A deep seismic sounding investigation of lithospheric heterogeneity and anisotropy beneath the Iberian Peninsula

35

Tectonophysics, 221(1993) 35-51 Elsevier Science Publishers B.V., Amsterdam

A deep seismic sounding investigation of lithospheric heterogeneity and anisotropy beneath the Iberian Peninsula ILIHA DSS Group * Reporters: J. Di’az, J. Gallart, D. Cbrdoba, L. Senos, L. Matias, E. Suririach, A. Him and P. Maguire

(Received July 14, 1991; revised version accepted August 24, 1991)

ABSTRACT ILIHA DSS Group, 1993. A deep seismic sounding investigation of lithospheric heterogeneity and anisotropy beneath the Iberian Peninsula. In: J. Badal, J. Gallart and H. Paulssen (Editors), Seismic Studies of the Iberian Peninsula. Tectonophysics,

221: 35-51.

The dimensions of the Iberian Peninsula, the facility of firing large charges in the surrounding waters, the well-known and relatively uniform geology, and prior knowledge of the crustal structure, resulted in it being chosen as a study area for the investigation of the seismic structure of the lower lithosphere via refraction-wide-angle reflection seismic profiling. The Iberian Lithosphere Heterogeneity and Anisotropy experiment (ILIHA), with a star-shaped arrangement of six long-range DSS profiles, was carried out in October 1989. The models derived from a first interpretation of the recorded data are presented. Three crustal profiles cover the same western and central part of the Hercynian Massif as the mantle profiles. The resulting interpretations all include a middle as well as a lower crustal layer above the mantle. The velocities of the layers in all three models are similar; however, the layer depths vary beneath the profiles. Interpretation of the mantle derived data suggests a layered lower lithosphere. One reversed line and an intersecting unreversed line indicate the layering penetrates to at least 90 km depth. The homogeneity of these layers contrasts strongly with the heterogeneous Hercynian surface geology. Velocities derived from reflected data from the deep layers suggest the constituent materials are either anisotropic or that the layers suffer a slight regional dip.

The ILIHA Project, DSS experiment

Correspondence to: Dr. J. Diaz, now at lnstituto de Ciencias de la Tierra “Jaume Almera”, C.S.I.C., Marti i Franquts s/n, 08028 Barcelona, Spain. * J. Diaz Cusi, A. Hirn and A. Nercessian (IPG, Paris); D. Cordoba, J. Tellez and A. Udias (Univ. Complutense, Madrid); J. Gallart (CSIC Jaume Almera, Barcelonak L. Senos (INMG, Lisboa); L. Matias and L.A. Mendes Victor (Univ. Lisboa); A. River0 and A. Lopez Arroyo (IGN, Madrid); E. Suriiiach (Univ. Barcelona); A. Rota and S. Figueras (Servei Geologic de Catalunya); M. Catalan (I.H.M. San Fernando); B. Jacob (DIAS, Dublin); J. Ansorge, St. Mueller and R. Freeman (ETH, Zurich); P. Maguire (Univ. Leicester); C. Prodehl and F. Hauser (Univ. Karlsruhe); M. Demartin (IGL, Milano); H. Thybo (Univ. Kobenhavn).

0040-1951/93/$06.00

The ILIHA (Iberian Lithosphere Heterogeneity and Anisotropy) Project was designed to be a large-scale test for vertical and lateral heterogeneity and seismic anisotropy in the subcrustal lithosphere over the Hercynian domain of the Iberian Peninsula. The upper mantle, over a depth range from 30 to 100 km, has been investigated by two methods: (1) analysis of earthquake surface and body waves recorded by the broadband stations of the permanent Spanish and Portuguese observatories, and also the portable NARS array (Paulssen, 1990); and (2) a large-scale Deep Seismic Sounding (DSS) experiment in which six profiles along different

