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 ( 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.

14

S. M. JONES ETAL.

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.

16

S. M. JONES ETAL.

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 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 ( 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

125

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

126

B. W. H. HENDRIKS & P. A. M. ANDRffiSSE N

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

128

B. W. H. HENDRIKS & P. A. M. ANDRIESSE N

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

DENUDATION HISTOR Y O F NORTHERN SCANDINAVI A

129

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

132

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

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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(eds) Petroleum Geology o f North West Europe. Graham & Trotman , London , 1131-1148 . FALEIDE, J.I. , VIGNES, E. & GUDLAUGSSON , S.T. 1993 . Late Mesozoic-Cenozoi c evolutio n o f th e south western Barent s Se a i n a regiona l rift-shea r tectonic setting . Marine an d Petroleum Geolog\, 10, 186-214 . FLEISCHER, R.L. , PRICE , P.B . & WALKER , R.M . 1975 . Nuclear Tracks in Solids: Principles and Applications. Universit y o f Californi a Press , Berkeley , CA. GALBRAITH, R.F . 1981 . O n statistica l model s fo r fission track counts. Mathematical Geology, 13 (6), 471-478. GALLAGHER, K . 1995 . Evolvin g temperature historie s from apatit e fission-track data. Earth and Planetary Science Letters, 136 , 421-435 . GALLAGHER, K . & BROWN , R . 1997 . Th e onshor e record o f passiv e margi n evolution . Journal o f th e Geological Society, London, 154 , 451-457 . GALLAGHER, K . & BROWN , R . 1999 . Denudatio n an d uplift a t passive margins: the record o n the Atlantic Margin o f souther n Africa . Philosophical Transactions o f the Roval Society o f London, Part A, 357 , 835-859. GALLAGHER, K. , BROWN , R . & JOHNSON , C . 1998 . Fission trac k analysi s an d it s application s t o geological problems . Annual Review o f Earth an d Planetary Sciences, 26, 519-572. GEOLOGICAL SURVEY S O F FINLAND , NORWA Y AN D SWEDEN. 1987 . Geological Ma p o f Northern Fennoscandia, 1:1000000 . Geologica l Survey s of Finland, Norway and Sweden , Helsinki. GLEADOW, A.J.W . & DUDDY , I.R . 1981 . A natura l long-term trac k annealin g experiment fo r apatite . Nuclear Tracks, 5 , 169-174 . GLEADOW, A.J.W . & BROWN, R.W. 2000. Fission-track thermochronology an d th e long-ter m denudational response t o tectonics . In : SUMMERFIELD , M.A . (ed.) Geomorphology an d Global Tectonics. Wiley, Chichester, 57-75. GLEADOW, A.J.W. , DUDDY , I.R. , GREEN , P.P . & HEGARTY, K.A . 19860 . Fissio n trac k length s i n the apatite annealing zone an d the interpretation of mixed ages . Earth an d Planetary Science Letters, 78, 245-254. GLEADOW, A.J.W. , DUDDY , I.R. , GREEN , P.P . & LOVERING, J.F . 1986£ . 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. GRIST, A.M . & RAVENHURST , C.E . 1992 . A step-by step laborator y guid e t o fissio n trac k thermo chronology a t Dalhousi e University . In: ZENTILLI , M. & REYNOLDS , PH . (eds ) Lo w Temperature Thermochronology. Shor t Cours e Handbook , Mineralogical Associatio n o f Canada , 20 , 189-201. GUDMUNDSSON, A . 1999 . Postglacia l crusta l doming , stresses an d fractur e formatio n wit h application to Norway. Tectonophysics, 307 , 407-419 . HANSEN, K. , PEDERSEN, S. , FOUGT, H. & STOCKMARR , P. 1996 . Post-Sveconorwegia n denudatio n an d cooling histor y o f Evj e area , souther n Setesdal ,

