NE-trending growth folds

The vast majority of growth folds within the NE Faroe–Shetland Basin display a NE trend (Fig. 2). The axial traces of these anticlines and monoclines vary in length significantly, with the largest extending over 70 km. The fold amplitudes range up to 3000 m, and both symmetrical and asymmetrical forms are present. The ages of the seismic horizons used in evaluating the timing of formation of the structures largely follows the stratigraphic schemes of STRATAGEM partners (2002) and Davies & Cartwright (2002), Ritchie et al. (2003) and Johnson et al. (2005). The growth folds are generally only clearly observed at top Palaeogene lavas and younger stratigraphic levels, as the presence of these lavas and also contemporaneous igneous intrusive complexes, effectively mask the deeper structure of the basin (e.g. Ritchie et al. 2003, figs 5, 7, 8 & 9). Though the structure and nature of the pre-Cenozoic succession is largely obscured in this area, the main trend of the growth anticlines is parallel with the pervasive NE Caledonian/pre-Caledonian structural grain that dominates this part of the margin (e.g. Duindam & van Hoorn 1987; Hitchen & Ritchie 1987; Stoker et al. 1993; Dean et al. 1999; Roberts et al. 1999; Coward et al. 2003), perhaps suggesting a link between pre-existing structural architecture, i.e. the presence of deeply buried NE-trending tilted fault blocks, and subsequent growth fold development. NE-trending anticlines, comprising folds A–D and the Pilot Whale Anticline (Fig. 2) have previously been described by Ritchie et al. (2003) and are considered representative of growth folds within this part of the basin. These were mainly developed in early to mid Miocene times, though some also show early Pliocene growth (Fig. 3). There is some biostratigraphical evidence from well 214/04-1 of a late Eocene unconformity (separating upper Eocene from lower Oligocene/upper Eocene strata) as documented by Davies & Cartwright (2002). There is also limited supporting seismic evidence for growth fold development at the time (Davies & Cartwright 2002, fig. 5). However, the widespread effects of any compression on the Eocene and Oligocene succession remain difficult to assess due to seismic masking by pervasive polygonal faulting. However, there are signs of limited fold growth during this time associated with the development of Anticline C (Fugloy Ridge) (e.g. Boldreel & Andersen 1998; Ritchie et al. 2003; Johnson et al. 2005).

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Fig. 3.

Summary chart of the timing/duration of Cenozoic fold growth development within the Faroe–Shetland Basin and adjacent areas within the NE Atlantic margin (modified from Ritchie et al. 2003 and Johnson et al. 2005). Abbreviations: C10, Early Pliocene Unconformity; C30, Late Eocene Unconformity; IEUa, b and c, Intra-Eocene unconformities; IMU, Intra-Miocene Unconformity; INU, Intra-Neogene (Early Pliocene) Unconformity; L3, Mid Eocene (late Lutetian); LE, Early Eocene (latest Ypresian) Unconformity; MMU, Mid Miocene Unconformity; TBF, Top Balder Formation; TPL, Top Palaeogene lavas; TPU, Top Palaeogene Unconformity.

NNE-trending growth folds

The NNE-trending Pilot Whale Anticline is one of only two significant growth folds with this orientation in NE Faroe–Shetland Basin (Fig. 4). The axial trace of this slightly asymmetrical anticline extends for about 20 km and lies close to the extreme NE margin of the Faroe–Shetland Basin (Fig. 2). The fold is clearly imaged at top Palaeogene lavas level, with an amplitude of approximately 1400 m (Fig. 4). The anticline is defined by the top Palaeogene lavas and the diagenetic Opal A–Opal C/T transformation horizon (Ritchie et al. 2003; Johnson et al. 2005). This diagenetic horizon has a late Miocene age ascribed to its formation by Davis & Cartwright (2002) (Fig. 3). The relatively undeformed mid Pliocene to Recent succession becomes thinner over the crest of the anticline possibly suggesting relative sea level rise and passive burial of the structure.

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Fig. 4.

Seismic profile A–A′ across the Pilot Whale Anticline within the NE Faroe–Shetland Basin. See Figure 2 for location of profile. Abbreviations: INU, Intra-Neogene (Early Pliocene) Unconformity and TPL, Top Palaeogene lavas. Note that Opal A–Opal C/T transformation horizon is a diagenetic reflector of late Miocene age (Davies & Cartwright 2002).