0 1993 - Elsevier Science Publishers B.V. All rights reserved

36

azimuths sampled the lower lithosphere over a central region in the Hercynian core of Iberia. First results of this latter survey will be presented here. The conventional geometry of refraction seismic experiments normally provides at best a twodimensional velocity-depth model. To ascertain the significance of variations in such models with respect to lateral heterogeneity it is necessary to implement an intersecting array of profiles over the study region. Patterns of crossing profiles are essential for studies of the mantle as there is evidence that seismic anisotropy may be important in the lower lithosphere. Anisotropy is a key to understanding the continental lithosphere as it may be related to its mode of formation, to its internal deformation, its present stress pattern, and/or the relative or absolute plate motion. Separating the two effects of lateral heterogeneity and anisotropy is difficult and requires a large data set. Also strongly heterogeneous crustal segments are to be avoided as their influence would probably dominate the wave propagation, making

lLIHA

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it more difficult, and perhaps impossible, to extract the more subtle signature of anisotropy. The Hercynian domain of Iberia provides one of the best opportunities in Europe to study the lateral heterogeneity and the anisotropic properties of continental lithosphere in a large, old and globally stable block. The dimensions of the Iberian Peninsula are ideal for the design of long-range profiles along different azimuths. The surrounding waters enable the firing of high efficiency sea-shots, while crossing profiles can be designed so that the turning points for rays penetrating the mantle occur under a common region in the centre of the homogeneous Hercynian core. The experiment was designed to avoid, where possible, complex lithospheric structure beneath the margins of Iberia which has suffered a number of different tectonic events: Late TriassicLate Jurassic rifting to the west, Eocene colhsion with Europe to the north, Neogene rifting of the Valencia Trough to the east, and the Miocene Betic orogen and Neogene Alboran Sea extension to the south.

Fig. 1. Geotectonic map of the Iberian Peninsula displaying ILIHA DSS experiment lines and shotpoints. 1 = Hercynian basement; 2 = Deformed Mesozoic cover of the Pyrenean realm; 3 = Mesozoic aulocogenic area; 4 = Mesozoic external zones of the Betic realm; 5 = Internal units (Mesozoic and Palaeozoic) of the Betic realm; 6 = Intermediate Flysch units of the Gibraltar arc; 7 = Mesozoic of the African margin; 8 = Undeformed Mesozoic cover; 9 = Tertiary basins; 10 = Badajoz-Ckdoba and South Iberian shear zones.

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The array of four reversed and two unreversed, long range profiles undertaken during ILIHA is shown in Figure 1. The dimensions of the Peninsula are such that the profiles provide observations of waves travelling to a nominal distance of about 800 km and likely to penetrate to more than 150 km below the surface. The Spanish and Portuguese navies were in charge of the sea shooting. Shots of 1000 kg were fired using dispersed charges (Jacob, 1975) each being divided into ten packets of 100 kg, distributed along a 250 m line at a nominal depth of about 90 m. However, at some shotpoints total charges of between 500 and 900 kg only were detonated. Ten shots were fired successfully, and 3 failed due to technical problems or bad weather conditions. Unfortunately, the two planned shots off the north coast of Spain (Fig. 1, shotpoint A) had to be cancelled just before the field work started. Five land-shots, from large quarry blasts at points G and P (Fig. l), were also recorded out to distances of 250-300 km, thus improving the crustal control along the lines. The data have been digitized at five processing centres and a final complete data base in SEG-Y format will be available in the near future. In this presentation only part of the data are shown (from profiles B-X, B-G, C-F, and D-A) together with a preliminary interpretation. This will be developed in more detail at a later date once the complete data set is available.

Samples

of Hercynian

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37

PENINSULA

Shot B-2, profile B-X Shot B-2 was fired 35 km offshore Viana do Castelo and recorded along the NW-SE profile B-X (Fig. 2a). Between 40 and 100 km the first arrivals display an apparent velocity slightly greater than 6.0 km. s-l, with a near zero delay time. They are interpreted as resulting from the Pg-wave through the outcropping Hercynian basement, which is characterized by velocities increasing up to 6.1 km. SC’.

DISTANCE (km) “RED i 6.0 km/s

crustal structure

All the sea-shots have been recorded to at least 250 km enabling the derivation of one- and two-dimensional crustal models. The Hercynian crust is well sampled from sea-shots at B, C and D. In addition, the quarry blasts from land-shotpoints G and P, in fairly central positions, provided complementary data. In this section, data will be presented from sea-shots at B (B-2) on profile B-X, and C (C-l and C-2) on profile C-F, and a land-shot at G (G-2) on profile B-E (Fig. 1) which all variously sample the Hercynian crust.

Fig.

2. Profile

Record

section.

B-X.

Crustal

(b) Synthetic

arrivals

from

seismograms.

diagram.

shot

B-2.

