Crustal structure of the Goban Spur rifted continental margin, Ne Atlantic

Geophys. J . Int. (1993) 119, 1-19

Crustal structure of the Goban Spur rifted continental margin, NE Atlantic

Susan J. Horsefield,’.* Robert B. Whitmarsh: Robert S. White’ and Jean-Claude Sibuet3 ’ Bullard Laboratories, Department of Earth Sciences, University of Cambridge, Madingley Rise, Madingley Road, Cambridge CB3 OEZ, UK ’ lnstirute of Oceanographic Sciences Deacon Laboratory, Wormlev, Godalrning, Surrey GU8 5UB, UK IFREMER Centre de Brest, B. P. 70, 29263 Plouzant, France

Accepted 1994 February 15. Received 1993 August 13; in original form 1992 June 22

S U M M A R Y We present results that constrain the crustal structure across a typical ‘non-volcanic’ rifted continental margin, the Goban Spur in the NE Atlantic. Traveltime and amplitude modelling of wide-angle ocean-bottom seismograph profiles shows that the continental crust thins in a seaward direction to about 7 km before it breaks to allow a new oceanic spereading centre to develop. The crustal-velocity structure, the stretching factors and the density distribution inferred from gravity anomalies are consistent with bulk pure shear stretching for the development of the rifted continental margin.

The seismic-velocity structure of the thinned continental crust suggests that thinning was accompanied by only limited igneous intrusion and extrusion in the upper crust. The adjacent oceanic crust is abnormally thin (5.4km). This is explained most readily by accretion at a very slow spreading centre (approx 11 mm yr-’ full rate), although additional cooling by conductive heat loss from the parent mantle as it rose slowly beneath the stretched and thinned continental lithosphere, together with minor igneous intrusion into, and surface flows onto, the adjacent continental crust may also have reduced the oceanic igneous thickness immediately adjacent to the continent-ocean transition.

Key words: continental margin, Goban Spur, ORS, ocean-continent boundary, seismic structure.

1 INTRODUCTION

The Goban Spur continental margin is located south-west of Britain (Fig. l), and was formed by break up between North America and Europe. It was created by crustal thinning mostly during the last rifting phase during the early Cretaceous to middle Albian (de Graciansky et al. 1985; Masson, Montadert & Scrutton 1985). There are only a few well-developed tilted fault blocks (Sibuet et al. 1985), and the crests of these blocks are largely eroded, preventing calculation of the extensional factor from the block geometry.

The Goban Spur continental margin is about 200 km long and is characterized by a 2-D structure. The Western Approaches Margin (WAM) deep seismic-reflection profile (Fig. l), recorded jointly in 1985 by BIRPS and ECORS

*Now at: Esso Expro UK Ltd, Mailpoint 15, Esso House, Ermyn Way, Leatherhead, Surrey KT22 8UY. UK.

(Peddy et al. 1989) and shot perpendicular to this margin, is thus repesentative of the whole margin. It was shot to image the structure of the thinned continental crust in order to test theories concerning the formation of this margin. Beyond the shelf edge it is only possible to image the layered lower crust clearly above the first seabed multiple. However, a careful examination of potential primary reflectors located below the first multiple has been made by Pinet et al. (1991), who propose that the layered lower crust extended as far landward as the shelf. These lower-crustal reflections, less continuous than below the continental shelf, have been interpreted as post-extension remnants of the lower-crustal reflectors observed on the continental shelf. As such, they would thus predate the extensional episode that led to the opening of the North Atlantic at this latitude.

Variations of the thickness of the lower and upper crust suggest that the continental crust on this margin has been stretched by pure shear. This is also confirmed by heat-flow data acquired along the WAM profile that follow the

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2 S. J . Horsefield et al.

16" 14" 12" 50"

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Figure 1. Seabeam bathymetric map of the Goban Spur and vicinity from Sibuet et al. (1985) showing location of seismic refraction profiles and position of instruments (open circles with dots). Inset shows general location on continental margin south-west of the UK and Ireland. Locations of DSDP holes 548-551 indicated by solid circles. Depths contoured in 200 m intervals.

predicted trend for pure shear rifting, provided the upper-crustal radiogenic heating is low (Louden, Sibuet & Foucher 1991). However, in this area of great extension, White (in White & Peddy 1990) suggests that the adiabatic decompression of upwelling mantle is likely to have generated a significant amount of melt. With that hypothesis, the lower-crustal layering could be caused, at least in part, by igneous rocks intruded during rifting.

The purpose of the seismic-refraction survey carried out on the Goban Spur margin in 1987 and reported here was to define the crustal structure and the extent of igneous rocks in order to complement the information provided by deep seismic-reflection data and to constrain better the mode of formation of the rifted continental margin.

2 THE EXPERIMENT

The geophysical data set used here was collected over the Goban Spur in March-April 1987 aboard the R/V Le Suroit as part of the Rkframarge project. It includes 500km of four-channel normal incidence seismic-reflection profiles, magnetic and bathymetric data, and three airgun wide-angle seismic-refraction lines. In this paper we describe the results of traveltime and amplitude modelling of the seismic data,

together with 2-D modelling of the observed gravity and magnetic fields.

