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

Abstract

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 4.9-5.2 and 5.3-5.6 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.

Full text

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 2008.
[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 4.9–5.2 and 5.3–5.6 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, B12405, doi:10.1029/2008JB005675.
1. 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 accompa-
nied 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 sub-
basalt 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
al. [1988] concluded that the crust within Rockall Trough is
continental in character, although Smythe [1989] had some
doubt th at this conclu sion was sup ported 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 conti-
nental 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 Klingelho¨fer
et al. [2005] found 24-km-thick continental crust. Between
Lousy Bank and George Bligh Bank, the crust thins to 14 km
[Klingelho¨fer 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
JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 113, B12405, doi:10.1029/2008JB005675, 2008
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1
Geological Surve y of Denmark and Greenland, Copenhagen,
Denmark.
2
Faroese Earth and Energy Directorate, To´rshavn, Faroe Islands.
Copyright 2008 by the American Geophysical Union.
0148-0227/08/2008JB005675$09.00
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seismic experiment was carried out in 2004. 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 multichan-
nel 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 volca-
nism occurred in the Faroe-Rockall area, which is attributed
to the arrival of the Iceland plume in that region [White,
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 [Klingelho¨fer 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.
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1992]. On the Faroe Islands, the basalts are divided into
lower, middle and upper formation. The lower formation
was e xtruded 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 subaer-
ial 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 contine nt-ocean transition
zone, interpreted as magmatic underplating, for example
at Lousy Bank [Klingelho¨fer 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 two-
ship experiment carried out in 2004. The Danish ship
Esvagt Connector was used for the deployment and recov-
ery of the ocean bottom seismometers (OBS), while the
Norwegian commercial seismic vessel Polar Princess fired
the air gun array and collected coincident multichannel-
seismic data along the refraction seismic lines.
[
10] A total of 40 OBS were available for the experiment.
The Institut Franc¸ais 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.
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 [Klingelho¨fer 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.
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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 i s a 39 5-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 w ere
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 calcu-
lated. 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 dis-
turbed 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 pro-
duced from the data.
3.2. 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 sedi-
mentary column and partly within the basalt sequences were
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.
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primarily defined by t he coincident MCS data. Late r,
velocities within layers were optimized by using the inver-
sion algorithm in RAYINVR.
3.3. 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 seismo-
grams. Phase names in the records reflect the later interpre-
tation 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 wide-
spread occurrence of basaltic sills and lava flows on the
Faroe-Rockall Plateau . Reflection seismic images of sub-
basalt structure are often degraded owing to poor penetra-
tion, 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 v olcanic 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 signal-
to-noise ratio decreases significantly for d eeper 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
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.
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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. Fi rst, 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.
4.1. Velocity Models
4.1.1. 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
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).
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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 con-
strained along portions of the line.
4.1.2. 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 cons ist of lava flows or vol canic debris, pos sibly
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).
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).
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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 [Klingelho¨fer et
al., 2005].
[
28] The remainder of the crust on line B has a three-
layered 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 under-
plated 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). Veloc-
ities of 8.0 km/s are determined for the mantle.
4.2. 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
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).
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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 withi n the sedimentary units and
