Bristol Channel Basin, as a result of the injection of intra-formationa l san d bodie s int o th e hydraulic fractures , thereb y preserving the m a s sedimentary dyke s and sills. Furthermore, a later set of sub-horizontal, bedding-parallel, hydraulic fractures, which pertain to the exhumation phase, have been preserved i n outcrop, as a network of satin spa r veins , withi n evaporite-ric h horizons . The orientatio n an d spatia l organizatio n o f th e satin spa r vein s i s consisten t wit h th e tectoni c compression o f th e Bristo l Channe l Basi n during th e exhumatio n tha t wa s initiate d i n Mid-Cretaceous time.

Empirical observations and discussion: exhumed Atlantic margin and borderland basins Many Atlantic margin and borderland basins are characterized b y exhumatio n durin g Cenozoi c time (Fig. 10 ) Although prior exhumation events may have occurred during the evolution of these basins, in terms of top-seal assessment, Cenozoic exhumation is most critical, as it generally occur s in thes e basins afte r th e initia l migratio n o f hydrocarbons int o traps. Shale an d evaporit e cap-rock s for m th e main regional seal s t o hydrocarbo n accumulation s i n exhumed basin s o f th e Atlanti c margi n an d borderlands (Fig . 11) . Shal e cap-rock s o f Jurassic-Cretaceous ag e ar e prevalen t i n th e Celtic Sea , Inner Moray Firth, West of Shetland s and Barents Sea basins, and mixed evaporite and shale seal s o f Triassi c ag e ar e encountere d i n three basin s (Slyn e Trough , Eas t Iris h Se a an d Southern Nort h Sea) . Th e prodigiou s Zechstei n evaporite sea l i s cap-roc k t o a n estimate d

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ultimate recover y (EUR ) o f >150TC F (trillio n cubic feet) of gas in the southern Permian Basin, including th e gian t Groninge n accumulatio n (97 TCP) an d c . 3 5 TCP containe d i n 3 5 accumulations i n the UK sector o f the Souther n North Se a Basin (Glennie 1997 ) (Fig . 11) . There ar e a number o f hydrocarbo n trappin g and top-sea l configuration s observe d i n thes e exhumed basins . I n th e Celti c Se a Basin , a maximum ga s colum n o f 91 m i n th e Kinsal e Head Ga s Fiel d i s retaine d b y 46 m o f Gaul t claystone i n a compressiona l anticlina l flexur e (Fig. 12) . Thi s basin-centre d accumulatio n experienced c . 900 m o f exhumatio n durin g early Cenozoic tim e and has been overprinted by a compressiona l deformatio n tha t i s poorl y constrained bu t probabl y post-Paleocen e i n ag e (Murdoch e t al 1995) . Loca l evidenc e suggests that th e maximu m buria l dept h o f th e Gaul t claystone wa s i n th e rang e o f 1700-1800m , which ma y no t hav e bee n sufficien t t o achiev e shale embrittlemen t befor e th e applie d defor mation associate d wit h exhumatio n an d com pressional inversion . I n thi s scenario , i t i s postulated tha t fracture development wa s inhibited a s the Gaul t claystone responded by plastic flow to Cenozoic deformation. This hypothesis is consistent wit h th e experimenta l result s o f Bolton e t a l (1998) , whic h suggest s tha t underconsolidated clayey sediments, undergoing shear, defor m b y bul k volum e loss , whic h reduces permeabilit y an d result s i n weakl y developed deformatio n fabric s tha t hav e littl e impact o n th e hydrologica l propertie s o f th e claystone. All four generic leakage mechanisms (tectonic breaching, capillar y leakage , hydrauli c leakag e and diffusion ) operat e i n bot h continuousl y subsiding basin s an d exhume d basi n settings . However, th e critica l aspec t o f tra p leakag e i n exhumed basin setting s i s a lower probability of trap replenishment, as a result of the 'switchin g off o f hydrocarbo n generatio n durin g regiona l uplift. I n such cases, top-seal failure (induce d by tectonic breachin g o r hydraulic leakage) durin g exhumation may result in the catastrophic loss of a pre-existin g hydrocarbo n fil l wherea s post exhumation thes e traps ca n onl y be replenished from a curtaile d hydrocarbo n budget , whic h consists primaril y o f remigratin g oi l an d gas . This i s consisten t wit h th e observatio n tha t a number o f hydrocarbo n accumulation s i n

Fig. 10 . Uplifte d basin s o f th e northeaster n Atlanti c margi n an d borderlands . Man y o f thes e basin s ar e characterized by Cenozoi c exhumatio n events, which hav e occurred afte r the initia l migratio n of hydrocarbons into traps. HB, Hatton Bank; RB, Rockall Bank; BAF, Barr a Fan; SSF , Sul a Sgeir Fan; JM , Jan Mayen; VMH , V0ring Marginal High; BJF, Bj0rn0y a Fan .

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Fig. 11 . Regional seals to hydrocarbon accumulation s i n exhumed basins of the Atlantic margin and borderlands . Shales an d evaporites ar e the most common cap-roc k lithologie s in this setting. MMG, Mercia Mudstone Group; MBOR, millio n barrel s o f oil recoverable .

exhumed basin settings along the Atlantic margin are characterize d b y underfille d traps (Fig . 13) . Empirical observation als o indicates tha t hydro carbon accumulation s i n exhume d basin s ar e characterized b y hydrostaticall y pressure d o r modestly overpressure d reservoirs , wherea s significantly overpressure d reservoir s ar e com mon i n basin s tha t hav e experience d relativel y continuous subsidence (Fig . 14) . This suggest s a close causal relationship between regional uplift , hydrocarbon remigratio n an d dissipatio n o f overpressures. Exhumation ma y als o hav e positiv e impli cations fo r th e capillar y an d hydraulic retentio n capacity of mudrock seals. Increased mechanica l compaction as a result of burial results in reduced interconnected pore-throa t size s an d increasin g shear strength and tensile strength for a claystone rock. For example, whe n a claystone is exhume d it retain s th e tensil e strengt h o f it s maximu m burial depth and , consequently , a highe r pore fluid pressur e wil l b e require d t o induc e hydrofracturing tha n fo r a clayston e i n a continuously subsidin g basi n a t th e sam e depth . Leak-off test s (LOTs ) an d formatio n integrit y tests (FITs) from a subset of the Atlantic margin basins suppor t thes e observation s (Fig . 15) .

These dat a indicat e that , fo r an y give n buria l depth, the minimum horizontal stres s or fractur e pressure (define d by the lowe r envelop e o f LOT pressures) i s highe r i n exhume d basin s tha n i n those basin s characterize d b y continuou s sub sidence. Critically , ther e ar e a numbe r o f FIT s performed o n exhume d claystone s o f Carbon iferous, Triassic and Jurassic age, which indicate that thes e sea l rock s hav e ver y hig h tensil e strengths, appropriat e t o thei r maximu m buria l depth befor e exhumation . Thi s suggest s tha t post-exhumation top-sea l integrit y i s relativel y high i n man y o f th e Atlanti c margi n an d borderland basin s under low differential stresses . However, the anomalously high shear strength of exhumed mudstone s ma y resul t i n th e develop ment o f dilatan t shea r fracture s (unde r lo w confining pressures ) i f shal e embrittlemen t ha s been achieve d befor e exhumation . Th e absenc e of seismi c chimney s acros s majo r ga s accumu lations i n man y of thes e basin s (e.g . Celti c Sea , East Iris h Se a Basin , Souther n Nort h Sea ) suggests tha t dynami c leakag e throug h th e top seal i s no t occurrin g a t th e presen t da y an d that pore-fluid pressure s ar e belo w th e top-sea l capillary an d hydrauli c leakag e threshold s fo r these accumulations.

