A refraction seismic transect from the Faroe Islands to the Hatton-Rockall Basin

Transcription

1 JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 113,, doi: /2008jb005675, 2008 A refraction seismic transect from the Faroe Islands to the Hatton-Rockall Basin Thomas Funck, 1 Morten S. Andersen, 1 Judith Keser Neish, 2 and Trine Dahl-Jensen 1 Received 7 March 2008; revised 3 July 2008; accepted 28 July 2008; published 11 December [1] The crustal structure of the Faroe-Rockall Plateau was studied by a 790-km-long refraction seismic transect consisting of two intersecting lines. The air gun shot spacing was 200 m, and the signals were recorded by 77 ocean bottom seismometers. A P wave velocity model was developed from forward and inverse modeling of the wide-angle seismic data and incorporation of coincident multichannel reflection seismic data. Continental crust with velocities ranging from 5.6 to 6.8 km/s can be traced from the Faroe Islands, across the banks to the SW of the Faroes and into the Hatton-Rockall Basin. The thickness of the subvolcanic crust is up to 25 km on the banks but is as little as 8 km in the channels between the banks. The thinning in the channels may be related to NW-trending shear zones extending from major lineaments in NE Rockall Trough. Basalt layers are found along the entire transect with a total thickness of up to 4 km. Two layers with velocities of and km/s are thought to represent Paleogene flood basalts that can be correlated from the Faroe Islands to George Bligh Bank. Close to George Bligh Bank, an 80-km-wide and up to 9-km-thick body with velocities of 6.5 km/s is interpreted as intrusion. A 5-km-thick, high-velocity lower crustal layer (7.3 km/s) extends from the area of the intrusion into the northern Hatton-Rockall Basin. At the northern flank of Lousy Bank the transition zone to oceanic crust was encountered. Citation: Funck, T., M. S. Andersen, J. Keser Neish, and T. Dahl-Jensen (2008), A refraction seismic transect from the Faroe Islands to the Hatton-Rockall Basin, J. Geophys. Res., 113,, doi: /2008jb Introduction [2] The NW European continental margin (Figure 1) is characterized by a long history of rifting in the Rockall Trough and the Faroe-Shetland Trough, during which the Faroe-Rockall Plateau was separated from the British Isles prior to the final opening of the northeast Atlantic between east Greenland and the Faroe-Rockall Plateau [e.g., Knott et al., 1993]. The final breakup in the Tertiary was accompanied by massive volcanism that formed the North Atlantic Igneous Province. Paleogene basalts are attributed to the Iceland plume [White, 1992] and are found in most parts of the Faroe-Rockall Plateau, the Rockall Trough and the Faroe-Shetland Trough. Poor seismic data quality for subbasalt and intrabasalt arrivals is not uncommon in this area and this is true for both reflection and refraction seismic data. Fliedner and White [2003] summarize some of the problems that are associated with the seismic imaging of subbasalt series. [3] There is some controversy on the crustal composition within Rockall Trough, which has to be partly attributed to poor deep-crustal seismic data and to ambiguous velocity functions derived from those data. Bott et al. [1979] assumed oceanic crust in Rockall Trough. Later Roberts et 1 Geological Survey of Denmark and Greenland, Copenhagen, Denmark. 2 Faroese Earth and Energy Directorate, Tórshavn, Faroe Islands. Copyright 2008 by the American Geophysical Union /08/2008JB005675$09.00 al. [1988] concluded that the crust within Rockall Trough is continental in character, although Smythe [1989] had some doubt that this conclusion was supported by the data. Joppen and White [1990] compiled a crustal velocity profile and concluded that these velocities are consistent with either oceanic crust or with thinned continental crust heavily intruded by syn-rift igneous rocks. Later publications, in particular the RAPIDS data, favored a continental affinity of the crust [Hauser et al., 1995; O Reilly et al., 1995; Morewood et al., 2005]. [4] Published refraction seismic data show clear continental crust on the Faroe-Rockall Plateau, both for the Hatton and Rockall Bank [Morgan et al., 1989; Keser Neish, 1993; Vogt et al., 1998] and beneath the Faroe Islands [Bott et al., 1976; Richardson et al., 1998, 1999]. However, the area between the Faroes and the Hatton and Rockall banks is less well studied. This part of the Faroe- Rockall Plateau is segmented into a number of shoal banks and intermittent channels (Figure 2). The banks to the SW of the Faroe Islands include the Faroe Bank, Bill Bailey Bank, Lousy Bank, and George Bligh Bank. Roberts et al. [1983] suggested a continental affinity of the banks on the basis of plate reconstructions. Refraction seismic evidence for this is only available for Lousy Bank, where Klingelhöfer et al. [2005] found 24-km-thick continental crust. Between Lousy Bank and George Bligh Bank, the crust thins to 14 km [Klingelhöfer et al., 2005]. [5] To obtain a more complete image of the crustal structure of the bank area SW of the Faroes, a wide-angle 1of25

2 Figure 1. Physiographic map of the study area. The elevation model is shaded by artificial illumination from the southeast. Red solid lines show the location of the two refraction seismic profiles of this study. Dashed lines show the location of other seismic experiments that are discussed in the text: lines AMP-D, AMP-E, and BANS-1 [Klingelhöfer et al., 2005]; isimm line 5 [Smith et al., 2005; White et al., 2008]; lines 87 3 [Keser Neish, 1993]; and RAPIDS lines 21, 13, and 14 [Vogt et al., 1998]. Filled circles indicate the location of DSDP wells at sites 116 and 117 [Laughton et al., 1972]. Elevation models: onshore GTOPO30 (U.S. Geological Survey); offshore [Smith and Sandwell, 1997]. Abbreviations are as follows: BBB, Bill Bailey Bank; EB, Edoras Bank; FB, Faroe Bank; GBB, George Bligh Bank; HRB, Hatton-Rockall Basin; LB, Lousy Bank; and RB, Rosemary Bank. seismic experiment was carried out in The transect consists of two lines extending from the southern Faroe Islands, across Faroe Bank, Bill Bailey Bank, Lousy Bank, the NW flank of George Bligh Bank, and into the Hatton- Rockall Basin (Figure 2). In addition, coincident multichannel seismic (MCS) data were acquired along these lines. The objective of the experiment was to determine the nature of the underlying crust in the banks area and to identify the processes that have shaped the present morphology of the banks. What has caused the segmentation of the Faroe- Rockall Plateau adjacent to the NE Rockall Trough into some 100-km-wide banks and what is the nature of the intermittent channels? How much was the banks area modified by Paleocene flood basalts [Waagstein, 1988] and magmatic underplating during the opening of the NE Atlantic, and how did the rifting in Rockall Trough affect the present crustal configuration of the banks? 2. Geological Setting [6] The Faroe-Rockall Plateau extends from the Faroe Islands in the NW to the Hatton and Rockall Banks in the SW and is separated from the British and Irish continental shelf by the Rockall Trough and the Faroe-Shetland Trough (Figure 1). Similar to the discussion of the nature of the crust in Rockall Trough (see above), there is also some debate on the age of Rockall Trough [e.g., Corfield et al., 1999; Shannon et al., 1999; Morewood et al., 2004] but it is generally agreed that rifting occurred in several episodes. On the basis of plate reconstructions, the rifting may have started in end-carboniferous to Early Permian time [Knott et al., 1993]. The last episode of rifting occurred in the Late Cretaceous [Knott et al., 1993] and may have continued into the Eocene as suggested by plate tectonic modeling [Cole and Peachey, 1999]. [7] Opening of the modern NE Atlantic occurred between east Greenland and the Faroe-Rockall Plateau; the age of the continent-ocean boundary was determined to follow very shortly after magnetic chron C25n between 56 and 55.5 Ma [Tegner et al., 1998; Storey et al., 1998; Holbrook et al., 2001]. Just prior to and during breakup, extensive volcanism occurred in the Faroe-Rockall area, which is attributed to the arrival of the Iceland plume in that region [White, 2of25