basalts mostly result from the lack of good reverse ray
coverage in these shallow layers. However, additional con-
straints on the layer geometry from the coincident MCS data
compensate for this shortcoming. Velocities within the high-
velocity 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 be-
neath the Faroe Islands from this data set. The base of the
intrusion (line B, 130–210 km) is mapped in detail by
reflections, while the midcrustal boundary along the re-
mainder 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 inde-
pendently 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 well-
determined velocities within the basalts may be caused by
anisotropy. Layer boundaries match within ±1 km.
4.3. Gravity Modeling
[
34] To verify the consistency of the velocity models
(Figures 10 and 11) with available gravity data, two-
dimensional gravity modeling was performed using the
algorithm of Talwanietal.[1959]. Grav ity 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. How-
ever, at the NE end of line B, the misfit can be reduced
substantially when a high-density lower crustal layer is
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).
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introduced (Figure 16) similar to the western flank of
Lousy Bank on lines A and AMP-E [Klingelho¨fer 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.
4.4. Comparison With Other Studies
[
36]LineAMP-E[Klingelho¨fer et al., 2005] crosses
Lousy Bank exactly at the intersection of lines A and B
(Figure 2). This offers the opportunity to check how
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 [Klingelho¨fer 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, low-
velocity zone.
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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,
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
[Klingelho¨fer et al., 2005], and line BANS-1 [Klingelho¨fer 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.
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line A divides the sequence into two layers with velocities of
5.6–5.8 km/s and 6.1–6.2 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 attrib-
uted to some complications with the overlying basalt se-
quence. 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 tw o ba salt 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 recog-
nized 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 baseme nt 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
[Klingelho¨fer 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 i ntrusions, 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 Ban k 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 distin-
guish two different crustal domains: the continent ocean
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 nt
rms
,ms c
2
Direct wave 837 0.031 0.413
P
S
(all sediments) 1067 0.042 0.353
P
S
P (all sediments) 111 0.054 0.411
P
B
(all basalts/subbasalts) 8606 0.050 0.362
P
B
P (all basalts/subbasalts) 2958 0.089 0.844
P
c1
2347 0.068 0.553
P
c1
P 1476 0.112 1.243
P
c2
1514 0.090 0.836
P
c2
P 398 0.117 1.187
P
c3
5353 0.105 1.045
P
c3
P 871 0.076 0.894
P
c4
32 0.149 2.305
P
m
P 3509 0.140 1.187
P
n
1676 0.127 0.682
All phases 30755 0.091 0.747
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 nt
rms
,ms c
2
Direct wave 2693 0.037 0.421
P
S
(all sediments) 1583 0.038 0.292
P
S
P (all sediments) 367 0.058 0.350
P
B
(all basalts/subbasalts) 5850 0.046 0.312
P
B
P (all basalts/subbasalts) 3349 0.093 0.863
P
c1
1712 0.091 0.822
P
c1
P 1041 0.110 1.350
P
c2
3729 0.085 0.614
P
c2
P 1644 0.106 1.040
P
c3
2491 0.103 0.751
P
c3
P 3049 0.140 1.448
P
c4
206 0.101 0.706
P
m
P 5512 0.168 1.909
P
n
2246 0.191 1.267
All phases 35472 0.112 0.931
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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 12
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lower crust. Below, these features of the crust are discussed
in some more detail.
5.1.1. 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; Garie´py 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 Rock-
all Trough suggests that the Lewisian-Islay basement ter-
rane 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 Green-
land vary between 5.4 and 5.6 km/s [Chian and Louden ,
1992; Chian et al., 1995] and upper crustal velocities of
5.4 k m/s off west Greenlan d 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
[Klingelho¨fer 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.
5.1.2. Continent-Ocean Transition
[
47] Crustal velocities at the northeastern end of line B
increase to 5.8 –6.2 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 bound-
ary (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 conti-
nental 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 inter-
mediate 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 t ransition 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,
whic h 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 transi-
tion zone. Line AMP-E is perpendicular to the margin and
covers the entire marginal magnetic high, where
Klingelho¨fer 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 enou gh seaward to allow a distinction
between clear oceanic crust and a continent-ocean transition
zone.
5.1.3. 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 remain-
der 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
.
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normal oceanic crust (6.6– 7.6 km/s) [Whit e 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 13
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B12405