The growth of the Pilot Whale Anticline is, we believe, closely associated with the development of the Pilot Whale Diapirs (Haflidason et al. 1996) and the associated seabed mud mounds and subsurface injection features that can be observed from seismic data (Fig. 4). This association of diapirs, seabed mounds and injection features is similar to phenomena described from the Norwegian margin around the Vema and Naglfar domes and the Helland Hansen and Modgunn arches, and in the North Viking Graben area (e.g. Hjelstuen et al. 1997; Hovland et al. 1998). The morphological expression of the Pilot Whale mud mounds ranges from almost circular features, to complex groups with irregular perimeters. They rise to at least 120 m above the sea bed and have been described as comprising sediments that range in age from Pliocene to Oligocene. The development of the mud mounds, diapirs and intrusive features is considered here to have been triggered initially by the growth of the Pilot Whale Anticline from early Pliocene times onwards. Before mobilization and injection, the Eocene and Oligocene succession within the NE Faroe–Shetland Basin is inferred to have comprised mainly smectite-rich, under-compacted, low-density mudstones that were subsequently mantled by a seismically well-layered, mainly fine-grained Pliocene and younger succession. It is suggested that the growth of the Pilot Whale Anticline may have facilitated fracturing and breaching of the Eocene and younger successions, and that the developing anticline may have acted as a focus for the migration of fluids and the mobilization of this under-compacted, overpressured, low-density succession. The interpreted presence of v-brights which are considered to indicate gas or water charged sandstone ring dykes (Huuse & Mikelson 2004) within the Eocene interval suggests that the natural buoyancy of this succession may have been significantly assisted by gas, fluid and sediment injection.

Ritchie et al. (2003) speculated, from strain ellipse analysis, that the generation of the NNE-trending Pilot Whale Anticline may have been caused by sinistral strike-slip fault movement along the Magnus Lineament or Transfer Zone (Fig. 2). This hypothesis is similar to that promulgated by Doré & Lundin (1996) for the NNE- to north-trending Cenozoic anticlines and domes observed within the Norwegian margin. However, it is now considered more likely that the Pilot Whale Anticline formed as a result of buttressing of basin fill against Palaeozoic and older basement of the Møre Marginal High which includes the basic igneous plutonic mass that forms the core of the Brendan Volcanic Centre (Fig. 2). Regionally, the NNE-trending Pilot Whale Anticline represents the most southerly of a series of NNE- to north-trending domes and anticlines that are extensively developed along the Norwegian margin (Fig. 1), though it is separated from them by a considerable expanse in the Møre Basin where no anticlines have been reported. To the SW of the Pilot Whale Anticline, the compressional folds follow mainly NE trends within the Faroe–Shetland Basin.

NE- to NNE-trending mud anticlines and diapirs

Linear NE- to NNE-trending mud anticlines with axial traces varying between 5 and 20 km in length have been described by Lamers & Carmichael (1999, fig. 13) around the southern half of Quadrant 214 (Fig. 2). Those authors believe that Upper Cretaceous and lowermost Paleocene (T10 stratigraphic interval of Ebdon et al. 1995) argillaceous rocks within the Faroe–Shetland Basin were mobilized during a period of ‘instability’ in early Paleocene (T20 to T30 intervals) times, to form the basement-detached, linear mudstone-cored anticlines. Speculatively, it is conceivable that a distinct phase of compression in early Paleocene times was the direct trigger for the formation of these structures, initiating their mobilization and development into discrete linear mudstone walls or anticlines. Such compression may have been associated with inversion of postulated underlying mainly NE-trending extensional Mesozoic faults. The effects of compression are described elsewhere within the Faroe–Shetland area during Paleocene times, with for example, broad inversion of the Judd sub-basin (Roberts et al. 1999) and also fold growth associated with the initiation of the Wyville Thomson Ridge (Fig. 3) (e.g. Boldreel & Andersen 1993; Johnson et al. 2005).