3 Figure 2. Bathymetric map of the study area. Lines show the location of the refraction seismic experiment. Positions of ocean bottom seismometers are marked by open circles, and numbers indicate the station number. The filled circle shows the location of the Lopra well [Chalmers and Waagstein, 2006]. Other seismic experiments discussed in the text are marked by dashed lines: lines AMP-D, AMP- E, and BANS-1 [Klingelhöfer et al., 2005]. Bathymetric data are from Smith and Sandwell [1997]; the contour interval is 200 m. Abbreviations are as follows: BBB, Bill Bailey Bank; FBC, Faroe Bank Channel; GBB, George Bligh Bank; LB, Lousy Bank; RB, Rosemary Bank; and WTR, Wyville- Thomson Ridge. 1992]. On the Faroe Islands, the basalts are divided into lower, middle and upper formation. The lower formation was extruded between 59 and 56 Ma, the two higher formations during chron C24r at 55 Ma [Waagstein, 1988; Waagstein et al., 2002]. Recently, Passey and Bell [2007] introduced a new stratigraphic division, in which the upper, middle and lower basalt series roughly correspond to the Enni, Malinstindur and Beinisvør* formations, respectively. The total thickness of the basalts on the Faroes is >5 km [Waagstein, 1988; Richardson et al., 1998]. [8] The continental margin to the NE of the Faroe-Rockall Plateau displays features that are typical of volcanic margins. Seaward dipping reflectors, interpreted as subaerial basalt flows, are found on many reflection seismic profiles [e.g., Smythe, 1983; Morgan et al., 1989; Barton and White, 1997a, 1997b]. Refraction seismic data provide evidence for thick lower crustal layers with high seismic velocities (>7.2 km/s) in the continent-ocean transition zone, interpreted as magmatic underplating, for example at Lousy Bank [Klingelhöfer et al., 2005], Hatton and Edoras banks [Fowler et al., 1989; Barton and White, 1997a; Vogt et al., 1998]. 3. Wide-Angle Seismic Experiment 3.1. Data Acquisition and Processing [9] The refraction seismic experiment was part of a twoship experiment carried out in The Danish ship Esvagt Connector was used for the deployment and recovery of the ocean bottom seismometers (OBS), while the Norwegian commercial seismic vessel Polar Princess fired the air gun array and collected coincident multichannelseismic data along the refraction seismic lines. [10] A total of 40 OBS were available for the experiment. The Institut Français de Recherche pour l Exploitation de la Mer (Ifremer) and the German GEOMAR Research Centre for Marine Geosciences provided 15 and 25 instruments, respectively. The Ifremer OBS were equipped with a hydrophone and with three-component 4.5-Hz geophones. The same applies to 16 of the GEOMAR OBS, whereas the remaining instruments only had a hydrophone component. 3of25

4 Figure 3. (top) Record section with computed travel times and (bottom) raypath diagram for the hydrophone of OBS 48 (line B) located in the center of Lousy Bank. Horizontal scale in the record section is shot-receiver distance (offset), and the vertical scale is the travel time using a reduction velocity of 6.5 km/s. A triangle indicates the receiver location. See text for description of phases. The horizontal scale of the raypath diagram is distance along the velocity model (Figure 11). LVZ, low-velocity zone. On the second line (line B), two of the hydrophone systems were upgraded with additional three-component geophones. [11] The experiment consisted of two lines (Figure 2). Line A is a 395-km-long transect that runs from the southern Faroe Islands, across the northern flanks of the Faroe and Bill Bailey banks and across the center of Lousy Bank. Line B is 390 km long and extends from Lousy Bank, across the northwesternmost George Bligh Bank into the Hatton-Rockall Basin. Along line A, 40 OBS were deployed at sites 1 through 40. Along line B, 39 OBS were deployed at sites 41 through 79. The nominal station spacing was 10 km. No data were retrieved for sites 19 and 69. [12] The seismic source was a tuned air gun array that consisted of 40 guns ranging in size from 0.7 to 4.9 L; the total volume was 95 L (5800 cubic inches). The array was fired every 200 m. For navigation, a differential Global Positioning System (GPS) was used. Water depths along the lines were obtained from the echo sounder onboard the vessel Polar Princess. [13] After recovery of the OBS, the data were dumped to disk, corrected for OBS clock drift, converted to SEGY format, and debiased. Travel time picks of the direct wave were used to recalculate the position of the instruments at the seafloor, from which shot-receiver ranges were calculated. The maximum distance between deployment and recalculated position was 378 m. For the display of the record sections (Figures 3 9), a band-pass filter from 5 to 24 Hz was applied. In addition, some records were disturbed by signals with a frequency of 5 Hz that was removed by a notch filter. On many records a deconvolution improved the recognition of seismic phases. [14] The coincident reflection seismic lines were collected with an 8.1-km-long streamer with 648 channels, the shot spacing was 25 m and the volume of the air gun array was 76 L. Time-migrated record sections were produced from the data Methodology [15] The goal of the analysis of the refraction seismic data was to obtain a two-dimensional velocity model for the sediments, basalts, crust, and uppermost mantle along the two lines. Even though some shear wave energy was identified in the record sections, the modeling was limited to the P waves. Both lines were shot along great circle arcs that define the baselines for the velocity models. The recalculated OBS positions at the seafloor were projected onto this baseline. [16] The P wave velocity models were developed using the program RAYINVR [Zelt and Smith, 1992; Zelt and Forsyth, 1994]. Initially, forward models were developed from top to bottom (seafloor to mantle) by fitting the observed travel times. Layer boundaries within the sedimentary column and partly within the basalt sequences were 4of25

