Late Cretaceous–Paleocene tectonic development of the NW Vøring Basin

Marine and Petroleum Geology 20 (2003) 177–206 www.elsevier.com/locate/marpetgeo

Late Cretaceous – Paleocene tectonic development of the NW Vøring Basin Shicun Rena,b,*, Jan Inge Faleidea, Olav Eldholma, Jakob Skogseidc, Felix Gradsteind a

Department of Geology, University of Oslo, P.O. Box 1047 Blindern, N-0316 Oslo, Norway b Norsk Hydro ASA, P.O. Box 7190, N-5020 Bergen, Norway c Norsk Hydro ASA, N-0246 Oslo, Norway d Paleontological Museum, University of Oslo, P.O. Box 1172 Blindern, N-0318 Oslo, Norway Received 18 June 2001; received in revised form 31 December 2002; accepted 11 January 2003

Abstract The Late Cretaceous – Paleocene rifting in the NW Vøring Basin is characterized by four main fault complexes and pronounced uppercrustal structural segmentation. The fault complexes are linked by accommodation zones, which separate fault systems of different polarities and thick from thinner coeval sedimentary successions. Structural and stratigraphic analyses suggest that the early rift phase (,81 to 65 Ma) was characterized by large-scale normal faulting, along-margin segmentation and varying structural styles; whereas the late rift phase (, 65 to 55 Ma) was associated with continued extension, regional uplift, intrusive igneous activity and subsequent erosion. The rifting ended with breakup at ,55 Ma accompanied by massive, but gradually waning extrusive igneous activity over the next 3 Myr. The mode of rifting appears to have changed from brittle to more ductile extensional deformation from the early to late rift phase. The changing rift rheology is probably related to the arrival of the Iceland mantle plume and initiation of associated igneous activity. Hence, the NW Vøring Basin provides an example of complex interaction of structural and magmatic relationships during rifting and breakup. q 2003 Elsevier Ltd. All rights reserved. Keywords: Continental margin; Offshore Norway; Rifting; Structure

1. Introduction The Vøring continental margin off Norway (Fig. 1), characterized by massive intrusive and extrusive activity during late rifting and breakup, is commonly considered a volcanic end-member of rifted passive continental margins (e.g. Eldholm, Skogseid, Planke, & Gladczenko, 1995). Due to the large amount of high-quality seismic reflection data, numerous wide-angle velocity experiments and several deep boreholes, the margin is considered a natural laboratory to study the tectono-magmatic processes and their interplay during rifting, breakup and volcanic margin formation. However, the relationship of the tectonic events leading to breakup and the voluminous breakup magmatism has been a matter of contention; both non-extensional (Hinz et al., 1987; Mutter, Buck, & Zehnder, 1988) and extensional (e.g. Ren, * Corresponding author. Present address: Norsk Hydro ASA, P.O. Box 7190, N-5020 Bergen, Norway. Tel.: þ 47-55995459; fax: þ 4755996639. E-mail address: [email protected] (S. Ren). 0264-8172/03/$ - see front matter q 2003 Elsevier Ltd. All rights reserved. doi:10.1016/S0264-8172(03)00005-9

Skogseid, & Eldholm, 1998; Skogseid, Pedersen, Eldholm, & Larsen, 1992a) settings have been discussed. Previously, we have documented a Late Cretaceous– Paleocene rift episode in the northern Fenris Graben – Gjallar Ridge area (Fig. 1), inferred to be active , 20 Myr prior to breakup at , 55 Ma (Ren et al., 1998). Nonetheless, questions remain about the timing and sequence of tectonomagmatic events in the outer Vøring Basin. In this study, we focus on these topics by interpreting structural elements and seismic stratigraphy of the NW Vøring Basin based on recently available stratigraphic information. In particular, we address the timing of the onset of rifting and the subsequent vertical motion history as an input to a tectonomagmatic model for the area.

2. Geological setting The Vøring margin comprises three geological provinces: The Vøring Marginal High, Vøring Basin and

178

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

Fig. 1. Main tectono-magmatic features on the Vøring margin based on this study, Blystad et al. (1995), Eldholm et al. (2002) and Fichler et al. (1999). Main study area within box. The area west of the Fles Fault Complex experienced significant Late Cretaceous– Paleocene deformation. Filled and open fault symbols refer to Late Jurassic–Cretaceous and Late Cretaceous –Paleocene structures, respectively. BFZ, Bivrost Fracture Zone; BL, Bivrost Lineament; CJMFZ, Central Jan Mayen Fracture Zone; FSE, Faeroe-Shetland Escarpment; GFZ, Gleipne Fracture Zone; GL, Gleipne Lineament; LVM, Lofoten–Vestera˚len Margin; MOR, mid-ocean ridge; RAZ, Rym accommodation zone; SL, Surt Lineament; VE, Vøring Escarpment.

Trøndelag Platform (Fig. 1), reflecting several periods of lithospheric extension since the end of the Caledonian Orogeny. Regional rifting episodes occurred in Carboniferous to Permian, in late Middle Jurassic to Early Cretaceous, and in Late Cretaceous to Early Cenozoic times (e.g. Blystad et al., 1995; Brekke, 2000; Brekke, Dahlgren, Nyland, & Magnus, 1999; Bukovics, Cartier, Shaw, & Ziegler, 1984). The Vøring Basin, bounded by the Bivrost and Jan Mayen lineaments of Blystad et al. (1995), developed by regional subsidence subsequent to the Late Jurassic – Early Cretaceous extension episode (Skogseid, Pedersen, & Larsen, 1992b). However, the definition and the timing of the various tectonic events remain a matter of debate. For example, Brekke et al. (1999) suggested three main tectonic phases: (1) extension in late Middle Jurassic to late Cenomanian, (2) extension and compression in late Cenomanian to early Paleocene, and (3) extension and compression from early Paleocene to the present. On the other hand, Dore´ et al. (1999) and Lundin and Dore´ (1997) divided the Late Jurassic –Early Cretaceous extension episode into two separate tectonic phases (Late Jurassic and Early Cretaceous), suggesting that the Vøring Basin mainly formed by subsidence following Early Cretaceous extension. In addition to Late

Cretaceous – Paleocene extension, a mid-Cretaceous extensional phase has been proposed (Dore´ et al., 1999; Lundin & Dore´, 1997). Structurally, the Vøring Basin can be divided into two sub-provinces by the Fles Fault Complex which is mainly a Late Jurassic – Early Cretaceous feature, but reactivated during the Late Cretaceous– Paleocene rifting (Fig. 1). The eastern sub-province comprises the Træna and Ra˚s basins, whereas two prominent structural highs, the Gjallar Ridge and Nyk High, separate, respectively, the Vigrid and Na˚grind synclines from the Fenris and Hel grabens in the western sub-province. In the NW Vøring Basin, the early phase of the Late Cretaceous – Paleocene rifting was characterized by lowangle normal faulting, subsidence and syn-rift sedimentation. Later, the arrival of the Iceland mantle plume initiated large igneous activity and subsequent magmatic underplating, which resulted in regional uplift and erosion in Paleocene time. The final rifting culminated with complete lithospheric separation accompanied by massive igneous activity at the Paleocene – Eocene transition (Eldholm, Thiede, & Taylor, 1989). The intrusive activity in the upper crust was most intense during the late Paleocene, whereas the later breakup flood basalt volcanism

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

constructed the Vøring Marginal High and up to 70 km wide extrusive complex between the Vøring Escarpment and the flood basalt boundary (Fig. 1). After breakup, the Vøring margin experienced thermal subsidence and modest sedimentation during a generally cooling climate, followed by Plio-Pleistocene uplift of Fennoscandia and the onset of Northern Hemisphere glaciations resulting in greatly increased erosion and sedimentation (Hjelstuen, Eldholm, & Skogseid, 1999).

3. Seismic interpretation and stratigraphy The seismic interpretation is based on the study of 10,500 km time-migrated, 8 s record length, multi-channel seismic (MCS) profiles (VB-87, VB-89, VB-90 and GVN92 surveys). Line spacings are , 2.5 to 10 km and , 4 to 12.5 km for dip and strike lines, respectively. The seismic

179

resolution generally decreases with depth, whereas the presence of igneous intrusive and extrusive complexes makes a detailed interpretation of the pre-mid-Cretaceous sections ambiguous. Most commercial wells on the Vøring margin have been drilled in the Trøndelag Platform –Halten Terrace region (Fig. 1), resulting in a well known post-Lower Triassic stratigraphy (e.g. Dalland, Worsley, & Ofstad, 1988; Gradstein & Ba¨ckstro¨m, 1996; Gradstein, Kaminski, & Agterberg, 1999). In addition to eight relatively shallow DSDP and ODP holes on the Vøring Plateau (Eldholm, Thiede, & Taylor, 1989), only a few shallow and five deep commercial wells have yet been drilled in the Vøring Basin proper (Figs. 1 and 2). Moreover, drilling samples are mostly cuttings with only few cores. In this study, we have focused on six regional seismic reflectors (Fig. 3). These are: base Pleistocene (BPl), base Pliocene (BP), top Eocene (TE), top Paleocene

Fig. 2. Depositional environment (left column) and simplified lithology (right column) in five commercial wells. Red lines indicate correlation of three stratigraphic markers associated with the Late Cretaceous–Paleocene rifting: TPa, top Paleocene; TC, top Cretaceous and IMC, intra middle Campanian.

