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Geophys. J . Int. (1993) 1l3,701-726 The ocean-continent boundary in the Gulf of Lion from analysis of expanding spread profiles and gravity modelling G. P. Pascal,'" A. Mauffret2 and P. Patriat3 'Instirut Frangais du Pe'frole,BP 31 1, 92506 Rueil-Malmaison, France 'Laboratoire de Ge'odynamique, Tectonique et Environnement, University P. et M. Curie, Place Jussieu, 75252 Paris Cedex 05, France 'Instirut de Physique du Globe, University P. et M. Curie, Place Jussieu, 75252 Paris Cedex 05, France Accepted 1992 November 2. Received 1992 November 2; in original form 1991 November 5 SUMMARY Two-ship multichannel seismic profiles, deep penetration (ECORS) and conventional seismic lines (LIGO surveys) are used to study the crustal structure of the Gulf of Lion (Western Mediterranean). 11 full ESPs (Expanded Spread Profiles) with total shot-receiver ranges up to 60 km were shot in 1981 perpendicular to the margin of the Gulf of Lion and in 1988 a deep MCS seismic profile (ECORS-CROP program) was performed parallel to the ESPs. These ESPs were analysed by matching traveltime and amplitude variations in both the x - f and z-p domains. The resulting P-wave velocity/depth model has the following features, (a) beneath the continental slope of the Provenqal margin a rapid rise of the Moho from 20 to 14 km and the existence of an anomalous 7.2-7.4 km s-' velocity layer, (b) from the base of the slope to the extensive salt-domes domain a 5-6 km thin crust which does not appear typically oceanic in nature, (c) quite typical oceanic crust up to the Sardinian margin. Gravity modelling is consistent with the seismic results. The OCB (ocean-continent boundary) could be placed north of that postulated by previous authors, where the data indicate a remarkably narrow transition between continental and 'oceanic' crust, or south where a typical oceanic crust, which correlates well with the domain of the salt domes and of large magnetic anomalies, has been determined. A very prominent reflector is clearly seen, at the base of the continental slope, on the seismic reflection profiles and corresponds to the top of an 7.2-7.4kms-' velocity layer. The high-velocity layer is 2-3 km thick where the crust is thinnest and has a limited lateral extent seawards. This anomalous crustal structure could be the result of extremely thinned and possibly broken up, continental crust underplated and intruded by partial melt, or could represent serpentinized peridotite material. Important questions about the evolution of the Gulf of Lion cannot be addressed using these new results alone without addition of other constraints. Nevertheless a two-stage mechanism of drifting and rifting of this part of the Western Mediterranean Sea is proposed. Key words: crustal structure, ESP, Gulf of Lion, Mediterranean, ocean-continent boundary. INTRODUCTION Plate kinematic models suggest that the African and European plates have been converging since anomaly 34 'Now at: University of Brest, Avenue Le Gorgeu, 29283 Brest cedex, France. (Olivet et af. 1984). Despite convergence between the two plates, the Western Mediterranean Sea is the result of a late Oligocene-early Miocene extension. The Proven~al Basin is placed in the northwestern Mediterranean Sea and includes the Gulf of Lion to the west and the Ligurian Sea to the east. It opened, after a rifting episode from Oligocene to earliest Miocene, as the result of 70 1 702 G. P. Pascal, A. Mauffret and P. Patriat profiles (ECORS, Etude de la CroQte Continentale et OcCanique par Reflexion et RCfraction Sismiques) as well as a large collection of conventional multichannel seismicreflection profiles (mainly LIGO surveys), and gravity data across the Gulf of Lion. These results are compared to a previous interpretation of Le Douaran, Burrus & Avedik (1984), which relied only on refraction arrivals. In light of the foregoing seismic and gravity analysis we are able to propose a new ocean-continent boundary (OCB) along the studied seismic lines. PREVIOUS WORK Figure 1. Location map of the Provenpl basin. The volcanism in Sardinia is omitted. NPF?: hypothetic extension of the North Pyrenean Fault; GLP2 and Si: GLP2 and Sirocco wells. The positions of the ECORS NW, NS, ZZ, CROP, and Ligo 4 seismic profiles are indicated. Large numbers: ESPs of the present study; thin number: ESPs of the Le Douaran et al.3 paper (1984). the southeastward drifting of Corsica and Sardinia (Le Pichon et af. 1971; Rehault, Boillot & Mauffret 1984) in early Miocene time. A probable emplacement of oceanic crust (Burrus 1984; Le Douaran, Burrus & Avedik 1984; Rehault et af. 1984) is coeval with this opening of the northwestern Mediterranean. The Gulf of Lion is in line with the southeast Basin (Fig. 1) where up to 10km of Mesozoic sediments have been accumulated (Curnelle & Dubois 1986). The opening of the Bay of Biscay and the eastward drift of the Iberian block (Arthaud, Ogier & Seguret 1981; Olivet, Auzende & Beuzart 1983) affected the region, mostly by left-lateral motion along the North Pyrenean fault zone. The faults (Cevennes and Durance Faults) which limit the Southeast Basin and other faults of the same orientation (Nimes Fault) acted first with a strike-slip component during the Pyrenean Orogeny, then with a normal component during the subsequent extension (Arthaud & Seguret 1981; Bergerat 1982; Mauffret & Gennesseaux 1989) which formed the grabens onshore and beneath the continental shelf of the Gulf of Lion. The southwestern edge of the Provenqal Basin is close to another Oligocene-Miocene rift basin, the Valencia Trough, where recent seismic and gravity studies (Pascal et af. 1992; Torne et af. 1992) have demonstrated that, except in the northeastern part of the basin where transitional to oceanic type crust is present, the Valencia Trough is underlain by a thinned continental-type crust. At the southeastern boundary of the Provenqal Basin, in Sardinia, volcanic rocks ranging from Oligocene to Quaternary cover the western side of the island and a large NS oriented rift extending along the entire western part of the island (Cherchi & Montadert 1982). This study presents a detailed analysis of the velocity and crustal structure of the Gulf of Lion margin based on eleven ESPs (Expanding Spread Profiles), deep-penetration seismic The study area has been the subject of several geophysical and geological studies over the last decades. Northwest of the Cevennes fault, in the Massif Central, the crust is 30 km thick, but the Moho rises up to 24 km in the Southeast Basin which underwent Mesozoic and Cenozoic riftings (Sapin 8c Him 1974). The crust is about 20km thick beneath the continental part of the Gulf of Lion and as thin as 8 km at the base of the continental slope. In the presumed oceanic domain the crust is 6-7 km thick (Le Douaran et af. 1984). A long-range seismic profile in the Western Mediterranean Basin (Hirn, Steinmetz & Sapin 1977) indicates a 7.7 km s-l velocity at a depth of about 11km which marks the boundary between the crust and the top of the uppermost mantle. At the southeastern edge of the Provenpl Basin, below Sardinia, the Moho rapidly deepens from 20-25 km depth at the margin to 30 km beneath the island (Egger et al. 1988). Le Douaran et af. (1984) proposed placing the northwestern