5 Figure 4. (top) Record section with computed travel times and (bottom) raypath diagram for the vertical geophone of OBS 44 (line B) located at the northeast flank of Lousy Bank. Horizontal scale in the record section is shot-receiver distance (offset), and the vertical scale is the travel time using a reduction velocity of 6.5 km/s. A triangle indicates the receiver location. See text for description of phases. The horizontal scale of the raypath diagram is distance along the velocity model (Figure 11). B, basalt layer. primarily defined by the coincident MCS data. Later, velocities within layers were optimized by using the inversion algorithm in RAYINVR Seismic Data [17] For this survey, a total of 266 record sections were available for the data analysis, counting only the geophone/ hydrophone components of stations from which data were retrieved. A selection of record sections is briefly discussed below to get an idea of the main features in the seismograms. Phase names in the records reflect the later interpretation of the velocity model. P S1 through P S11 label P wave refractions within sedimentary layers and P S1 P through P S11 P label the reflections from the base of each of these sediment layers. P B1 through P B5 define refractions within the main basalt layers and within subbasalt layers; P B1 P through P B5 P are the corresponding reflections from the base of these layers. The underlying basement is divided into four crustal layers with the corresponding refractions P c1 through P c4 ; the intracrustal reflections are P c1 P through P c3 P. P m P and P n denote the Moho reflection and mantle refraction, respectively. [18] The data quality is generally good given the widespread occurrence of basaltic sills and lava flows on the Faroe-Rockall Plateau. Reflection seismic images of subbasalt structure are often degraded owing to poor penetration, scattering, and attenuation of high frequencies [Chironi et al., 2006]. The lower frequencies used in refraction seismic experiments are less influenced by these effects [Fliedner and White, 2001]. Basalt layers often overlie sedimentary rocks or other volcanic layers with lower velocities, from which no refractions can be observed. Such low-velocity zones are very characteristic for our data set and are illustrated in Figure 3, where the low-velocity zone beneath basalt layer 2 creates a typical time delay before deeper arrivals are observed. On some records, the signalto-noise ratio decreases significantly for deeper crustal phases (Figure 3), which is in particular true for OBS 74 through 79 in Hatton-Rockall Basin, where two low-velocity zones are observed. High noise levels there may also be related to strong bottom currents in the basin [Hitchen, 2004]. [19] Despite these local complications, there are a number of excellent record sections that allow the correlation of seismic energy up to offsets of 160 km. An example for this is OBS 44 located at the northern flank of Lousy Bank. The record (Figure 4) displays a high-amplitude P m P phase at offsets between 45 and 80 km and a strong mantle refraction (P n ). No evidence for a high-velocity layer at the base of the crust is observed on this record, in contrast to OBS 66 at the NW flank of George Bligh Bank (Figure 5). Here, two reflections P c3 P and P m P are observed that mark the top and the base of the high-velocity lower crustal layer. These reflections are not as distinct as the P m P on OBS 44 owing to the reduced velocity contrast across the layer 5of25

6 Figure 5. (top) Record section with computed travel times and (bottom) raypath diagram for the vertical geophone of OBS 66 (line B) located northwest of George Bligh Bank. Horizontal scale in the record section is shot-receiver distance (offset), and the vertical scale is the travel time using a reduction velocity of 6.5 km/s. A triangle indicates the receiver location. See text for description of phases. The horizontal scale of the raypath diagram is distance along the velocity model (Figure 11). boundaries. Other examples for record sections with P c3 P and P m P reflections are shown close up in Figures 6 and 7. High-amplitude midcrustal reflections (P c2 P) are observed to the NE of OBS 66 (Figure 5). This feature is limited to an 80-km-wide zone on line B and indicates a crustal structure that is different from the remainder of the two lines. [20] The variable topography along the lines with the banks and intermittent channels results in undulating travel time curves as seen on OBS 4 (Figure 8), where arrival times from shots in the channels at offsets of 30 and 130 km are delayed compared to the shots on Faroe Bank (offset 60 km). This variable topography is also the reason for the asymmetric shape of the travel time curves to either side of the OBS. The phase velocity of the P c1 to the SW of OBS 4 is for example much lower than to the NE (4.7 km/s versus 5.8 km/s), which requires careful interpretation. Some record sections lack the midcrustal P c2 refraction, as seen on OBS 18 for observations to the SW (Figure 9). Here, only a P c1 and a P c3 phase are observed, with phase velocities of 5.5 and 6.4 km/s, respectively. 4. Results [21] Below, the P wave velocity models for lines A and B are presented. First, the models are described, then an account of the resolution and model uncertainties is given followed by two-dimensional gravity modeling to check for the consistency of the model with the gravity data. Finally, the velocity models are compared with other data from the region Velocity Models Line A [22] The P wave velocity model for line A is shown in Figure 10. Sediments and sedimentary rocks with velocities between 1.6 and 4.3 km/s are found in the channels between the banks, where they are up to 2 km thick. At these shallow levels, velocities close to 4 km/s and higher may indicate a substantial amount of volcanic and volcaniclastic material. The detailed geometry of the sedimentary layers was obtained from correlation with the coincident MCS data. [23] Below the sedimentary units, the basalt and subbasalt sequence is divided into three layers that can be correlated across the entire line. The upper two layers are interpreted as volcanic sequence on the basis of their velocities between 4.9 and 5.6 km/s. The third layer is a low-velocity zone (LVZ) and, hence, no refractions were observed that could have determined the velocity within the LVZ. This layer may have a volcanic composition, but it may as well contain some sedimentary rocks that predate the volcanism on the Faroe-Rockall Plateau. The total thickness of all three layers varies between 2 and 5 km assuming a velocity of 5.2 km/s in the LVZ. This velocity is not constrained by the seismic data, but velocities of 5.0 km/s were found on line B in 6of25

7 Line B [26] Hatton-Rockall Basin at the SW end of line B is characterized by up to 2-km-thick sedimentary units with velocities <3.7 km/s (Figure 11). Between George Bligh Bank and Lousy Bank, some layers with velocities between 3.9 and 4.7 km/s can be seen close to the seafloor. They may consist of lava flows or volcanic debris, possibly interbedded with sedimentary rock. The basalt and subbasalt sequence is up to 6 km thick with velocities up to 5.5 km/s. Similar to line A, the lowermost layer of the sequence is a LVZ along most of the line. However, in the Hatton-Rockall Basin, where the higher-velocity basalt cover is absent, refractions indicate a velocity of 5.0 km/s in the layer above the upper crust. [27] The crust along line B is laterally divided into two zones. To the NE of Lousy Bank crustal velocities are increased compared to the remainder of the line or to line A. The upper layer is 2 km thick and has velocities of 5.8 to 6.2 km/s, velocities in the 7.5-km-thick lower layer range from 6.9 to 7.0 km/s. White et al. [2008] interpret increased velocities in the continent-ocean transition of this volcanic margin as heavily intruded continental crust. A few internal Figure 6. (top) Record section with computed travel times and (bottom) raypath diagram for lower-crustal reflections for the vertical geophone of OBS 60 (line B) located southwest of Lousy Bank. Horizontal scale in the record section is shot-receiver distance (offset), and the vertical scale is the travel time using a reduction velocity of 6.5 km/s. A triangle indicates the receiver location. See text for description of phases. The horizontal scale of the raypath diagram is distance along the velocity model (Figure 11). Hatton-Rockall Basin at depths >5 km in a layer that may be the southward continuation of the LVZ (Figure 10). [24] The underlying basement is divided into three layers with a velocity structure that is typical for continental crust. The upper crust is characterized by velocities of 5.4 to 5.6 km/s and the layer thickness varies between 2 and 5 km. Midcrustal velocities range from 6.05 to 6.3 km/s. The maximum thickness of the midcrustal layer is 6 km at the Faroe Bank. Beneath Bill Bailey Bank and the Faroe Bank Channel there is no seismic evidence for these midcrustal velocities. The lower crust (6.55 to 6.8 km/s) is also very variable in thickness; between the banks the thickness is as low as 4 km whereas up to 18-km-thick lower crust exists beneath Bill Bailey Bank and Faroe Bank. The total thickness of all three crustal layers varies between 8 and 24 km, the overlying basalt sequence not included. [25] To the SW of Lousy Bank, some deep reflections (P c3 P and P m P) were observed defining the top and the base of a high-velocity lower crustal layer (Figure 7). The velocity within this layer is set to 7.4 km/s and the layer thickness is 6 km. Mantle velocities of 8.0 km/s are constrained along portions of the line. Figure 7. (top) Record section with computed travel times and (bottom) raypath diagram for lower-crustal reflections for the hydrophone of OBS 34 (line A) located on Lousy Bank. Horizontal scale in the record section is shot-receiver distance (offset), and the vertical scale is the travel time using a reduction velocity of 6.5 km/s. A triangle indicates the receiver location. See text for description of phases. The horizontal scale of the raypath diagram is distance along the velocity model (Figure 10). 7of25