Fig. 12 . Kinsale Hea d Ga s Field, a compressional inversio n structur e i n the exhumed Celti c Se a Basin, (a) Depth structur e ma p on top main reservoir , indicatin g tha t th e gas-water contact (GWC at — 2967 feet) is coincident with the spill point to the north of the structure. C.I., contour interval, (b) NNW-SSE seismic line, illustrating reverse faulting on the southern limb of the anticline, (c) Type log for the 'A Sand-Gault Claystone reservoir-top-seal couplet at the Kinsale Head Field. A maximum gas column of 91 m, in the Greensand reservoir, is retained in situ, without apparent leakage, b y 46 m of Gault Claystone. Maximum burial depth of the Gault claystone, in the Kinsal e Head area , i s estimated to have been <2km . As a result, it is postulated that Cenozoic exhumatio n of the Gault Claystone occurre d befor e shal e embrittlement , thereb y inhibiting the formation o f dilatant fracture s (see Fig . 6) . (Dat a fro m Tabe r e t al. 1995. )

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Fig. 13. Underfilled traps in exhumed Atlantic margin basins. Of the five major gas accumulations from the Celtic Sea to the Barents Sea (Kinsale Head, South Morecambe, Comb, Victory, Sn0hvit) only the Kinsale Head Field is full t o structura l spill point, at the present day.

Fig. 14 . Comparison o f initial reservoir pressure s v. depth fo r hydrocarbon accumulation s in exhumed basins v. non-exhumed basins . Accumulation s i n exhume d basin s ar e characterize d b y hydrostaticall y pressure d o r modestly overpressured reservoirs, whereas significantly overpressured reservoirs are common in basins that have experienced relativel y continuou s subsidence. SNS , Souther n North Sea Basin; EISB, East Irish Se a Basin; ST, Slyne Trough ; BS , Barent s Se a Basin ; TVDSS , tru e vertica l depth , sub-sea . (Dat a compile d essentiall y fro m Spencer e t al 1986 ; Abbott s 1991 ; Pooler & Amory 1999. )

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479

Fig. 15. Formation pressures (repeat formation tests (RFTs) and drill stem tests (DSTs)), formation integrity tests (FITs) an d leak-of f test s (LOTs ) fo r basins offshore Ireland . Dat a indicate that, for an y given burial depth, th e minimum horizonta l stres s o r fractur e pressur e (define d b y th e lowe r envelope o f LO T pressures) i s higher i n exhumed basins than in those characterized b y continuous subsidence.

Conclusions A numbe r o f effectiv e regiona l cap-rock s ar e recognized i n the petroleum system s of exhumed Atlantic margin basins. However, the observation of underfille d traps , clos e t o hydrostaticall y pressured reservoir s an d breache d trap s wit h residual oil and gas shows, suggests that top-seal behaviour exercise s a critica l contro l o n por e pressure evolutio n an d hydrocarbon distribution and redistribution in exhumed basin settings. The following ar e th e principa l conclusion s o f thi s review. (1) Depressurizatio n o f reservoir s durin g exhumation i s a consequenc e o f faul t reactivation, fracturin g o f th e cap-roc k an d th e large-scale remova l o f overburde n durin g regional uplift . A s a result , exhume d basin s

typically sho w limite d overpressurin g excep t where residual overpressure s hav e been retaine d by extremel y low-permeabilit y aquitard s o r exceptionally ric h sourc e rock s (wit h sub nanodarcy permeabilities ) o r wher e overpres sures have been generate d after regiona l uplift . (2) Roc k por e dilatanc y an d flui d volum e reduction, cause d b y a lowering o f temperatur e during exhumation, may result in underpressured reservoirs, whic h ar e commo n i n th e exhume d Laramide basin s o f th e US A an d Canada . However, wit h th e exceptio n o f on e reporte d occurrence i n th e Barent s Se a Basin , under pressuring ha s no t bee n reporte d i n exhume d Atlantic margin basins, although it is anticipated. (3) Post-exhumatio n overpressur e generatio n is primaril y drive n b y tectoni c compression , aquathermal pressurin g an d th e evolve d

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hydrodynamic regim e a s disequilibriu m com paction, mineral transformation and hydrocarbon generation mechanism s ca n no longer contribute to overpressure generatio n i n the exhumed basin setting. (4) Lithofacie s i s a major contro l o n top-sea l efficiency i n exhume d basi n settings . Th e juxtaposition o f a halit e directl y abov e th e hydrocarbon-bearing reservoi r offer s th e mos t favourable conditio n fo r th e retentio n o f hydro carbons i n exhume d traps . However , th e empirical evidenc e fro m th e Atlanti c margi n suggests tha t mudrock s ca n for m efficien t top seals in exhumed basins under certain conditions. (5) Th e behaviou r o f an y cap-roc k litholog y during exhumatio n i s dependen t upo n th e physical and mechanica l characteristic s of the cap-rock a t th e tim e o f exhumatio n an d th e timing an d condition s o f th e associate d defor mation relativ e t o th e timin g o f hydrocarbo n emplacement. (6) Mudrock s may form effective cap-rock s in exhumed basin s whe n th e deformatio n associ ated wit h exhumatio n occur s befor e embrittle ment and the cap-rock exhibit s ductile behaviour. Where exhumatio n occur s post-embrittlemen t the shal e cap-roc k wil l facilitat e hydrocarbo n leakage throug h th e developmen t o f extensiv e fracture networks . (7) Syn-exhumation top-sea l efficienc y (flui d retention capacity ) is a major exploration risk in exhumed basi n settings , althoug h post-exhumation top-sea l integrit y i n thes e basin s ma y b e relatively high . Thi s suggest s tha t a majo r exploration risk factor in exhumed basin settings pertains t o th e limite d hydrocarbo n budge t available post-regiona l uplif t an d th e efficienc y of th e remigration process .

The authors woul d like to thank A. Carr and J. Gluyas for thei r helpful reviews o f this manuscript. W e would also lik e t o than k J . Kipp s (Statoi l (UK ) Ltd), who draughted mos t of the figures in this paper. Finally , the authors woul d lik e t o than k Statoi l Exploratio n (Ireland) Ltd for their sponsorship o f the color printing costs.