DENUDATION HISTORY O F NORTHERN SCANDINAVIA Central Sout h Norway . Norges Geologiske Unders0kelse Bulletin, 431, 49-58 . HURFORD, AJ . & GREEN , P.P . 1983 . Th e zet a ag e calibration o f fission-trac k dating . Isotope Geoscience, 1 , 285-317. KETCHAM, R.A., DONELICK , R.A . & CARLSON , W.D . 1999. Variabilit y o f apatit e fission-trac k annealin g kinetics: III . Extrapolatio n t o geologica l tim e scales. American Mineralogist, 84 , 1235-1255 . KETCHAM, R.A. , DONELICK , R.A . & DONELICK , M.B. 2000. AFTSolve : a progra m fo r multi-kineti c modeling o f apatit e fission-trac k data . Geological Materials Research, 2, 1-32. LARSON, S.E. , TULLBORG , E.-L. , CEDERBOM , C . & STIBERG, J.-P . 1999 . Sveconorwegia n an d Caledo nian forelan d basin s i n th e Balti c Shiel d reveale d by fission-track thermochronology. Terra Nova, 11, 210-215. LASLETT, G.M. , GREEN , PR , DUDDY , LR . & GLEADOW, AJ.W . 1987 . Therma l annealin g o f fission tracks i n apatit e 2 . A quantitative analysis . Chemical Geology, 65, 1-13 . LEHTOVAARA, J . 1976 . Apatit e fissio n trac k datin g o f Finnish Precambria n intrusives . Annales Academiae Scientiarum Fennicae, 117 , 1-94 . LIDMAR-BERGSTROM, K . 1993 . Denudatio n surface s and tectonics in the southernmost part of the Baltic Shield. Precambrian Research, 64 , 337-345. 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. LUTZ, T.M . & OMAR , G . 1991 . A n invers e method o f modeling thermal historie s fro m apatit e fission track data . Earth an d Planetary Science Letters, 104, 181-195 . NAESER, C.W. 1981 . The fading o f fission tracks in the geologic environment—dat a fro m dee p dril l holes . Nuclear Tracks, 5 , 248-250. OLESEN, O. , TORSVIK , T.H. , TVETEN , E. , ZWAAN , K.B., L0SETH , H . & HENNINGSEN , T . 1997 . Basement structur e o f th e continenta l margi n i n the Lofoten-Lopphave t area , norther n Norway : constraints fro m potentia l fiel d data , on-lan d structural mappin g an d palaeomagneti c data . Norsk Geologisk Tidsskrift, 77 , 15-30 . REEMST, P . 1995 . Tectonic modelling o f rifted continental margins —basin evolution and tectono-magmatic development of the Norwegian and N W Australian margin. Ph D thesis , Vrij e Universiteit Amsterdam .

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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. ROBERTS, D . & GEE, D.G . 1985 . An introduction to the structure of the Scandinavian Caledonides. In: GEE , D.G. & STURT , B.A . (eds ) Th e Caledonide Orogen—Scandinavia an d Related Areas. Wiley , Chichester, 55-68 . ROHRMAN, M . 1995 . Thermal evolution o f th e Fennoscandian region from fission track thermochronology—an integrated approach. Ph D thesis , Vrije Universitei t Amsterdam. ROHRMAN, M . & VA N DER BEEK, P . 1996 . Cenozoi c postrift doma l uplift o f North Atlanti c margins : a n asthenospheric diapiris m model . Geology, 24 , 901-904. SRIVASTAVA, S.P . & TAPSCOTT , C.R . 1986 . Plat e kinematics o f the North Atlantic. In: VOGT, PR. & TUCHOLKE, B.E . (eds ) Th e Western North Atlantic Region, M . Geologica l Societ y o f America , Boulder, CO, 379-404. 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. STUWE, K. , WHITE , L . & BROWN , R . 1994 . Th e influence o f erodin g topograph y o n steady-stat e isotherms. Applicatio n t o fissio n trac k analysis . Earth an d Planetary Science Letters, 124 , 63-74 . SUMMERFIELD, M.A . & BROWN , R.W . 1998 . Geo morphic 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 in Fission-Track Geochronology. Kluwer, Dordrecht, 269-284 . VAN DE R BEEK , PA . 1995 . Tectonic evolution o f continental rifts —inferences from numerical modelling and fission track thermochronology. PhD thesis, Vrije Universitei t Amsterdam . WAGNER, G.A . 1968 . Fissio n trac k datin g of apatites . Earth an d Planetary Science Letters, 4 , 114-145 . WAGNER, G.A . & VA N DEN HAUTE, P. 1992 . FissionTrack Dating. Kluwer , Dordrecht. ZECK, H.P. , ANDRIESSEN , P.A.M. , HANSEN , K. , JENSEN, PK. & RASMUSSEN , B.L . 1988 . Paleozoi c paleo-cover o f th e souther n par t o f th e Fennoscandian shield—fissio n trac k constraints . Tectonophysics, 149 , 61-66 . ZIEGLER, P . 1987 . Evolutio n o f th e Arctic—Nort h Atlantic Borderlands . In : BROOKS , J . & GLENNIE , K. (eds ) Petroleum Geology o f North West Europe. Graham & Trotman, London , 1201-1204 . ZWAAN, K.B. , FARETH , E . an d GROGAN , P.W . 1998 . Geologisk kart over Norge, berggrunnskart Troms0, M 1:25 0 000 . Norwegia n Geologica l Survey, Trondheim .

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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

LATE NEOCENE EVOLUTION O F MID-NORWEGIAN MARGI N

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. )

LATE NEOCEN E EVOLUTION OF MID-NORWEGIAN MARGI N

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) .