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 t he 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 intru-
sion is mapped by numerous wide-angle reflections
(Figure 13). According to the potential field data analysis
of Edwards [2002], the Hatton-Rockall Basin is character-
ized 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 recog-
nized 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 in trusions 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 Hat-
ton-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 [Klingelho¨fer
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.
5.1.4. 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 indi-
cate a Selandian to Ypresian age [Riisager et al., 2002;
Abrahamsen, 2006]. Offshore, two layers with velocities of
4.9–5.2 km/s and 5.3 –5.6 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
).
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 [Klingelho¨fer 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. Abbrevia-
tions are LVZ, low-velocity zone; HVZ, high-velocity zone.
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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.
B12405 FUNCK ET AL.: SEISMIC TRANSECT AT FAROE-ROCKALL PLATEAU
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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 thick-
ness 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 co nsist 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 w ithin 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 possibil-
ity. 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 reflec-
tivity 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 otherw ise 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 (3.1–3.8 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 (4.1–5.0 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].
5.2. Tectonic Development of the Banks Area
[
59] One intriguing result of the refraction seismic exper-
iment 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
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
[Klingelho¨fer et al., 2005]. For location of these lines, see
Figure 1. Abbreviations are LVZ, low-velocity zone; Sed.,
Sedimentary layer.
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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 Dore´, 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 Islan ds
toward Lousy Bank. In particular it can be seen that the
distance of individual banks from the Hebrides shelf (indi-
cated 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 linea-
ment 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 suggeste d lineaments by Kimbelletal.[200 5]
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 indic ates NW–SE trending lineame nts in Rockall
Trough to the north and south of George Bli gh Bank,
although the one to the north is rather weak.
[
62] Crustal thinning associated with continental trans-
form 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 [Klingelho¨fer et al., 2005], rifting there went into an
advanced stage and a comparison with transform margins
seems to be instructive.
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 [Klingelho¨fer et al., 2005]. Abbreviations are RB, Rosemary Bank;
WTR, Wyville-Thomson Ridge.
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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 [Klingelho¨fer et al., 2005].
Abbreviations are as follows: DIC, Darwin Igneous Complex; RB, Rosemary Bank; SS, Sigmundur
Seamount; and WTR, Wyville-Thomson Ridge.
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[63] In the eastern equatorial Atlantic, a transform conti-
nental 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 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 two-
way travel time and are shown as solid lines. Abbreviations are LVZ, low-velocity zone; v, velocity.
Figure 21. Subset of migrated record section from reflection seismic line coincide nt 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.
B12405 FUNCK ET AL.: SEISMIC TRANSECT AT FAROE-ROCKALL PLATEAU
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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 mini-
mum thic kness 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 [Klingelho¨fer 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 fr om 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
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 [Klingelho¨fer 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.
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to 25 km) but thins to as little as 8 km in the channels
between the banks (Figures 10 and 11), not including the
overlying basalt layers. These channels are subparallel to
known transfer zones and lineaments in the region. Dextral
shear was interpreted along the Wyville-Thomson lineament
[Tate et al., 1999] and is consistent with the southwestward
widening of the NE Rockall Trough. We therefore suggest
that the channels between the banks were formed by
continental tran sform f aults, similar to observed crustal
thinning along portions of the Dead Sea transform [Ben-
Avraham, 1985] or across continental transform margins
[Todd et al., 1988; Edwards et al., 1997].
[
67] Two main basalt layers with velocities o f 4.9 to
5.6 km/s can be correlated from the Faroe Islands up to
George Bligh Bank; the total thickness varies between 1
and 4 km with the maximum occurring in the channel between
Lousy Bank and Bill Bailey Bank. The thickening of the
basalts in the channels indicate that the channels were formed
prior to the Paleogene volcanism. Other volcanic layers are
found on top of the two main basalt layers and a low-velocity
zone beneath the basalts may contain additional material of
volcanic rock or, alternatively, may represent subbasalt sed-
imentary rock. Locally, the total thickness of the basalts may
be as high as 6 km. The two main basalt layers cannot be
correlated south of George Bligh Bank, where the seismic line
becomes more distal to the Faroe-Hatton volcanic margin.
Boldreel and Andersen [1994] suggest that the basalts on the
southern end of line B are subaqueous.
[
68] High-velocity lower crustal layers are observed at the
NW side of Lousy Bank and in a 4- to 5-km-thick layer that
extends from George Bligh Bank into the northern Hatton-
Rockall Basin (Figures 10 and 11). This indicates a sub-
stantial addition of melt to the crust, either as magmatic
underplating or by intrusion into the lower crust as sills. An
80-km-wide and up to 9-km-thick intrusion NW of George
Bligh Bank is possibly associated with the high-velocity
lower crust. Tomographic results 70 km to the south of line
B indicate lower crustal velocities of 7.3 km/s [White et al.,
2008] but interpretations of other deep-seismic data in
southern Hatton-Rockall Basin [Keser Neish, 1993; Vogt
et al., 1998] lack an underplated layer. However, a review of
these data indicate that a high-velocity lower crust cannot be
excluded, which may suggest that the high-velocity zone is
not only restricted to the northernmost part of the Hatton-
Rockall Basin. Widespread intrusions in Hatton-Rockall
Basin [Edwards, 2002] may indeed be more compatible
with such an interpretation. If part of the lower crust in
southern Hatton-Rockall Basin has to be reinterpreted as
magmatic underplating or sill intrusion, then the thickness
of the continental crust in the basin has to be adjusted
downward, which would result in larger stretching factors of
the crust than previously thought.
[
69] Acknowledgments. We thank the captains, crew and scientists
on board Polar Princess and Esvagt Connector. Comments from Bob
White and an anonymous reviewer helped improve the manuscript. The
project was part of the Danish Continental Shelf Project. Funding was
provided by the Faroese Government and the Danish Ministry of Science,
Technology, and Innovation.
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