180

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

Fig. 3. Simplified lithology and sonic log of well 6607/5-2 correlated to seismic reflectors and sequences in the NW Vøring Basin. Lithostratigraphy from Dalland et al. (1988). BPl, base Pleistocene; BP, base Pliocene; MM, Middle Miocene; IO, intra Oligocene; TE, top Eocene; TPa, top Paleocene; TC, top Cretaceous; IEM, intra early Maastrichtian; ILC, intra late Campanian; IMC, intra middle Campanian; IS, intra Santonian.

(TPa), top Cretaceous (TC), and intra middle Campanian (IMC). In addition, middle Miocene (MM), intra Oligocene (IO), intra early Maastrichtian (IEM), intra late Campanian (ILC) and intra Santonian (IS) reflectors have been interpreted locally (Fig. 3). The stratigraphic and lithological control is achieved by correlation to the five deep commercial wells in the province (Figs. 2 and 4). Because these wells are located on structural

highs and provide relatively poor basin-wide stratigraphic control, the interpretation is also based on seismic character and regional correlations of Blystad et al. (1995) and Skogseid and Eldholm (1989). Because of the distance to the drilled sites, the stratigraphy in the Hel Graben is principally based on a shallow borehole, 6707/ 04-J-01 (Fig. 1), as well as correlation of seismic character to the Nyk High and Gjallar Ridge.

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

181

Fig. 4. Commercial wells projected onto seismic profiles across (a) Nyk High, Na˚grind Syncline and Utgard High; (b) northern Gjallar Ridge, Vigrid Syncline, Vema Dome and Nyk High. Most wells are located on the line, and well 6607/5-1 is 5.2 km to SE of the line. Location in Fig. 1. Reflector code in Fig. 3.

Fig. 5a and b show typical basin sequences and main structural elements on two dip lines across the study area. The Late Cretaceous – Paleocene normal faults are well expressed along the western Hel Graben and Nyk High (Fig. 5a), and in the Gjallar Ridge – Fenris Graben province and in the Fles Fault Complex, on each side of the Vigrid Syncline, respectively (Fig. 5b). The profile in Fig. 5c represents a tie between wells in the Vema Dome –Nyk High region and northern Gjallar Ridge. Stratigraphic markers associated with the Late Cretaceous – Paleocene rifting are the intra middle Campanian, top Cretaceous and top Paleocene reflectors (Fig. 2). The intra middle Campanian reflector marks the upper boundary of a prominent high-amplitude seismic package (Kittilsen, Olsen, Marten, Hansen, & Hollingsworth, 1999) of Santonian to middle Campanian age. The top Cretaceous reflector represents a regional erosional unconformity, and is onlapped by the Paleocene sediments. Finally, reflector IF (inner flows), marks the top breakup lavas landward of the Vøring Escarpment, and we recognize abundant sills and/or low-angle dykes within the Cretaceous and pre-Cretaceous sequences (Fig. 5).

4. Main structural elements Structurally, the study area is dominated by four main fault complexes in the vicinity of the northern Gjallar Ridge, Hel Graben, Nyk High and North Vema Dome, and the Rym accommodation zone located at the central part of the Rym Fault Zone of Blystad et al. (1995) (Figs. 1 and 6). In the northern Gjallar Ridge and Nyk High areas, our structural map is more detailed, but generally consistent with the regional map of Blystad et al. (1995). Pronounced differences exist, however, in the Hel Graben and Vema Dome areas. Moreover, it appears to be little evidence to relate the Rym accommodation zone with the Surt Lineament. Our interpretation of the seismic, magnetic and gravity data does not support a deep crustal lineament extending across the entire outer Vøring Basin. The structural mapping and gravity data suggest, on the other hand, the Gleipne Lineament as a diffuse structural lineament landward of the Gleipne Fracture Zone (Figs. 1 and 6). The lineament was also, based on gravity and magnetic data, inferred by Fichler et al. (1999). The fracture zone (originally called Lofoten Fracture Zone) was first

182

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

Fig. 5. Typical profiles across NW Vøring Basin: (a) Vøring Escarpment (VE) –Hel Graben–Naglfar Dome–Nyk High; (b) Vøring Escarpment –Fenris Graben– northern Gjallar Ridge–Vigrid Syncline–Fles Fault Complex; (c) Vigrid Syncline–Vema Dome– Nyk High. Location in Fig. 1. Reflector code in Fig. 3.

mapped by Hagevang, Eldholm, and Aalstad (1983), as a left-lateral offset of the oldest sea floor spreading anomalies, 24A – B. In the Vøring Basin, the number of faults and fault displacements decrease significantly towards the Gleipne Lineament along which some smaller landward dipping faults are present (Figs. 1 and 6). Furthermore, the lineament is associated with a distinct eastward shift of the lava front (Fig. 6). Hence, it may have governed the segmentation and topography of the Gjallar Ridge, and the distribution of the Paleocene lavas.

We divide the study area between the Gleipne and Bivrost lineaments, into ridges and sub-basins on either side of the Rym accommodation zone. In the south, the NE trending Gjallar Ridge is separated from the basin fill in the Fenris Graben by a boundary fault (Ga, Figs. 6 and 7). Regionally, the ridge comprises a northeastward narrowing fault complex of seaward dipping listric normal faults with a characteristic zigzag pattern, and locally by interlocking crooked X shapes (Morley, 1995) (Figs. 5a, 6 and 8a). The easternmost faults (Ge-f) appear older than the crestal faults (Ga-d), and

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206 Fig. 6. Structural map of Late Cretaceous–Paleocene structures. The outline of the Utgard High from Blystad et al. (1995). MCS data coverage in thin black lines, other interpreted MCS profiles, are shown in subsequent figures. GFZ, Gleipne Fracture Zone; Rym AZ, Rym accommodation zone.

183

184

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

Fig. 7. Stack of interpreted seismic profiles in the northern Fenris Graben/Gjallar Ridge region showing changes of stratigraphic and structural styles. Location and fault annotation in Fig. 6. Reflector code in Fig. 3. MCD, middle crustal dome.

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

Fig. 7 (continued )

185

186

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

the paleotopographic high is shifted westward (Fig. 6), implying a westward propagation of fault activity with time. The onlapping Upper Cretaceous– Paleocene Vigrid Syncline strata pinch out along the eastern ridge flank (Figs. 5a and 8c). Only a thin Paleocene sequence, 150– 450 m, is present above the thick Fenris Graben basin fill. The structural and stratigraphic styles change along the Gjallar Ridge. In the southwest, the faults are low-angle, and detach at , 5.5 s (, 7 km) above a middle crustal dome structure (profiles G1-5, Fig. 7; Figs. 6 and 9) which is associated with a gravity high. The dome has been interpreted as a metamorphic core complex caused by isostatic unroofing of the middle– lower crust in response to low-angle detachment faulting, analogical to the nonvolcanic West Iberian margin core complexes (Ren et al., 1998). Note that the base of our mid-crustal dome (MCD; Fig. 7) corresponds to the top of the core complexes proposed by Lundin and Dore´ (1997) to be the result of middle– upper crustal anatexis induced by magmatic underplating. The fault blocks are strongly rotated and deeply eroded reflecting repeated activation since mid-Campanian time until the latest Cretaceous. Locally, a very low-angle fault cuts the older structures, and the top Cretaceous unconformity appears as a flattened erosional surface (Fig. 8a and b). Syn-rift sedimentation is well preserved in profile G1 (Fig. 8b), but absent on top of the Gjallar Ridge

(e.g. profiles G3 and G4). The middle crustal dome disappears towards the Rym accommodation zone corresponding to a shift in trend of the paleotopographic high (Fig. 6). In this region, most faults become high-angle and apparently undetached with decreased displacement (profiles G6-10, Fig. 7). Nonetheless, we observe several phases of syn-rift sedimentation and increased basin fill along the boundary fault (Fig. 8d). The northern part of the Vema Dome is a topographic high at middle Oligocene level over a core of tilted and nearly symmetric fault blocks with abundant intrusions (Figs. 9 and 10). In map view, some of the faults appear as semi-circular features (Fig. 6), while cross-sections show that the highly reflective Santonian-middle Campanian sequence is absent on the hanging wall and that several small-offset faults have formed antithetically against the fault plane (Fig. 9). Most faults cut both Upper Cretaceous and Paleocene strata, whereas other only extend into the Santonian-middle Campanian sequence. At shallow levels the Vema Dome is also characterized by extensive shale diapirism, in places piercing the sea floor. Diapirs are found at several stratigraphic levels within the Quaternary and Pliocene sediments (Hjelstuen, Eldholm, & Skogseid, 1997; Hovland, Nygaard, & Thorbjørnsen, 1998). We divide the study area north of the Rym accommodation zone into the Hel Graben –Naglfar Dome and Nyk

Fig. 8. Seismic examples in the northern Fenris Graben/Gjallar Ridge region showing (a) erosional surface and unconformities both at top Paleocene (TPa) and near top Cretaceous (TC); (b) syn-rift sedimentation (shaded) above intra middle Campanian reflector (IMC); (c) onlap of sediments onto TC, IMC and intra Cenomanian (ICen) reflectors; (d) significant growth along the Fenris Graben boundary fault. IF, inner flows. Location in Fig. 7.