LATE NEOCEN E EVOLUTION OF MID-NORWEGIAN MARGI N

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

LATE NEOCEN E EVOLUTION O F MID-NORWEGIAN MARGI N

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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and th e M0r e Basin ; th e influenc e o f basemen t structural grain , an d th e particula r rol e o f th e M0re-Tr0ndelag Faul t Complex . Marine an d Petroleum Geology, 16, 443-465. GRADSTEIN, F . & BACKSTROM , S.A . 1996 . Cenozoi c biostratigraphy an d palaeobathymetry , norther n North Se a an d Haltenbanken . Norsk Geologisk Tidsskrift, 76 , 3-32 . GUDMUNDSSON, A . 1999 . Postglacia l crusta l doming , stresses an d fractur e formatio n wit h application to Norway. Tectonophysics, 307, 407-419 . HAFLIDASON, H., AARSETH, I., HAUGEN, J.-E., SEJRUP , H.P., L0VLIE , R . & REITHER , E . 1991 . Quaternar y stratigraphy o f th e Drauge n area , mid-Norwegia n shelf. Marine Geology, 101, 125-146 . HAFLIDASON, H., SEJRUP , H.P., BRYN , P . & MIENERT , J. 2000 . Glid e plane s an d slid e frequenc y i n th e Storegga Slid e Comple x of f mid-Norway . Geoscience, 2000, 16. HENDRIKSEN, S . & VORREN , T . 1996 . Lat e Cenozoi c sedimentation an d uplif t histor y o n th e mid Norway continenta l shelf . Global an d Planetary Change, 12 , 171-199. HENRIET, J.P. ; MIENERT , J . 1998 . Ga s Hydrates: Relevance to World Margin Stability and Climatic Change. Geologica l Society , London , Specia l Publications, 137 . JAPSEN, P . & CHALMERS , J.A . 2000 . Neogen e uplif t and tectonics around the North Atlantic: overview. Global and Planetary Change, 24, 165-174 . KING, E.L., SEJRUP , H.P., HAFLIDASON, H., ELVERHOI, A. & AARSETH , I . 1996 . Quaternar y seismi c stratigraphy o f th e Nort h Se a Fan : glacially-fe d gravity flo w aprons , hemipelagi c sediments , an d large submarin e slides . Marine Geology, 130 , 293-315. KRISTOFFERSON, Y. , WINTERHALTER , B . & SOLHEIM , A. 2000 . Shel f progradatio n o n a glaciate d continental margin, Queen Maud Land, Antarctica. Marine Geology, 165, 109-122 . LABERG, J.S. , VORREN , TO. , DOWDESWELL , J.A. , KENYON, N.H . & TAYLOR , J . 2000 . Th e And0y a Slide an d th e And0y a Canyon , north-easter n Norwegian-Greenland Sea . Marine Geology, 162, 259-275 . LABERG, J.S. , VORREN , V.O. , MIENERT , J. , KENYON , N.H., EVANS , D. , HENDRIKSEN , S . & DOWDESWELL, J.A . 1999 . Th e Traenadjupe t Slide area offshor e Norway . In : MARTINSEN , O.J . & DREYER, T . (eds ) Sedimentary Environments Offshore Norway —Palaeozoic t o Recent. Norwegian Petroleum Society, Extended Abstracts. 223-226. LARTER, R.D . & BARKER , P.P . 1991 . Neogen e interaction o f tectonic an d glacia l processe s a t th e Pacific margi n o f th e Antarcti c Peninsula . In : MACDONALD, D.I.M . (ed. ) Sedimentation, Tectonics and Eustasy: Sea-level Changes at Active Margins, Internationa l Associatio n o f Sedimentologists, Specia l Publications , 12 , 165-186. LlDMAR-BERGSTROM, K. , OLLIER , C.D . & SULEBAK ,

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LATE NEOCENE EVOLUTION O F MID-NORWEGIAN MARGI N MANGERUD, J. , JANSEN , E . & LANDVIK , J.Y . 1996. Late Cenozoi c histor y o f th e Scandinavia n an d Barents Se a ic e sheets . Global an d Planetary Change, 12, 11-26. MCNEILL, A.E. , SALISBURY , R.S.K. , 0STMO , S.R. , LIEN, R . & EVANS , D . 1998 . A regiona l shallo w stratigraphic framewor k of f mi d Norwa y an d observations o f dee p wate r 'specia l features' . OTC paper 8639 presented at the 1998 Offshore Technology Conference, Houston , TX, May 4-7 . MUIR WOOD , R . 1988 . Extraordinar y deglaciatio n reverse faultin g i n norther n Fennoscandia . In : GREGERSEN, S . & BASHAM , RW . (eds ) Earthquakes at North Atlantic Passive Margins: Neotectonics and Postglacial Rebound. Proceedings of the NATO Advanced Research Workshop, Vordingbord, Denmark, 9-13 Ma y 1988. Kluwer , Dordrecht, 141-173 . POOLE, D.A.R . & VORREN , T.O . 1993 . Miocen e t o Quaternary palaeoenviroment s an d uplif t histor y on the mid-Norwegian shelf. Marine Geology, 115, 173-205. 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. Rns, F. & FELDSKAAR, W. 1992. On 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 . & TELLERAAS , E . (eds ) Structural and Tectonic Modelling and its Application to Petroleum Geology. Elsevier , Amsterdam , 163-185. ROKOENGEN, K. , RlSE , L. , BRYN , P. , FRENGSTAD , B.,

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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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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

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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 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.

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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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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

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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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185

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P. JAPSEN ETAL.

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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('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 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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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 (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 )

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(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.

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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).

SCOTLAND'S DENUDATIONA L HISTORY

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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 (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) .

DEPRESSURIZATION OF HYDROCARBON-BEARING RESERVOIRS

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

475

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 .

476

A. CORCORAN & A.G. DOR E

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