The configuration of the three seismic-refraction lines and the instrument locations are shown in Fig. 1. Line 3 was positioned to coincide with the WAM deep seismic- reflection profile across the margin.

Several digital ocean-bottom seismographs (DOBS) from the Institute of Oceanographic Sciences Deacon Laboratory (Kirk, Langford & Whitmarsh 1982; Peal & Kirk 1983) with hydrophone and internal three-component geophone pack- ages were deployed on each line. The source was an array of eight 16 1 (1000 cubic inches) airguns fired every 120 s, giving a shot spacing of approximately 300m. The data were recorded at 75 samples per second for a time of up to 60s. Shot-receiver ranges were determined using the technique outlined in the Appendix of Whitmarsh, Avedik & Saunders (1986). Ranges were adjusted until the recorded traveltimes of the waterwave arrivals matched the calculated (ray- traced) time-distance curve.

Two sets of four-channel reflection profiles were recorded along the two margin-parallel refraction lines (Lines 1 and 2) to provide control on the sediment structure and depth to acoustic basement: one was recorded during wide-angle shooting using the large airgun source; the other used a

Goban Spur continental margin 3

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Figure 2. Line drawing of the section of the WAM profile along Line 3 (model ranges at bottom). Faults delineating tilted blocks are shown by heavy lines and the DOBS instrument positions are marked by triangles. Location of DSDP holes indicated by sea-bed arrows.

higher-frequency watergun source. The data from the different channels were summed and filtered in real time, then displayed on an electrostatic plotter, so no further processing was possible. The WAM line (Fig. 2) was used to provide control on the sediment structure and depth to basement along Line 3.

3 MODELLING RESULTS

The DOBS data were traveltime and amplitude modelled at Cambridge using a program based on Maslov asymptotic ray theory developed by Chapman & Drummond (1982), restructured by D. G. Lyness and modified to run on SUN Sparcstations. Initial models were set up by using: (a) the reflection profiles, sonobuoy refraction (Sibuet 1987) and Deep Sea Drilling Project (DSDP) results (de Graciansky et al. 1985) for the structure to acoustic basement and sediment velocities, and, (b) 1-D slope-intercept traveltime solutions for an estimate of the deeper structure. Lines 1 and 2 were modelled first to provide control at the intersections with Line 3; the intersections are marked on the ray-traced models in Figs 3-8.

The modelling strategy follows that outlined by Peirce & Barton (1991) except that shallow sedimentary arrivals (4.) were not modelled. Apart from the fact that most P, phases are not visible on the record sections, the determination of precise sediment velocities is not as important in modelling the deep crustal structure as ensuring that the time delays resulting from travel through the sediments are accurate. The revised strategy is thus:

(1) the traveltimes and amplitudes of deep sediment layers (P,) and near-offset upper crustal arrivals Pg are modelled around each instrument and the velocity structure is interpolated between DOBSs.

(2) The traveltimes of lower crustal and upper mantle phases (PmP, Pg and P n ) are modelled by altering the velocity structure and geometry down to the Moho until a satisfactory fit is obtained.

(3) After obtaining a reasonable traveltime fit for arrivals at all the instruments, the velocity gradients and the velocity contrasts are altered (keeping the average velocities, and hence traveltimes, unchanged), until a good amplitude fit is obtained with the main amplitude features of the data. We

concentrate on modelling the first-arriving energy, the mantle reflections and the amplitudes in the vicinity of the crust-mantle triplication. In a few places discontinuous later-arriving energy is visible (e.g. 140-160 km, Fig. 6). This energy may be caused by short reflectors in the crust, or by interbed multiples, and in view of the poor amplitude control we have not attempted to reproduce such minor features in the synthetics.

During the ray tracing care was taken to specify increments of ray slowness giving good ray coverage across the entire model, ensuring that rays reach the surface of the model no further than 2km apart (the distance between synthetic traces). This ensures that the algorithm calculating the synthetic seismograms has sufficient information to compute the amplitude variations with range.

Modelling results are discussed below together with the main features of the record sections. For each instrument the synthetic seismograms are shown together with the observed data and the ray-traced model. The data and synthetics are plotted with a reduction velocity of 8 km s-' and on both data and synthetics the same linear range scaling of trace amplitudes has been applied to enhance the small-amplitude arrivals at large ranges. The data have been bandpass filtered (4.5-11.5 Hz) and the positive excursions are displayed as a variable area plot with noisy traces removed. The bandpass filter was applied to attenuate high- and low-frequency noise. Outside this frequency band, the airgun source produced little energy. For clarity, the ray-traced models are displayed with a much reduced number of rays, the ray-parameter increment between rays varying through the model, but being the same for any one phase.