(a)

(c) Ray-tracing

38

Three later phases can be correlated in the record section. The first, P*, is seen beyond a distance of 50 km with increased amplitude after 70 km. It is attributed to a reflection from a discontinuity at the base of the upper crust at a depth of 14 km, where the velocity increases to 6.3 km. s-’ (Fig. 2). The second reflected phase, P,P, can be correlated from 80 km to the end of the record section, with high amplitudes at around 120 km. It corresponds in the model to waves reflected at the top of the lower crust. This is located at 23 km depth and is characterized by an increasing velocity gradient from 6.7 km * s-l at the top to 6.9 km . s-l at the Moho. There is a weak wide-angle PmP reflection from the Moho which can be correlated between 100 and 190 km distance, and which indicates the crust/mantle boundary lies at 31 km depth. A high-gradient zone of about 2 km thick has to be introduced on top of the Moho discontinuity to fit the observed PmP amplitudes.The Pn phase, refracted in the upper mantle, is well defined from this high-energy sea-shot, and displays an apparent velocity of about 8.0 km . s- ’ (see also Fig. 5a). In an experiment carried out previously in Galicia (Cordoba et al., 1987, 1988), one shot was situated very close to B-2 and was recorded on a SW-NE profile (shot A on the profile La Guardia-Ribadeo, in: Cordoba et al., 1987). This profile showed a similar upper crust, but with no clear indication of a discontinuity at about 15 km depth. A conspicuous PmP phase was correlated beyond 70 km towards the northeast, in contrast to the weak arrivals observed on the present profile beyond 100 km towards the southeast. The Galicia profile Pn phase displayed an apparent velocity of 8.1 km * s-l to the northeast. Shots C-l-C-2, profile C-F The seismic image of a well defined threelayered Hercynian crust, obtained from the northwestern part of profile B-X, can also be inferred from shots C-l and C-2 recorded on profile C-F. Analogous seismic phases can be correlated in the record section (Fig. 3a). For this profile the first arrivals show an increasing re-

ILIHA

DISTANCE VWD

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(km)

i 6.0 km/s

10

DISTANCE

Fig. 3. Profile (a) Record

C-F.

section.

Crustal

arrivals

(b) Synthetic

,lrm,

from shots C-l and C-2.

seismograms.

(c) Ray-tracing

diagram.

duced time delay of up to 1 s to a distance of about 40 km. This delay results from ray-paths crossing the Neogene sedimentary cover present in the westernmost part of the profile (Fig. 1). The Pg phase is correlated from 40 to 100 km with an apparent velocity of about 6.0 km. s-l. Examination of P* (Fig. 3a) results in a discontinuity at the base of the upper crust at a depth of 12 km, with the mid-crustal layer having a velocity of 6.3 km. s-l. A lower crustal layer, at 21 km depth under profile C-F, has an average velocity of 6.8 km . s-l. The reflected phase from

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shotpoint C is 28 km. A similar thin crust has already been reported for the Tagus Valley (Mendes-Victor et al., 1980). To match the crustal thickness of 30 km about 150 km inland, as reported by Caetano (19831, and of 31 km at 250 km, derived from a profile recorded from shotpoint G towards the west-southwest (Banda et al., 19811, a gentle eastward downdip of the Moho is introduced in the model (Fig. 3bl. To be consistent with the apparent Pn velocity of 8.0 km . s-l, the true velocity in the mantle is then taken as 8.1 km. s-l. Shot

40h-1m

THE IBERIAN

150

170

190

1

210

k

230

OlSTANCE CL.1

Fig. 4. Profile B-G. Crustal arrivals from shot G-2. (a) Record section. (b) Synthetic seismograms. (c) Ray-tracing diagram.

its top, P,P, is less apparent in this record section than in those from shots B-2 and G-2 (see Figs. 2a, 4a). This effect may be related to the lower frequency content of shots C-l and C-2. Such frequency dependence was illustrated by the Grupo de Trabajo de Perfiles sismicos profundos (1983) from three sea-shots, slightly shifted with respect to each other, which were recorded along the same profile on the southern Spanish Meseta. In contrast to the phase from shot B-2, PmP from shotpoint C is a dominant phase with a critical distance of about 85 km. The Moho depth under