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

187

Fig. 9. Seismic example and interpretation showing the complicated structure on the Vema Dome. Note that the intra middle Campanian reflector (IMC) is not observed on the hanging wall of the curved fault plane. Location in Fig. 6. Reflector code in Fig. 3.

High – Na˚grind Syncline structural provinces. The Hel Graben –Naglfar Dome is characterized by: (1) two normal fault systems (Figs. 11 and 12), (2) unconformities of middle Campanian and latest Cretaceous age (Figs. 11 and 13), (3) thick middle Campanian to latest Cretaceous basin fill increasing in thickness to the northeast, (4) inverted, lens-shaped lower Maastrichtian and Paleocene sequences (Figs. 5b and 11), and (5) pervasive sill intrusions (Figs. 5b, 11 and 12). The most prominent fault system is found in the northern flank of the Hel Graben, consisting of a series of major WSW-to-W trending and widely spaced listric normal faults (Ha – g, Fig. 6) east of the Vøring Escarpment. Fault Ha, characterized by rollover of the intra middle Campanian reflector, is the main northern boundary fault of the Hel Graben (Figs. 11 and 12). In contrast to the low –angle faults of Gjallar Ridge, the faults display a classic and idealized convex map pattern. A seaward dipping fault, Hi (Fig. 6), divides the fault system. South of the east trending part of Hi, we observe simple half-graben geometries, and the fault displacement at the intra middle Campanian horizon is distributed over four well-developed curved fault planes (Ha– d), reaching

a maximum throw of , 0.55 s twt (0.9 km, profiles H3-4, Fig. 11; Fig. 12b). To the north, the fault system becomes more complex and Hi changes trend bounding a small elongate basin, and cross-cuts a landward dipping fault Hh (profiles H5-9, Fig. 11; Fig. 6). Moreover, fault Hi is associated with significant flank uplift and erosion of part of the Paleocene sequences (Fig. 12c). Farther north, a , 7 km wide fault block between faults Hf and Hg (profile H8, Fig. 11) includes three Late Cretaceous sequences showing growth pattern against fault Hg (Fig. 13), indicating multi-phase fault movements. The other Hel Graben – Naglfar Dome fault system comprises a large number of small-offset faults dipping antithetically or synthetically to faults of the first system (Figs. 11 and 12). Most faults are present within the postintra middle Campanian sequences, while others only appear in Paleocene and lower Eocene sequences. The Nyk High, defined at the Upper Cretaceous level, is trending parallel to the Gjallar Ridge and bounded by a main boundary fault (Na) toward the Hel Graben, and locally by NW (Nc, Nf) and SE (Ne) dipping faults toward the Na˚grind Syncline (Figs. 6 and 14). The main faults (Na– f) are formed in the latest Cretaceous as documented by weak

188

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

Fig. 10. Seismic example and interpretation of a strike line across the Vema Dome and Nyk High. Note that the thickness of post-intra middle Campanian strata is greatest on the footwall of fault, Nc. Location and fault annotation in Fig. 6. Reflector code in Fig. 3.

onlap within the uppermost lower Maastrichtian sequence (profile N7, Fig. 14). The older Upper Cretaceous strata are, however, parallel and SE dipping. The faults strike in a zigzag pattern similar to the Gjallar Ridge, but show significant erosion of the footwall crests. Moreover, the thickness of the post-intra middle Campanian succession increases southeastward, particularly over fault Nc (profile N4, Fig. 14; Fig. 10). It may be explained by progressive faulting and subsequent erosion from Na to Nc resulting in the preservation of sediments only on the footwall of fault Nc. Some small-offset normal faults southwest of the high appear present only in the Santonian to intra middle Campanian sequences (Fig. 6). In terms of structural and depositional style, the Nyk High consists of two segments, and the narrow transition between the segments is marked by a northward shift of paleotopographic high locations from fault Nc to Na (profiles N4-5, Fig. 14; Fig. 15b). The wide southern segment is characterized by relatively fewer faults and a collapse feature caused by antithetic faulting toward Nc (profiles N2-3, Fig. 14; Fig. 15a). Furthermore, the high is buried by Paleocene sediments shallowing to the northeast with onlap of a very thin Paleocene sequence on the eastern

flank (profiles N1-5, Fig. 14). The northern segment is structurally more complex with presence of several SE dipping normal and reverse faults (profiles N6-9, Fig. 14; Fig. 15c). The northern end is again deeply buried and the faults show collapse patterns and reduced displacement (profile N10, Fig. 14). The Rym accommodation zone, which offsets the Gjallar Ridge and the north Vema Dome fault systems, is expressed as a narrow and elongate topographic high at early Eocene to Oligocene levels (Fig. 16). Its deeper features are, however, almost entirely masked by breakup lavas (Fig. 6). Nevertheless, the zone separates fault systems of different polarities in the Gjallar Ridge and Hel Graben, as well as the NE thickening of post-middle Campanian sequences in the Hel Graben from the thinner successions in the Vigrid Syncline. Figs. 16 and 17 show complex structures along strike of the accommodation zone. Two listric normal faults of opposite polarities are observed on both sides of the lavas in profile R1 (Fig. 16). To the west, the seaward dipping Fenris Graben boundary fault (Ga) appears as a triangular hanging wall feature where the Santonian– Campanian sediments are missing (Fig. 17a). Similarly, the Hel Graben boundary fault

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

189

Fig. 11. Stack of interpreted seismic profiles in the Hel Graben–Naglfar Dome region showing changes of stratigraphic and structural styles. Location and fault annotation in Fig. 6. Reflector code in Fig. 3.

190

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

Fig. 11 (continued )

(Ha) cuts the section twice, but dips landward, farther east. A narrow latest Cretaceous dome feature is mapped east of the lava boundary (Fig. 17b). A similar dome feature is observed in southern Hel Graben (profile R7, Fig. 16). The thickness of the post-intra middle Campanian sequences increases largely northeast of the accommodation zone, reaching , 1.5 s twt (, 2.3 km, Figs. 16, 17b and c). Small-offset normal faults, formed antithetically or

synthetically against the main graben faults are most abundant in the southern Hel Graben (Fig. 16).

5. Basin subsidence and paleobathymetry From the available lithologic and biostratigraphic information in the scientific and commercial boreholes

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

191

Fig. 12. Seismic examples from the Hel Graben–Naglfar Dome region illustrating major Late Cretaceous landward dipping faults (Ha–Hd), late Paleocene seaward dipping fault (Hi) and small-offset faults. Location in Fig. 11.

(Figs. 1 and 2), we have reconstructed the paleobathymetry of the NW Vøring Basin, and calculated burial curves for the decompacted sediments in three of the commercial wells (Fig. 18). However, the Cenozoic coverage is sketchy due to mostly open hole commercial drilling with limited sample return. Thus, the Cenozoic part of the curves should be considered with caution. The stratigraphy and paleobathymetry are consistent with zonations and paleowater depth models derived from detailed micropaleontological studies of a large number of commercial wells on the adjacent continental shelf (Gradstein & Ba¨ckstro¨m, 1996; Gradstein et al., 1999).