ocean-continent boundary (OCB) between ESPs 205 and 206 which show, according to their interpretation, a change in the upper crustal velocity from 5.3-5.4 km s-l to 5.7 km s-l and which correspond to a transition from a magnetic quiet zone to an area of high magnetic anomalies. Magnetic anomalies show a complex pattern and have been interpreted as the expression of sea-floor spreading (Bayer, Le Moue1 & Le Pichon 1973; Burrus 1984). Alternatively the OCB has been placed (Le Cann 1987) at the northwestern limit of the salt dome area which coincides with the zone of large magnetic anomalies in the central basin. According to Le Cann (1987), the shape of the boundary between continental and oceanic domains is likely to be represented by lineaments expressed by salt-wall alignments and easily recognized from the bathymetry. Various reconstructions, using different poles, of the eastward drifting of Corsica and Sardinia have been proposed (Auzende, Bonnin & Olivet 1973; Burrus 1984; Olivet ef al. 1982; Rehault et al. 1984). According to palaeomagnetic data (Montigny, Edel & Thuizat 1981; Vigliotti & Kent 1990) and aeromagnetic survey data (Galdeano & Ciminale 1987), Sardinia and Corsica underwent a 30" counterclockwise rotation during the early Miocene. Thermal asymmetry is indicated by the increase in heat flow from the Gulf of Lion to the northwest Sardinian margin (Burrus & Foucher 1986). The greatest heat flow values correlate with the highest Bouguer anomaly suggesting that intrusions from the mantle could be responsible for both observations. On the Provenqal side the heat flow and the rates of synrift and postrift subsidence are, Crustal structure of the Gulfof Lion at a first approximation, compatible with the predictions of the uniform extension model although an excess of subsidence is found in the centre of the basin (Burrus 1989); this excess of depth appears to be a general characteristic of the Western Mediterranean basin, as well as the Ligurian Sea (Jemsek et al. 1985) and the Tyrrhenian Sea (Hutchinson et al. 1985). In 1988 October a seismic survey was carried out in the Gulf of Lion within a French (Ecors)-Italian (Crop: Crosta Profonde) joint project. de Voogd et al. (1991) presented the preliminary results of the seismic profiles. Across the shelf and slope, several packages of S to S E dipping reflectors, that affect the upper and lower crust, correspond most probably to the deep roots of NW verging Pyrenean thrusts which have been recognized by onshore geological survey, or drilled in the industrial offshore boreholes (Arthaud et al. 1981). Few half grabens bounded by listric faults are observed and these only occur across the upper shelf. The zone where most of the crustal thinning occurs is characterized by prominent lower-crustal and possibly Moho reflections gently dipping towards the continent. In the deep-water domain, a deep depression followed by a flat basement are observed, whereas the basement topography is rough on the Sardinia margin. A crustal or Moho reflection dips abruptly from the ocean beneath the slope of the Sardinian margin. This margin is a starved one, with a slope much steeper than in the Gulf of Lion. Much more detailed work is in progress, especially the migration of the deep part of the ECORS profile. DATA ACQUISITION A N D REDUCTION The data used in this study consist of 11ESP (expanding spread profiles), some nearby industrial CDP lines and the ECORS-CROP multichannel seismic reflection lines (Fig. 1). The shooting ship in the 1981 ESP experiment was equipped with a 6 X 1000in3 airgun array and the recording ship towed a 2.5 km streamer with 48 channels (Le Douaran et al. 1984). An ESP experiment is a multiple-fold wide-angle reflection/refraction data set acquired by two ships steaming apart from a common mid-point, one shooting and the other recording (Stoffa & Buhl 1979). In this experiment the two vessels started from initial positions 65 km apart at the end points of each ESP profile, moved toward each other crossing a mid-point at separation less than 1km and moved away steaming on reciprocal tracks at a constant speed of 5.4knots, so that a common fixed-central-reference point was maintained. This common mid-point geometry is important to reduce the effects of dipping interfaces on determining interval velocities (Diebold & Stoffa 1981). In that manner, we obtained two profiles in one experiment, one at approaching ranges and another at departing ranges. As the two profiles are almost symmetrical, we used only one for the interpretation. The ECORS seismic reflection data were acquired by CGG (Compagnie G6nCrale de Geophysique) and the 2 0 k m of CROP data by OGS (Osservatorio Geofisico Sperimentale), both in 1988 October. A large airgun array was recorded by a 3000m long, 120 channel streamer; the distance between shots was 50 m, resulting in 3000 per cent coverage (de Voogd et al. 1991). Water-bottom multiples were successfully removed in the deep-water area, revealing 703 the top of the basement. A post-stack F-K migration was performed on the northern half of the NS and NW lines. Line drawings of the profiles can be found in de Voogd et al. (1991). As the structure below the ESPs 203 to 215 is quite laterally homogeneous, all the interpretation was made with a 1-D model. The velocity-depth models were constructed using x-t traveltime forward modelling and z-p inversion modelling techniques (Diebold & Stoffa 1981). This iterative modelling scheme incorporates near-vertical incidence reflections and wide-angle reflections and refractions to derive a complete traveltime solution. The modelling was further improved by matching computed reflectivity seismograms in the x-t domain and z-p domain. We were not able to reinterpret all the ESPs shot along the ECORS-CROP seismic profile. ESP 208 or 209 were not available and we could not reprocess the data. Both approaches were used to construct the velocitydepth models for the sedimentary layers. The first approach started with predictive deconvolution of the near-offset data and F-K filtering designed to eliminate water-borne energy. The measured traveltimes for a reflection were matched with traveltimes calculated by ray tracing through models in which the interval velocity for a given layer was initially varied over a wide range (Limond & Patriat 1975). The thickness of the layer was determined for each trial velocity from the zero-offset traveltime. This process was repeated until the variances of both the traveltimes and traveltime errors were minimized. Random errors will occur in picking and digitizing the different arrivals, but, in general, these errors are believed to have a minor effect on the final velocity structure. Since poor correlations of phases due to the interference of reflected and refracted waves from different interfaces can be responsible for larger errors, particular care was taken in identifying the different phases. Only one layer was optimized at a time. In the second approach, a preliminary estimate of the structure was constructed using the z-sum inversion technique developed by Diebold & Stoffa (1981). For this purpose, the data were first transformed from the offset traveltime ( x - t ) to the intercept time-ray parameter ( r - p ) domain by slant stacking the X-t data (Stoffa et al. 1981). Once in the z-p