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Index Page numbers i n italic, e.g . 70, refer t o figures. Pag e number s in bold, e.g. 6, signify entrie s i n tables . admittance 16-1 7 Albatross field 414 initial reservoir pressur e 478 Alligin field 415 Altmark-Brandenburg Basi n 70 And0ya Slid e 147 Anton Dohrn Seamoun t 31 4 apatite fission-track (AFT) analysis 5-7 , 6 Cenozoic uplif t an d denudatio n of southern Norway 222 factors affectin g accurac y and precision o f timing estimates 341-343, 342 Paleocene initiation o f Cenozoic uplif t i n Norway 60-61 quantifying Atlanti c exhumatio n 331-332 analytical problems 337, 338-339 approaches base d o n heat flow 334 availability o f palaeotemperature constraint s over a range of depths 33 9 basic assumption s 332-333 , 332 calibration of system response 33 5 depth resolutio n versus thermal resolutio n 34 4 Early Tertiar y palaeo-therma l effect s i n NW England 344-346, 345 factors affectin g th e accurac y of the magnitud e of exhumation 332-339 factors affectin g the precision of the magnitud e o f exhumation 339-340 heating rate s 334-335 identifying th e appropriat e unconformit y 335 inappropriate system kinetics 335-33 6 integration o f results fro m differen t method s 337-338 latest result s 344-34 9 limitations o f this technique 343-344 Mesozoic palaeotherma l episodes i n Ireland an d Central Irish Se a Basin 347-349 nature of palaeotemperature constraint s 339-340 new results fro m Central Englan d 346-34 7 non-linear palaeogeotherma l gradients 333-334, 333 onset versus duration of cooling o r exhumation 344 palaeo-surf ace temperature 334 reburial 344 thermal history resolution 343-344 timing o f exhumation 340-341 , 340, 347 Scandinavian denudation age and track length statistics 12 0 annealing 119-12 0 inverse modellin g 120-12 1 procedure 12 0 thermal an d denudational history o f Ireland 377-380 , 378-379, 382-383, 384^386 analytical details 38 0 contouring method 387 evaluation of geothermal gradient 387-388 modelling procedures 380-387 sampling 38 0 thermotectonic development of souther n Sweden 177-180 dating method 17 2

forward modellin g 172-17 6 geological an d geomorphological modellin g constraints 176 modelled thermal histories 174-175 modelling result s 173, 176-177 timing o f Norwegian denudatio n 28-31 , 33 Ardmore field 421 Armorican Block 86 , 87 Askeladden field 414 initial reservoir pressur e 478 Asland field 419 Assynt 276, 277 Asta Graben 23 7 Atlantic, North see North Atlanti c Atnajakka fluvia l valle y 759 , 160 Ballycotton field 421, 477 Baltic Shiel d 68 , 69 Barents Se a 118, 403,412-416, 414, 423, 474 cap-rocks 476 Barra, Isle of 276, 277 Barra Fan 314, 315, 317,318 Beatrice field 418 Ben Wyvis 273 Bennachie 27 3 Bill Bailey's Ban k 314 Binney field 419 Bivrost Fault Zone 14 7 Bjarmeland Platfor m 41 4 Bjorn Drif t 31 4 Bj0rnoya Basin 414 Blake field 418 Bothnia, Gulf o f 11 8 Bristol Channe l 86 , 372 Britain bathymetric settin g o f Atlantic margi n 31 4 Cenozoic denudatio n 18-19 , 19 mass balance calculation s 19-21 , 20, 21 mass balance results 21-22 correlation wit h Alpine tectonics 95-98 , 97 Early Tertiar y palaeo-therma l effect s i n NW Englan d 344_346, 34 5 Late Neogen e developmen t o f Antlantic margi n 313-314, 327 airgun profiles 327 , 378, 379, 320 basal unconformity 316-322 deep-water sedimentar y respons e 324-325 mechanics o f change 326-32 7 prograding wedges 322-324 setting 314-31 6 stratigraphy an d sedimentation 375 , 316-326, 377 summary even t stratigraphy 325-326, 325 timing of change 32 6 Mesozoic sedimentar y basin evolutio n 86-88 , 86 new results fro m Centra l Englan d 346-34 7 post-Variscan tectonic events 87 southern, Cenozoic inversion an d uplift 85-86 , 88 fault inversio n 88-89

486

INDEX

Britain continued inversion by bulk deformation an d uplift 9 0 Variscan Thrust Fron t 89 , 90 Tertiary landscap e evolutio n 98-9 9 Variscan structura l framework 87 Weald Basi n inversio n 90-91, 92 forward mode l 94-95, 94, 95, 96 model 91-93 , 9 3 Brona Basin 403, 474 Brygge Formation 141 , 141, 142, 143 Buchan 276, 277 Buchan field 418 Cairngorms range 276 Caithness 273 , 276 Caithness Ridg e 41 8 Calder field 419 Cape Wrat h 27 6 Captain field 418 Cardigan Ba y Basin 372, 474 Celtic Se a Basin cap-rocks 47 6 Kinsale Head ga s field 477 Cenozoic denudation o f Britain an d Ireland 18-19 , 19 mass balanc e calculation s 19-21 , 20, 21, 21-22 Cenozoic evolutio n o f the Faroe Platform 291 , 308-30 9 denudation depth 30 8 mass balance 306-308 denudation estimate s 30 6 sediment yiel d estimates 30 7 structural element s and regional setting 291-294 , 292 distribution o f Paleocene basal t 297, 298 seismic section s 293, 295, 303, 305 seismic unit s 294 stratigraphic correlatio n 294-298 , 296 structural evolution 30 7 tectonics, denudatio n an d depositio n Early and Mid-Miocene time 304-305 Eocene depositio n 303-30 4 Eocene-Recent denudation of Paleocene platea u basalts 299-303 , 300, 301, 302 Late Miocene-Earl y Pliocene tim e 305 Late Pliocen e tim e 305-30 6 Oligocene tim e 304 Paleocene time 298-299 , 299 Pleistocene tim e 306 Cenozoic evolutio n of the North Se a 229-230, 235-236, 244, 26 3 correlation wit h global climat e and sea level 221 correlation wit h regional tectoni c event s 219-221 database 21 6 depocentres 275 Eocene tim e 248-250, 249 Miocene tim e 254-258, 255 Neogene incise d valley s 25 6 Oligocene tim e 250-251, 252, 254 palaeogeographical developmen t 216-219 , 218 palaeo-water depth 240-243, 240, 241, 245, 249, 257 Paleocene time 244-248, 245, 246 Plio-Pleistocene time 257, 258-259, 258, 260 possible vertica l movemen t mechanism s 259-261 intra-plate stress 262 Late Paleocene-Early Eocene vertica l movement s 261 Oligocene-Miocene domal uplif t o f souther n Norway 26 2 Plio-Pleistocene glacia l erosio n an d uplift 26 2