The prominent feature of the Cretaceous Vøring Basin is the 7– 10 km thickness of mud-prone sediments, derived mainly from land areas in the northeast, north and west of the present basin (Brekke et al., 1999). Subsidence after Late Jurassic – Early Cretaceous rifting, possibly amplified by renewed modest mid-Cretaceous extension and a 250 – 350 m sea level rise (Hay, 1995), created the Cretaceous seaway between Greenland and Norway resulting in accommodation space for the widespread mudstones. At the onset of the Cenomanian, the region had become a gently subsiding shallow basin passing through neritic environments (Fig. 18). This setting persisted through

192

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

Fig. 13. Seismic example and interpretation of a strike line across the Hel Graben–Naglfar Dome showing rift unconformities at the top Cretaceous and top Paleocene levels, and significant growth against fault Hg. Location and fault annotation in Fig. 6. Reflector code in Fig. 3.

the Santonian followed by rapid Campanian deepening, which created considerable sediment accommodation space in a middle to upper bathyal marine setting documented by cosmopolitan deep water agglutinated foraminifer assemblages and Atlantic-type diatom blooms (Gradstein et al., 1999). Thick mudstones, siltstones, and gravity flow sands were deposited in the basin during early and middle Campanian time. However, the paleo-water depth suggests that tectonic subsidence outpaced sedimentation. The situation reversed in the late Maastrichtian with rapidly shallowing environments lasting through Paleocene times. This created a hiatus in the three wells shown in Fig. 18, implying subaerial or shallow water conditions resulting in erosion of the thick Campanian sequence. Noting that the wells are located on structural highs, the reversal of vertical motion is interpreted as a combined effect of local tectonism and the Paleocene arrival of the Iceland mantle plume inducing thermal and dynamic regional uplift (Skogseid et al., 2000). Although the stratigraphic and paleobathymetric data indicate uplift and erosion, local structures such as the Hel Graben did not experience significant erosion (Figs. 11 and 13). This model is also consistent with the history of vertical motion and late Paleocene to early Eocene depositional conditions inferred from the upper and lower volcanic series drilled at ODP Site 642 on the Vøring Marginal High (Eldholm et al., 1989). Campanian –Paleocene sediments are missing at the conjugate areas in East Greenland, where Santonian

sandstones in the Hold-with-Hope group are overlain unconformably by Tertiary basalts and sedimentary rocks (Kelly, Whitham, Koraini, & Price, 1998). Quantitative thermal history modelling of apatite fission track data from SW Clavering Ø in East Greenland reveals a midCretaceous cooling episode which may reflect denudation resulting in large amounts of siliciclastic detritus transported into the proto-Vøring Basin (Johnson & Gallagher, 2000). 6. Timing of Late Cretaceous –Paleocene rifting The seismic stratigraphic and structural interpretation suggest that the intra middle Campanian reflector (IMC) marks the lower boundary of the Late Cretaceous– Paleocene rift episode which may be divided in three tectono-magmatic phases. 6.1. Onset of rifting and early rift phase (, 81 to 65 Ma) A major onlap surface, the intra Cenomanian horizon, on the eastern flank of the Gjallar Ridge (Fig. 8c) has been interpreted in terms of an early Cenomanian onset of regional extension in the Vøring Basin by Bjørnseth et al. (1997). Similarly, Brekke (2000) suggested that the rifting began in latest Cenomanian to earliest Turonian time, but with renewed or accelerated subsidence reaching a maximum during Campanian time. He proposed that the Gjallar Ridge started to rise near

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

the Cenomanian– Turonian transition as demonstrated by the thinning and onlap of the post-Cenomanian strata towards the tilted late Cenomanian units on the eastern flank, and the late Cenomanian –early Turonian eastward tilting of the ridge has a counterpart on the conjugate basin flank along the southern Trøndelag Platform. Recent studies of wells 6607/5-1, 6707/10-1, and 6706/ 11-1 indicate, however, a latest Turonian age for the initial stage of tectonism (Brekke, Sjulstad, Magnus, & Williams, 2001) probably with reference to the same onlap surface. On the other hand, Dore´ et al. (1999) and Lundin and Dore´ (1997) have suggested that the onlap represents a separate mid-Cretaceous extensional event. Although they indicate Cenomanian to Paleocene faulting on the Gjallar Ridge, they dated the onset of renewed rifting to late Maastrichtian time. We observe growth on normal faults immediately above IMC in the northern Fenris Graben – Gjallar Ridge area. Furthermore, middle –upper Campanian strata onlap IMC on the eastern flank of the Gjallar Ridge. Similar growth patterns also occur in the Hel Graben and Vema Dome areas. In particular, IMC forms a rollover structure against the Hel Graben boundary fault, Ha (profile H2, Fig. 11). Moreover, the Vøring Basin wells indicate a significant increase in paleo-water depth prior to the deposition of the syn-rift middle –upper Campanian sediments (Figs. 2 and 18) probably resulting from the initial subsidence due to extension of the lithosphere. These observations provide us with compelling indications of new, or renewed, basin-wide extension and tectonism. Thus, we place the onset of the Late Cretaceous– Paleocene rift episode in middle Campanian time, , 81 to 76 Ma according to the Gradstein et al. (1994) time scale. Note that the correlation of the intra middle Campanian and intra late Campanian reflectors across the Fenris Graben –Gjallar Ridge boundary fault remains uncertain. Dore´ et al. (1999), for example, interpret shallower levels of these reflectors, a fact we ascribe to their younger, late Maastrichtian, age of the rift onset. However, we attribute the multi-phase syn-rift sedimentation and increased basin fill along the boundary fault as well as minor growth below IMC on the central Gjallar Ridge faults to a separate, but poorly defined Cenomanian extensional event (Fig. 8d). Thus, there is general agreement that the onlap of the post-Cenomanian strata on the eastern flank of the Gjallar Ridge represents an extensional event which marks the onset of a period of increased tectonism and subsidence. In addition, we point out that many inferences about pre-Campanian rifting arrive from onlap patterns, which may only image basin subsidence following a period of rifting expressed by brittle faulting. We note that our study area is located on the western flank of the main Late Jurassic– Early Cretaceous rift zone (Skogseid et al., 2000). Moreover, we do not recognize unambiguous evidence of Cenomanian faulting in the seismic data, whereas the Gjallar Ridge onlap

193

surface clearly predates the onset of Campanian faulting. Therefore, we consider the Cenomanian relative uplift of the Gjallar Ridge and associated onlap patterns, simply as the flexural response during the thermal cooling phase of the Late Jurassic –Early Cretaceous rifting, and/or the more poorly delineated mid-Cretaceous rift event, consistent with the symmetrical basinward tilting of the basin flanks (Brekke, 2000). Except for the Nyk High, the main structural features were developed during the main Late Cretaceous rift phase. Extensional deformation in the Gjallar Ridge, Hel Graben and North Vema dome areas yielded accommodation space for a greatly expanded post-IMC succession in the Hel Graben and progressively developed the Rym accommodation zone marking a lateral change in rift polarity (Fig. 6). As extension and differential subsidence continued towards the end of the period, intense local tectonism created the Nyk High resulting in faulting, block rotation, rift flank uplift and erosion (Figs. 14 and 15). A slight thickening of the lower Maastrichtian sequence on the western flank of the Nyk High implies that the main period of faulting and block rotation began in the early Maastrichtian lasting for a period of less than , 5 Myr (Fig. 5a). 6.2. Late rifting-regional uplift (, 65 to 55 Ma) Although extension continued, the late rift phase was characterized by regional uplift, subsequent erosion and redeposition, contributing an additional source for the Paleocene sediments in the Vøring Basin. The uplift led to an erosional surface above sediments of early Maastrichtian age in the Gjallar Ridge, Nyk High, and Utgard High wells, whereas upper Maastrichtian sediments remain in the Vema Dome well. In the seismic data, the uplift is documented by the near top Cretaceous erosional unconformity over the entire Vøring Basin, and deep local erosion such as into Campanian on the Nyk High (profile N9, Fig. 14; Fig. 15). Whether the unconformity marks subaerial or shallow water erosion is not resolved, but it is likely that parts of the structural highs were periodically emergent. Whether the uplift of the intra-basinal structural highs is related to local syn-tectonic effects, rising hot mantle material, or both, is not obvious. Nonetheless, the Maastrichtian sequence is thicker, and presumably, more complete, in the Hel Graben, suggesting that the regional, pre-breakup uplift component may be constrained to the Paleocene epoch. We date this event to the late Paleocene and infer a maximum uplift near the incipient plate boundary, gradually diminishing in magnitude eastward in the Vøring Basin (Skogseid et al., 2000). This led to extensive and deep erosion in the west leaving predominantly upper Paleocene depocentres of variable thickness between the intra-basinal highs (Skogseid et al., 1992a) and a thin cover of Late Paleocene sands on top of the highs (Fig. 2).

194

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

Fig. 14. Stack of interpreted seismic profiles across the Nyk High showing changes of stratigraphic and structural styles. Location and fault annotation in Fig. 6. Reflector code in Fig. 3.

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

Fig. 14 (continued )

195

196

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

Fig. 15. Seismic examples on the Nyk High illustrating: (a) a collapse feature northwest of fault Nc; (b) the narrowest fault zone; (c) a complicated faulting pattern of both seaward (faults Ha, Nd) and landward (e.g. fault Nf) dipping faults and collapse features. Note that paleotopographic high locations have shifted from fault Nc on profile (a) to fault Na on profile (c). Location in Fig. 14.