domain, discrete picks for each ray parameter were performed on the postcritical arrivals and a velocity-depth function was derived using the z-sum inversion technique. This preliminary solution was then refined by modelling the seismic traveltime trajectories in both the x-t and z-p domains until a final solution that satisfied the data in both domains was obtained. For the deep structure, however, we could not use the z-p stack data for all the ESPs, since the r-p stacking did not enhance deep arrivals and did not result in increased penetration and resolution at depth in the velocity profile. The synthetic z-p seismograms, except for ESPs 201 and 202, were computed using the approach developed by Dietrich (1988). In this approach, the generalized reflection and transmission matrix is used to compute the seismograms in the r-p domain. The same parameters used for x - t seismograms were incorporated in this scheme. The x-t reflectivity seismograms included not only primary P and S arrivals but every possible multiple and converted arrival. Synthetic seismograms were calculated following Kennett’s Crustal structure of the Gulf of Lion 0 RAY PARAMETER (slkm) 0.4 0.6 RAY PARAMETER (slkm) 0.4 0.6 0 0.2 0 0.2 1 - z2 w 3 EL 4 6 H 5 k6 z7 8 9 0 1 5 2 v t7 a 9 ESP 2 0 3 Figure 2(b). Upper panel shows data from ESP203 transformed to the domain of intercept time (T) and ray parameter (p). The input data were 50m equi-spaced seismograms from 0 to 60km. The stacking was performed for ray parameters from 0.02 to 0.67 s km-' at 0.002 s km-' increment. Lower panel shows that r-p traveltimes corresponding to the final velocity model. The labels correspond to those in Fig. 2(a). velocities increase from 5.7 to 6.6 km s-' (phases Tuc and Tmc). The match between synthetics and observed amplitudes corresponding to the crustal arrivals can be seen on Fig. 2(c). In the same way 7-p curves exhibit well developed post-critical crustal arrivals corresponding to the three layers (Fig. 3b). To fit the major amplitude peak at about 30 km offset we associated a moderate gradient velocity layer, in which the velocity increases from 7.1 to 7.3 km sC1. The gradient needs to be approximately 0.225s-' to model correctly the amplitudes and the positions of the post-critical Moho reflection and of the head wave. The Moho is modelled as a single velocity step from 7.3 to 8.2 km sC1 at a depth of 14.8 km. ESP 205 is located above a less perturbed basement (Fig. 20b). Sedimentary cover is not very different from that of the previous ESP (Figs 4a,b,c and d). The first amplitude peak at 7-8 km range is produced by the gradient zone from 2.9 to 3.8 km depth, for which velocities increase from 2.5 to 3.2kms-'. The second amplitude peak is produced by another gradient zone from 3.8 to 4.4 km in which velocities 705 increase from 3.5 to 4.2 km sC'. Crustal arrivals (phases Tuc and Tmc) are mapped clearly in both domains as post-critical arrivals (Figs 4a and b). The phase Tuc corresponds to a reflection arrival in the 5.6 km s-' constant velocity layer. The phase Tmc is due to energy turning in a moderate-gradient velocity layer in the middle crust. A clear arrival with apparent velocity of 8.1 kms-' is seen over the distance range from 40 to 50 km. The phase is identified as the Pn wave. To model the strong Pn arrival at the appropriate range and the pre-critical PmP from about 20 km offset a velocity gradient at the base of the lower crust explains well both traveltimes and amplitudes of the PmP phase. However, the actual magnitude of the amplitudes of the Pn phase does not match well the observed phase. Stronger head wave would need a high velocity gradient at the base of the crust, but would reduce the precritical Moho reflection. Sedimentary cover of ESP 206 (Figs 5a,b and c) is very similar of that of the other ESPs. A sudden increase in amplitude of the arrivals at about 7 km offset arises from a gradient velocity layer of 4.0 to 4.6 km s-' over a 1.6 km depth range which corresponds to the salt layer. Very clear reflections at small offsets are observed and correspond to the evaporite and salt layers. Seismic phases T1 to T3 correspond to the same arrivals as described previously. Middle and lower crustal arrivals have been modelled as turning rays coming from moderate-gradient velocity layers. As can be observed in the x - t section of Fig. S(a) a very thin layer of 7.2 km s-' mean velocity was introduced in the model to fit the amplitude peak observed at about 25-30 km offset as well as the gravity modelling (see next section). ESP 207 is quite perturbed by the extensive salt domes, as seen on the CDP seismic section (Fig. 16). Below the sea-floor reflection (Figs 6a and b) high gradient layers were modelled from 2.5 to 3.9 km in which the velocity increases from 2.0 to 3.1 kms-'. Branches T1 and T2 correspond to turning rays in moderate gradient velocity sedimentary layers. The branches Tuc and Tmc mapped clearly in the x-t section and less visible on the 7 - p domain because of the perturbation caused by the salt domes, correspond to seismic reflections in the crust. The first one is produced by a gradient from 9.6 to 11.9 km depth for which velocities increase from 5.5 to 6.0 km s-', and which corresponds to an inflexion point in the r-p curve (Fig. 6b). The second one is due to a constant velocity layer at 11.9-14.7 km depth for which the velocity is 6.8 km s -'. To fit the strong Moho reflection we used a single-velocity step from 6.8 to 8.1 kms-' at a depth of 14.7 km. For the two last ESPs synthetic seismograms have been computed but are not shown in this paper. ESP 212 like ESP 211, was shot parallel to the seismic section Lig 4 (Fig. 1). Below the sea-floor reflection (Fig. 7) high-gradient velocity layers were modelled, in which the velocity range from 1.8 to 3.1 km s - ' . Branches T1 and T2 provide the velocity structure from about 3.6 to 6.3 km depth. It comprises two constant velocity layers, in which the velocity increases from 4.1 to 4.7 km s-I. A difference with ESP 203 is the absence of the 5.1 kms-' velocity sedimentary layer. Upper and middle crustal reflections have been interpreted as corresponding to rays travelling in constant velocity layers. The first layer is characterized by a 5.7 km sC1 velocity, the second one by a 6.2 km s-' velocity, G. P. Pascal, A. Mauffret and P. Patriat 706 @ ) O F F S E T (km) 3.0 30 40 40 h v) Y -I 1 1 1 Figure 2(c). Upper panel: comparison between observed record section of ESP 203 and reflectivity seismogram for the final velocity model presented in Table 1. The amplitudes of both the sections are scaled proportional to the distance. In the synthetic section all the phases are computed. Consequently some phases are not present in the observed section due to poor signal-to-noise ratio or strong interference between phases. Lower panel: comparison between observed record section of ESP 204 and reflectivity seismogram for the final velocity model. Crustal structure of the Gulf of Lion Table 1. (Continued.) Table 1. ESP 201 Layor no 1 2 3 4 5 8 7 8 9 ESP202 1 2 3 4 5 6 7 8 9 10 11 12 13 ESP203 1 2 3 4 5 6 7 6 9 10 11 12 13 ESP204 1 2 3 4 5 6 7 6 9 10 11 12 13 14 ESP205 1 2 3 4 5 6 7 8 9 10 11 12 13 ESP206 1 2 3 4 5 6 7 8 9 10 11 12 13 14 707 (d 1 .oo Vtop (km/s) 1.500 2.550 3.300 3.700 4.300 5,000 6.100 7.200 