regional profile s 23 9 regional settin g in Early Oligocene tim e 25 3 regional settin g in Late Paleocene-earliest Eocen e time 247 regional settin g in mid-Miocene tim e 25 5 seismic line s 246, 252, 258, seismic mappin g 236-240, 238 stratigraphy 273, 238 structural ma p 237 study are a 23 6 tectonic modellin g and subsidence analysi s 242, 243-244, 244 Cenozoic uplif t o f North Atlantic 2, 2 2 Paleocene initiation 45^46, 60, 63-64 active mechanisms of surface uplift 54-6 0 Bouguer gravity map 46 chalk overburial 53 changing in-plane stress 54-57, 55, 56 climate an d eustasy 46-47 gravity 6 3 heat flow 63 isostatic adjustments caused b y erosion o f existing topography 48, 49-51, 49, 50 isostatic and erosional response t o sea-level fal l 51-54,57,52 lithospheric delamination 58-60, 58, 59, 62 magmatic underplatin g 57-58, 5 7 mechanism o f surface uplift 61-6 3 passive mechanism s of surface uplift 49-5 4 pre-Quaternary geologica l ma p 54 sediment structur e and fission tracks 60-61 tectonics 47^4 9 thermal an d topographic response o f lithosphere 57, 58, 59, 62 Paleocene permanen t uplift 22-2 3 Paleocene-Eocene dynamic support 23 Oligocene-Recent dynami c support 23-24 timing and mechanisms 27-28, 40 denudation o f the eastern Atlanti c margi n 28-31 Neogene uplif t o f Norway 31-35 proposed mechanism s 28 Central Channel High 86, 89 Central Englis h Channel 86, 86, 88 Central Graben , North Sea 68-69, 68, 69, 70 Central Grabe n Dome , Nort h Sea 67-68, 68, 70 , 74 Mesozoic structur e and evolution 75 chemical denudation , definition 4 Cheshire Basi n 419 Clair field 475, 416-417 Claymore field 418 Clyde field 419 coherence analysi s 35, 35 Coll, Isl e of 276, 277 compaction analysi s 6 Cornubian Massi f 87, 87 Corona Basi n 475 Corona Ridg e 292 Corrib field 420 capillary sea l capacit y 46 3 initial reservoir pressur e 478 underfilled trap s 47 8 cosmogenic nuclid e 6 cosmogenic nuclid e studie s erosion o f glaciated passive margins 155 , 162-164, 763 complex exposur e an d shielding histories 157-15 8 erosion depth s 157

INDEX surface exposur e datin g an d erosion rate s 155-15 6 suface exposur e datin g and erosion rates 156, 15 7 Cowal peninsula 273, 276, 277 Cromarty field 418 Crossans field 419 crustal uplift, definitio n 4 Cullins 27 6 Dalradian belt 27 7 Danish Basin 73 , 185 Danish Megablock 67-68 , 68, 69, 71 Jurassic-Cenozoic rotation 7 0 Darwen field 419 Darwin volcanic centr e 292 Deemster Basi n 419 Deemster Platform 41 9 Dellen meteorit e impac t 104 Denmark Late Cenozoi c erosio n comparison wit h other studie s 19 8 estimates of missing sectio n 193-195 , 795 magnitude 196, 195-197, 797 timing o f maximum buria l an d subsequent erosio n 197-198 denudation, definition 4-5 Donegal Basin 372, 403 Donegal Fa n 314, 317 Douglas field 419 Dragon field 427 drill stem tests (DSTs) 47 9 Dutch Bank Basin 41 8 East Faroe Grabe n 415 East Faro e Hig h 292 East Faroes wedge 314, 315, 317 East Iris h Basi n 372 East Irish Se a Basin 417^22, 479, 423, 474 cap-rocks 476 East Midland s Shel f 7 6 East Orkne y Basin 41 8 East Rockall wedg e 374, 375 East Shetlan d Basi n 23 7 East Shetlan d Platfor m 23 7 Egersund Basi n 403, 474 Egret field 478 Elgin 273 erosion, definitio n 4 erosion o f glaciated passiv e margin s 153-155 , 75 4 cosmogenic nuclid e studies 155 , 162-164, 76 3 complex exposur e and shielding histories 157-15 8 erosion depth s 15 7 suface exposur e dating and erosion rates 155-156 , 756, 75 7 landscape surfac e reconstruction i n northern Swede n 158 fluvial inclusion during interglacial periods 158-161, 759 glacial erosio n b y selective linear erosio n 16 1 localized subglacia l modification o f preglacia l upland surface s 161-16 2 lowering o f preglacial uplan d surfaces by nonglacial processe s 162 preglacial fluvia l drainag e system 158 uneven sedimen t productio n 164-16 5 Erris Basin 403

487

Erris Ridge 420 Erris Trough 372, 474 Eubonia Basi n 479 exhumation, definitio n 3-4, 5, 210-212 exhumation of the North Atlantic margin 1-3, 348, 349 basins 403 mechanisms accelerated burial befor e uplif t 35 0 comparisons wit h other region s 350-351 regional events in NW Europe 349-35 0 repeated cycles o f burial an d exhumation 35 0 overview 5, 1 0 continental margin record 7- 9 mechanisms 7, 8 petroleum exploratio n 9-1 0 techniques 5-7, 6 quantification usin g apatite fission-track analysis and vitrinite reflectance dat a 33 1 analytical problems 337, 338-339 approaches based o n heat flow 334 availability o f palaeotemperature constraint s ove r a range o f depths 33 9 basic assumptions 332-333, 332 calibration o f system response 33 5 depth resolution versu s thermal resolutio n 34 4 Early Tertiary palaeo-thermal effects i n NW England 344-346, 345 factors affectin g accurac y an d precision o f AF T timing estimates 341-343 , 342 factors affectin g th e accuracy of the magnitud e of exhumation 332-339 factors affectin g th e precision of the magnitude of exhumation 339-34 0 heating rates 334-335 identifying th e appropriate unconformit y 33 5 inappropriate syste m kinetic s 335-336 , 336, 337 integration o f results from differen t method s 337-338 latest result s 344-34 9 limitations of these technique s 343-34 4 Mesozoic palaeotherma l episodes i n Ireland an d Central Iris h Se a Basin 347-34 9 nature o f palaeotemperature constraint s 339-34 0 new result s from Centra l England 346-347 non-linear palaeogeothermal gradients 333-334 , 333 onset versus duration of cooling o r exhumation 344 palaeo-surface temperatur e 33 4 reburial 344 thermal history resolution 343-34 4 timing from AF T data 340-341, 340, 347 Faroe Bank 374 Faroe Ban k Knoll 292 Faroe Island s 293, 374 Faroe Platform , Cenozoi c evolutio n 291, 308-309 denudation dept h 308 mass balance 306-308 denudation estimates 30 6 sediment yield estimate s 30 7 structural element s an d regional settin g 291-294 , 292 distribution of Paleocene basal t 297, 298 seismic section s 293, 295, 303, 305 seismic unit s 294 stratigraphic correlation 294-298 , 296 structural evolution 307 tectonics, denudatio n an d depositio n

INDEX Faroe Platform , Cenozoi c evolutio n continued Early an d Mid-Miocene time 304-305 Eocene depositio n 303-30 4 Eocene-Recent denudation of Paleocene platea u basalts 299-303, 300, 301, 302 Late Miocene-Earl y Pliocene tim e 305 Late Pliocen e tim e 305-30 6 Oligocene tim e 30 4 Paleocene tim e 298-299 , 299 Pleistocene tim e 30 6 Faroe Shel f 314, 403 Faroe-Shetland Basin 415 Faroe-Shetland Channel 293, 314, 315 Faroe-Shetland Escarpment Slid e 146 Farsund Basin 70 , 403, 474 Fastnet Basi n 37 2 Fastnet Hig h 42 1 Feni Ridg e 314, 317 Fennoscandian Borde r Zon e 170 Fennoscandian Hig h 7 0 Finnmark Platfor m 41 4 Flett Basin 292, 415 Flett Ridge 292, 415 flexural rigidity 34-35 Foinaven Basi n 292 , 293, 295, 415 Foinaven field 415 formation integrit y test s (FITs ) 476, 479 Formby field 419 Formby Point Faul t 41 9 Foula sub-basin 415 Friesland Dome 68, 69, 70 , 71 Fr0ya Basin 40 3 Fugloy Basi n 292, 293, 295, 303 Gaick 27 7 Gardar Drift 31 4 gas reserves and resources 401^02 Barents Se a 414, 416 dissolved i n Earth's crus t 449—45 0 East Iris h Se a Basin 417^22, 419 effects o f exhumation 412, 413 Inner Moray Firt h Basi n 417, 418 North Celti c Se a Basin 421, 424 release fro m groundwater 450-452 efficiency 452-45 3 resource estimatio n 409—41 2 risk analysis 402-409, 404 Slyne-Erris Basin 420 , 422-^24 solubility in water 447-448, 448, 449 West Shetlan d Basi n 415, 416-417 Geikie Escarpment 32 0 geomorphological analysis 6 George Blig h Bank 314, 315 Glelprie Fault Zone 14 7 Godred Crove n Basi n 41 9 Goliath field 414 Goteborg 17 0 Grampian Highland s Massi f 41 8 Great Gle n 2 73 Great Gle n Fault 274, 276 grus 10 8 Halibut Horst 273, 418 Halibut Shel f 41 8 Haltenbanken 14 0