Judd Anticline

The Judd Anticline/Monocline is interpreted to form an important, generally west-trending structure (Smallwood & Kirk 2005, fig. 15) that extends for at least 33 km within the Judd sub-basin (Figs 5 & 6(i)). However, the west-trending ‘Judd Anticline’ of Smallwood (2004, fig. 13) refers to a broad structural feature that comprises an amalgamation of the Judd, Westray and South Judd anticlines as described in this paper. Seismic reflection evidence suggests that there are at least two main phases of fold growth associated with the development of this anticline. Considerable depositional thinning of the lower to middle Eocene (latest Ypresian to late Lutetian) succession within the NW limb of the Judd Anticline (Fig. 6(i)) is interpreted to be indicative of the first pulse of deformation. The onset of this deformation is marked by the development of the LE (early Eocene i.e. latest Ypresian) Unconformity . This was followed by a more protracted phase of fold growth during mid Eocene times (Fig. 3). Significant attenuation of the lower to middle Eocene succession occurs on the SE limb of the anticline in particular, at the prominent L3 (late Lutetian) Unconformity . In addition, the overlying middle Eocene to Oligocene succession is also deformed, though the age of this deformation cannot be accurately ascertained as the upper part of the interval is absent through erosion over the crest of the Judd Anticline due to the development of the Top Palaeogene Unconformity (TPU) . Notwithstanding this uncertainty, it is suggested that this latter phase of folding occurred sometime during the late Oligocene. This interpretation supports that of Smallwood (2004), who recognized several phases of deformation including Eocene (latest Ypresian and late Lutetian), Oligocene and ?mid Miocene times, associated with NE–SW orientated compressional stress.

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Fig. 6.

(i) Seismic profile B–B′ across the Judd Anticline, (ii) Seismic profile C–C′ across the South Judd and Westray anticlines and (iii) Seismic line D–D′ across the South Judd, Judd and Westray anticlines. See Figures 3 and 5 for age of reflectors and location of profiles, respectively. Abbreviations: INU, Intra-Neogene (Early Pliocene) Unconformity; L3, Mid Eocene (late Lutetian); LE, Early Eocene (latest Ypresian) Unconformity; MMU, Mid Miocene Unconformity; TBF, Top Balder Formation; TPU, Top Palaeogene Unconformity.

Westray Anticline

The term Westray Anticline is informally introduced here for a generally asymmetrical, NW-trending growth fold within the Judd sub-basin (Figs 5 & 6(ii)). The axial trace of this anticline extends for approximately 45 km, and its amplitude increases markedly towards the SE where it merges with the eastern deformational margin of the ‘Judd Anticline’ as defined by Smallwood (2004, figs 12 & 13). At least three main phases of growth have been recognized for the Westray Anticline. Like the Judd Anticline, a significant onlap and thinning of the lower to middle Eocene (latest Ypresian–late Lutetian) succession occurs on the NE limb of the Westray Anticline (Fig. 6(ii)) and is considered to mark a distinct phase of fold growth activity during latest early Eocene (latest Ypresian times) times (Fig. 3). This was followed by a more gradual and protracted phase of growth until mid Eocene (late Lutetian) times. The lower to middle Eocene interval is absent on the SW flank of the structure, possibly due in part to erosion associated with the development of the L3 (late Lutetian) Unconformity. There is also evidence for depositional thinning of the upper part of the middle Eocene to Oligocene succession across the Westray Anticline. Here for example, a distinct debrite flow with chaotic seismic facies onlaps the SW flank of the structure. This suggests that the fold had significant penecontemporaneous relief, and exerted considerable influence on the distribution of the debrite. Rejuvenation of fold growth occurred during development of the TPU (Fig. 3) and was associated with onlap of Miocene to lower Pliocene sediments, particularly on the NE limb of the anticline. This was followed by a further phase of growth associated with formation of the Neogene Unconformity (INU) , with onlap of Pliocene to Recent sediments onto the SW margin of the Westray Anticline. Indeed the seabed itself appears to be slightly domed over the Westray Anticline (Fig. 6ii), although this may be caused in part by differential subsidence on its flanks due to the effects of sediment loading and compaction rather than recent compression.

