Download The ocean-continent boundary in the Gulf of Lion from analysis of

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. Olivet provided an unpublished structural map
and helped us with this manuscript by providing helpful
suggestions. The authors are grateful to P. Stoolweg who
made the gravity data reduction.
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