Hamilton field 419 Hammerfest Basi n 403, 414 Hampshire-Dieppe Basi n 86, 88-89 Hano Bay 73 , 7 3 Harlingen gas field 465 Harris, Isl e of 273, 276, 277 Harstad Basi n 414 Hatton Bank 314, 403 Hatton Drif t 31 4 Hatton Rockall Basin 314 Hebridean Basins 403, 474 Hebrides Shel f 314, 317 Hebrides Terrac e Seamoun t 314 Helmsdale Faul t 276 Helsingborg 77 0 Helvick field 421 Heron field 478 Highland Boundar y Fault 274, 276 Horda Platfor m 237, 403, 423 hydrocarbon system s in exhumed basins 401^02, 424-425, 443 abnormal por e pressure s 472-475 commercial implication s 431 depressurization 457-459, 479-480 abnormal formatio n pressure s 45 8 capillary leakag e 459-462, 460 capillary sea l capacit y 463 dilatant shea r fracture s 46 4 hydraulic leakage 461, 462^-65 molecular transport 461, 465 tectonic breachin g 459, 460 top-seal leakag e mechanism s 459^65 diagenesis 433 diagenetic sequenc e 440 near-surface fluid s 435—43 6 empirical observation s 475-479 formation pressure s 479 initial reservoir pressure s 47 8 regional cap-rock s 476 underfilled trap s 47 8 examples fro m N W European margin 412, 423 Barents Sea, western 412-416, 414, 423 East Iris h Se a Basin 417^22, 419, 423 Horda Platfor m 423 Inner Moray Firth Basi n 417, 418, 423 North Celti c Sea Basin 427, 423, 424 North Se a Basin 42 3 Slyne-Erris Basin s 420, 422-424, 423 West Shetlan d Basi n 475, 416-417, 423 exhumation processe s 433, 469-470 changing hydrodynamic regime 470—472 , 477 chemical processes 43 5 physical processe s 43 4 stress 470, 473 fracturing 440-44 1 porosity an d permeability 447 key characteristic s 412, 473 natural gas 447, 453 dissolved ga s i n crus t 449—45 0 efficiency o f gas releas e 452—45 3 formation pressure s 479 hydrodynamic effects 451-452 , 470-472. 477 hydrostatic effect s 450—15 1 initial reservoi r pressures 478 solubility o f gas i n water 447^149, 44 8 underfilled trap s 478 uplift-related ga s releas e 450—452, 45 1

INDEX oil-bearing reservoir s 441-443 fluid-fill history 44 2 physical properties o f cap-rocks 465-466 lithology, porosity an d permeability 466 porosity-depth behaviour 46 8 strength, ductilit y an d brittleness 461-469, 467 porosity trend s 436-437 causes o f porosity chang e 436 clay distributio n 43 9 effect o f erosion o n pore pressure evolution 43 8 effective stres s 437 erosion rate 437 percentage oi l residue 439 secondary porosit y and sandstone framewor k stability 437-43 9 prospect resourc e estimation 409—410, 409 formation volum e factor 411 gross roc k volum e 410 net-to-gross ratio , porosity an d hydrocarbon saturation 410-411 recovery facto r 411-412 prospect risk analysi s 402-405, 404 probability o f oil versus gas 406-407 probability o f reservoir 40 5 probability of source and charge 405-406 probability of trap and seal 407^09 Iceland Basin 314, 317 Iceland (mantle ) Plume 7, 13-14 , 27, 58 influence i n the Faroe-Shetland area 294, 309 influence i n the Irish Se a basin are a 35 6 North Se a basin subsidence 261, 263 Paleocene uplif t o f British Isles 30 9 present-day dynami c suppor t 14-1 5 estimates fro m bathymetr y 15-16 estimates fro m gravit y 16-18 , 17 role i n Cenozoic surfac e uplift 61-63 , 62, 212, 214 Iceland-Faroe Rise 314 Indefatigable field 478 Inner Mora y Firt h Basi n 403, 417, 418, 423 cap-rocks 47 6 inversion, definition 5 Ireland 37 2 Cenozoic denudatio n 18-19 , 1 9 mass balance calculations 19-21 , 20, 21 mass balance result s 21-22 denudation sinc e Triassic time 388-389, 393 Triassic time 388 Jurassic time 389-390, 389 Cretaceous tim e 390-391, 390 Paleocene-Eocene time 391, 391 Oligocene-mid-Miocene time 391-392, 392 Mesozoic palaeotherma l episode s 347-349 post-Variscan thermal and denudation history 371-373, 396 eastern and northern offshor e flanks 377 Irish landmas s 373-375, 373, 37 4 palaeothermal an d thermochronological data 377-388, 378-379, 382-383, 384-386 southern offshore flank 375-376 western offshore flank 376-377 total denudationa l efflu x o f Irish landmas s 392-396 , 394, 395 Irish Platform 420 Irish Se a Basin 40 3 Mesozoic palaeotherma l episode s 347-349 regional setting and stratigraphy 356-357

489

sonic velocit y analysis of Tertiary denudatio n 355-356, 365-368 , 365, 366 curve-fitting, geologica l erro r an d derivation of exhumation 357, 358, 359,360-361, 361-364 methods 357-36 4 results 362 , 363, 364-365, 364 theoretical background and mechanical contro l over velocity 357-36 1 Irish Shel f 31 4 Jan Ma y en Fault Zone 14 7 Jan Mayen Ridg e 40 3 Judd High 415 Kai Formation 141 , 141, 142, 14 3 Karsatjakka mountai n block 759 Keys Basin 419 Keys Faults 419 Kilfenora Horst 420 Kinsale Head ga s field 421, 475, 478 capillary sea l capacity 463 initial reservoi r pressur e 478 underfilled trap s 478 Kish Bank Basin 372 Klibreck 27 7 Knoydart 276 Labadie Bank High 421 Laggan field 415 Lake Distric t Massi f 41 9 Lambda field 419 leak-off test s (LOTs ) 476, 479 Leman field 478 Lennox field 479 Lofoten Basi n 74 0 Lofoten Islan d 779, 133 Lomre Terrace 23 7 Loppa Hig h 474 Lome, Firt h of 27 7 Lousy Ban k 374 Lower Saxon y Basin 70 Loyal field 475 Magnus Basin 23 7 Main Porcupine Basi n 372 Malin Shel f 374 Malin Trough 372 Mal0y Fault Block s 23 7 Mamock field 478 mantle plume 7, 13-14 , 27, 58 influence i n the Faroe-Shetland area 294 , 309 influence i n the Iris h Se a basin area 35 6 North Sea basin subsidence 261, 263 Paleocene uplift o f British Isle s 309 present-day dynami c support 14-1 5 estimates from bathymetr y 15-16 estimates fro m gravit y 16-18 , 77 role in Cenozoic surfac e uplif t 61-63 , 62, 212, 214 mantle upwelling 27-28 uplift o f southern Norway 35-3 7 Marulk Basin 23 7 mass balance analysi s 6 Cenozoic denudatio n of Britain an d Ireland 19-21 , 20, 21 results 21-22