The regional uplift and associated erosional unconformity make it difficult to resolve when the movement on many individual faults ceased. On the North Vema Dome, the faulting appears to have continued beyond breakup into Eocene time (Fig. 9). On the other hand, most faults on the central Gjallar Ridge and Nyk High end at the top Cretaceous erosional unconformity which is overlain either by a thin upper Paleocene cover or lower Eocene and Pliocene sediments, respectively (Figs. 7 and 14). However, in the Hel Graben – Naglfar Dome area, most faults active during the early phase continue into the Paleocene and/or lower Eocene strata. In addition, there are numerous smalloffset faults within the post-intra middle Campanian sequences, in particular in the Paleocene and lower Eocene

sequences. These faults are either structurally connected to the older normal faults, or appear as separate features. Thus, we propose that the area was undergoing extension and faulting from middle Campanian time to breakup, with localized, modest adjustment after the onset of sea floor spreading in the Lofoten Basin. Landward of the flood basalt cover, the structural style changed from the early to the late rift phase. The late phase was characterized by generally small displacements and reduced frequency. Moreover, the deformation appears most intense in the Hel Graben – Naglfar Dome and adjacent areas. For example, the Hel Graben seaward dipping fault (Hi) shows significant flank uplift and part of the Paleocene succession has been eroded. In addition, continuous faulting along the Hel Graben faults

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206 Fig. 16. Stack of interpreted seismic profiles across the Rym accommodation zone (Rym AZ) showing changes of stratigraphic and structural styles. Location and fault annotation in Fig. 6. Reflector code in Fig. 3.

197

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

Fig. 16 (continued )

198

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

199

Fig. 17. Seismic examples across the Rym accommodation zone (RAZ) showing (a) a complicated fault pattern and lava flows in two branches; (b) an updomed feature southwest of the RAZ; (c) thickening of post-intra middle Campanian sediments toward northeast. Location in Fig. 16.

led to further footwall uplift and erosion of the Nyk High. In contrast, with the exception of the boundary fault (Ga) which offsets the top Cretaceous unconformity in the west, the thin upper Paleocene cover and onlapping Paleocene sequences along the eastern flank of the Gjallar Ridge are unfaulted (Figs. 7 and 8c). The late rift phase was also associated with abundant intrusive igneous activity. Numerous sills and/or low-angle dykes appear as strong seismic reflections of good continuity terminating abruptly within the transparent and weakly layered pre-Tertiary sequences. The intrusions appear below the middle Campanian reflector, but become shallower in the stratigraphy and more abundant on the central Naglfar and Vema domes (Figs. 5b and 9). In addition, the intrusions are fewer beneath the Gjallar Ridge and Nyk High, whereas they become abundant and thicker in the basinal areas, i.e. Hel Graben, Vigrid and Na˚grind synclines. Most intrusions follow the stratification, and commonly cut and/or follow the fault planes, suggesting a contemporaneous or later time of emplacement than the faulting. Note that the intrusions cut both the early and late

rift phase faults, but they appear to follow early phase fault planes only. The sills are particularly abundant and thick in the Hel Graben, forming prominent complexes at , 5 km depth (Berndt, Skogly, Planke, & Eldholm, 2000) (Figs. 11 and 13). The sill intrusions in the Vøring Basin have only been sampled in well 6607/5-2 on the Utgard High which penetrated three sills (Fig. 3). Cuttings of the two lower sills (90 and estimated 150 m thick) show that they are coarse-grained gabbroic mafic rocks with only little secondary alteration (Munz, Iden, & Stray, 1997), geochemically distinct from both the upper and lower volcanic series drilled at ODP Site 642 on the Vøring Marginal High. Although the ages of the two sills are unknown, the thermal basin history modelling from vitrinite measurements indicates a high heat flow associated with mafic intrusion in the Paleocene (Ren et al., 2001). The large amount of intrusives locally may have caused the observed pre-breakup doming and inversion of the Hel Graben (Fig. 5b), and amplified the rift flank uplift beneath the northern Vema Dome (Fig. 9).

200

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

Fig. 18. (a) Paleobathymetric trend and (b) burial curves of three commercial wells using decompacted sediment thicknesses for the Late Cretaceous– Cenozoic succession in the NW Vøring Basin.

With reference to the initiation of regional magmatism elsewhere in the North Atlantic Volcanic Province at 63 –60 Ma (e.g. Saunders, Fitton, Kerr, Norry, & Kent, 1997), the onset of the Vøring Basin intrusions may have occurred in late early Paleocene time, continuing through breakup.

igneous intrusions as sills and low-angle dykes; an inner flow complex present within a 10– 40 km wide zone landward of the Vøring Escarpment (IF, Figs. 4b, 5a, b, 7, 11 and 16); various tuff units including a prominent tuff marker, the top Paleocene (TPa) reflector; and several volcanic vents at or near TPa (Figs. 4 and 5) (e.g. Skogseid et al., 1992a).

6.3. Breakup phase (, 55 to 52 Ma) The breakup phase was marked by regional volcanism, producing voluminous mafic melts at very high rates. The excess volcanism was most intense in the first 1– 1.5 Myr after breakup, gradually changing into waning, smallvolume melts of variable compositions (Eldholm, Tsikalas, & Faleide, 2002). The breakup events produced the Vøring Marginal High underlain by expanded oceanic crust and intruded and underplated continental crust separated by a continent – ocean boundary some tens of kilometres west of the Vøring Escarpment (Eldholm et al., 1989). In the Vøring Basin, we recognize

7. NW Vøring Basin rift evolution At the onset of the Late Cretaceous –Paleocene rift episode, the area between Norway and Greenland was an epicontinental sea covering a region of highly attenuated crust and greatly accentuated lithospheric relief reflecting several previous post-Caledonian rift episodes. The Late Jurassic – Early Cretaceous rifting, in particular, was accompanied by considerable crustal extension and thinning forming grabens within older half-graben basins (Faleide, Va˚gnes, & Gudlaugsson, 1993; Skogseid et al., 1992a,

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

2000). In mid to early Late Cretaceous times, the Vøring Basin experienced uniform thermal subsidence although there is evidence for modest mid-Cretaceous extension (Dore´ et al., 1999; Lundin & Dore´, 1997). In the NE Atlantic realm, coeval mid-Cretaceous (Aptian – Albian?) extension is also observed on the Lofoten – Vestera˚len margin (Tsikalas, Faleide, & Eldholm, 2001), onshore East Greenland (Surlyk, 1990; Whitham, Price, Koraini, & Kelly, 1999), and in the SW Barents Sea (Faleide et al., 1993). The Late Cretaceous – Paleocene rift episode was initiated in middle Campanian time, and lasted until continental separation near the Paleocene –Eocene transition, leaving a more than 300 km wide, segmented rift zone along the eventual axis of continental separation (Skogseid et al., 2000). Although our study area comprises the eastern rift segment between the Gleipne and Bivrost lineaments, we stress that the rift continues below the lavas onto the conjugate Greenland margin. In fact, stretching distributions derived from crustal thinning and basin subsidence estimates indicate that the extension increases significantly towards the continent – ocean boundary (Skogseid, 1994; Skogseid et al., 2000), hence considerable Paleocene brittle structural deformation may be hidden beneath the lava flows. However, a changing rift rheology due to plume arrival may have given rise to a change in mode of rifting from brittle to more ductile.

201

The Vøring Basin rift evolution is also characterized by rift segmentation and varying across- and along-rift structural styles as exemplified by the five structural subprovinces in Fig. 19. In both a local and regional sense, the Vøring rift system has many features similar to those observed in continental rift systems such as the East African Rift (e.g. Ebinger, 1989; Ebinger, Jackson, Foster, & Hayward, 1999; Morley, Nelson, Patton, & Munn, 1990; Nelson, Patton, & Morley, 1992; Rosendahl et al., 1986). Among these are: (1) rift basins bordered on one or both sides by relatively long (tens of kilometres) normal fault systems that largely control basin morphology, (2) fault systems composed of high-angle normal faults occurring in a 10– 15 km wide zone, (3) curvature of the bounding faults, which commonly terminate by splaying along strike, (4) throws that are greatest near the central part of spoonshaped basins, and decrease towards the basin tips, (5) tilted or rotated fault blocks, (6) block asymmetry on dip lines, (7) presence of axial highs where full-graben morphology is observed, and (8) thickening of stratigraphic units towards bounding faults. As in other continental rift zones developed in stable cratonic lithosphere (e.g. Baikal), the East African rift system is characterized by narrow rift zones (, 100 km) bounded by high-angle border faults and small degrees of extension, in contrast to wide rifts (. 1000 km) developed

Fig. 19. Main tectono-magmatic events and relative vertical motion in the NW Vøring Basin during the Late Cretaceous–Paleocene rift episode: I, early rift phase; II, late rift phase; III, breakup phase. MCD, middle crustal dome; Rym AZ, Rym accommodation zone.