G-2, profile G-B

Records from the quarry blast at G (Fig. 11, located near Toledo and recorded to the northwest towards shotpoint B (Fig. 4a), display features which are comparable to those from seashots B-2, C-l and C-2 discussed above. First arrivals between 5 and 40 km have an apparent velocity of about 5 km . s- ’ across recording stations located in the Neogene Tagus Basin. A recording gap between 60 and 80 km due to instrument failure hinders a complete study of the Pg phase between 40 and 60 km observed at 1.2 s reduced time and at 0.8 s after 80 km. However, such a change in delay time of 0.4 s has been documented in a previous profile from the same shotpoint but oriented slightly further northeast (Surifiach and Vegas, 1988). It was interpreted as resulting from the transition from the Tagus Basin to the Hercynian basement of the Central System (Fig. 1). This geological contact is located along profile G-B at 80 km from shotpoint G. A reflected phase from the upper crust, P*, occurring about 0.6-0.3 s after the Pg arrivals is observed to a distance of 140 km. It can be interpreted as resulting from a velocity change from 6.0 to 6.3 km. se1 at 15 km depth. An alternative interpretation, introducing a low velocity layer of 5.6 km . s- ’ between 7 and 11 km depth, can also be proposed, consistent with previous models for the Central Meseta (Banda et al., 1981; Surifiach and Vegas, 1988). The difference between these two models may not be resolved from the amplitude of the present data.

40

Strong reflections originating from the top and bottom of the lower crust are well defined after distances of 120 km and 100 km, respectively. The observed amplitude pattern of the PmP phase is consistent with that on the nearby profile, its critical distance being not less than 85 km. The PmP arrival times indicate that the crust under the central part of profile B-E is 34 km thick. This value is also consistent with the Pn traveI times to a distance of 160-220 km. This phase has an apparent velocity of 8.0 km . sp ‘. D~c~ssion of crustal st~cture

The crust under the Hercynian Massif which outcrops in the western half of the Iberian Peninsula has been sampled by refraction and wide-angle reflection data along different segments of the ILIHA profiles. Several common features, consistent with previous seismic models, can be discerned in the three profiles discussed. First arrivals can be separated into two refracted phases, Pg and Pn, well defined for tens of kilometres, which constrain the depth to refractor and velocities at the basement and Moho levels. Three main boundaries separating the upper, middle, and lower crust, and the mantle, are consistently inferred from the associated reflected seismic phases. The veiocity distribution within the crust is similar for all three profiles. In the northwest part of the Hercynian Massif, near shotpoint B, the first discontinuity, where the velocity reaches 6.3 km - s-l, is identified at 14 km depth. Similar results have been obtained from previous profiles sampling areas in GaIicia to the northeast of shotpoint B (Cordoba et al., 1988). Comparable depths are also obtained beneath the central part of the Massif from shotpoint G, as well as from shotpoint C situated southeast of Lisbon. The top of the lower crust, characterized by a velocity increase to 6.7-6.8 km - s-l is well defined from wide-angle reflections and is located at about 21-23 km depth beneath all three profiles. The central part of the Hercynian Massif has a thick crust with respect to that close to the Atlantic margin. The Moho is at 28 and 30 km

ILIHA

DSS GROUP

depth close to shotpoints C and B, respectively, and reaches 34 km under the Central System beneath the profile from shotpoint G. The thickening primarily seems to affect the lower crust. It was attributed by Suriiiach and Vegas (1988) to the Cretaceous-Miocene shear-zone activity of the Central System in which rotations of brittle upper crustal segments, together with ductile deformation and thickening of the deep crust, resulted in elevated topography, contemporaneous with the opening of the Bay of Biscay. The thinner crust under sea-shotpoints B and C is similar to that beneath other parts of the Hercynian range, for example beneath the Armorican-Central Massif of France (Perrier and Ruegg, 1973; Sapin and Prodehl, 1973). It may be associated with the opening of the Atlantic as has previously been suggested (Mueller et al., 1973; Mendes-Victor et al., 1980). The similari~ in velocity/depth structure of the upper 20 km of the crust is notable considering the distances between the three profiles and the different terranes sampled. Beneath shotpoint C is the South Portuguese continental terrane, accreted to Iberia along the Hercynian subduction and shear zone. Beneath shotpoint B is the Central Iberian autochthon, massively intruded by Palaeozoic plutons and later subjected to shearing in the case of the Central System under shotpoint G. However, examined in more detail, there are different crustal responses between the profiles, for example that of the Moho providing extremely clear short range PmP reflections from shotpoint C, and very indistinct ones from B. In contrast, there is a clear PmP arrival from shotpoint B to the northeast (Cordoba et al., 1987). Such variations imply that there may be differences in the nature of the interfaces, which in turn could be related to their different locations within the Hercynian domain. Since there are also significant variations in the Moho response for sea-shots of different spectral content, care must be taken in the interpretation of fine differences in the reflection character. The immediate sub-Moho velocity cannot be better constrained due to the lack of strictly reversed profiles. However, values of 8.0-8.2 km. s-’ are consistent for the range of Pn apparent