6.400 Vbot (kmk) 1s o 0 2.600 3.500 3.650 4.450 5.000 6.100 7.400 1.41 2.11 2.41 2.64 3.06 7.44 7.99 1.060 1.520 2.130 2.690 2.950 3.500 4.960 5.460 6.990 7.910 18.040 19.800 1.500 2.100 2.400 3.200 3.500 3.700 4.400 5.150 5.350 5.600 6.200 7.100 8.400 1 so0 2.250 2.650 3.200 3.500 3.700 4.550 5.150 5.450 5.800 6.200 7.300 1.44 1.64 2.33 2.66 2.63 3.12 3.79 3.99 4.54 4.65 8.12 6.61 1.770 2.010 2.300 2.600 3.500 3.900 4.450 5.740 8.200 7.940 9.300 12.500 15.100 1.500 1.900 2.300 2.450 3.050 3.700 4.300 4.700 5.100 5.300 5.700 8.200 7.200 8.100 1.500 2.100 2.300 2.650 3.250 3.700 4.450 4.700 5.100 5.300 5.700 6.200 7.400 2.36 2.60 2.65 3.23 3.67 3.69 4.14 4.69 4.67 5.53 6.00 7.04 7.75 2.100 2.350 2.750 3.260 3.740 4.300 5.400 6.610 8.090 9.700 11.250 13.M 14.750 1.500 1.600 2.150 2.630 3.040 3.700 4.200 4.700 5.050 5.300 5.700 6.600 7.100 6.200 1.500 2.100 2.260 2.830 3.040 3.700 4.400 4.700 5.150 5.300 5.700 6.600 7.300 2.60 3.05 3.40 3.60 4.11 4.41 4.91 5.51 6.01 6.66 7.35 7.75 6.15 1.500 1.500 2.000 2.300 3.200 4.200 4.400 4.700 5.100 5.300 5.600 6.650 7.250 3.15 3.55 3.70 4.47 4.79 5.34 5.64 6.39 7.15 7.74 6.27 6.39 1.500 2.100 2.400 2.800 3.300 4.300 4.600 4.800 5.100 5.300 5.900 6.900 7.300 3.23 3.50 3.60 4.00 4.47 4.62 5.17 5.62 6.32 6.96 7.51 6.13 8.32 Depth (km) 0.750 1.310 2.490 3.060 4.000 4.600 17.900 19.800 2.360 2.760 2.940 3.630 4.440 5.650 6.630 8.220 10.340 12.000 13.&10 14.200 2.420 2.690 3.040 3.310 4.060 4.360 5.570 7.100 6.360 10.090 11.870 13.710 15.100 13 0 0 2.300 2.500 3.500 4.400 4.700 4.950 5.300 5.600 8.550 7.100 8.100 1.500 2.000 2.250 2.600 3.100 4.000 4.450 4.650 5.100 5.300 5.450 6.450 7.100 8.100 ESP207 1 2 3 4 2.550 2.800 3.250 3.450 3.920 4.450 5.250 7.380 9.060 9.620 11340 14.710 1.so0 2.000 2.280 2.700 3.100 3.700 4.450 4.700 5.100 5.300 5.500 8.600 6.100 1.500 2.000 2.260 2.700 3.100 4.200 4.600 5.000 5.100 5.300 6.000 6.600 3.40 3.65 4.04 4.19 4.50 4.75 5.1 1 6.03 6.70 6.90 7.71 6.52 ESP211 1 2 3 4 5 6 7 8 9 10 11 12 13 2.450 2.640 3.310 3.610 4.200 5.160 6.070 7.400 10.160 11.200 13.500 14.600 1s o 0 1.600 2.400 3.100 3.500 4.100 4.300 4.750 5.350 5.700 6.600 7.100 6.200 1.500 2.000 2.700 3.100 3.500 4.300 4.550 4.850 5.350 5.950 6.900 7.300 3.27 3.65 4.01 4.33 4.56 5.01 5.41 5.96 7.00 7.35 6.04 6.34 ESP212 1 2 3 4 5 6 7 8 9 10 11 12 2.270 2.610 2.660 3.190 3.840 5.260 6.300 7.720 9.590 10.470 13.340 1.500 1.600 2.300 2.500 2.900 4.100 4.700 5.350 5.750 6.240 7.100 6.000 1S O 0 2.100 2.500 2.700 3.050 4.100 4.700 5.350 5.750 6.240 7.300 3.03 3.37 3.58 3.61 4.10 4.89 5.32 5.85 6.50 7.05 7.60 2.700 2.900 3.410 3.770 4.300 4.800 5.840 7.330 8.970 11.410 12.070 14.160 1.500 1.BOO 2.250 2.900 3.450 4.300 4.100 4.500 5.050 5.500 6.600 6.600 6.000 1.500 2.100 2.500 2.900 3.700 4.300 4.100 4.700 5.050 6.200 6.600 7.000 3.60 3.60 4.25 4.50 4.80 5.03 5.54 6.19 6.84 7.66 7.86 6.50 2.720 3.070 3.500 4.090 4.650 5.660 7.670 9.310 11.210 14.950 1.500 1.900 2.250 3.000 4.300 4.100 4.400 5.000 5.400 6.700 8.000 1.500 2.200 2.650 3.800 4.300 4.100 4.900 5.150 6.300 6.900 3.63 3.85 4.30 4.65 4.91 5.41 6.26 6.90 7.46 6.56 5 6 7 8 9 10 11 12 13 ESP215 1 2 3 4 5 6 7 6 9 10 11 12 13 ESP216 1 2 3 4 5 6 7 6 9 10 11 as found in the ESP 203, and indicate a continental crust type. The deeper crustal structure was modelled in a similar way to ESP 203 (phase TIC). ESP 211 is quite different from ESP 212 (Figs. 8a,b and c). Below the sea-floor reflection high-gradient velocity layers were modelled, in which the velocity ranges from 1.8 to 3.5 km s-'. A sudden increase in P velocity from 3.5 to 4.1 kms-' marks the top of the salt layers. Branch T1 provides the velocity structure from about 4.2 to 6 km depth. 708 G. P. Pascal, A . Mauffret and P. Patriat Figure 3(a). Upper panel: observed record section of ESP 204. Lower panel: calculated traveltimes curves. Upper right inset is the location map for ESP 204. Cross marks ESPs midpoint location. The main arrivals are labelled: Rsf: sea-floor reflection; T1: rays turned in the 4.4 km s-' layer, T2: rays turned in the 4.7 km s-' layer; Tuc: rays turned in the upper crust; Tmc: rays turned in the mid-crust: PmP: Moho reflection critical distance. It comprises two velocity gradient layers, in which the velocity increases from 4.1 to 4.6 km s-I. As for the previous seismic section the 5.1 km s-' velocity layer is absent. Upper and middle crustal reflections (Tuc and Tmc) have been interpreted as turning rays in moderate gradient velocity layers, very similar to those determined for ESP 204. The deeper crustal structure was modelled in a similar way to ESP 204. ESPs 215 and 216 (Figs. 9a,b and 10) are located in the deeper part of the basin. Both are quite perturbed by the presence of large salt domes visible on the CROP profile. Below the sea-floor reflection and very thin sedimentary layers whose velocities vary from 1.9 to 3.6 kms-', salt layers are prominent. The main feature of these two seismic sections is the presence of a velocity inversion in the salt layers. Classical salt velocities are recognized: 4.3, 4.1 and 4.5 to 4.8 kms-I. The slight disruption of the arrivals between 8 and 15 km, seen in both x-f and 5-p domains (Figs. 9a,b and 10) has been interpreted as a low-velocity zone (LVZ) located just above the 5.1 km s-' velocity layer. The other main common feature is the presence of gradient velocity layers in the crust (phases Tmc and Tlc). These branches mapped clearly in both domains as post-critical arrivals (Figs 9a,b and 10) and correspond to turning rays coming from moderate-gradient velocity layers. These velocity gradients appear to be related to those of a classical oceanic crust. ESP 201 is first characterized by a poor signal-to-noise ratio (Figs l l a and b). Sedimentary layers are quite thick ESP 2 0 4 Figure 3(b). Upper panel shows data from ESP 204 transformed to the domain of intercept time (7)and ray parameter (p).The input data were 50m equi-spaced seismograms from 0 to 60km. The stacking was performed for ray parameters from 0.02 to 0.67 s kmat 0.002 s km-' increment. Lower panel shows the 7-p traveltimes corresponding to the final velocity model. The labels correspond to those in Fig. 3(a). ' (about 4 km) and velocities increase from 2.6 to 5 km s-', sometimes with moderate velocity gradients (Fig. 20a). The first sedimentary layers show very clear reflections without distortion; so our assumption of 1-D modelling was accurate for these layers, more or less parallel to the sea floor. The crustal arrivals were modelled in a 2-D model, taking into account the variable topography of the basement and using both ray tracing and synthetic seismograms. High amplitudes from range of 35 km correspond to strong reflections at the base of the thick continental layer (velocity of 6.1 km s-I). The lower crustal arrivals from 50 km offset are modelled by a low gradient layer having a mean velocity of 7.3 km s-I. The Moho is modelled here as a single velocity step from 7.4 to 8.4 km s-', the latter being perhaps a too high velocity. Both the magnitude and the shape of the amplitude high from the Moho reflection are quite well matched. Because of the low signal-to-noise ratio and the fact that the ESP was shot perpendicular to the margin, the error in the amplitudes is greater than for the other lines. It is possible that a layered Moho would provide an equally good