490

INDEX

mass balanc e analysi s continued mechanical denudation , definitio n 4 methane rates o f generation 45 2 solubility in brine 448, 449 Millom field 419 Mizzen Basi n 42 1 modelling apatite fisson track (AFT) analysis 177-18 0 AFTsolve program 120-121 , 172 dating metho d 17 2 forward modellin g 172-17 6 geological an d geomorphological modellin g constraints 17 6 modelled therma l historie s 174-175 modelling result s 173 , 176-17 7 thermotectonic developmen t o f southern Swede n 172-177 erosion o f souther n Scandinavi a chronostratigraphic even t definition 190-191 erosion model 191-19 3 heat-flow mode l 19 1 model calibratio n an d results 192, 193 , 194 model descriptio n 190-19 3 palaeo-surface temperature s 191 , 19 1 Neogene uplif t o f souther n Norway coherence analysi s 35 , 35 estimating flexural rigidity 34-35 parameter value s 31-35, 31 tectonic uplif t an d isostati c reboun d 36 North Se a Basin lithospher e 78 , 80-81, 81 new mode l o f Cenozoic subsidenc e an d marginal dome uplif t 7 9 previous models o f Mesozoic-Cenozoic evolution 78-79 synrift an d post-rift basin evolution 78 North Se a tectonics an d subsidence analysi s 242, 243-244, 244 Paleocene initiatio n of Cenozoic uplif t passive margi n formation 54-57 , 55 , 56 thermal an d denudationa l histor y o f Irelan d 380-387 Weald Basi n inversion 91-93, 93 forward mode l 94-95 , 94, 95, 96 Moine Thrus t 27 4 Moray Firt h 277 , 474 Moray Firt h Graben 68 , 68, 70 M0re Basi n 140, 474 M0re-Tr0ndelag Fault Complex 147 M0re-Tr0ndelag Fault Zone 104, 113 Morecambe field 419 initial reservoir pressur e 47 8 Mormond Hil l 27 7 Morven 276, 277 Morvern 273, 276 Mounth 273 Muddus plains 10 8 compared with Norwegia n Palaei c relief 109-11 0 Mull, Isle o f 273, 276, 277, 278 Munken Basin 29 2 Myrsilde field 414 natural gas see gas reserves an d resource s Naust Formation 141 , 141, 142-143, 142, 143, 145, 145, 146 net uplift , definitio n 4

Nordkapp Basi n 403, 414 Norewegian-Danish Basin 403 North Atlantic anomalous topography 7 5 Cenozoic uplift 2 , 22, 27-28, 40 denudation of the eastern Atlanti c margin 28-31 Neogene uplif t o f Norway 31-3 5 Oligocene-Recent dynamic suppor t 23-24 Paleocene-Eocene dynamic suppor t 2 3 Paleocene permanen t uplif t 22-2 3 proposed mechanism s 28 free-air gravit y anomal y ma p 1 4 Neogene evolutio n of mid-Norwegian margi n 139-141, 147-149 , 147 north of Storegga Slid e Complex 142-143 , 143, 143-147 stratigraphy 141-147 , 141, 146 Paleocene initiatio n of Cenozoic uplif t 45-46 , 60, 63-64 active mechanisms of surfac e uplift 54-6 0 Bouguer gravity map 46 chalk overburia l 5 3 changing in-plane stress 54-57, 55, 56 climate an d eustasy 46^7 gravity 6 3 heat flow 63 isostatic adjustment s caused by erosion o f existing topography 48, 49-51, 49, 50 isostatic an d erosional respons e t o sea-leve l fal l 51-54,57,52 lithospheric delamination 58-60, 58, 59, 62 magmatic underplatin g 57-58, 57 mechanism o f surface uplif t 61-6 3 passive mechanisms of surfac e uplift 49-5 4 pre-Quaternary geological ma p 54 sediment structur e and fission tracks 60-61 tectonics 47^49 thermal an d topographi c response o f lithosphere 57, 58, 59, 62 seismic tomograph y 37—4 0 P-wave map s 38, 39 topographic an d bathymetric map 3 North Buchan Graben 418 North Celti c Sea 86, 90 North Celtic Sea Basin 372, 403, 421, 423, 424, 474 North Minch 27 7 North Porcupin e Basin 372 North Se a 67 bathymetry 210 Cenozoic depocentre s 275 Cenozoic evolutio n 229-230, 244 correlation wit h globa l climat e an d se a level 221 correlation wit h regiona l tectonic events 219-221 database 21 6 palaeogeographical developmen t 216-219, 218 Paleocene time 244-248, 245, 246 Cenozoic stratigraph y 273 central North Se a Dome 68-7 1 Danish Megabloc k 71 Jurassic-Cenozoic lithosphere models 78 , 80-81, 81 new mode l o f Cenozoic subsidenc e and marginal dome uplif t 7 9 previous model s o f Mesozoic-Cenozoic evolutio n 78-79 synrift an d post-rif t basi n evolution 78 Jurassic-Cenozoic uplift-subsidenc e reversa l 75 Cenozoic accumulatio n 77

INDEX Mesozoic structur e and evolution of Central Graben Dome 7 5 pre-Quaternary geology 7 6 structure and evolution of North Sea Basin 75-77 migration o f clinoform breakpoint s 216 seismic profile s 277 , 220 South Swedis h Dome 71-75 tectonic impact o n sedimentary processes 235-236, 263 Eocene tim e 248-250, 249 intra-plate stress 26 2 Late Paleocene-Early Eocene vertica l movements 261 Miocene tim e 254-258, 255 Neogene incised valley s 256 Oligocene tim e 250-251, 252, 254 Oligocene-Miocene domal uplift o f southern Norway 262 palaeo-water depth 240-243, 240, 241, 245, 249, 257 Plio-Pleistocene glacial erosio n an d uplift 26 2 Plio-Pleistocene time 257, 258-259, 255, 260 possible vertica l movemen t mechanisms 259-262 regional profile s 23 9 regional settin g in Early Oligocen e tim e 253 regional settin g in Late Paleocene-earliest Eocen e time 247 regional settin g in mid-Miocene tim e 255 seismic line s 246, 252,25 8 seismic mappin g 236-240, 238 stratigraphy 238 structural map 237 tectonic modelling an d subsidence analysis 242, 243-244, 244 upper crustal configuration 67-68 , 6 8 North Sea Basin 403, 423 North Se a Fan 740 , 374, 317 North Westray Ridge 47 5 Northern Highlands Massif 475 Northern Scandes 103-110 , 705, 112 , 77 8 Jurassic-Cretaceous denudation 125-132 , 726-727, 128-129, 130-131 Tertiary denudatio n 132-13 3 Norway 1 , 28 AFT analysis 30, 33 Bouguer gravity anomaly map 29 Cenozoic uplif t an d denudation 209-210, 229-230 climatic an d eustatic chang e 212-214 constraints on magnitude 222-224 constraints o n timing 22 4 definition 210-21 2 differences betwee n souther n Norway and Shetland Platform 228-229 hypothetical mode l 226-228 , 227 isostatic respons e t o localized denudatio n 224—226 isostatic uplif t respons e 21 2 study rationale 214-21 6 tectonic mechanisms 212 isostasy 3 4 mean elevation 30, 32 Neogene uplif t 31-3 4 calculated tectoni c uplif t an d isostatic rebound 36 coherence analysis 35, 35 estimating flexural rigidity 34-35 mechanisms 35-3 7 parameter value s 31 Oligocene-Miocene domal uplift 26 2