8 Figure 8. (top) Record section with computed travel times and (bottom) raypath diagram for the vertical geophone of OBS 4 (line A) located at the eastern flank of the Faroe Bank Channel. Horizontal scale in the record section is shot-receiver distance (offset), and the vertical scale is the travel time using a reduction velocity of 6.5 km/s. A triangle indicates the receiver location. See text for description of phases. The horizontal scale of the raypath diagram is distance along the velocity model (Figure 10). crustal reflectors (P c1 P) are observed in that portion of line A, which would argue against normal oceanic crust as interpreted at the NW end of line AMP-E [Klingelhöfer et al., 2005]. [28] The remainder of the crust on line B has a threelayered structure similar to line A. The upper crust is 1 to 4 km thick (velocities 5.45 to 5.90 km/s); the lower crust is 3 to 13 km thick (velocities 6.45 to 6.80 km/s). Midcrustal velocities range from 6.1 to 6.2 km/s with exception of a segment to the west of George Bligh Bank, where higher velocities of 6.5 to 6.7 km/s are observed. This velocity increase correlates with a thickening of the midcrustal layer from 3 to 9 km. This high-velocity midcrustal body is interpreted as an intrusion and will be discussed later. At the SW end of the line, midcrustal velocities of 6.2 km/s are not observed. The total thickness of the continental crust varies between 18 km at Lousy Bank and 7.5 km beneath the Hatton-Rockall Basin. [29] A 5.5-km-thick lower crustal layer with a velocity of 7.25 km/s extends beneath George Bligh Bank and the Hatton-Rockall Basin. This layer is consistent with underplated mafic crust [White et al., 1987] or heavily intruded lower crust [White et al., 2008]. The high-velocity lower crust shown at the NE end of the line is not based on seismic data but on gravity modeling (see section 4.3). Velocities of 8.0 km/s are determined for the mantle Model Resolution and Uncertainty [30] Tables 1 and 2 summarize the formal error analysis for individual phases on both lines. The normalized c 2 is based on assigned pick uncertainties of 35 to 250 ms depending on the quality of each individual travel time pick. The pick uncertainties are graphically indicated in Figures 12 and 13. With these uncertainties, a normalized c 2 of 0.75 (line A) and 0.93 (line B) was obtained, slightly below the optimum value of 1, for which we see two possible explanations: (1) the pick uncertainties were slightly overestimated, or (2) the additional information extracted from the coincident MCS data (in particular layer geometry of the sedimentary units and basalts) has improved the fit for the upper part of the model. The root-mean square misfit between calculated and picked travel times is around 100 ms. [31] The diagonal values of the resolution matrix of the velocity nodes are a good indicator to distinguish between poor and well-resolved parts of a model. Values >0.5 indicate well-resolved model parameters [Lutter and Nowack, 1990]. Figures 10 and 11 show the resolution values for lines A and B, respectively. The three upper crustal layers are characterized by resolution values >0.5 with a few exceptions. These include the boundary zones of the model where the ray coverage is reduced (in particular the NE end of line A close to the Faroe Islands), the thin 8of25

9 Figure 9. (top) Record section with computed travel times and (bottom) raypath diagram for the vertical geophone of OBS 18 (line A) located between Bill Bailey Bank and Faroe Bank. Horizontal scale in the record section is shot-receiver distance (offset), and the vertical scale is the travel time using a reduction velocity of 6.5 km/s. A triangle indicates the receiver location. See text for description of phases. The horizontal scale of the raypath diagram is distance along the velocity model (Figure 10). upper crustal layer on top of the intrusion on line B, and a small portion of the lower crust in the channel between Lousy Bank and Bill Bailey Bank. Resolution within the basalt and subbasalt sequence is generally good apart from the low-velocity zones. Within the sedimentary layers, resolution values are variable and often <0.5. Low-resolution values within the sedimentary units and basalts mostly result from the lack of good reverse ray coverage in these shallow layers. However, additional constraints on the layer geometry from the coincident MCS data compensate for this shortcoming. Velocities within the highvelocity lower crustal layer are well-resolved in the central part of line B, and also on line A resolution values are >0.5 in some portions of the layer. [32] The plots with the ray coverage (Figures 12 and 13) provide information on how well individual portions of the models are sampled by rays and there is a close correlation with the resolution plots (Figures 10 and 11). In addition, it can be seen where layer boundaries are constrained by reflections. The Moho as well as the high-velocity lower crustal layer are almost continuously sampled by P c3 P and P m P reflections, providing good control on the crustal thickness and the thickness of the high-velocity zone. No P m P reflections were observed at the NE end of line A, preventing the determination of the crustal thickness beneath the Faroe Islands from this data set. The base of the intrusion (line B, km) is mapped in detail by reflections, while the midcrustal boundary along the remainder of the two profiles is less densely covered by P c2 P reflections. [33] Figure 14a shows the velocity models for lines A and B at their intersection. Both models were developed independently from each other and provide therefore an idea on the absolute error of the velocities and of the depth of boundary layers. Velocities are considered to be correct within ±0.1 km/s. The difference of 0.2 km/s for the welldetermined velocities within the basalts may be caused by anisotropy. Layer boundaries match within ±1 km Gravity Modeling [34] To verify the consistency of the velocity models (Figures 10 and 11) with available gravity data, twodimensional gravity modeling was performed using the algorithm of Talwani et al. [1959]. Gravity data were extracted from the satellite-derived free-air gravity [Sandwell and Smith, 1997] and density models were obtained from conversion of the P wave velocities to density using the curve shown in the work of Ludwig et al. [1970]. The calculated gravity for the density models shows a good agreement with the general pattern of the observed gravity (Figures 15 and 16). Most misfits are probably related to deviations from the two-dimensionality. However, at the NE end of line B, the misfit can be reduced substantially when a high-density lower crustal layer is 9of25

10 Figure 10. (top) P wave velocity model along line A. Numbers indicate velocity in km/s. Unconstrained velocities in low-velocity zones are shown in parentheses. The outer perimeter of the model with no ray coverage is omitted. Red circles mark the location of the OBS used for the modeling; the gray circle shows the location of OBS 19 with no data recovery. Position of intersection with line B and line AMP-E [Klingelhöfer et al., 2005] is marked at the top. (bottom) Diagonal values of the resolution matrix of the P wave velocity model. Abbreviations are FBC, Faroe Bank Channel; LVZ, lowvelocity zone. introduced (Figure 16) similar to the western flank of Lousy Bank on lines A and AMP-E [Klingelhöfer et al., 2005]. [35] The lithostatic pressure at the base of the density models is shown in Figures 15 and 16. In average, the pressure at 40 km depth is some 1160 MPa and deviations from the average are <35 MPa and <20 MPa on lines A and B, respectively. This indicates that the crust is isostatically balanced Comparison With Other Studies [36] Line AMP-E [Klingelhöfer et al., 2005] crosses Lousy Bank exactly at the intersection of lines A and B (Figure 2). This offers the opportunity to check how 10 of 25