202

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

above collapsing orogenic belts or areas with higher thermal gradients (e.g. Aegean, Basin and Range). These variations can be explained by the thermomechanical state of the lithosphere at the time of rifting (Buck, 1991); hot, weak lithosphere develops broad rift zones bounded by high-angle border faults; and cold, strong lithosphere develops narrow rift zones bounded by low-angle border faults. The rifted Vøring continental margin, part of the . 300 km wide segmented NE Atlantic rift system, differs therefore, from the East African Rift, i.e. the border fault systems are commonly broader (, 13 to 27 km) and made up of listric and in places, low-angle detachment faults. However, because the stretching factor, estimated from the structural restoration and subsidence analysis, is , 1.6 in the study area (Ren et al., 1998; Skogseid, 1994), we have established a structural evolution model by using the terminologies of the East African rift (Rosendahl, 1987) associated with mechanics of brittle rock deformation in a low mean stress environment. In this model, the key to understand the rift morphology and structure pertains to how half-grabens link together by accommodation zones (for terminology, see Peacock, Knipe, & Sanderson, 2000). These are generally

10– 20 km wide zones of varying geometry which transfer, or accommodate, extension between individual faults and rift segments along the strike of the rift. By applying this concept, we interpret the Bivrost and Gleipne lineaments as accommodation zones and introduce two new zones within the NW Vøring Basin rift segment (Fig. 20). From seismic stratigraphic and structural interpretation, we have divided the syn-rift history into two phases. In the early phase, from middle Campanian to the Cretaceous– Tertiary transition, the Gjallar Ridge, Hel Graben and North Vema Dome experienced extensive listric normal faulting and subsequent syn-rift subsidence. The Gjallar Ridge developed along low-angle detachment structures that updome thick Cretaceous sequences and sole out at medium-to-deep intra-crustal levels, accompanied by the formation of a core complex (Fig. 19; Ren et al., 1998). During this period, thick units of Campanian to Maastrichtian age were deposited in basinal areas and formed a Campanian depocentre in the Hel Graben. In early Campanian to early Maastrichtian time, the Gjallar Ridge and Hel Graben fault zones had formed non-overlapping opposing-polarity half-grabens linked by a synthetic/

Fig. 20. A tectonic model for the Late Cretaceous– Paleocene structuring in the NW Vøring Basin. (a) In middle Campanian to early Maastrichtian time, a HRAZ and a LRAZ formed between Hel Graben (HG) and Gjallar Ridge (GR), and between Hel Graben (HG) and Vema Dome half-grabens, respectively. (b) In Late Maastrichtian time, the LRAZ extended further northeastward after the development of the Nyk High (NH) half-graben. (c) Local stress field in the latest Maastrichtian and/or earliest Paleocene time causing main activity along the Rym accommodation zone (RAZ) which in (d) is inferred by a slight rotation of the local stress field, resulting in strike– slip motion along the RAZ. The associated depression was later ponded by Early Tertiary breakup lavas (shaded).

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

synthetic high relief accommodation zone (HRAZ), as defined by Rosendahl (1987), at the northern end of the Rym accommodation zone (Fig. 20a). The northernmost part of the Gjallar Ridge boundary fault (Ga) shows major displacement, suggesting that the fault continues farther north and that the subsidence caused by the fault is transferred to the Hel Graben fault zone. Thinning of the Campanian –Maastrichtian strata towards the Rym accommodation zone suggests that the HRAZ may represent a remnant of pre-rift rocks that have been excluded from significant syn-rift subsidence by the particular arrangement of the bounding fault system. At the same time, the Hel Graben and North Vema Dome fault zones developed as overlapping and opposing-polarity half-grabens linked by a low relief accommodation zone (LRAZ) along which the subsidence is greatest on the Hel Graben side (Fig. 20a). In late Maastrichtian time, when the Nyk High experienced extensive faulting, fault block rotation and rift flank uplift, the LRAZ continued towards northeast and formed an antiformal welt between the half-grabens in the Hel Graben and Nyk High (Fig. 20b). However, the subsidence along the boundary faults (Ha, Na) was asymmetric. Alternating stratal thickness relationships, progressing from the lower Campanian – Maastrichtian to the upper Maastrichtian sequences correspond to alternating movement on the boundary faults that control individual half-graben development. This model of LRAZ formation is comparable to the ‘subsidence teeter – totter’ of Rosendahl et al. (1986). The strike of the faults in the Hel Graben and southwest of the Nyk High deviates as much as 408 to the regional NW – SE extension direction (Figs. 6 and 20b). Moreover, the North Vema Dome is underlain by a highly curved fault complex, although part of the faults trends along the regional direction (Fig. 20b). In many rift environments, there is a tendency for faults to zigzag, or the influence of anisotropy and heterogeneity of the pre-rift structural and/or basement fabric may cause some rift features to form at an acute angle to the regional extension direction (e.g. Christopher, 1998; Daly, Chorowicz, & Fairhead, 1989; Dore´ et al., 1997; Smith & Mosley, 1994). As extensional faulting occurs in a relatively low mean stress environment, mechanical anisotropies tend to be reactivated. Tension tests in layered rocks by Youash (1969) imply that fabrics, even striking at quite high angles (608) to the extension direction, may be activated (Morley, 1995). The development of the Hel Graben and North Vema Dome had created a local stress field deviating , 50 to 608 to the regional extension direction (Fig. 20b), however a slight clockwise rotation of the regional stress field may have occurred in the latest Maastrichtian and/or earliest Paleocene time resulting in local strike – slip motion in this area (Fig. 20c). A small strike – slip basin similar to the polygenetic basin type described by Nilsen and Sylvester (1999), locally bounded by the northernmost Gjallar Ridge boundary fault (Ga), was formed on top of the Rym

203

accommodation zone and later ponded by lavas (Fig. 20d). The development of the pull-apart terrain on top of the initial HRAZ has obscured its nature as a high. In fact, the central part of the HRAZ was not completely covered by lavas which form two separate flow branches (Fig. 20d). Hence, the mild strike –slip motion caused minor folding and slight thickening of the Paleocene strata towards the Rym accommodation zone, a relief which subsequently allowed a limited, narrow cover of lavas (Figs. 16 and 17). During the late rift phase, the area experienced regional uplift accompanied by intrusive igneous activity and subsequent erosion causing the regional hiatus (Fig. 18). Extensional faulting, observed as reactivation of the Fenris Graben boundary fault and the small-scale normal faults in the Hel Graben –Naglfar Dome, continued onto breakup, but with small displacements and reduced frequency. The NW Vøring Basin rift evolution shows that the East African Rift architecture may provide an adequate analogue explaining the rift segmentation and the mode of rifting during the early rift phase. The pre-existing basement grains and/or structural lineaments may act as high-relief accommodation zones separating rift segments by strike – slip motion. For example, the first-order Jan Mayen and Bivrost lineaments separate the Møre, Vøring and Lofoten – Vestera˚len basins, respectively, and the Lofoten – Vestera˚len margin is divided into three main rift segments by high-relief accommodation zones (Tsikalas et al., 2001). In the study area, the half-graben linkage and structural evolution is largely controlled by the Rym HRAZ, whereas the Gleipne and Bivrost HRAZ provide boundary conditions. In contrast, we document a change in the mode of rifting during the late rift phase. The extending rift became affected by the arrival and impingement of the Iceland Plume inducing dynamic and thermal uplift in mid-to-late Paleocene. As the hot asthenosphere migrated into regions of thinned lithosphere, decompressional partial melting accelerated, culminating with massive subaerial volcanism during breakup. The change in thermal regime in the continental lithosphere caused by the plume arrival may have modified the rheology of the continental crust (Buck, Martinez, Steckler, & Cochran, 1988; Pedersen & Skogseid, 1989), thus contributing to a change from brittle to more ductile extensional deformation from the early to late rift phase. Furthermore, we cannot rule out that part of the Paleocene brittle deformation presently is hidden beneath the breakup lavas, i.e. below the inner flow complex and on the easternmost Vøring Marginal High.