fit, as seen on the CDP section (Fig. 20a). Crustal structure of the Gulf of Lion Figure 4(a). Upper panel: observed record section of ESP 205. Lower panel: calculated traveltimes curves. Upper right inset is the location map for ESP 205. Cross marks ESPs midpoint location. The main arrivals are labelled: Rsf; sea-floor reflection; T1: rays turned in the 4.3 km s-' layer; T2: rays turned in the 4.7 km s-' layer; T3 corresponds to rays travelling trough the 5.1 km s-' layer; Tuc: rays turned in the upper crust; Tmc: rays turned in the mid-crust. The main difference between ESPs 201 and 202 (Figs 12a and b) is the presence, in ESP202, of a 5.3 kms-' sedimentary layer just above the basement. This layer is restricted to the deep margin and is present at the base of the depression between the locations of the two ESPs (Fig. 20a), but does not extend to the upper margin as shown by the modelling. High amplitudes from an offset of 30km correspond to reflections in a thin 5.8 km s-I velocity layer and the still-thick 6.3 km s-l velocity layer. Another amplitude high at about 45 km with high-velocity arrivals of about 7.1 km s - I is seen on the section (Figs. 12a and b). At last a weak Moho arrival with about 8.4 km sC1 velocity is observed. A low gradient layer with a mean velocity of 7.2 km s - ' could match both traveltimes and amplitudes. It is clear that more precision on such data is difficult, although we used quite detailed methods of interpretation. In our opinion the lower crustal layer could represent a gradient velocity at the base of the crust or the presence of a weak laminated lower crust, especially below ESP 201. GRAVITY MODEL To constrain the crustal structure along the profile between ESP 201 and ESP 215 we computed the gravity effect of the seismic structure deduced from the ECORS-CROP profile and ESP data set, and compared it with the observed free-air gravity anomaly (Finetti & Morelli 1973). The gravity anomaly map (Fig. 13) shows that the modelled 709 ESP 2 0 5 Figure 4(b). Upper panel shows data from ESP 205 transformed to the domain of intercept time ( 7 ) and ray parameter @). The input data were 50m equi-spaced seismograms from 0 to 60km. The stacking was performed for ray parameters from 0.02 to 0.67 s km-' at 0.002 s km-' increment. Lower panel shows the r-p traveltimes corresponding to the final velocity model. The labels correspond to those in Fig. 4(a). profile is perpendicular to the long-wavelength gravity anomaly trend. This fact allows us to assume a 2-D model, which is reasonable, particularly for the deeper structure. There is a good agreement between the velocity-depth solutions derived from ESPs and the ECORS seismic profile in terms of thickness and geometry of the sedimentary layers. The bathymetry as well as the sedimentary layers were digitized on the ECORS profile from ESP 201 to ESP 207. The velocity-depth solutions were used to assign the density of the sedimentary layers using velocity-density relationships (Nafe & Drake 1963). In our calculations we assumed a 1030kgmP3 for the water layer. We did not introduce all the sedimentary layers; so we assumed a first layer with a uniform density of 2330kg11-~ which represents the 5.1-5.3 km s- layer. To introduce the topographic variations of the basement along the profile, we digitized the corresponding reflector on the ECORS profile, after depth conversion of the seismic line. Then, according to the velocities determined from the ESPs, we assumed a density of 2730 kg mP3 for the upper crust and a density of 2770 kg mP3 for the crystalline crust at the north of ESP 204. ' G. P. Pascal, A . Mauffret and P. Patriat 710 OFFSET ( k m ) 40 6 h a v) v t- 10 12 14 Figure 4(c). Comparison between observed record section of ESP 205 and reflectivity seismogram for the final velocity model presented in Table 1. The amplitudes of the both sections are scaled proportional to the distance. In the synthetic section all the phases are computed. Consequently some phases are not present in the observed section due to poor signal-to-noise ratio or strong interference between phases. The anomalous layer which represents the high-velocity zone at the base of the lower crust had a density of 3130 kg mP3 and the upper mantle a density of 3300 kg m-3 (Fig. 14). Once the gravity effect of the water, sediments and upper-mid-crust layers were removed, we defined an initial crust-mantle boundary geometry by smoothly joining the base of the crust defined at the ESPs. We did not introduce the 7.2 km s - ' anomalous layer in a first trial; the Moho reflector was defined as the line joining the Moho depth at ESP 202 to the top of the 7.2 km s-' layer of the ESP 203 (Fig. 14a). A second modelling was then performed according to the seismic model (Fig. 14b). The bottom of the crust was, if necessary, modified to match the observed free-air anomaly, and the geometry of the transition zone between ESP 203 and ESP 204. Since the morphology and thickness of the sediments are well known from seismic data, it is not to be expected that gravity will contribute to a better definition of their geometry. Although gravity can contribute in better defining their densities, we take their geometry and density from seismic results. Therefore, the main contribution of the gravity modelling has been to allow us to infer the importance of the 7.2 km s-' anomalous layer at the base of the crust. Figure 14 shows the model that better fits the gravity and seismic data. It is clear that the second modelling (Fig. 14b) fits the observed free-air anomaly quite well. However, the ECORS profile crosses a particular anomaly which is limited in its lateral extent, as demonstrated in Fig. 13 east and west of the ECORS line, the modelling results are certainly different, as demonstrated in the lower panel of the Fig. 13. As the crustal velocities are not well determined for the two lines located east and west of the ECORS seismic line, we are not able to model the gravity anomalies. DISCUSSION Representative velocity-depth profiles, extracted from the models, appear in Fig. 15 and in Table 1 to illustrate the changing crustal structure from continent (ESP 201) to the deep basin (ESP 215). The depth section of the ECORS NW seismic profile is also shown for comparison (Fig. 16). On the MCS line a striking feature is the presence of a 5.3 km s-' layer which appears to represent sediments just above basement and must be related to a post-rift or a syn-rift sequence. In the previous interpretation (Le Douaran et al. 1984) this layer was included as pre-rift sedimentary basement. The velocity profiles between the basement and the Moho present a variety of forms. Below the so-called tilted blocks of the slope (ESPs 201 and 202) continental crust is undoubtedly present. A transition zone appears between ESP 203 and 204 where crustal velocities change laterally from 6.2 to 6.6 km s - ' . The ECORS NW Crustal structure of the Gulf of Lion 711 RAY PARAMETER (s/krn) - v ) ' Y F k a Figure 4(d). Comparison between ESP 205 7-p section (upper panel) and synthetic r-p seismograms (lower panel) computed from the final velocity model. Note that the gains for the data and the synthetics are not the same therefore the amplitudes are not matched in absolute terms. seismic line shows that a large depression in the basement