11 Figure 11. (top) P wave velocity model along line B. Numbers indicate velocity in km/s. Unconstrained velocities in low-velocity zones are shown in parentheses. The outer perimeter of the model with no ray coverage is omitted. Red circles mark the location of the OBS used for the modeling; the gray circle shows the location of OBS 69 with no data recovery. Position of intersection with line A, line AMP-E [Klingelhöfer et al., 2005], and line BANS-1 [Klingelhöfer et al., 2005] is marked at the top. (bottom) Diagonal values of the resolution matrix of the P wave velocity model. Abbreviations are B, Basalt; HVZ, high-velocity zone. consistent our velocity models are with other data. Figure 14b shows that line AMP-E has a thin underplated layer at the intersection, while the high-velocity zone (HVZ) on line A has just pinched out. This discrepancy is of minor nature because the thinning and pinch out of the HVZ is not well resolved on either line. Thickness and velocities of the lower crust are almost identical, only the Moho depth is offset by 2 km. However, this is within the limits of the general uncertainty of the models. [37] Major differences become obvious in the top part of the velocity profiles. While line AMP-E has a 10-km-thick upper crustal layer with velocities between 5.6 and 6.4 km/s, 11 of 25

12 Table 1. Number of Observations, n, RMS Misfit Between Calculated and Picked Travel Times, t rms, and Normalized c 2 for Individual Phases on Line A Phase n t rms,ms c 2 Direct wave P S (all sediments) P S P (all sediments) P B (all basalts/subbasalts) P B P (all basalts/subbasalts) P c P c1 P P c P c2 P P c P c3 P P c P m P P n All phases line A divides the sequence into two layers with velocities of km/s and km/s, and a combined thickness of 6.5 km. The overall velocity range is not too dissimilar and the difference in upper crustal thickness can probably be attributed to some complications with the overlying basalt sequence. Line AMP-E shows a 4-km-thick basalt layer on top of Lousy Bank with a high-velocity gradient from 4.4 km/s at the top to 5.6 km/s at the base of the layer. In contrast, line A shows two basalt layers with velocities of 4.9 km/s and 5.6 km/s with an underlying LVZ of unknown composition (see discussion in section 5.1.4). This LVZ was not recognized on line AMP-E, which can be due to the data quality or some possible anisotropy within the basalts. However, it is interesting to note that the top of the basement on line AMP-E lies exactly in the middle between our 5.6-km/s basalt layer and our 5.6-km/s basement. This may indicate that it was difficult on the AMP data to distinguish between two refraction branches with a similar velocity that were slightly offset by a LVZ. [38] Figure 17 compares the results from line B with other deep-crustal seismic data from the Hatton-Rockall Basin, including line 87 3 [Keser Neish, 1993], isimm line 5 [White et al., 2008], and RAPIDS line 13 [Vogt et al., 1998] (for location, see Figure 1). The Moho depth on all four lines varies between 17 and 20 km and the velocities in the two upper crustal layers on line 87 3 and our line B are very similar (5.6 to 5.8 km/s in the upper layer and 6.5 km/s underneath). The most significant variations occur in the lower crust. Line B indicates a 5-km-thick HVZ at the base of the crust (7.3 km/s), which is absent on line 87 3 and RAPIDS line 13, while the tomographic results on isimm line 5 indicate lower crustal velocities of up to 7.3 km/s. [39] The HVZ on line B is well constrained by reflections from its top and the base as well as by a few refractions (Figure 13). The HVZ can be correlated farther to the north, where line BANS-1 (Figure 1) reports underplating [Klingelhöfer et al., 2005] at the intersection with line B as indicated in Figure 17. The fit at the cross point is not too well, which is partly related to the limitations in the BANS-1 data set and that the cross point is located right at the edge of an intrusion. If the models for line 87 3 and RAPIDS line 13 are right, this would suggest that the HVZ is restricted to the northern part of the basin. [40] Edwards [2002] inferred from potential field data that the rifted continental crust in the Hatton-Rockall Basin is characterized by numerous intrusions, similar to the intrusion that line B shows close to George Bligh Bank (Figure 11). Hence, the existence of a high-velocity lower crustal layer in association with all these intrusions is not implausible and we therefore want to have a closer look at the two southern lines. [41] The data set along line 87 3 includes MCS and sonobuoy data as well as expanding spread profiles [Keser Neish, 1993]. The crustal section shows the same Moho depth beneath Hatton Bank as underneath Hatton-Rockall Basin, which would result in an isostatic imbalance owing to the deeper sedimentary units in the basin. Replacing some of the lower crust with a denser underplated layer can reduce some of this isostatic imbalance. [42] Comparison between line B and RAPIDS line 13 (Figure 17) shows that the HVZ on line B and the lower crustal layer on the RAPIDS line are very similar in depth, both extending from ca. 15 km down to a depth of 20 km. This raises the question, if the lowermost crustal layer on the RAPIDS line could be modeled with higher velocities. Vogt et al. [1998] show no raypath distribution but the record sections indicate that lower crustal velocities are not controlled too well since refractions through the lower crust never become a first arrival. Record sections show strong reflections from the top of the lower crust. The velocity contrast across this boundary is only 0.1 km/s, which should not give such a strong reflection and, indeed, the amplitudes of the synthetic seismograms do not match this reflection very well. Replacing the 6.8-km/s lower crustal layer with higher velocities would increase the velocity contrast and would result in a better fit with the amplitudes. In addition, Edwards [2002] noticed an isostatic imbalance along the RAPIDS profile between Hatton Bank and the Hatton- Rockall Basin. This imbalance would be reduced by higher-density material in the lower crust. 5. Discussion 5.1. Crustal Composition [43] Results from the two refraction seismic lines distinguish two different crustal domains: the continent ocean Table 2. Number of Observations, n, RMS Misfit Between Calculated and Picked Travel Times, t rms, and Normalized c 2 for Individual Phases on Line B Phase n t rms,ms c 2 Direct wave P S (all sediments) P S P (all sediments) P B (all basalts/subbasalts) P B P (all basalts/subbasalts) P c P c1 P P c P c2 P P c P c3 P P c P m P P n All phases of 25

13 transition zone at the NE end of line B, and continental crust of slightly variable structure along the remainder of the lines with a major intrusion close to George Bligh Bank (Figures 10 and 11). In addition, a basalt sequence can be correlated across both lines while large segments of the profiles are characterized by a high-velocity zone in the Figure of 25