South Judd Anticline

The South Judd Anticline is considered to form a short, slightly asymmetrical NW-trending growth fold and occurs within the SW part of the Judd sub-basin (Figs 5 & 6(iii)). The axial trace of the anticline extends for approximately 18 km. In contrast to the Westray Anticline, its amplitude increases towards the NW, where it merges with the east-trending Judd Anticline. There is only evidence for one major phase of deformation associated with the South Judd Anticline, though its exact timing is not easily defined due the truncation of key stratigraphic units. Generally, the parallel-bedded Palaeogene (and possibly even the Miocene to early Pliocene) succession as a whole appears deformed by a single episode of compression (Fig. 6(iii)). However, the stratigraphic succession has been truncated over the crestal part of the South Judd Anticline by a composite unconformity that combines the TPU and INU (and possibly others too). The main age of the deformation is clearly post-Eocene and is probably of late Oligocene age; similar to the proposed growth on the parallel Westray Anticline (Fig. 3).

In terms of local controls on the style and distribution of the fold structures within the Judd sub-basin, the Westray and South Judd anticlines appear to be closely spatially associated with the inferred traces of the Westray and Judd lineaments of Rumph et al. (1993) and from BGS unpublished information (Fig. 5). These NW-trending lineaments possibly originated as Proterozoic Laxfordian-style shear zones, similar to those described from the Lewisian Complex on mainland Scotland and the Outer Isles (e.g. Park et al. 2002). Offshore, these lineaments are considered to have been reactivated as Permo-Triassic, Cretaceous and Cenozoic transfer zones (e.g. Earle et al. 1989; Rumph et al. 1993; Dean et al. 1999), partitioning the NE-trending sub-basins of the Faroe–Shetland Basin. Transfer zones typically have components of strike-slip movement associated with them (e.g. Lister et al. 1986) and there is a possibility that Westray and South Judd folds represent positive flower structures. A more plausible explanation, however, is derived from the observation that the anticlines have a close spatial association with footwall blocks or ramps of the Judd and Westray highs (Fig. 5), and that the folds probably developed as a consequence of buttressing of the sedimentary succession against these relatively rigid structural highs.

The east-trending Judd Anticline has a rather anomalous orientation with regard to the mainly NE-trending Caledonian structural grain prevalent within the Faroe–Shetland Basin. A plausible explanation could be that the anticline formed in response to north–south compression during Eocene times, as a result of ridge-push associated with seafloor spreading at the Aegir Ridge to the north of the Faroe Islands (Boldreel & Andersen 1993) which persisted until Chron 12 times (earliest Oligocene) (Lundin & Doré 2005). This view is supported by Smallwood & Kirk (2005), who believe that inversion was caused by north–south compressional stress and was concentrated between the Westray Fault and the Judd and Rona highs.

The variation in the phases of fold growth of the Judd, South Judd and Westray anticlines (Fig. 3) may suggest that they should be considered as distinct (though generically-linked) structural features, and not as a single compressional structure as described by Smallwood (2004).

Wyville Thomson Ridge

The Wyville Thomson Ridge forms a large symmetrical, WNW-trending anticline, with an axial trace that extends for more than 200 km (Figs 7 & 8). The fold has an amplitude and wavelength of approximately 2 km and 40 km, respectively, at the level of the TPL (Fig. 8). A number of deformational phases are considered to have contributed to the evolution of the Wyville Thomson Ridge. On the NE flank of the anticline, the lower and middle parts of the Eocene succession (between the Intra Miocene Unconformity (IMU) and an intra-Eocene Unconformity (IEUb) and possibly also the upper Paleocene to Eocene lavas) thin towards the ridge, suggesting a long-lived episode of Paleocene and Eocene growth (Johnson et al. 2005) (Fig. 3). The Eocene to Oligocene interval above the IEUb event is strongly attenuated; the angular break interpreted to represent a composite unconformity that combines the effects of the C30 (late Eocene) , TPU IMU unconformities (e.g. Johnson et al. 2005). On the basis of regional evidence, the development of the IMU is regarded to be the most significant of these unconformities and considered to be associated with an early to mid Miocene phase of growth on the Wyville Thomson Ridge (e.g. Johnson et al. 2005; Stoker et al. 2005a, b) (Fig. 3). According to Boldreel & Andersen (1993) and Tate et al. (1999), the Wyville Thomson Ridge represents part of a system of ramp-anticlines, which include the Ymir and Munkagrunnur ridges (see below) that formed above a crustal detachment that was active during latest Paleocene times and onwards as a result of mainly north–south compression associated with the spreading on the Aegir Ridge. However, a minor component of the fold growth may be due to strain partitioning associated with regional sinistral transpression along the NW European margin during Palaeogene times (e.g. Imber et al. 2005).