The data, synthetics and model are shown in Figs 3-8. The quality of the data is good and it is relatively easy to recognize first and second arrivals, even out to ranges of 150 km, because the trace-to-trace coherence is high. The source waveform has a predominant frequency of 8 Hz and is relatively long (-0.5 s ) as a result of using only one size of airgun in the source array. This means that there is a significant amount of interference between phases arriving at similar times (see, for example, the triplication at 100 km in the Fig. 3 model). However, this waveform is also relatively easy to recognize even in the presence of noise,


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Figure 4. Ray tracing of final model of Line 2 for DOBS 2015. Plotting and display details as for Fig. 3. (a) Observed data. (b) Synthetic seismograms including DOBS end water-bottom multiple for the Pg phase. (c) Ray-traced final model showing ray paths for water-bottom multiple. The intersection with Line 3 (at 94 km model range) is marked.


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8 S. J. Horsefield et al.

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and the amplitudes are higher than would have been the case if other, smaller airguns had replaced some of the 16 litre guns in order to tune the array. The models presented here reproduce the traveltimes generally to within 50 ms of those observed, and in the worst case, to better than 100 ms.

3.1 Line 1 (Fig. 3) The structure of this line is less well constrained than planned because only one of the two DOBSs recorded data (Fig. 1).

DOBS 2014

Amplitude modelling allows us to determine the velocity gradient in the upper crust in the vicinity of the DOBS, and this same velocity structure is then assumed to apply beneath the basement across the whole model (Fig. 3). The triplication of Pg, PmP and Pn arrivals between 70 km and 125 km in the model then constrains the velocity structure down to the Moho. The Moho here lies 14-16 km below sea level, the constrained part being indicated by a heavier line on the ray diagram. Traveltimes and amplitudes are well matched by the model. Note especially the much lower amplitude of the Pn phase compared to the Pg phase, and its particularly low amplitude between 115 km and 135 km. The amplitude variation of PmP and Pn provides a constraint on the velocity structure of the upper mantle.

The final model shows -2 km sediments overyling 10 km of thinned continental crust. The structure of the thinned crust is similar to that found on the Biscay margin (Ginzburg et al. 1985; Whitmarsh et al. 1986). Although the line is unreversed, the arrivals from this DOBS constrain the depth and shape of the Moho over a 90 km section on the profile. This zone includes the intersection with Line 3. The Moho has a small degree of topography and the crust thins slightly towards the south.

3.2 Line 2 (Figs 4 and 5)

This profile was shot landward of Line 1 and was recorded by DOBSs 2015 and 4005 (Fig. 1). The water-bottom multiples (bottom-surface reflections at the shot end, or surface reflection at the receiver end) arrive only a few seconds after the direct phases. For these two instruments the water-bottom multiple has also been modelled. Only those multiples from the DOBS end of the profiles are shown on Figs 4 and 5.

DOBS 2105

The Pg-PmP-Pn triplication between 60 km and 155 km is particularly well matched by the model synthetics (Fig. 4). Features of the amplitude variations that allow us to constrain the velocity structure include:

(1) the two regions of low-amplitude Pg arrivals at 50-60km and 90-105km (modelled by changes in the velocity gradient of the crust).

(2) High amplitudes of PmP phase between 95 and 160 km (constraining the velocity contrast across the Moho).

(3) Very low amplitude of Pn phase compared to the

other phases (constraining the velocity gradient of the upper mantle).

DOBS 4005

This record section has a lower signal-to-noise ratio than that from DOBS 2015 and the low-amplitude arrivals are harder to pick in places (Fig. 5) . Nevertheless, most of the arrivals are fitted to within 50 ms traveltime. Particular amplitude features modelled to constrain the velocity structure are:

(1) low-amplitude Pg arrivals at 100-125 km and 150-160 km (constraining two changes in the velocity gradient of the crust).

(2) High-amplitude Pg arrivals between the termination point at 45 km and 80 km (constraining the lower crustal velocities and crustal thickness).

(3) Low-amplitude of Pn (constraining the velocity contrast across the Moho).

The final model has a relatively simple structure, with a few kilometres of sediment overyling a 20km thick crust with typically continental velocity structure. There is no detectable topography on the Moho.

The thickness of crust determined by our modelling relates to a plane perpendicular to the Moho since the PmP phase is the main arrival we model. This plane is not vertical because, as will be shown later, the margin is still thinning at this location. Thus lower-crust Pg arrivals are seen on the record sections which have travelled longer, deeper paths out of the plane of the section than those modelled (see Fig. 5a at 30-40 km, 7.5 s).

3.3 Line 3 (Figs 2, 6, 7 and 8)

Line 3 intersects Lines 1 and 2 at 125 km and 159km respectively (Fig. 1). The results of modelling Lines 1 and 2 were incorporated without alteration into the model for Line 3 at the intersection points.

DOBS 201 7 east

The data exhibit good signal-to-noise ratios, except noticeably between 100 and 120km model range where arrivals are obscured by noise (Fig. 6). This DOBS lies in only 11OOm of water and so the water-column multiples follow shortly after the direct arrivals. This results in primary second arrivals being obscured, especially the near-offset PrnP arrivals (170-200 km), making it impossible to determine crustal thickness beyond 190 km model range. Particular amplitude features modelled to constrain the structure are:

(1) two zones of lower-amplitude Pg arrivals at 120-145 km and 175-190 km (two changes in velocity gradient through the crust).