14 lower crust. Below, these features of the crust are discussed in some more detail Continental Crust [44] Continental crust with velocities between 5.5 and 6.8 km/s can be correlated from the Faroe Islands across the banks into the Hatton-Rockall Basin (Figures 10 and 11). This interpretation is in agreement with other geophysical data [Bott et al., 1974, 1976; Richardson et al., 1998] and isotopic evidence [Hald and Waagstein, 1983; Gariépy et al., 1983; Holm et al., 2001] that the Faroe Islands are underlain by continental crust. Lateral thickness variations of the continental crust are substantial, between 24 km at the banks and 8 km in the intermittent channels and basins. The variable amount of crustal extension along the profiles will be discussed later. [45] Crustal contamination of the lavas around the Rockall Trough suggests that the Lewisian-Islay basement terrane boundary crosses the trough and passes between George Bligh and Rockall Banks [Hitchen et al., 1997]. This may explain the slightly lower velocities south of George Bligh Bank (Figure 11), where upper and lower crustal velocities decrease by 0.1 km/s compared to the area to the north. However, this variation lies within the velocity uncertainty. Lewisian basement is composed predominantly of Archean granitoid gneisses, extensively reworked during the Early Proterozoic [Park, 1994]. Upper crustal velocities of 5.6 km/s north of George Bligh Bank are actually lower than what laboratory measurements indicate for granite gneiss at a pressure of 200 MPa (6.0 km/s) [Christensen, 1996]. In order to explain the lower velocities, Hughes et al. [1998] suggest that the original continental crust may have been intruded with basalts during the Tertiary igneous episode. Other explanations for the reduction of the velocity are conceivable, such as intense fracturing of the uppermost 2 to 4 km of the brittle crust by rift-related extension. Alternatively, the upper crust is perhaps not composed of gneiss but consists of a different rock type. Measurements on a foliated granite from the Lewisian complex show a pronounced seismic anisotropy at a pressure of 200 Ma, with velocities between 5.9 and 6.3 km/s [Hall and Simmons, 1979], which is higher than the observed velocity of 5.6 km/s. However, velocities within the widespread granitic intrusions offshore southern Greenland vary between 5.4 and 5.6 km/s [Chian and Louden, 1992; Chian et al., 1995] and upper crustal velocities of 5.4 km/s off west Greenland are probably related to granite [Funck et al., 2007]. [46] Lower crustal velocities of 6.6 km/s north of George Bligh Bank (Figures 10 and 11) are compatible with other data from Lewisian basement on the Hebrides shelf [Klingelhöfer et al., 2005; Keser Neish, 1993]. According to the compilation of Holbrook et al. [1992], these lower crustal velocities fall within the range of mafic granulite. Crustal velocities south of George Bligh Bank are only slightly lower than north of the bank, which is why the composition is probably similar Continent-Ocean Transition [47] Crustal velocities at the northeastern end of line B increase to km/s and 6.9 km/s in the upper and lower crust, respectively (Figure 11). While these velocities fall within the typical range of layers 2 and 3 in oceanic crust [White et al., 1992], the observation of intracrustal reflections in that zone (P c1 P) would argue against normal oceanic crust. Kimbell et al. [2005] discuss the difficulties of defining the exact location of the continent-ocean boundary (COB) at the volcanic Faroe-Hatton margin and suggest a nominal COB at the landward limit of the marginal magnetic high (Figure 18). [48] The problems of defining the COB are related to the difficulty in distinguishing between heavily intruded continental crust and oceanic crust. White et al. [2008] observed a transition zone at the Faroe-Hatton margin in which the velocities lie between continental crust and oceanic crust observed further seaward and they suggest that the intermediate velocities correspond to intruded continental crust. On the basis of reflection seismic data, White et al. [2008] argue that the high-velocity lower crust in this continent ocean transition zone corresponds to continental crust intruded by sills. As discussed above, gravity modeling suggests the presence of a high-density lower crust at the NE end of line B (Figure 16). However, the orientation of the line almost parallel to the margin makes it difficult to find reflections from sills that crosscut continental fabric, which is the argument used by White et al. [2008] to distinguish between underplating and intrusions. [49] With the NE end of line B located almost parallel to the landward limit of the marginal magnetic high (Figure 18), we interpret this initial velocity increase in the crust related to intrusions in the continent-ocean transition zone. Line AMP-E is perpendicular to the margin and covers the entire marginal magnetic high, where Klingelhöfer et al. [2005] interpret the crust to be of oceanic character with a sharp transition to continental crust at the landward limit of the high. However, the line does probably not extend far enough seaward to allow a distinction between clear oceanic crust and a continent-ocean transition zone Intrusion and High-Velocity Lower Crust [50] At the NW flank of George Bligh Bank, the velocity model for line B (Figure 11) shows velocities of 6.5 km/s at much shallower levels (6 km depth) than along the remainder of the profile. One explanation for this anomalous velocity profile could be that it represents oceanic crust. Velocities of 6.5 km/s lie close to layer 3 velocities in Figure 12. Ray coverage of the model of line A with every tenth ray from point-to-point ray tracing. The upper part of each of the four panels shows the observed data, indicated by vertical bars, with heights representing pick uncertainty; calculated data are indicated by solid lines. A reduction velocity of 6.5 km/s has been applied for the travel times. The lower part of the panels shows the raypaths. LVZ, low-velocity zone. (a) Reflections and refractions in the basalt layers, as well as basement reflections. (b) Refractions P c1 and P c2 in the upper and middle crustal layer, as well as the reflection P c1 P from the interface between the two layers. (c) Middle crustal reflections (P c2 P) and lower crustal refractions (P c3 ). (d) Reflections from the Moho discontinuity (P m P) and from the top of the high-velocity lower crustal layer (P c3 P), as well as mantle refractions P n. 14 of 25

15 normal oceanic crust ( km/s) [White et al., 1992] and there are observations that velocities can be even lower (e.g., 6.4 km/s off eastern Canada [Funck et al., 2004]). However, this interpretation is rejected for two reasons. First, this part of line B lies landward of the positive magnetic anomaly to the NW of the banks that Kimbell et Figure of 25

16 Figure 14. Comparison of P wave velocities at (a) the intersection of lines A and B, and at (b) the intersection of lines A and AMP-E [Klingelhöfer et al., 2005]. For all lines a 30-km-wide section adjacent to the intersection is shown. Numbers indicate velocities in km/s. Horizontal scale is distance along the respective velocity models. Abbreviations are LVZ, low-velocity zone; HVZ, high-velocity zone. al. [2005] interpret as COB. Second, a low-velocity zone (LVZ) is observed beneath the 6.5-km/s layer, which is not compatible with velocity profiles in oceanic crust [White et al., 1992]. [51] We therefore interpret the 8-km-thick layer with velocities between 6.5 and 6.7 km/s as an intrusion. In this scenario, the LVZ represents lower crustal rock that can be correlated underneath the intrusion. The base of the intrusion is mapped by numerous wide-angle reflections (Figure 13). According to the potential field data analysis of Edwards [2002], the Hatton-Rockall Basin is characterized by a number of circular intrusions in the NE and linear intrusions in the SW. Close to the intrusion on line B, Edwards [2002] interpret the twin circular positive gravity anomalies over George Bligh Bank (Figure 19) as intrusion. The intrusion on line B close to George Bligh Bank does not stand out as a circular anomaly on the potential field maps (Figures 18 and 19). However, the intrusions recognized and modeled by Edwards [2002] are also shallower than the intrusion on line B. In addition, the thick basalt layers on line B may mask the signature of the intrusion. [52] The age of the intrusions is not determined. Edwards [2002] noticed that the circular intrusions in the northern part of the Hatton-Rockall Basin are similar to known Late Cretaceous or Tertiary igneous intrusions to the north and west of Scotland described by Hitchen and Ritchie [1993]. Several Paleogene intrusions are also known from the conjugate margin in east Greenland [Holm and Prægel, 2006]. [53] Our velocity model for line B (Figure 11) shows a high-velocity lower crustal layer in the northernmost Hatton-Rockall Basin. Such lower crustal high-velocity zones (HVZ) are commonly referred to as underplated igneous material. At the continental margin beneath Edoras Bank and Hatton Bank, the HVZ is thought to be related to the Iceland mantle plume [Barton and White, 1997a; Vogt et al., 1998; Fowler et al., 1989]. Alternatively, the HVZ could represent continental crust intruded by sills [White et al., 2008]. In any case, the HVZ indicates addition of melt into the crust or below the base the crust. With that, the major intrusion observed on line B may originate from the same source as the HVZ. [54] While a 3- to 6-km-thick HVZ is observed beneath the Hatton-Rockall Basin and at the western side of Lousy Bank, no lower crustal high-velocity layer was detected on line A northeast of Lousy Bank (Figures 10 and 11) despite the proximity to the volcanic Faroe-Hatton continental margin. Absence of a high-velocity lower crustal layer underneath the adjacent NE Rockall Trough [Klingelhöfer et al., 2005] is consistent with a restriction of lower crustal high-velocity zones to the northwesternmost part of Lousy Bank, Bill Bailey Bank and Faroe Bank Basalts [55] The Faroe Islands are characterized by Paleogene flood basalts exceeding a thickness of 5 km [Waagstein, 1988]. Palaeomagnetic dating of the onshore basalts indicate a Selandian to Ypresian age [Riisager et al., 2002; Abrahamsen, 2006]. Offshore, two layers with velocities of km/s and km/s can be correlated along the Figure 13. Ray coverage of the model of line B with every tenth ray from point-to-point ray tracing. The upper part of each of the four panels shows the observed data, indicated by vertical bars, with heights representing pick uncertainty; calculated data are indicated by solid lines. A reduction velocity of 6.5 km/s has been applied for the travel times. The lower part of the panels shows the raypaths. HVZ, high-velocity zone. (a) Reflections and refractions in the basalt layers, as well as basement reflections. (b) Refractions P c1 and P c2 in the upper and middle crustal layer, as well as the reflection P c1 P from the interface between these two layers. (c) Middle crustal reflections (P c2 P) and lower crustal refractions (P c3 ). (d) Reflections from the Moho discontinuity (P m P) and from the top of the high-velocity lower crustal layer (P c3 P), as well as mantle refractions (P n ) and refractions in the HVZ (P c4 ). 16 of 25