8. Summary and conclusions A dense grid of seismic profiles in the NW Vøring Basin provides structural, stratigraphic and magmatic information which we interpret in terms of a Late Cretaceous – Paleocene rift episode culminating with

204

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

breakup and volcanic margin formation near the Paleocene – Eocene transition. The rifting began in the middle Campanian time, and was characterized by (1) regional large-scale normal faulting and along-margin varying structural styles during the early rift phase, , 81 to 65 Ma, and (2) continued extension, regional uplift, intrusive igneous activity and subsequent erosion in the late rift phase, , 65 to 55 Ma. The mode of rifting changed from brittle, in the early rift phase, to more ductile extensional deformation in the late rift phase. The rifting ended with breakup at , 55 Ma accompanied by massive eruption of lavas, most intense near breakup and gradually waning over the next 3 Myr. The NW Vøring Basin, bounded by the Bivrost and Gleipne lineaments, comprises four separate fault complexes in the northern Fenris Graben – Gjallar Ridge, North Vema Dome, Hel Graben and Nyk High areas. The fault complexes are linked by the Rym accommodation zone. Both the Gjallar Ridge and Nyk High are deformed by seaward dipping listric normal faults and show rift flank uplift, whereas landward dipping normal faults and numerous Paleocene –Eocene small-scale faults dominate in the Hel Graben. In addition, the Gjallar Ridge is characterized by low-angle detachment faulting soling out at middle crustal levels, locally accompanied by core complex development, whereas the normal faults on the North Vema Dome are highly curved. The Rym accommodation zone which appears to offset the Gjallar Ridge and Nyk High fault complexes, separates fault systems of different polarities in the Gjallar Ridge and Hel Graben, as well as northeast thickening post-middle Campanian sequences in the Hel Graben from thinner coeval successions in the Vigrid Syncline. For the early rift phase we propose a simple tectonic model for the development of the NW Vøring Basin rift segment (Fig. 20). In middle Campanian to early Maastrichtian times, the Gjallar Ridge and Hel Graben fault zones developed as non-overlapping opposingpolarity half-grabens linked by the Rym HRAZ. Another LRAZ formed between the overlapping and opposingpolarity Hel half-graben and the North Vema Dome. In late Maastrichtian time, when the younger Nyk High halfgraben began to develop, the LRAZ progressed farther northeast and led to a ‘subsidence teeter– totter’ setting between the Hel and Nyk High half-grabens. In the latest Maastrichtian and/or earliest Paleocene time, faulting in the Hel Graben and North Vema Dome half-grabens created a local stress field trending , 50 to 608 to the regional extension direction which induced a slight rotation of the regional stress field resulting in strike –slip motion along the Rym accommodation zone and formation of a small strike –slip basin later ponded by lavas. Although the East African Rift model appears consistent with NW Vøring Basin syn-rift evolution prior to Paleocene time, we infer that a change from brittle to more ductile extensional deformation may have occurred during the late rift phase.

Hence, we interpret the change in mode of rifting as the combined effects of rift rheology due to plume arrival and the fact that most accentuated expressions of brittle deformation may be hidden beneath the lavas erupted during breakup.

Acknowledgements We gratefully acknowledge the Norwegian Petroleum Directorate and Statoil for providing access to multichannel seismic and well data. This study has been supported by the University of Oslo Passive Margin Research Group through a research grant from Statoil. We thank F. Tsikalas and many of our Statoil colleagues, in particular V.B. Larsen, for data, support and constructive discussions. We are also grateful to referee, T. Dore´ an anonymous referee, and D. G. Roberts for their very constructive comments.

References Berndt, C., Skogly, O. P., Planke, S., & Eldholm, O. (2000). High-velocity breakup-related sills in the Vøring Basin, off Norway. Journal of Geophysical Research, 105, 28443–28454. Bjørnseth, H. M., Grant, S. M., Hansen, E. K., Hossack, J. R., Roberts, D. G., & Thompson, M. (1997). Structural evolution of the Vøring Basin, Norway, during the Late Cretaceous and Palaeogene. Journal of the Geological Society, London 154, 559–563. Blystad, P., Brekke, H., Færseth, R. B., Larsen, B. T., Skogseid, J., & Tørudbakken, B. (1995). Structural elements of the Norwegian continental shelf. NPD Bulletin, 8, 45. Brekke, H. (2000). The tectonic evolution of the Norwegian Sea Continental margin with emphasis on the Vøring and Møre Basins. In A. Nøttvedt, et al. (Eds.), Dynamics of the Norwegian Margin (pp. 327 – 378). Geological Society London Special Publication, 167. Brekke, H., Dahlgren, S., Nyland, B., & Magnus, C. (1999). The prospectivity of the Vøring and Møre basins on the Norwegian Sea continental margin. In A. J. Fleet, & S. A. R. Boldy (Eds.), Petroleum Geology of Northwest Europe: Proceedings of the Fifth Conference (pp. 261 –274). London: Geological society. Brekke, H., Sjulstad, H. I., Magnus, C., & Williams, R. W. (2001). Sedimentary environments offshore Norway: an overview. In O. J. Martinsen, & T. Dreyer (Eds.), Palaeozoic to recent (pp. 7–37). NPF Special Publication, 10. Buck, W. R. (1991). Modes of continental lithosphere extension. Journal of Geophysical Research, 96, 20161–20178. Buck, W. R., Martinez, F., Steckler, M. S., & Cochran, J. R. (1988). Thermal consequences of lithospheric extension: pure and simple. Tectonics, 7, 213–234. Bukovics, C., Cartier, E. G., Shaw, N. D., & Ziegler, P. A. (1984). Structural and development of the mid-Norway continental margin. In A. M. Spencer (Ed.), Petroleum geology of the North European margin (pp. 407 –424). London: Graham & Trotman. Christopher, H. D. (1998). Basement-controlled transfer zones in an area of low-magnitude extension, eastern Basin and Range Province, TransPecos, Texas. In J. E. Faulds, & J. H. Stewart (Eds.), Accommodation zones and transfer zones: the regional segmentation of the Basin and Range Province (pp. 75–88). Geological Society of America, Special Paper, 323.

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206 Dalland, A., Worsley, D., & Ofstad, K. (1988). A lithostratigraphic scheme for the Mesozoic and Cenozoic succession offshore mid- and northern Norway. NPD Bulletin, 4, 65. Daly, M. C., Chorowicz, J., & Fairhead, D. (1989). Rift basin evolution in Africa: the influence of reactivated steep basement shear zones. In M. A. Cooper, & G. D. Williams (Eds.), Inversion tectonics (pp. 749–762). GSA Bulletin, 93. Dore´, A. G., Lundin, E. R., Fichler, C., & Olesen, O. (1997). Patterns of basement structure and reactivation along the NE Atlantic margin. Journal of the Geological Society, London, 154, 85– 92. Dore´, A. G., Lundin, E. R., Jensen, L. N., Birkeland, Ø., Eliassen, P. E., & Fichler, C. (1999). Principal tectonic events in the evolution of the northwest European Atlantic margin. In A. J. Fleet, & S. A. R. Boldy (Eds.), Petroleum geology of Northwest Europe: Proceedings of the Fifth Conference (pp. 41–61). London: The Geological Society. Ebinger, C. J. (1989). Geometric and kinematic development of border faults and accommodation zones, Kivu-Rusizi Rift, Africa. Tectonics, 8(1), 117–133. Ebinger, C. J., Jackson, J. A., Foster, A. N., & Hayward, N. J. (1999). Extensional basin geometry and the elastic lithosphere. Philosophical Transactions Royal Society London, 357, 741 –765. Eldholm, O., Skogseid, J., Planke, S., & Gladczenko, T. P. (1995). Volcanic margin concepts. In E. Banda, M. Talwani, & M. Torn (Eds.), Rifted ocean–continental boundaries (pp. 1–16). NATO ASI series volume, Dordecht: Kluwer. Eldholm, O., Thiede, J., & Taylor, E. (1989). Evolution of the Vøring volcanic margin. In O. Eldholm, J. Thiede, & E. Taylor (Eds.), (pp. 1033 – 1065). Proceedings of Ocean Drilling Program (ODP), Scientific Results, 104. Eldholm, O., Tsikalas, F., & Faleide, J. I. (2002). Continental margin off Norway 62–758N: Palaeogene tectono-magmatic segmentation and sedimentation. In B. Bell, & D. Jolley (Eds.), The North Atlantic igneous Province: stratigraphy, tectonics, volcanic and magmatic processes (pp. 39– 68). Geological Society London Special Publication, 197. Faleide, J. I., Va˚gnes, E., & Gudlaugsson, S. T. (1993). Late Mesozoic– Cenozoic evolution of the southwestern Barents Sea in a regional rift-shear tectonic setting. Marine and Petroleum Geology, 10, 186–214. Fichler, C., Rundhode, E., Olesen, O., Sæther, B. M., Ruesla˚tten, H., Lundin, E., & Dore´, A. G. (1999). Regional tectonic interpretation of image enhanced gravity and magnetic data covering the midNorwegian shelf and adjacent mainland. Tectonophysics, 306, 183–197. Gradstein, F. M., Agterberg, F. P., Ogg, J. G., Hardenbol, J., van Veen, P., Thierry, J., & Huang, Z. (1994). A Mesozoic time scale. Journal of Geophysical Research, 99, 24051–24074. Gradstein, F. M., & Ba¨ckstro¨m, S. A. (1996). Cenozoic biostratigraphy and palaeobathymetry, northern North Sea and Haltenbanken. Norsk Geologisk Tidskrift, 76, 3–32. Gradstein, F. M., Kaminski, M. A., & Agterberg, F. P. (1999). Biostratigraphy and paleoceanography of the Cretaceous Seaway between Norway and Greenland. Earth Science Reviews, 46, 27–98. Hagevang, T., Eldholm, O., & Aalstad, I. (1983). Pre-23 magnetic anomalies between Jan Mayen and Greenland-Senja fracture zones in the Norwegian Sea. Marine Geophysical Researches, 5, 345 –363. Hay, W. W. (1995). Cretaceous paleoceanography. Geologica Carpathica, 46, 257 –266. Hinz, K., Mutter, J. C., Zehnder, C. M., Alsop, J., Buhl, P., Meyer, H., & Roeser, H. A. (1987). Symmetric conjugation of continent –ocean boundary structures along the Norwegian and East Greenland margins. Marine and Petroleum Geology, 4, 166–187. Hjelstuen, B. O., Eldholm, O., & Skogseid, J. (1997). Vøring Plateau diapir fields and their structural and depositional settings. Marine Geology, 144, 33 –57.