(2) High-amplitude combined Pg and PrnP arrivals at 105-130 km, and strong PrnP at 138-152 km and close to 160 km (constrains lower crustal velocities, velocity contrast across the Moho and topography of the Moho).

(3) Low-amplitude Pn arrivals especially between 105 and


120 km (constrains the velocity structure of the upper mantle).

The depth and shape of the Moho (130-190km) were determined mainly by the PrnP phase and the deeper Pg diving rays. The portion of the Moho constrained by arrivals is indicated in Figs 6 and 9 by a bold line.

DOBS 2016 west

This DOBS was located close to DSDP site 551 (de Graciansky et al. 1985), which drilled basalt, and is located above suspected extruded basalts (Scrutton 1985). Features of particular importance whose amplitudes are modelled include (Fig. 7):

(1) high-amplitude Pg arrivals from 90 km to 100 km. (2) High-amplitude PmP arrivals from 62km to 72km

(3) Low-amplitude of Pn with respect to other phases, and from 77 km to 85 km.

especially between 41 km and 55 km.

Crust with a seismic-velocity structure typical of a very slow spreading oceanic ridge, formed at less than 15 mm yr-' full spreading rate (White, McKenzie & O'Nions 1992) is present only to the west of 95 km. The structure appears typical of stretched continental crust to the east of 110 km.

Shots to the east of DOBS 2016 constrain tightly the velocity, thickness and extent of the extruded basalts. The basalts are located between model ranges 108-120 km, have a thickness of 1.5 km, lie above thinned continental crust and exhibit seismic velocities of 5.55-5.65 kms-'.

DOBS 5001

This DOBS was located on oceanic crust just west of the escarpment that Masson et al. (1985) define as the ocean-continent boundary (see Fig. 2). The quality of the data recorded was not as good as for the other instruments as is evident from Fig. 8. However, it was possible to use this data to constrain the velocity structure of the oceanic crust between 40-95km, the structure of the basalt knoll and extrusives, the velocities of the continental crust, and the location of the Moho between 105 km and 150 km.

3.4 The Final Model

The best-fitting model is shown in Fig. 9 together with velocity-depth profiles at locations of particular interest. The main features of the final model are:

(1) the sedimentary sequence is divided into three main layers, divided laterally into two blocks on either side of the ocean-continent transition. The sediment velocities overly- ing the oceanic crust range from 1.85-3.2 km sC1 with a localized fourth layer exhibiting velocities of 3.6-3.8 km s-' at model ranges of 90-100 km. East of the ocean-continent transition the velocities range from 2.0-3.4 km s-' for post-rift sediments to 3.6-3.8 km s-' for the deepest layer, which probably represents syn-rift deposits. The maximum sediment thickness of 3 km is attained in the half-graben at 150 km model range.

(2) The oceanic crust is modelled with a 2.5 km thick layer 2 having velocities of 4.9-5.8 km s-*. There is a 0.5 km thick transition to layer 3, which is 3.5 km thick and exhibits

-- 1 5 9 1 5 9

kmls kmls 1 5 9

kmls

Model Range km

Figure 9. The final velocity-depth model for Line 3. Intersections with Lines 1 and 2 are marked. Velocity-depth profiles have been plotted across the oceanic crust (at 75 km), extruded basalts and thinned continental crust (at 115 km), and slightly thinned continental crust (at 180 km). Filled triangles show locations of velocity profiles, aligned at the 5 km s-' marks on the seismic velocity scales. Dotted lines represent velocity contours at 5.4, 5.8, 6.0, 6.2, 6.4, 6.6, 6.8, 8.2 and 8.4 km s-'.


velocities of 6.6-6.85 km s- '. A typical velocity-depth profile is shown at 75 km model range on Fig. 9.

(3) The top of the continental basement is modelled with a velocity of 5.0 km s-', The velocity gradient (at 190 km model range, see Fig. 9) in the upper crust is 0.15 s-' down to 6.2 kms-' at lOkm depth. The velocity gradient in the lower crust is 0.05 s-' down to a velocity of 7.0 km sC1 just above the Moho. The continental basement is modelled by two linear gradient zones, the inflection point being the 6.2 km s-' contour. This occurs at approximately the same depth as that at which the basement faults sole out and may represent the 'brittle-ductile' boundary at the time of rifting. The 'upper' and 'lower' crust above and below this boundary show the same degree of thinning across the margin, with no differential stretching. This is consistent with pure shear stretching of the crust.

(4) The continental crust thins from a Moho depth of 28km below sea-level at 190km to 12km at 100km. The base of the oceanic crust is relatively flat at 11.5 km depth below sea-level. PmP and Pn phases are observed on all record sections and have been modelled for every instrument. The region of the Moho constrained by the modelling is indicated by the bold line in Fig. 9.