17 Figure 15. Two-dimensional gravity modeling for line A. (top) Lithostatic pressure at a depth of 40 km. (middle) Observed (gray line) and calculated gravity (dashed line). (bottom) Densities in the model, given in kgm 3. Figure 16. Two-dimensional gravity modeling for line B. (top) Lithostatic pressure at a depth of 40 km (model A, dashed line; model B, solid line). (middle) Observed (gray line) and calculated gravity (model A, dashed line; model B, solid line). (bottom) Densities in the model, given in kgm 3. Model A is with a high-velocity lower crustal layer at the northeast end of the profile, and model B is without such a layer. 17 of 25

18 Figure 17. Comparison of P wave velocity profiles from line B with other data from the Faroe-Rockall Plateau. Velocities are specified in km/s. Solid lines indicate layer boundaries; dashed lines represent velocity contours. Profiles are taken from line 87 3 [Keser Neish, 1993], isimm line 5 [White et al., 2008], line B (Figure 11), RAPIDS line 13 [Vogt et al., 1998], and line BANS-1 [Klingelhöfer et al., 2005]. For location of these lines, see Figure 1. Abbreviations are LVZ, low-velocity zone; Sed., Sedimentary layer. entire length of line A (Figure 10) and have velocities that are compatible with those of basalts from the Faroe Islands [Kern and Richter, 1979; Boldreel, 2006]. The total thickness of these layers varies between 1.5 and 4 km. On top of these continuous layers, an additional but discontinuous layer is found with variable velocities between 3.7 and 4.9 km/s and a maximum thickness of 800 m. Also these portions of line A are interpreted to consist mainly of basaltic rock. Velocities at the low end may indicate intercalated nonvolcanic sedimentary rock. The coincident reflection seismic record (Figure 20) shows for example a series of fairly continuous and high-amplitude reflections within a 3.9-km/s layer between Bill Bailey Bank and Faroe Bank. The high amplitudes may be explained with a high impedance contrast between volcanic units and sedimentary rocks. [56] Below the two continuous basalt layers on line A (Figure 10), a low-velocity zone (LVZ) was identified, which could represent sedimentary rocks that predate the volcanism on the Faroe-Rockall Plateau, or which could partly consist of basalts. Boldreel [2006] noticed a decrease of seismic velocity at the transition from the subaerially extruded basalts to the subaqueous basalts at the Lopra well on the Faroe Islands (Figure 2). Velocities within the hyaloclastites of the well are as low as 4.5 km/s [Christie et al., 2006]. At basements highs (essentially the individual banks), possible sedimentary layers should be rather thin and we assume that at least part of the LVZ consists of hyaloclastites. [57] The channels between the banks form basement lows and thicker prevolcanic sediments sequences are a possibility. The coincident reflection seismic data have to be treated with some caution since it is not always easy to distinguish between primary energy and multiples. However, the reflectivity in the LVZ in the channel between Bill Bailey Bank and Faroe Bank (Figure 20) appears to be real and is characterized by numerous continuous reflectors with rather high amplitude. This reflection character is very similar to the overlying two basalt units with the exception that the vertical spacing of reflectors in the LVZ is larger. This can be interpreted in two ways, either the LVZ consists of thicker basalt flows compared to the overlying layers, or it could consist of sedimentary rocks and the reflectors represent some sills. [58] SW of George Bligh Bank, the otherwise good continuity of the two main basalt layers on line A and B ends. Locally, up to two LVZ can be found in the Hatton- Rockall Basin (Figure 11). The upper LVZ ( km/s) is located below a series of high-amplitude reflectors between 2.2 and 2.4 s two-way travel time (TWT) (Figure 21). These reflectors resemble those at DSDP site 116 (for location see Figure 1), which were identified as Oligocene chalk [Laughton et al., 1972; Smith et al., 2005]. The lower LVZ ( km/s) is located beneath some sills between 3.0 and 3.3 s TWT. Paleocene basalts were drilled at DSDP site 117 (for location see Figure 1) some 120 km to the south of line B at the western flank of Rockall Bank [Laughton et al., 1972] Tectonic Development of the Banks Area [59] One intriguing result of the refraction seismic experiment is the very variable thickness of the continental crust on line A (Figure 10). Within the banks, the continental crustal thickness varies between 20 and 24 km, while the crust is as thin as 8 km in the intermittent channels. There is 18 of 25