topography is present between ESP 204 and 205. In this depression a chaotic layer could be interpreted as syn-rift series (Burrus 1984), which may overlie a thinned continental crust. We observe a slight decrease in depth of the acoustic basement from ESP205 to 207. The last puzzling feature is the fact that the lower crust below the continental slope consists of a 2-3 km thick high-velocity layer (7.2 to 7.4 km s-') with a low velocity gradient. The reflector (reflector S as in the west Galicia margin?) at the top of this layer is very prominent on the ECORS NW and NS profiles and on other industrial seismic lines (LIGO surveys for example). We will discuss later its origin. Figure 17 shows the model proposed by Le Douaran et al. (1984) and the new one from the present study. The differences are quite important, especially in the crustal layers and around the previously postulated OCB. They can be summarized as follows: (1) no large depth discrepancies are observed in the Figure 5(a). Upper panel: observed record section of ESP206. Lower panel: calculated traveltimes curves. Upper right inset is the location map for ESP 206. Cross marks ESPs midpoint location. The main arrivals are labelled: Rsf: sea-floor reflection; T1: rays turned in the 4.3 km s-' layer; T2:rays turned in the 4.7 km s-' layer; T3 corresponds to rays travelling through the 5.1 km s-' layer; Tuc: rays turned in the upper crust. sedimentary layers for velocities up to 5.1 km s-'. The geometry of the so-called tilted blocks was extracted from the refraction data by Le Douaran er al. (1984); in our interpretation the depths and slopes were measured from the ECORS-reflection seismic line. (2) The 5.3 kms-' layer exhibits much difference in the two models and it follows that its interpretation is different. In the new model this layer is shallower and extends up to ESP 207; consequently it is related to a post-rift sequence, partly to syn-rift series. It was previously proposed that this layer may represent ante-rift sediments included in the acoustic basement. (3) The palaeo-oceanic ridge near ESP 208, as proposed by Le Douaran et al. (1984), is not supported by the ECORS-CROP transect (de Voogt et al. 1991). The velocity/depth solution of ESP 215 infers the same conclusion. (4) The thick 6.2 km s-l granitic layer is replaced by a 6.7 km sfl layer below ESP 204 in our interpretation. This occurs in the region where the original crust is now thinnest. ( 5 ) The last main difference between the two models is the presence of a high-velocity layer in the lower crust, whose thickness is greatest between ESPs 203 and 204. The top of this layer corresponds to a dipping reflector clearly seen on the multichannel seismic lines, similar to those observed on the Canadian margin by Keen & de Voogt (1988) which are coincident with the high velocity zone of Reid & Keen (1990). According to the free-air gravity map (Fig. 13) the ECORS profile crosses a feature which is 712 G. P. Pascal, A. Mauffret and P. Patriat Figure 6(a). Upper panel: observed record section of ESP207. Lower panel: calculated traveltimes curves. Upper right inset is the location map for ESP207. Cross marks ESPs midpoint location. The main arrivals are labelled: Rsf: sea-floor reflection; T1: rays turned in the 4.3 km s-' layer, T2: rays turned in the 4.7 km s layer; Tuc: rays turned in the upper crust; Tmc: rays turned in the mid-crust. ESP 2 0 6 Figure 5(b). Upper panel shows data from ESP 206 transformed to the domain of intercept time ( r ) and ray parameter (p). The input data were 50m equi-spaced seismograms from 0 to 60km. The stacking was performed for ray parameters from 0.02 to 0.67 s km I at 0.002 s km ' increment. Lower panel shows the r-p traveltimes corresponding to the final velocity model. The labels correspond to those in Fig. 5(a). RAY PARAMETER (S/KM) w Figure S(c). Synthetic 7-p seismogram for ESP 206 computed from the final velocity model. Note the S waves between 2 and 4s seen on both observed and synthetic 7-p sections. relatively isolated in the Gulf of Lion, although it appears an important one on the seismic profile. The prominent reflector, which is placed at the top of the high-velocity layer (ESP 203) is clearly seen between 4" and 6"E on unpublished industrial lines shot parallel to the E C O R S profile and is never seen on seismic profiles in the Ligurian Sea. A second objective of this work was to investigate the extent of thinned continental crust and, if possible, to define the location of the OCB (ocean-continent boundary). Thickness of the crust is not a discriminant of crustal type, just as the seismic velocities. According to White (1984) high velocity gradients (0.4-0.6 s p l ) in the lower crust may be indicative of uniformly stretched continental crust; lowvelocity gradients are typical of either oceanic crust or of stretched continental crust that has been heavily intruded by igneous rocks. So one method of differentiating between oceanic and thinned continental crust devoid of igneous intrusion appears to be the nature of the lower crust. White (1984) gathered North Atlantic velocity structures and found that, for crust over SOMyr old, they fall within a quite narrow range. Whitmarsh, Miles & Mauffret (1990) made this comparison with success on the continental margin of Iberia. In our case the oceanic crust is younger, as the creation of the ProvenGal Basin occurred during the Miocene, around 19-21 Myr. The comparison between ESPs 215, 216, and possibly 207 and the model (Fig. 18) is very conclusive and we are able to confirm that these seismic sections are located o n an oceanic area. However, we cannot discriminate easily for ESPs 204 to 206. The O C B could be placed at three locations: Crustal structure of the Gulf of Lion RAY PARAMETER (s/krn) 713 OFFSET ( k m ) YF ESP 2 0 7 Figure 6(b). Upper panel shows data from ESP 207 transformed to the domain of intercept time (7)and ray parameter (p). The input data were 50m equi-spaced seismograms from 0 to 60km. The stacking was performed for ray parameters from 0.02 to 0.67 s km-' at 0.002 s km ' increment. Lower panel shows the T - p traveltimes corresponding to the final velocity model. The labels correspond to those in Fig. 6(a). (1) between ESPs 203 and 204. W e observe a drastic lateral change in seismic velocities of the crust and a very thin crust. With these arguments it appears that a reasonable OCB could be placed between these ESPs. In this area probably the mantle was unroofed and the 7.2-7.4 k m s - ' layer corresponds either to serpentinization of the peridotites (Boillot ef al. 1989) or t o a characteristic underplating. Such landward dipping reflections have been observed at the presumed O C B of the Atlantic Canadian margin with a true dip of 28" (Keen & d e Voogd 1988). T h e reflector identified beneath the Gulf of Lion has the same disposition but the true dip is only of 10". However, the position of the Canadian O C B is controversial and could be placed much further seaward (Tucholke, Austin & Uchupi 1990). An argument contrary to this first O C B position is based on the velocity-depth sections of ESPs 204 t o 206 which are not characteristic of typical oceanic crust although a wedge of high velocities (7.4 km can be found within the oceanic crust near the continental margin (Reid & Keen 1990). ( 