(5) The upper mantle is modelled by a 2km zone of high-velocity gradient (0.1 s - ' ) immediately underneath the Moho, beneath which the gradient decreases to 0.1 s-'. This gradient change was necessary to model the amplitude changes of the Pn phase for DOBSs 2014 (Line 1) and 2017 East (Line 3). and is consistent with the data from all other DOBSs.

3.5 Resolution of the Model

The final model is well constrained both by the crossing ray paths to the three DOBS on Line 3 and by the intersections of Lines 1 and 2 with Line 3. The resolution of the model was estimated by varying the velocities, and the depths of interfaces and gradient changes, to determine the effect on traveltimes and amplitudes. The thickness of the upper continental crust and the depth to the oceanic layer 2/layer 3 boundary can be varied by only a few hundred metres before the fits become unacceptable (deviation from the observed traveltimes greater than 150 ms). Similarly the lower crustal and upper mantle velocities can be varied only by *O.l kms-' and the depth to the Moho only by &1 km before the fits become unacceptable. This resolution is approximately the wavelength of the 8 Hz source when travelling through the lower crust or upper mantle. Although in detail the crust is likely to contain considerable small-scale heterogeneity not constrained by our model, the sensitivity of the final model to perturbations gives confidence that the overall average velocities and depths to the major boundaries are well represented by our model.

4 CONVERSION TO T W O - W A Y TRAVELTIME

The final velocity-depth model for Line 3 was converted to two-way traveltime (twtt) to enable a direct comparison to be made with the WAM deep reflection line. The result is

shown in Fig. 10(a). Features to note are:

(1) exaggeration of the apparent relief of the sedimentary basins on a time section, compared to a depth section, due to the relatively low sedimentary seismic velocities.

(2) Location of the Moho at 9-10s twtt across the whole margin.

(3) Pull-up of the Moho reflection beneath basement highs (see Fig. 10a at 105, 140 and 160 km).

This twtt model is overlain onto a plot of the WAM data in Fig. lO(c). Features of particular note are:

(1) The good fit of the extent of the extruded basalts determined from the wide-angle data between 110-120 km with the extent suggested by the reflection profiles. The continental basement can be seen on the WAM line beneath the prism of volcanic rocks.

(2) The Moho determined from the wide-angle arrivals is located at the base of the packets of lower crustal reflectors seen in the data between 110-135 km. Reflectors to the east of this are obscured by water-bottom multiples so a similar comparison cannot be made in this region. Across the ocean-continent transition zone a good fit is seen between the modelled Moho and the seismic-reflection expression of the base of the crust. It is interesting to note that the main velocity contrast mbdelled by the wide-angle data occurs at the base of the zone of lower crustal reflectivity. True mantle velocities are seen only below this.

5 GRAVITY MODELLING

The seismic model (Fig. 9) was also converted to a density model using a straight-line approximation of the mean velocity-density relationship of Nafe & Drake (1957) between 2 and 7 km s-l (see also Barton 1986; Ludwig, Nafe & Drake 1970; Nafe & Drake 1962). For density p in Mg mP3 and seismic velocity V, in km sC1 the relationship is given by

p = 1.6 + O.19Vp

The water column was assigned a density of 1.03 Mg m-3. Densities for the sediments were taken from DSDP results and varied between 2.0-2.25 Mg mP3. The continental and oceanic crust were divided into 11 layers (increments of 0.2 km s-' in velocity) and the equation listed above used to calculate density. The mantle was assigned a density of 3.30 Mg m-3 (Sibuet et al. 1990).

The free-air gravity anomaly calculated from the seismic model was compared with observed anomalies from RRS Shackleton cruise 4/77 along the same track (Fig. l l a ) . It was necessary to adjust only slightly the densities of a few mid-crustal layers (by at most 0.02Mgm-3) to achieve a good fit to the observed free-air anomaly (Fig. llc).

The fit is good, to within 10mga1, especially between 80-160 km, where the greatest changes occur in sediment thickness and Moho depth. The margin is close to isostatic equilibrium (see Fig. l l b ) with out-of balance stresses of less than 140 bars (14 MPa) that can be sustained by the crust.


ENE

t - 0 150 200

I I I I

rFigure ' O C 1 0 50 100

wsw (4 0 4

4

C 1

11 - 11

12 - 12 13 - 13 14 - 14 15 15 -

0 50 100 153 200

30

1 6.6 I

8-01 3c

35 I I I I J . 1 1 1 ) I I I I I I I I I I I I I I I I I I I l m ' l " . ' ] . " 1 1 " " 1 . " " ~ ~ ~ ~ ~ ~ ~ , , ~ , , ~ , ~ ~ , ~ ~ , , , , ~ , , ~ ~ ] , , , , 35

distance km 0 10 20 30 40 50 60 70 80 90 100 110 1x1 130 140 150 160 170 180 190 200 210

Figure 10. Traveltime conversion of the final model for Line 3. (a) Two-way traveltime model derived from (b), the velocity-depth model. (c) Interfaces and velocity contours from (a) superimposed on a plot of the WAM profile. Water column shown in blue, sediments shown in green, oceapic crust and the basalt flows in red, and the continental crust in yellow. Water-column multiples obscuring later primary reflections are marked.