491

palaeic surfac e 108-109 , 70 9 compared wit h Swedish Muddus plains 109-11 0 Paleocene initiatio n of Cenozoic uplif t 45-^-6 , 60, 63-64 active mechanisms of surface uplift 54-6 0 Bouguer gravity map 46 chalk overburial 53 changing in-plane stress 54-57, 55, 56 climate and eustasy 46-47 gravity 63 heat flow 63 isostatic adjustment s cause d b y erosion o f existing topography 48, 49-51, 49, 50 isostatic an d erosional respons e t o sea-level fal l 51-54,57,52 lithospheric delamination 58-60 , 58 , 59, 62 magmatic underplating 57-58, 5 7 mechanism of surface uplift 61-63 passive mechanisms of surface uplift 49-5 4 pre-Quaternary geologica l ma p 54 sediment structur e and fission tracks 60-61 tectonics 47^49 thermal and topographic respons e o f lithosphere 57, 58, 59, 62 temperature-depth plot 3 7 timing of denudation 28-31 topographic cross-section s o f south Norwegian dome 225 topographic relie f ma p 47 topography, Moho and Bouguer gravity 48, 49, 50 Nyk Slid e 74 6 offshore sedimentar y response analysi s 6 Ogham Platform 479 oil reserves an d resources 401^4-0 2 Barents Se a 474, 416 East Irish Sea Basin 417-422, 479 Inner Moray Firt h Basin 417, 478 North Celtic Se a Basin 427, 424 resource estimatio n 409^12 risk analysis 402^09, 404 Slyne-Erris Basin 420, 422^24 West Shetland Basin 475, 416-417 Orcadian Basin 274 Orkney Islands 273, 478 Orkney-Shetland Platform 475, 478 Oryx field 479 Oskarshamn 77 0 Outer Hebrides Island s 27 3 partial annealin g zone (PAZ) 119-120, 125 Peach Slid e 375, 378 Pembroke Ridg e 427 Peterhead Ridg e 47 8 petroleum exploratio n 9-10 Phoenix field 478 Porcupine Abyssa l Plain 374 Porcupine Ban k 374 Porcupine Seabigh t 374 prospect resourc e estimatio n 409^410 , 409 formation volum e factor 411 gross rock volume 410 net-to-gross ratio , porosity an d hydrocarbon saturation 410-411 recovery factor 411^412

492

INDEX

prospect risk analysis 402-405 , 404 probability o f oil versus ga s 406-407 probability o f reservoir 40 5 probability o f source an d charge 405^406 probability o f trap and seal 407^09 Purbeck-Wight Disturbance 86 , 88, 89 Ranger field 419 Rannock Moor 27 7 Rathlin/Ulster Basin 372 Rayleigh-Taylor instability 58, 61, 62 repeat formatio n tests (RFTs ) 479 Rhum, Isle of 273 Ribban Basin 40 3 Ringk0bing Fy n High 18 5 Rockall Ban k 314, 315, 403, 420 Rockall Plateau 314, 317 Rockall Troug h 314, 315, 317, 318, 319, 320, 372, 420 Rona Ridge 292, 415 Rosemary Ban k 314 Ross field 418 Sand0y Ridge 41 5 Scandinavia, landform s an d uplift 103 , 104, 113-114, 117-118, 134-13 5 analysis o f map s comparison betwee n Norwegian Palaei c surfac e and Swedish Muddus plains 109-11 0 grus saprolites an d landforms 10 8 major shap e o f Northern an d Souther n Scande s 103-110,705 Palaeic surfac e o f southern Norway 108-109 , 10 9 palaeosurfaces forme d b y etching an d planatio n 106-109 plains wit h residual hill s 108 , 108 sub-Cambrian peneplai n 106-107 , 10 6 undulating hill y relie f 107-108 , 10 7 zone o f incised valley s 10 9 analysis of profile s 110 , 777 Scandinavian dome s 11 2 apatite fisson track (AFT ) analysis age an d track length statistic s 12 0 annealing 119-12 0 inverse modellin g 120-12 1 procedure 12 0 results 122-123, 123-124 , 72 4 geological settin g 118-11 9 hinge lin e 112-113 , 773 interpretation of AFT data an d inverse modellin g 124-125 correlation wit h offshore geology 13 4 Jurassic-Cretaceous denudation o f Northern Scandes 125-132 , 126-127, 128-129, 130-131 Lofoten an d Vesterale n islands 13 3 pattern of denudation 133-13 4 post-orogenic cooling 12 5 Tertiary denudatio n o f Northern Scande s 132-13 3 methods 10 3 Neogene uplif t induce d by Souther n Swedis h Dom e 183-186, 198 , 202-203 burial anomal y and missing sectio n 189-19 0 chronostratigraphic event definition 190-19 1 denudation surface s 198-200 , 799, 200-201 erosion estimate d fro m basi n modelling 190-19 3 erosion estimate d fro m soni c dat a 186-19 0 erosion mode l 191-19 3 heat-flow mode l 19 1