2 ) Between ESPs 205 and 206 as postulated by Le S C ' ) Figure 7. Upper panel: observed record section of ESP 212. Lower panel: calculated traveltimes curves. Upper right inset is the location map for ESP 212. Cross marks ESPs midpoint location. The main arrivals are labelled: Rsf; sea-floor reflection; T1: rays turned in the 4.3 km s-' layer; T2: rays turned in the 4.7 km s ' layer; TIC: rays turned in the lower crust. Douaran et al. (1984) and Burrus (1984) which placed the O C B at this position because they observed an important change in velocity of the basement and the possible presence of syn-rift sediments landward of E S P 2 0 5 The E C O R S seismic lines as well as the industrial lines (Fig. 16) confirm the presence, landward of ESP205, of a huge depression which may b e filled by syn-rift sediments (de Voogd 1991); however, the present study does not show any significant difference in the velocity structure from ESP 204 to 206. (3) From ESP207. This zone corresponds to very prominent salt domes and large magnetic anomalies which contrast with the landward quiet zone (Le Cann 1987; Mauffret 1976). T h e velocity structures of ESPs 207,215 and 216 show a n 'oceanic' thickness and the velocity gradients in the crust are very similar to those of typical oceanic crust, although the thickness of layer 3 appears thinner than normal. This observation is not surprising; Duschenes, Sinha & Louden (1986) found such results in the Tyrrhenian Sea. In this interpretation, the O C B coincides with a belt of large salt domes (Le Cann 1987). This coincidence is puzzling because the formation of the salt domes can be mostly attributed t o the basinward flow of the salt in response to the Pliocene-Quaternary loading of the Rhone deep-sea fan. T h e structure below the salt domes cannot be clearly interpreted o n the seismic records which are highly disturbed by artefacts caused by the salt diapirism, however, 714 G. P. Pascal, A. Mauffret and P. Patriat Figure 8(n). Upper panel: observed section of ESP 211. Lower panel: calculated traveltimes curves. Upper right inset is the location map for ESP 211. Cross marks ESPs midpoint location. The main arrivals are labelled: R s t sea-floor reflection; T1: rays turned in the 4.4 km s-' layer; '12: rays turned in the 4.7 km s-' layer, Tuc: rays turned in the upper crust; Tmc: rays turned in the mid-crust; PmP: Moho reflection critical distance. there is some weak evidence (de Voogd et al. 1991) that the large salt dome belt is underlain by a depression of the basement. No simple ridge and transform geometry is observed on ECORS and CROP profiles; in a similar way magnetic anomalies do not show well developed sea-floor spreading magnetic pattern. Now we examine briefly some aspects of the reconstruction of this part of the Mediterranean Sea. The extension may have been accomplished by pure shear stretching (McKenzie 1978), by intrusion of mantle material into the crust (Royden & Keen 1980) or by simple shear failure of the lithosphere (Wernicke 1985). The developmental history of this part of the Mediterranean Sea is still a matter of current contention. Numerous geophysical data have been collected in the Gulf of Lion for many years and some models of evolution have been proposed. We proposed only to delineate the history of the Western Mediterranean Sea (Valencia Trough, Gulf of Lion, Ligurian Sea) through its structure, as quantified by the last experiments. A two-stage history, which can explain most of the observed data, can be postulated if we extend the drifting history to the Valencia Trough. In the first stage, continental breaking started in the Valencia Trough and in the Gulf of Lion at the same time. Extension leads to an extended crust in the former domain with localized formation of oceanic crust northeast of the Valencia Trough (Torne et al. 1992; Pascal et al. 1992). In the Gulf of Lion a detachment process leads to the unroofed E S P 211 Figure 8(b). Upper panel shows data from ESP 21 1 transformed to the domain of intercept time (T) and ray parameter @). The input data were 50m equi-spaced seismograms from 0 to 60km. The stacking was performed for ray parameters from 0.02 to 0.67 s km-' at 0.002 s km-' increment. Lower panel shows the r-p traveltimes corresponding to the final velocity model. The labels correspond to those in Fig. 8(a). mantle feature observed along the ECORS profile. The ridge (Fig. 19) as observed on seismic sections of the industrial lines is correlated with the negative free-air and positive magnetic anomalies located south of ESPs 211 and 212 and could correspond to this episode of aborted rifting. At that time Balearic and Sardinia islands moved to the south to form a single line. In the second stage, sea-floor spreading started in the Ligurian Sea and the Gulf of Lion while tectonic activity ceased in the Valencia Trough. Sardinia and Corsica islands moved to the southeast as observed. CONCLUSION Eleven ESPs and a parallel multichannel seismic section provide detailed information on the crustal structure of the Gulf of Lion. The main P-wave reflected and refracted phases were modelled both in the x-r and 7 - p domains and for traveltime and amplitude. The resulting P-wave velocity-depth model has the following features: (a) beneath the continental slope of the ProvenGal margin a rapid decrease of the Moho depth from 20 km to 10 km and the Crustal structure of the Gulfof Lion 715 (QOFFSET ( k m ) 30 3 4,O 30 4,o h cn Y I- 1 1 1 RAY PARAMETER (S/KM) 3 3 - P 0 r2 0 w 0 P 0 0 ul m existence of an anomalous 7.2 km s-' layer, (b) a 5-6 km thin extended crust from the base of the slope, (c) a quite 'typical' oceanic crust up to the Sardinian margin. Gravity modelling confirms the seismic model and particularly the high-velocity layer. The ocean-continent transition could be located at a place where the crust is thinnest and where the continental 6.2 km s-' layer abuts on 6.6 km s-' velocity layer, or at the boundary between the 5-6 km thin extended crust and the typical oceanic crust. Though our foregoing analysis is only limited to seismic and gravity modelling and to a brief examination of the Figure 8(c). Upper panel: comparison between observed record section of ESP 211 and reflectivity seismogram for the final velocity model presented in Table 1. The amplitudes of the both sections are scaled proportional to the distance. In the synthetic section all the phases are computed. Consequently some phases are not present in the observed section due to poor signal-to-noise ratio or strong inferference between phases. Lower panel shows the synthetic r-p seismogram for ESP 211 computed from the final velocity model. See text for explanations of the computation. magnetic anomalies, it appears several similarities between the structure of the Gulf of Lion and those of the back-arc basins which also suggest similar origins. So we propose a two stage mechanism of rifting and drifting of the Western Mediterranean Sea. The first one, occurred in the Gulf of Lion and Valencia Trough which underwent rifting processes, the second one gave way to the drifting of the Corsica-Sardinia plate. Careful and detailed analysis of other geophysical data will still be necessary in further constraining the evolution of this part of the Mediterranean Sea. 716 G. P. Pascal, A . MauJfret