19 Figure 18. Magnetic anomaly map. Data are taken from Verhoef et al. [1996]. The magnetic anomalies are shaded by artificial illumination from the southeast. Bold solid lines show the location of the refraction seismic experiment. Positions of ocean bottom seismometers are indicated by filled white circles, and numbers indicate the station number. Solid lines show the bathymetry, and the contour interval is 500 m. Other seismic experiments discussed in the text are marked by dashed lines: lines AMP-D, AMP-E, and BANS-1 [Klingelhöfer et al., 2005]. Abbreviations are RB, Rosemary Bank; WTR, Wyville-Thomson Ridge. no evidence in the literature for major extension in a SW NE direction (approximately parallel to line A) other than some weak Paleocene extension [Lundin and Doré, 2005]. Hence, in order to explain the crustal thinning in the channels, other mechanisms need to be employed. Below, we propose a model, in which the crustal thinning in the channels is related to transform faults. [60] Figure 22 shows that the Faroe-Shetland Trough and the NE Rockall Trough widen from the Faroe Islands toward Lousy Bank. In particular it can be seen that the distance of individual banks from the Hebrides shelf (indicated by the 1000-m depth contour) increases toward the SW. By the division of NE Rockall Trough into segments, in which the width of the rift or trough is constant, the widening is associated with strike-slip movement across the segment boundaries. This concept is sketched in Figure 22, where shear zones with strike-slip movement are drawn parallel to a system of lineaments identified by Kimbell et al. [2005] who highlight the importance of the broadly NW trending lineaments on the development of the post- Caledonian basin architecture. Some of the lineaments are interpreted as pre-caledonian structures that were reactivated as transfer zones during phases of Mesozoic extension. Tate et al. [1999] demonstrate that the Wyville-Thomson lineament acted as a transfer zone and they also interpret a dextral offset across the lineament. [61] In the Faroe Bank Channel, the postulated shear zone could be a continuation of the Judd Lineament (Figure 22). The channel between Bill Bailey Bank and Faroe Bank appears to line up with the Wyville-Thomson Ridge. None of the suggested lineaments by Kimbell et al. [2005] coincides exactly with our shear zone between Lousy Bank and Bill Bailey Bank, but is close to the Ymir Ridge lineament. However, the gravity data (Figure 19) illustrates that the channel between these two banks can be lined up with the Sigmundur Seamount and the Darwin Igneous Complex. Hence, this trend may indicate an underlying lineament; a weak zone along which the igneous bodies were formed. The vertical gravity gradient map (Figure 19) also indicates NW SE trending lineaments in Rockall Trough to the north and south of George Bligh Bank, although the one to the north is rather weak. [62] Crustal thinning associated with continental transform faults can be substantial as evidenced by the Dead Sea transform that is connected with the northern Red Sea rift. About 105 km of displacement is documented along this transform [Garfunkel et al., 1980] and the sedimentary basins along the transform are up to 5 km thick in the Gulf of Aqaba [Ben-Avraham, 1985] and 14 km thick in the Dead Sea basin [Ginzburg and Ben-Avraham, 1997]. Mascle and Blarez [1987] show the evolution of such initial continental transform contacts into transform margins. With substantial crustal thinning over a 200-km-wide zone in NE Rockall Trough [Klingelhöfer et al., 2005], rifting there went into an advanced stage and a comparison with transform margins seems to be instructive. 19 of 25

20 Figure 19. (top) Free-air gravity anomaly map and (bottom) vertical gravity gradient map. Data are taken from Sandwell and Smith [1997]. The maps are shaded by artificial illumination from the northeast. Bold solid lines show the location of the refraction seismic experiment. Positions of ocean bottom seismometers are indicated by filled white circles, and numbers indicate the station number. Solid lines show the bathymetry, and the contour interval is 500 m. Other seismic experiments discussed in the text are marked by dashed lines: lines AMP-D, AMP-E, and BANS-1 [Klingelhöfer et al., 2005]. Abbreviations are as follows: DIC, Darwin Igneous Complex; RB, Rosemary Bank; SS, Sigmundur Seamount; and WTR, Wyville-Thomson Ridge. 20 of 25

21 Figure 20. Subset of migrated record section from reflection seismic line coincident to line A. Vertical scale is two-way travel time; horizontal scale is the distance along the velocity model of line A (Figure 10). Open triangles mark the location of OBS, and station numbers are indicated at the top. Filled triangles show the location of two velocity-depth profiles (solid lines on white background) extracted from the velocity model. Layer boundaries between the top of the basalt and the basement were converted to twoway travel time and are shown as solid lines. Abbreviations are LVZ, low-velocity zone; v, velocity. [63] In the eastern equatorial Atlantic, a transform continental margin developed off Ghana along the Romanche fracture zone. The continental crust at that margin thins over a 17-km-wide zone from 20 km to just 5 km before the transition to oceanic crust is reached [Edwards et al., 1997]. At the SW Newfoundland transform margin, Todd et al. [1988] observe a thinning of the continental crust from 20 km to 8 km across a 25-km-wide zone. The channels Figure 21. Subset of migrated record section from reflection seismic line coincident to line B. Vertical scale is two-way travel time; horizontal scale is the distance along the velocity model of line B (Figure 11). Open triangles mark the location of OBS, and station numbers are indicated at the top. A filled triangle shows the location of a velocity-depth profile (solid line on white background) extracted from the velocity model. The depth to basement from the refraction model was converted to two-way travel time and is shown as solid line. Abbreviations are LVZ, low-velocity zone; v, velocity. 21 of 25

22 Figure 22. Tectonic model to explain some of the crustal thinning in the channels between the banks southwest of the Faroe Islands. For details on the background elevation model, see Figure 1. Contour interval of the bathymetry is 500 m (solid lines). Thin white lines indicate lineaments and transfer zones [after Kimbell et al., 2005]. Bold white lines show the suggested segmentation of the shelf to either side of northeast Rockall Trough and the Faroe-Shetland Trough; the orientation of segments is kept parallel to the contours at the eastern side of the trough. Dashed white lines mark shear zones, with arrows indicating the sense of motion. Bold dashed lines show the location of the refraction seismic experiment, and dashed lines mark the position of lines AMP-D and E [Klingelhöfer et al., 2005]. The dotted blue line shows the location of the continent-ocean boundary [after Kimbell et al., 2005]. Abbreviations are as follows: ADLC, Anton Dohrn Lineament Complex; ADS, Anton Dohrn Seamount; BBB, Bill Bailey Bank; CL, Clair Lineament; FB, Faroe Bank; FBC, Faroe Bank Channel; FST, Faroe-Shetland Trough; GBB, George Bligh Bank; JL, Judd Lineament; LB, Lousy Bank; RB, Rosemary Bank; SHL, South Hatton Lineament; WTLC, Wyville-Thomson Lineament Complex; WTR, Wyville-Thomson Ridge; and YR, Ymir Ridge. between the banks on line A (Figure 10) have a full width between 45 and 90 km, which compares well with the reported half widths for the transform margins. The minimum thickness of the continental crust at the transform margin (5 to 8 km) is also comparable to the values obtained for the channels between the banks (8 to 11 km). In summary, the geometry of the channels is very similar to what is observed at the continental portion of transform margins. [64] The observed thickening of the basalts in the channels suggests that the Paleogene basalt flows filled in preexisting bathymetric lows. Hence, we assume that the postulated shear zones in the channels are associated with the NW-striking lineaments in NE Rockall Trough, which were reactivated during Mesozoic rifting [Kimbell et al., 2005]. Similar to line A, line AMP-D (Figure 2) in the NE Rockall Trough displays a variable crustal thickness [Klingelhöfer et al., 2005], which supports our interpretation that the lineaments may indeed have some control on the crustal thickness on the banks and in the interjacent troughs. [65] In this context it is interesting to note that transform faults are often bounded by asymmetric ridges [Basile and Allemand, 2002], which could suggest that the Wyville- Thompson Ridge is related to the postulated transform fault that offsets Bill Bailey Bank from Faroe Bank. However, Tate et al. [1999] interpret the ridge as a late Eocene to Oligocene-Miocene fault-propagation fold that developed from inversion above a crustal-scale detachment during N S compression. 6. Conclusions [66] The results from the refraction seismic experiment show that continental crust can be correlated from the Faroe Island, across the banks to the SW of the Faroes and into the Hatton-Rockall Basin. The crust is thickest at the banks (up 22 of 25