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Fig. 8.

Seismic profile E–E′ across the Wyville Thomson and Ymir ridges. See Figures 3 and 7 for age of reflectors and location of profile, respectively. Abbreviations: C10, Early Pliocene Unconformity; IEUa, b and c, Intra-Eocene unconformities; IMU, Intra-Miocene Unconformity; INU, Intra-Neogene (Early Pliocene) Unconformity; TPL, Top Palaeogene lavas. Note that Opal A – Opal C/T transformation horizon is a diagenetic reflector of late Miocene age (Davies & Cartwright 2002).

Ymir Ridge

The Ymir Ridge forms an asymmetrical, NW-trending and faulted anticlinal complex that extends for over 100 km (Figs 7 & 8). The fold amplitude and wavelength are approximately 1.4 km and 40 km respectively at the level of the TPL (Fig. 8). The Eocene succession on the SW flank of the Ymir Ridge displays growth folds and is cut by penecontemporaneous reverse faults. The timing of this deformation is difficult to define accurately, as the Eocene succession is strongly truncated by a significant composite unconformity surface that combines the effects of the C30 (Late Eocene) , TPU and IMU . As for the Wyville Thomson Ridge, the Miocene unconformity surface is interpreted to mark the most important of the unconformity surfaces and may have formed as a result of significant pulse of compression within the Faroe–Shetland region (e.g. Johnson et al. 2005) (Fig. 3).

Munkagrunnur Ridge

The Munkagrunnur Ridge forms a slightly asymmetrical elongate NNW-trending anticlinal feature (Keser Neish 2003; Smallwood 2005) that extends for at least 135 km (Figs 7 & 8), though the nature of its transition with the Faroe Platform to the NW is unclear. The nature and origin of the Munkagrunnur Ridge is poorly understood, but the results of potential field modelling suggest that it comprises a crystalline basement block, capped by less that 2 km of pre-Eocene strata and 1 km of folded Palaeogene lavas (Smallwood et al. 2001). Because these deformed lavas crop out at the seabed over a large proportion of the ridge (Keser Neish 2003), the age of any compressional deformation is difficult to assess, other that it must have occurred in latest Paleocene or later times. By analogy with the Wyville Thomson and Ymir ridges, this deformation may have occurred during Eocene, Oligocene and Miocene times.

Alpin Dome

The Alpin Dome forms a little surveyed, slightly asymmetrical, east-trending anticline that extends for approximately 50 km in the North Rockall Basin (STRATAGEM partners 2002; Stoker et al. 2005a) (Fig. 7). The timing of deformation is not fully understood as yet, but it appears that a major phase of compression may have occurred during late Eocene (C30) and mid Miocene (IMU) times on the SW flank of the anticline, with the possibility of an additional late Oligocene (TPU) phase which only significantly affected the NE flank (Stoker et al. 2005a) (Fig. 3). Speculatively a local causal mechanism for the formation of the anticline could be the buttressing of the post-Palaeogene lavas sedimentary section against the southern flank of the Sigmundur Seamount (Fig. 7).

Structural trends and controls

Essentially, there are six main orientations of growth fold within the study area: (1) NE-trending; (2) NNE-trending; (3) east-trending; (4) NW-trending; (5) NNW-trending; and (6) WNW-trending. The pre-existing structural architecture of the underlying sedimentary basins is likely to have exerted a considerable influence on the development of these fold trends.