6 MAGNETIC MODELLING

A magnetic model was constructed by taking the shapes of the oceanic layer 2 (top of oceanic basement down to the rapid transition in velocity from 5.8 to 6.6 km s- ' , see Fig. 9) and the extruded basalts from the seismic model. Both the production of the basalt flows and the onset of sea-floor spreading took place during the Cretaceous normal polarity interval (Scrutton 1985; Masson et al. 1985).

The magnetic anomaly was modelled to fit observed data from RRS Shackleton cruise 4/77 along a profile coincident with Line 3. The remanent magnetization was varied to achieve a good fit (Fig. 12). It was not necessary to alter the geometry of the extruded basalts from that determined by the seismic modelling. A good fit is achieved. with less than 20 nT deviation for a uniform magnetization of 5.5 A m-l.

7 DISCUSSION

7.1 Continental crust

The Line 3 Goban Spur structure shows that the continental crust thins steadily oceanward reaching a minimum

thickness of about 7 km before it merges into oceanic crust. The velocity structure of the thinned continental crust exhibits a smooth increase of velocity with depth accompanied by an overall decrease in velocity gradient. In Fig. 13 we compare our results from the Goban Spur with six other structures for thinned continental crust within the ocean-continent transition (OCT) at non-volcanic margins.

The seven structures in Fig. 13 can be characterized by the crustal thickness, by the velocity at the base of the crust, and by the approximate distance from the edge of the oceanic crust (see Table 1). The velocity in the lower crust along seismic-refraction lines shot on the adjacent continent, including the continental shelf, lies in the range 6.2 to 7.1 kms-' (Sapin & Prodehl 1973; Ginzburg et al. 1985; Cordoba, Banda & Ansorge 1987; Moreira et al. 1982; Mendes-Victor, Hirn & Veinante 1980). Thus if the structures in Fig. 13 are grouped according to the velocity at the base of the crust then one can distinguish structures where this velocity substantially exceeds (by 0.5 km s-' or mote) the velocity at the base of the adjacent unthinned crust (N Biscay Line 2, NB; Iberia Abyssal Plain Line 2, 2; and Galicia Bank Line 7, GB) and where it falls within the normal range (the rest).

14 S. .I. Horsejield et al.

7" x 1 0

-2

100

50

mgal 0

-50

0

2 - 200

b a r s kg I? 0 '1 -

i ! -200

,(C)

3 ' 100 2 m

5

10

k m l 5

20

25

30

1 1 -50 L U

t A , 2.0 2:l 2.25 2.65 2.54 2,57 2.61 2.65 2.69 0

5

1 D

distance km

Figure 11. Gravity modelling of density model obtained from final velocity model for Line 3. (a) Free-air anomaly calculated using densities derived from seismic velocities. The solid line joins gravity values calculated for the observed points (marked by squares). Data is from RRS Shackleton cruise 4/77 (Scrutton 1979, 1985). (b) Load anomaly across the adjusted model calculated for every observed point. (c) Calculated and observed free-air anomalies across the margin using adjusted densities. (d) Adjusted density model used to produce gravity profile shown in (c). Densities are in Mgm-3.

There is a tendency for the structures in the first group with high basal velocities to be associated with thin crust and to lie close to the point of break-up (edge of the oceanic crust) (Table 1) from which one might deduce that the basal velocities of 7.4-7.6 km spl are indicative of underplating in the vicinity of the OCT (White & McKenzie 1989). However, there is no indication of the presence of anomalously hot underlying mantle from a plume at the time of break up, and such high-velocity underplated crust is normally generated only when the parent mantle is several hundred degrees hotter than normal. So we consider it

unlikely that the high basal velocities are indicative of underplating at the time of continental rifting.

An alternative explanation for the 7.4-7.6 km s-' lower crustal velocities is that they represent serpentinized periodite or upper mantle; this explanation was suggested by Pinheiro, Whitmarsh & Miles (1992) for such velocities found under both thin continental and oceanic crust associated with the OCT itself. This hypothesis is supported by the recent discovery of serpentinized peridotite basement underlying Iberia Abyssal Plain sediments close to the edge of the oceanic crust (ODP Leg 149 Shipboard Scientific


400

300 - 200 - 100 -

0

-100 -

-200

Anomaly nT

400

- 300

- 200

- 100

(a)

- T I -L

-200

Model distance km

Figure 12. Modelling of magnetic anomaly along Line 3. (a) Calculated and observed magnetic anomalies. Solid line joins anomaly values calculated for the observed points (marked by squares). Data is from RRS Shackleton cruise 4/77 (Scrutton 1979, 1985). (b) Model producing profile in (a). The shaded area represents magnetized material. This comprises layer 2 of the oceanic crust and the basalts extruded onto the thinned continental crust.

Party, 1993). It also appears from Table 1 that the 7.4-7.6kms-' velocities are found only near the base of unusually thinned crust (4 km or less) although such thin crust is not necessarily accompanied by such velocities (cf Iberia Abyssal Plain Line 1). A thin crust would more easily permit access of sea-water to the upper mantle and consequent serpentinization.