model calibratio n an d results 792 , 193, 194 palaeo-surface temperature s 191 , 797 quantification o f late Cenozoic erosio n i n Denmark 193-198, 79 5 sonic dat a 19 0 timing of Cenozoic uplif t 200-20 2 velocity-depth trends 186-189 , 186, 187, 188 pre-Quaternary geolog y 184 sampling 12 1 South Swedis h Dome 113 , 118 study are a 77 9 uplift o f the Scande s 11 2 Vaner Basi n 11 3 Schiehallion field 415 Scotland 1 , 27 3 Cenozoic uplif t an d denudatio n of the Highland s 1 , 287, 282-283 denudation and landscape evolutio n 283-285 denudational history 271-272, 285-286 former cove r rocks in the Highlands 279-282 Highlands 1 , 30 Palaeozoic-present histor y of the Highlands 272-279 morphotectonic unit s 277 post-Devonian depth s of denudation 274 tectonics and denudation in Tertiary Igneou s Province 278 Tertiary uplif t 27 6 Senja Faul t Zone 147 Senja Ridg e 414 Seven Head s field 421 Shetland Islands 27 3 Shetland Isles 41 5 Shetland Spin e Fault 415 Siljan rin g meteorit e impac t 10 4 Skagerrak-Kattegat Platfor m 68, 68, 69, 170, 185 Skeivi Bank 292 Skua field 478 Skye, Isle of 273, 276, 278 Slyne Basin 40 3 Slyne Ridge 420 Slyne Troug h 372,474 cap-rocks 476 Slyne-Ems Basin 420, 422-124, 423 Smith Bank Graben 41 8 Smith Bank Hig h 418 Sn0hvit field 414 initial reservoir pressure 478 underfilled trap s 478 Sogn Grabe n 23 7 Solan Basin 415 Solan field 415, 416 Sole Pi t faul t zon e 7 6 Solway Basi n 372 sonic velocit y analysis 6 Irish Se a Basin 355-356, 365-368, 365 , 366 curve-fitting, geologica l erro r an d derivation of exhumation 357 , 358 , 359 , 360-361 , 361-364 methods 357-364 results 362, 363, 364-365, 364 theoretical background and mechanical control over velocity 357-361 Sorgenfrei-Tornquist Zon e 185, 211 S0rvestnaget Basi n 414 South Celtic Sea 86 South Celtic Sea Basin 372, 421

493

INDEX South Halibut Basi n 41 8 South Hewitt Fault 86 South Morecamb e fiel d 47 8 initial reservoi r pressure 47 8 South Smalan d Peneplain 71 , 72 , 106, 170, 199 South Swedis h Dom e 68, 68, 69, 71-75, 72 , 73, 74, 105, 106 see also Swede n Neogene uplif t 183-186 , 198 , 202-203 burial anomal y an d missing sectio n 189-19 0 chronostratigraphic even t definition 190-19 1 denudation surface s 198-200 , 799, 200-201 erosion estimate d fro m basi n modellin g 190-19 3 erosion estimate d fro m soni c dat a 186-19 0 erosion mode l 191-19 3 heat-flow mode l 19 1 model calibratio n an d results 792, 193, 794 palaeo-surf ace temperatures 191 , 797 sonic data 19 0 timing of Cenozoic uplif t 200-20 2 velocity-depth trends 186-189 , 186, 187, 18 8 Southern Ga s Basin 47 4 Southern Nort h Se a Basin, cap-rocks 47 6 Southern Scande s 103-110 , 105, 112, 118 Southern Scande s Dom e 68 , 73 Spitsbergen 403, 474 St George's Channel 86 St George's Channel Basi n 372, 421 Stappen Hig h 414 Sticklepath Fault 86 Stord Basi n 237, 474 Storegga Slid e 140 , 140 stratigraphy 14 1 Storegga Slid e Comple x 140 , 141-142, 144, 147 north sid e 142-143 , 143, 14 6 south sid e 143-147 , 146 Strathmore field 475 Sub-Cambrian Peneplai n 106-107 , 106, 170 Suilven field 415 Sula Sgei r Fan 314, 315, 317 surface uplift , definitio n 4 Sweden see also Sout h Swedis h Dom e landscape surfac e reconstruction 15 8 fluvial inclusion during interglacial period s 158-161, 759 glacial erosio n b y selectiv e linea r erosio n 16 1 localized subglacia l modificatio n of preglacia l upland surface s 161-16 2 lowering o f preglacial uplan d surfaces by nonglacial processes 162 preglacial fluvia l drainag e syste m 158 Mesozoic and Cenozoic thermotectoni c developmen t 169, 177-180 , 180-181 Cenozoic exhumatio n 179-18 0 dating method 17 2 faults an d lineaments 171-172 , 777 forward modellin g 172-17 6 geological an d geomorphological modellin g constraints 17 6 large-scale Phanerozoi c tectonis m 18 0 modelled therma l historie s 7 74-7 75 modelling o f AFT data 172-17 7 modelling result s 173, 176-17 7 palaeosurfaces 169-17 1 Triassic an d Jurassic exhumation 177-179 , 778

Muddus plains 108 compared wit h Norwegian Palaei c relie f 109-11 0 slope ma p 10 8 Tampen Slid e 146 Tampen Spu r 237 Teusajaure valle y 75 9 Tornquist Zone 67-68 , 68 , 69, 70, 76, 105 Torridop field 475 Traenabanken Slid e 744, 746 Traenadjupet Slid e 140, 146, 147 Triple Junction Dome 68 , 68, 70 , 71 Troms0 Basi n 474 Tr0ndelag Platfor m 47 4 Tynwald Basin 479 Tynwald Fault 47 9 Uer Terrace 237 uplift, definitio n 4 , 210-212 Utsira High 23 7 Vale o f Pewsey Faul t 8 9 Vaner Basin 11 3 Vanern, Lake 77 0 Vattern, Lake 77 0 Vattern Grabe n 71 , 72 Vealevuomus glacial valley 16 0 velocity-depth trends 186-190 , 186, 187, 188 baseline fo r Lower Jurassi c marine shal e 20 4 baseline fo r Lower Triassic Bunter shale 204-205 revised tren d for North Se a Chalk 203-20 4 Vestbakken volcani c provinc e 47 4 Vesteralen Island 779 , 13 3 Vestfjorden Basi n 40 3 Victory field 475 capillary sea l capacit y 46 3 underfilled trap s 478 Viking Graben 68 , 68, 70 , 237 vitrinite reflectance (VR) 6 quantifying Atlanti c exhumatio n 331-332 analytical problems 337, 338-339 approaches base d o n heat flow 334 availability of palaeotemperature constraint s over a range o f depths 33 9 basic assumption s 332-333, 332 calibration o f system response 33 5 depth resolution versu s thermal resolutio n 34 4 Early Tertiary palaeo-therma l effects i n NW England 344-346, 345 factors affectin g th e accurac y o f the magnitud e of exhumation 332-33 9 factors affectin g th e precision o f the magnitud e of exhumation 339-34 0 heating rates 334-335 identifying th e appropriate unconformit y 335 inappropriate syste m kinetics 336, 337 integration of results from differen t method s 337-338 latest result s 344-34 9 limitations o f this technique 343-344 Mesozoic palaeotherma l episode s in Ireland an d Central Iris h Se a Basin 347-349 nature of palaeotemperature constraint s 339-340 new result s from Centra l Englan d 346-347 non-linear palaeogeothermal gradients 333-334 , 333

494

INDEX

vitrinite reflectance (VR) continued onset versus duration o f cooling o r exhumation 34 4 palaeo-surf ace temperature 33 4 reburial 344 thermal and denudational history o f Ireland 377-380 thermal history resolutio n 343-344 timing o f exhumation 34 7 V0ring Basin 140 V0ring Plateau 14 0 Wardour-Portdown structure 89 Watchet-Cothelstone Fault 86 , 89 Weald Basi n 86, 90-91, 92, 403

forward mode l 94-95, 94, 95, 96 inverse model 91-93, 93 West Bank Hig h 418 West Shetlan d Basi n 403, 415, 423, 47 4 cap-rocks 476 West Shetlan d Platform 415 West Shetlan d Shelf 314, 32 1 West Shetlan d wedge 314, 315, 31 7 Westray Ridge 292 Westray Transfer Zone 29 2 Witchground Graben 23 7 Wyville-Thomson Ridge 314 Wyville-Thomson Ridge Complex 292