and P. Patriat Figure 9(a). Upper panel: observed record section of ESP215. Lower panel: calculated traveltimes curves. Upper right inset is the location map for ESP 215. Cross marks ESPs midpoint location. The main arrivals are labelled: Rsf: sea-floor reflection; TI: rays turned in the 4.3 k m s I layer; T2: rays turned in the 4.7 k m s - l layet; T3 corresponds to rays travelling through the 5.1 km s layer: Tmc: rays turned in the mid crust; TIC: rays turned in the lowcr crust: f r n f critical: Moho reflection critical distance. Figure 9(b). Upper panel shows data from ESP 215 transformed to the domain of intercept time ( r ) and ray parameter (p).The input data were 50m equi-spaced seismograms from 0 to 60km. The stacking was performed for ray parameters from 0.02 to 0.67 s km-' at 0.002 s km-' increment. Lower panel shows the r - p traveltimes corresponding to the final velocity model. The labels correspond to those in Fig. 9(a); LVZ: low velocity zone in the salt layers. Crustal structure of the Gulf of Lion 717 OFFSET ( k m ) Figure 10. Upper panel: observed record section of ESP 216. Lower panel: calculated traveltimes curves. Upper right inset is the location map for ESP 216. Cross marks ESPs midpoint location. The main arrivals are labelled: Rsf: sea-floor reflection; T1: rays turned in the 4.3 km s ' layer; T2: rays turned in the 4.7 k m s - ' layer; Tmc: rays turned in the mid crust; TIC: rays turned in the lower crust; f r n p critical: Moho reflection critical distance. Figure ll(a). Upper panel: observed record section of ESP 201. Upper right inset is the location map for ESP201. Cross marks ESPs midpoint location. Lower panel: enlargement of the x - i section from 30 km range, corresponding to crustal arrivals. Superimposed are traveltime interpretations of these phases. OFFSET (km) OFFSET (km) 40 50 I 60 I 70 40 5 5 7 7 - 9 5? 70 9 - h ln ln Y I- I- ll ll 13 13 15 15 Figure ll(b). Traveltimc curves corresponding to the final velocity model of ESP 201 and final best fitting of amplitudes of the last crustal arrivals (see text for explanations). 718 G. P. Pascal, A . Mauffret and P. Patriat o(4 OFFSET ( k m ) 20 10 30 50 40 Figure 12(a). Upper panel: observed record section of ESP202. Lower panel: calculated traveltimes curves. Upper right inset is the location map for ESP 202. Cross marks ESPs midpoint location. OFFSET (km) 20 5 30 I 40 I OFFSET (km) 50 20 5 7 30 I 40 I 7 9 9 h v) v I- 11 ll 13 13 15 15 Figure 12(b). Traveltime curves corresponding to the final velocity model of ESP 202 and final best fitting of amplitudes of the some crustal arrivals (see text for explanations). Crustal structure of the Gulfof Lion 1 , , , . I .,.. I , . . . I . . . .I , . , , I , , 5E : . , , ! . , , . , , . . , , . I,. ,,,. .I, 719 JOkm iOE Figure 13. Upper panel: Free-air gravity anomaly map at 5mGal interval (data from the Bureau Gravimetrique International). The central dashed line corresponds to the ECORS N W seismic section. The two other dashed lines are parallel to the first one and cross the Gulf of Lion east and west of the ECORS profile. Lower panel: free-air gravity anomaly profiles along the three dashed lines. From the left of the figure note that the two first lines are not too different in shape but in the amplitudes of the anomalies. The third line is very different and could indicate another crustal structure. 201 202 203 204 205 206 207 215 Figure 14. (a) Comparison of the observed free-air gravity anomaly on N W ECORS profile to calculated profile based on seismically constrained structure from ESPs201 to 215. See text for explanations. Densities are given in kg m-3. (b) Comparison of the observed free-air gravity anomaly on NW ECORS profile to calculated profile based on seismically constrained structure from ESPs 201 to 215. See text for explanations. In the lower part of (b) crustal structure of the Gulf of Lion deduced from seismic and gravity data. 720 G. P. Pascal, A . Mauffret and P. Patriat 0 54 L 6 5 - 3 10- 4 1 tI + I h 10 15- n - __ VELOCITY I 15 ESP202 (KM/S) ESP 203 ‘I, VELOCITY (KM/S) VELOCITY (KM/S) VELOCITY (KM/S) VELOCITY (KM/S) 0 - 5 H c 1 fj 10 0 15 I I ESP 212 VELOCITY (KM/S) 2 - 4 6 5- z E h 10n 1 5 ~ VELOCITY VLLOCllY (KMIS) (KM/S) 0 E 5 2 4 VELOCITY (KM/S) B 6 i r’ kX fj 10 0 I 4 15 ESP 216 I VCLOCITY (KM/S) VELOCITY (KM/S) Figure 15. The velocity-depth function results of the Gulf of Lion expanding spread profiles. For the location map of the ESPs see Fig. 1. Crustal structure of the Gulf of Lion Figure 16. Depth section of the ECORS NW (a and b) and Ligo 4 (c). For the location of these seismic lines see Fig. 1. The results of the ESP are superimposed on the sections. The correlation between these ESP 'logs' and the seismic section is generally excellent; i.e. the acoustic basement is placed at the base of the 5.3 km sC' layer, however, the prominent reflector ( T ? ) seen on Fig. 19 corresponds to the top of the 7.3kms layer below ESP 203 whereas it is correlated to the Moho (8.4 km sC1) below ESP 202. Note the large basement depression between ESP 204 and 205 where syn-rift sediments may have been accumulated (Burrus 1984). The volume of salt in the large domes [right side of b] is overestimated because we obtain a pull-down in-depth section although a pull-up is observed on the time section. ' 721 Figure 17. Upper panel (section a): Crustal section of the Gulf of Lion, after Le Douaran et al. (1984). Lower panel (section b): Crustal section from this study. Numbers in the different layers indicate P velocities. Locations of the ESPs are indicated by the common central point. 722 I VELOCITY (KWS) 1 VELOCITY (KMIS) ESP207 10' VELOCITY (KMIS) 1 ESP215 - ESP216 10' VELOCITY (KM/S) E%gure18. Interpretation of some velocity-depth profiles from Fig. 15 with respect to reference oceanic crust and thinned continentalcrust models based on synthetic seismogram modelling. Dots indicate the envelope of North Atlantic models for oceanic crust over 50Myr old (White 1W). Bold line in the upper right of each figure is a model for thinned continental crust (with sediments omitted) from North Biscay (Whitmarsh et ul. 1986). EsPs structures are relative to acoustic basement. Crustal structure of the Gulfof Lion Figure 19. Depth to moho and horizon T (solid line). Contour interval: 1. On-land contours after Sapin & Hirn (1974). Thin line: bathymetric contour, contour interval: 1 km. Cross and solid large number: position and number of the ESP with depth of horizon T ; first small number and depth of the Moho, second small number. The bright reflector of the ECORS NW line (Figs 16 and 20a) corresponds to the first number (12.5 km, horizon T, below ESP 203) whereas it has to be correlated to the Moho (19.8km) below ESP 202. 723 G. P. Pascal, A. Maufret and P. Patriat 724 D -. 0 *, . 4 , 0 0 4 8 0 I Figure 20(a). Migrated ECORS profile between ESPs 201 and 203. Superimposed velocity-depth function derived for these ESPs. ESP 204 ESP 205 ESP 206 Figure 20(b). Unmigrated ECORS profile between ESPs 204 and 206. Superimposed velocity-depth function derived for these ESPs. Crustal structure of the Gulf of Lion ACKNOWLEDGMENTS We thank SNEA(P) for providing the ESP’s magnetic tapes and our IFP colleagues for their assistance in processing the data. J. L. 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