A problem remains, however, with the hypothesis that the 7.4-7.6 km s-l material is serpentinized periodite alone. Continental crust as thin as 3.6 to 4.2 km implies stretching factors of 8 to 10. In theory, for a given asthenospheric potential temperature at the time of rifting, the greater the amount of stretching, the greater the volume of melt produced by adiabatic decompression (White & McKenzie 1989). In the following discussion we calculate the melt generated by adiabatic decompression using McKenzie & Bickle's (1988) method. Note that we have made a number of improvements to the calculations (Bown & White 1994). Rather than a value for the entropy of fusion of 250 J/kg/"C, we use a better estimate of 400J/kg/"C. We also take account of the reduction in mantle-source volume caused by the extraction of melt, and stop the mantle upwelling at the

base of the crust rather than allowing it to rise to the surface. The first and the last of these corrections tend to decrease the amount of melt formed for a given mantle temperature and stretching factor, whilst the other correction increases the calculated melt volume. The net effect is to give predicted mantle temperatures quite similar to those originally reported by McKenzie & Bickle (1988) and by White & McKenzie (1989).

The vertical resolution of seismic models is close to a wavelength which, for many wide-angle lower crustal experiments, is about 850 metres. A thinner layer would not be resolved. If the thickness of melt produced and underplated is 850 m this implies an asthenospheric temperature close to 1220 "C which is significantly less that the range of 1280-1320°C corresponding to the normal range of thickness for oceanic crust (White 1992; Bown & White 1994). However, it is now recognized that oceanic crust adjacent to the non-volcanic rifted margins of the NE Atlantic can be significantly thinner (often only 2.0-4.0 km thick) than the 7 km modal thickness (see below and White 1992; Whitmarsh et a/. 1993). Oceanic crust as thin as 2.0-4.0 km suggests mantle potential temperatures in the

16 S. .I. Horsejield et al.

0

2 n

Y

C

E v c

E 4 a3 v) Q D

0

D

Q

3 5 6

5 ;

8

10

Velocity (km.s-')

4 6 8

Figure 13. Velocity/depth structure for the profiles listed in Table 1. See Table 1 for locations and origins of the profiles. Depths are relative to the top of the non-sedimentary continental crust. AR = Tagus Abyssal Plain; GB = Galica Bank; GS = Goban Spur; NB = North Biscay; T1 = Tyrrhenian Sea; 1 = Iberia Abyssal Plain Line 1; 2 = Iberia Abyssal Plain Line 2.

range of 1210-1260°C at the time of break up. Thus we obtain the interesting result that the asthenosphere potential temperature at the time of continental rifting appears to have been some SO-100 "C cooler than the temperature subsequently reached during normal sea-floor spreading. We can only speculate at present on the reason for this but it seems likely that the subcontinental asthenosphere ascends so slowly during continental break up that sufficient heat is lost by conduction to inhibit, or at least to diminish, melt formation. We conclude that although small quantities of melt will have accompanied stretching by a factor of 8-10, particularly if volatiles were present in the mantle, so far they have not been detected seismically.

The second group of structures in Table 1, to which Goban Spur belongs, are characterised by the absence of 7.4-7.6 km s- ' material and (except for Iberia Abyssal Plain Line 1) by stretching factors of 4 to 5. The Goban Spur margin is unique in this group in having a circa 2 km thick approximately syn-rift layer of basaltic material. In other respects, however, the continental part of Line 3 across the Goban Spur is typical of thinned continental crust at margins with stretching factors of 5 or less and where there has not been appreciable syn-rift underplating at the base of the crust.

7.2 Oceanic crust

The igneous portion of the oceanic crust adjacent to the Goban Spur continental margin averages 5.4 km thick, which is significantly thinner than the mean crustal thickness of 7.0 + 0.8 km of normal oceanic crust formed away from the influence of hot spots, fracture zones or subduction zones (White et al. 1992). However, apart from its reduced thickness the crustal structure shows a distribution of layer 2 and layer 3 velocity gradients within the crust which is quite normal (Fig. 14), so it appears to represent an attenuated version of normal oceanic crust rather than a radically different crustal type.

We consider three possible explanations for the thin oceanic crust adjacent to the Goban Spur margin. The first and most easily dismissed is that it represents fracture zone crust that is commonly found to be abnormally thin (White et al. 1992). The reasons this is an unlikely explanation are: there is no evidence of fracture-zone offsets in the sea-floor spreading anomalies; the crustal velocities are normal whereas fracture-zone crust is commonly highly atypical, often lacking a normal seismic layer 3; and, perhaps most compellingly, thin oceanic crust is found widely along the Atlantic non-volcanic margins and it is unlikely that every seismic profile is located precisely over a fracture zone (Pinheiro et al. 1992; White 1992; Whitmarsh et al. 1993).