205

Hjelstuen, B. O., Eldholm, O., & Skogseid, J. (1999). Cenozoic evolution of the northern Vøring margin. GSA Bulletin, 111, 1792–1807. Hovland, M., Nygaard, E., & Thorbjørnsen, S. (1998). Piercement shale diapirism in the deep-water Vema Dome area, Vøring basin, offshore Norway. Marine and Petroleum Geology, 15, 191–201. Johnson, C., & Gallagher, K. (2000). A preliminary Mesozoic and Cenozoic denudation history of the North East Greenland onshore margin. Global Planetary Change, 24, 261 –274. Kelly, S. A. R., Whitham, A. G., Koraini, A. M., & Price, S. P. (1998). Lithostratigraphy of the Cretaceous Hold with Hope Group, NE Greenland. Journal of the Geological Society, London, 155, 993– 1008. Kittilsen, J. E., Olsen, R. R., Marten, R. F., Hansen, E. K., & Hollingsworth, R. R. (1999). The first deepwater well in Norway and its implications for the Upper Cretaceous play, Vøring basin. In A. J. Fleet, & S. A. R. Boldy (Eds.), Petroleum geology of Northwest Europe: Proceedings of the Fifth Conference (pp. 275–280). London: The Geological Society. Lundin, E. R., & Dore´, A. G. (1997). A tectonic model for the Norwegian passive margin with implications for the NE Atlantic Early Cretaceous to break-up. Journal of the Geological Society, London, 154, 545–550. Morley, C. K. (1995). Developments in the structural geology of rifts over the last decade and their impact on hydrocarbon exploration. In J. J. Lambiase (Ed.), Hydrocarbon habits in rift basins (pp. 1 – 32). Geological Society Special Publication, 80. Morley, C. K., Nelson, R. A., Patton, T. L., & Munn, S. G. (1990). Transfer zones in the East African rift system and their relevance to hydrocarbon exploration in rifts. AAPG Bulletin, 74, 1234–1253. Munz, I., Iden, K., & Stray, H (1997). Sill intrusions in the Vøring basin. Mid Norway: preliminary geochemical results. Abstract. Mid-Norwegian Volcanic Margin Conference, Ife, Oslo, Norway, 10–11 February 1997. Mutter, J. C., Buck, W. R., & Zehnder, C. M. (1988). Convective partial melting. 1. A model for the formation of thick basaltic sequences during the initiation of spreading. Journal of Geophysical Research, 93, 1031–1048. Nelson, R. A., Patton, T. L., & Morley, C. K. (1992). Rift-segment interaction and its relation to hydrocarbon exploration in continental rift systems. AAPG Bulletin, 76, 1153–1169. Nilsen, T. H., & Sylvester, A. G. (1999). Strike-slip basins, part 1. The leading edge, 18(10), 1146–1152. Peacock, D. C. P., Knipe, R. J., & Sanderson, D. J. (2000). Glossary of normal faults. Journal of Structural Geology, 22, 291 –305. Pedersen, T., & Skogseid, J. (1989). Vøring plateau volcanic margin: extension, melting and rifting. In O. Eldholm, J. Thiede, & E. Taylor (Eds.), (pp. 985 –991). Proceedings of Ocean Drilling Program (ODP), Scientific Results, 104. Ren, S., Aasgaard, K., Pedersen, T., Wangen, M., Munz, I. A., Faleide, J. I., & Eldholm, O. (2001). Nature of sill intrusions on the Utgard High, Vøring Basin. Abstract. Norwegian Geological Winter Meeting, Oslo, Norway, 8– 10 January 2001. Ren, S., Skogseid, J., & Eldholm, O. (1998). Late Cretaceous –Paleocene extension on the Vøring volcanic margin. Marine Geophysical Researches, 20, 343 –369. Rosendahl, B. R. (1987). Architecture of continental rifts with special reference to East Africa. Annual Review of Earth and Planetary Sciences, 15, 445–503. Rosendahl, B. R., Reynolds, D. J., Lorber, P. M., Burgess, C. F., Scott, M. D., Lambiase, J. J., & Derksen, S. J. (1986). Structural expressions of rifting: lessons from Lake Tanganyika, Africa. In L. E. Frostick, R. W. Renaut, I. Reid, & J. J. Tiercelin (Eds.), Sedimentation in the African rifts (pp. 29 – 43). Geological Society London Special Publication, 25. Saunders, A. D., Fitton, J. G., Kerr, A. C., Norry, M. J., & Kent, R. W. (1997). The North Atlantic Igneous Province. Geophysical Monography, 100, 45–93.

206

S. Ren et al. / Marine and Petroleum Geology 20 (2003) 177–206

Skogseid, J. (1994). Dimensions of the late Cretaceous – Paleocene Northeast Atlantic rift derived from Cenozoic subsidence. Tectonophysics, 240, 225– 247. Skogseid, J., & Eldholm, O. (1989). Vøring plateau continental margin, seismic interpretation, stratigraphy and vertical movements. In O. Eldholm, J. Thiede, & E. Taylor (Eds.), (pp. 993–1030). Proceedings of Ocean Drilling Program (ODP), Scientific Results, 104. Skogseid, J., Pedersen, T., Eldholm, O., & Larsen, B. T. (1992a). Tectonism and magmatism during NE Atlantic continental break-up, the Vøring margin. In B. C. Storey, T. Alabaster, & R. J. Pankhurst (Eds.), Magmatism and the causes of continental break-up (pp. 305 –320). Geological Society Special Publication, 68. Skogseid, J., Pedersen, T., & Larsen, V. B. (1992b). Vøring basin: subsidence and tectonic evolution. In R. M. Larsen, H. Brekke, B. T. Larsen, & E. Talleraas (Eds.), Structural and tectonic modelling and its application to petroleum geology (pp. 55–82). NPF Special Publication, 1. Skogseid, J., Planke, S., Faleide, J. I., Pedersen, T., Eldholm, O., & Neverdal, F. (2000). NE Atlantic continental rifting and volcanic margin formation. In A. Nøttvedt et al. (Eds.), Dynamics of

the Norwegian margin (pp. 295–326). Geological Society London Special Publication, 167. Smith, M., & Mosley, P. (1994). Crustal heterogeneity and basement influence on the development of the Kenya rift, East Africa. Tectonics, 12, 591–606. Surlyk, F. (1990). Timing, style and sedimentary evolution of Late Paleozoic–Mesozoic extensional basins of East Greenland. In R. P. F. Hardman, & J. Brooks (Eds.), Tectonic events responsible for Britain’s oil and gas reserves (pp. 107 –125). Geological Society London Special Publication, 55. Tsikalas, F., Faleide, J. I., & Eldholm, O. (2001). Lateral variations in tectono-magmatic style along the Lofoten–Vestera˚len volcanic margin off Norway. Marine and Petroleum Geology, 18, 807 –832. Whitham, A. G., Price, S. P., Koriani, A. M., & Kelly, S. R. A. (1999). Cretaceous (post-Valanginian) sedimentation and rift events in NE Greenland (71–778N). In A. J. Fleet, & S. A. R. Boldy (Eds.), Petroleum Geology of Northwest Europe: Proceedings of the Fifth Conference (pp. 325–336). London: The Geological Society. Youash, Y. (1969). Tension tests in layered rocks. Geological Society of America, 80, 469–482.