The NE-trending anticlines are by far the most numerous within the study area but are mainly restricted to the NE Faroe–Shetland Basin. Here, fold generation is inferred to be linked to buttressing of strata against pre-existing basement-involved architecture that has a strong inherited Caledonian or pre-Caledonian structural grain. If pre-existing structure is a key factor in the determination of the orientations of Cenozoic growth folds, then this implies that pre-existing Caledonian structure may be much less important in the SW Faroe–Shetland Basin and Wyville Thomson areas. In these areas, the presence of Palaeogene igneous centres and generally NW-trending shear zones may be more influential in the development of growth folds. NNE-trending Cenozoic reverse faults are rare, but a single example associated with the Pilot Whale Anticline in the NE Faroe–Shetland Basin formed either by buttressing against the Brendan Volcanic Centre, or by growth associated with left-lateral strike-slip fault movement on the NW-trending Magnus Lineament or Transfer Fault. East-trending faults are also rare, but the slightly arcuate Alpin Dome in the North Rockall Basin may have formed as a result in compression of strata against the southern flank of the Sigmundur Seamount. The similarly trending Judd Anticline within the Judd sub-basin developed during phases of NE to SW compression, though the degree to which the underlying structural architecture of the area was an influence on its formation remains unclear. The Munkagrunnur and Westray anticlines form NW- to NNW-trending growth folds, with the latter considered to represent part of a system of ramp anticlines formed during phases of north–south compression associated with the development of the Aegir Ridge. The Westray Anticline lies parallel to, and in close proximity, to the Westray Lineament or Transfer Fault and extends between basement blocks that form the north and south Westray High (Fig. 2). It is probably associated with buttressing against a step in the existing basement architecture although, alternatively, it may represent a flower structure associated with lateral shear along the Westray Lineament. Similarly, the NW-trending South Judd Anticline appears to be almost coincident with the Judd Lineament or Judd Transfer Fault. The most probable cause of formation is buttressing of sedimentary strata against the basement of the juxtaposed Judd High. The NW-trending Ymir Ridge and WNW-trending Wyville Thomson Ridge are considered to form part of a series of ramp anticlines similar to that described for the Munkagrunnur Ridge. The Wyville Thomson Ridge is unlikely to have a significant component of strike-slip movement associated with its formation, as the platform margin to the SE is not offset (Kimbell et al. 2005).

Regional deformation history

Within the context of the NE Atlantic margin as a whole, it appears that the earliest phase of Cenozoic fold growth development occurred in late Paleocene to early Eocene times, particularly on the Hatton High, Hatton Basin, Wyville Thomson and Ymir ridges (Fig. 3). During mid to late Eocene times, significant development of anticlinal structures occurred within the SW Faroe–Shetland Basin, Hatton High and Basin, respectively. In mid Eocene to early Oligocene times, major fold growth of north- to NNW trending domes occurred within the Norwegian margin, with a second major pulse prevalent during early Miocene times (e.g. Lundin & Doré 2002). These two pulses of activity resulted in the formation of margin-wide unconformities. The early Miocene event correlates with a major episode of fold growth observed within the NE Faroe–Shetland Basin, though here, there is also a significant slightly younger mid Miocene phase too (e.g. Ritchie et al. 2003). There are also indications of Miocene fold development in the SW Faroe–Shetland Basin (Smallwood 2004) and particularly around the Wyville Thomson and Ymir ridge area (e.g. Johnson et al. 2005; Stoker et al. 2005a, b). The effects of localized early Pliocene deformation is observed throughout the Faroe–Shetland Basin, along the Wyville Thomson and Ymir ridges and within the Hatton High and Basin areas. However, this phase of deformation should not be confused with the effects of regional Pliocene epeirogenic uplift and continental margin tilting as described by Stoker et al. (2005a, b) and Praeg et al. (2005). Some anticlines within the Faroe–Shetland and Wyville Thomson Ridge areas have a positive topographic expression at the seabed (Fig. 8) (e.g. Johnson et al. 2005 figs 3 & 6) suggesting that fold growth activity may still be occuring.

Regional mechanisms causing compression

There have been many models suggested to account for the distribution, orientation and age of growth folds long the NE Atlantic margin between Norway and Greenland in the north, and the Hatton–Rockall area in the south. These include the closing of the Tethys Ocean (Roberts 1989), strike-slip fault movement and ‘shuffling’ along NW-trending transfer zones/lineaments (Doré & Lundin 1996), a combination of seafloor spreading geometries and ridge-push (Boldreel & Andersen 1993), intraplate stress (e.g. Cloetingh 1990), the pulsing plume hypothesis (Lundin & Doré 2002), ridge-push and mantle drag forces (Mosar et al. 2002) and plume-enhanced asthenospheric flow (e.g. Kusznir 2006). As there is presently no obvious regional spatial and temporal pattern to the evolution of compressional structures throughout the NE Atlantic margin, it is considered likely that a combination of mechanisms might best account for this. However, progress regarding an assessment of the relative importance of individual mechanisms is, we believe probably substantially hindered by poor calibration (i.e. inadequate dating) of growth folds (and hence regional correlation).