The second, and most likely explanation, is that sea-floor spreading rates immediately after continental break up were very low, leading to the accretion of unusually thin oceanic crust (White 1992; Bown & White 1994; Whitmarsh et al. 1993). Very-slow spreading allows the upwelling mantle beneath the rift to cool down by conductive heat loss as it rises, and this reduces the amount of melt that is generated as it decompresses compared to rapidly upwelling mantle. At full spreading rates below 15mmyr-' the amount of melt is reduced considerably (Bown & White 1994). The full spreading rate of the oceanic crust adjacent to the Goban Spur immediately after break up was about 11 mm yr-', and the 25 per cent decrease in thickness is consistent with the decrease that would be expected at such a low spreading rate. Similar reduced-thickness oceanic crust is found elsewhere along the eastern North Atlantic off the Galicia, Iberia and Biscay continental margins, in the final stages of opening of the Labrador Sea, and in the Arctic Ocean (Pinheiro et al. 1992; White 1992; Bown & White 1994; Whitmarsh et al. 1993), where spreading rates were also low. A further reason for reduced melt generation on the Goban Spur profile is that the melt forming the oceanic crust immediately adjacent to the continental margin was derived from mantle that had been rising slowly beneath the continental lithosphere as it stretched over a long period prior to break-up. So this parent mantle would have had a considerable period of time to cool by conductive heat loss both vertically and laterally into the adjacent continental lithosphere prior to its decompression beneath the nascent oceanic rift, thus reducing the initial melt generation.

A third, subsidiary reason for reduced oceanic crustal thickness immediately adjacent to the continental margin is that some of the newly formed melt may have flowed on the surface or intruded laterally into the easily penetrated

0

2

4

Depth into 6

igneous crust km

8

10

12


Table 1. Velocity structures within the continental crust of the ocean-continent transition (OCT) at non-volcanic rifted margins.

Margin

Goban Spur

N. Biscay

Galicia Bank

Iberia AP

Iberia AP

Tagus AP

Approx. Crustal Basal Distance

Profile thickness velocity from

Line 3

Line 2

Line 7

Line 1

Line 2

Line AR

Tyrrhenian Sea Line T1

(km)

7.6

3.6

4.2+

3.4

3.6

6.8

6.4

(km s-l-) break-up1

7.0 252

7.4 1 53

7.4+ 15‘

( k m )

6.5 1 304*5

7.6 60‘a5

(6.0)‘ 0-405

(6.35)‘ ?

this paper; Scrutton 1985

Whitmarsh et al 1986

Whitrnarsh et al 1993

Whitrnarsh et al 1990 ODP Leg 149 Shipboard Scientific Party 1993

Whitrnarsh et al 1990 ODP Leg 149 Shipboard Scientific Party 1993

Purdy 1975 Pinheiro et al 1992

Duschenes et a1 1986

Velocity km/s 2 4 6 8 10

Source

1 Break-up points defined variously as,

2 foot of basement escarpment at 98 km (Fig. 9)

3 boundary recognised on MCS profiles and wide-angle seismic recordings

4 peridotite basement ridge 5 eastern (landward) edge of seafloor spreading model

6 velocity not based on synthetic seismogram modelling

t Figure 14. Seismic velocity structure of the oceanic portion of Line 3 adjacent to the Goban Spur continental margin compared to the envelope of normal oceanic crustal structure from profiles constrained by synthetic-seismogram analysis over normal oceanic crust older than 125 Ma in the North Atlantic (from White 1992).

continental crust as sills and dykes. This is a commonly observed occurrence on volcanic continental margins where melt emplacement rates are much higher (White 1992), but it would not be surprising if it also occurred on a more limited scale on ‘non-volcanic’ rifted margins. As already noted there is indeed evidence of lower crustal layering and of oceanward-derived basalt flows at the surface near to the ocean-continent transition on the Goban Spur. However, this mechanism would only transfer a relatively small volume of melt and would only operate on a limited scale immediately adjacent to the OCT, because once a fully-fledged oceanic-spreading centre had developed the melt intrusion and extrusion would be limited to the axial region.

8 ACKNOWLEDGMENTS

The authors thank the Master and crew of the R. V. Suroit and the RCframarge scientific party for their help at sea under difficult conditions. R. E. Kirk and M. R. Saunders operated the DOBSs at sea and played a major role in the data processing ashore. The participation of IOSDL in RCframarge was largely funded by IFREMER.

Modelling was carried out using a Maslov computer program written by D. G. Lyness modified by S. J. H. to run on the Bullard Laboratories’ SUN Sparcstations. The wide-angle airgun data were processed and plotted on the Bullard Laboratories’ CONVEX C120 using Merlin Geophysical’s SKS package and programs written by S. J . H. and C. Peirce. This work also foams part of a study of the


relationship between tectonics and magmatism supported at Cambridge by the Natural Environment Research Council. S. J. H. was funded by a scholarship from Shell International Petroleum Co. Ltd. Department of Earth Sciences, Cambridge contribution number 2324. Institute of Oceanog- raphic Sciences Deacon Laboratory contribution 94018.

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Crustal structure of the Goban Spur rifted continental margin, Ne Atlantic

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