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velocities and interface slopes considered in the models. The models derived from the ILIHA crustal data, together with those obtained previously in the same Hercynian tectonic province, have shown that there are no significant lateral inhomogeneities in gross crustal structure. The present study is concerned with the southwest part of the Hercynian Massif beneath which the crustal thickness does not vary significantly except for a marginal increase towards the east near shotpoint G. Although this crustal model has yet to be verified in more detail, the mantle phases discussed below are unlikely to be affected by variations in crustal structure. Thus, any implications concerning lower lithosphere heterogeneity and/or anisotropy, derived from examination of the recorded mantle phases are likely to be valid. However, final consideration concerning the significance of any observed azimuthal variation in mantle velocity and structure should only be made after all the relevant data are included in the interpretation.

BENEATH

THE

41

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Mantle structure A first glance at the record sections shows there is no one continuous single mantle phase suggestive of a homogeneous lower lithosphere. Two or three different phases with apparent velocities higher than the Moho refraction, Pn, relay behind each other from 150 to 700 km distance and are well separated. The energy is carried by each phase only over a limited distance range before jumping to the later phase with a slightly higher apparent velocity. In a sense the energy propagates with a group velocity slower than the individual phase velocities of the successive branches. The simplest type of model accounting for such a picture assumes multipathing in a stratified lower lithosphere. This image has been documented elsewhere in continental Europe, e.g., France (Hirn et al., 1973, 1975; Kind, 1975), the British Isles (Barnford et al., 1976; Faber and Bamford, 1979; Bean and Jacob, 19901, and Scandinavia (Guggisberg et al., 1984; Hauser and Stangl, 1990). The derivation of an appropri-

500

400 DISTANCE

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600

7

(km)

VRED = 8.0 km/s

Fig. 5. Profile

B-X.

Mantle

arrivals

from shot B-2. (a) Record

section.

(b) Synthetic

seismograms

and u(z) model.

9

42

ILIHA

ate velocity-depth structure has been undertaken via forward modelling. This has included trial and error travel time fitting and critical distance and amplitude modelling using synthetic seismograms computed with the reflectivity method (Fuchs and Mueller, 1971). The immediate sub-Moho layer velocity is constrained by the Pn refraction. Direct measurement of the velocity beneath other interfaces is not provided by these data, since head or refracted waves corresponding to prograde branches are not clearly identified. Arrival times, critical distances and relative amplitudes are read from the data and adjusted by the theoretical models; the average velocities in these layers are constrained by the travel time and the velocity contrast at the interfaces is constrained by the relative amplitude of reflected waves. With the characteristics of the survey, numerical simulation provides a 0.1 km. s-l resolution in velocities for a 0.3-0.5 s accuracy in time picks. Precise modelling of the subcritical amplitudes leads, in most cases, to the introduction of strong gradient of around 0.2 km . s-’ over a 2 km depth interval, instead of first order discontinuities between lay-

DSS GROUP

ers. The ILIHA observational scheme provided reversed and intersecting profile coverage enabling the preliminary models to be derived first via 1-D and then 2-D ray-tracing. These seismic models have been derived assuming a flat-Earth model. However, prior to their discussion below, the layer velocities have been corrected to accommodate curvature of the Earth and hence correspond to rock velocities. Profile B-X

Line B-X is the best example of a reversed profile (Figs. 5 and 6). It lies between Marbella in southern Spain (shot X-2) and Viana do Castelo in northwest Portugal (shot B-2). From the south it crosses the western end of the Betic Cordillera, the Guadalquivir sedimentary basin and, after 160 km it enters the Hercynian domain in which it remains to its northwest end (Fig. 1). The total length is 690 km. Shots X-2 and B-2 have been modelled separately using ray-tracing and the reflectivity method. Even though the profile cannot strictly

b)

DISTANCE (km) VRED = 8.0 km/s

Fig. 6. Profile

B-X.

Mantle

arrivals

from shot X-2. (a) Record

section.

(b) Synthetic

seismograms

and c(z) model.

ADEEPSEISMICSOUNDINGlNVESTIGATIONOFTHELlTHOSPHEREBENEATHTHElBERIANPENlNSULA

be considered as reversed because of the different segments of the interfaces that are explored by the two shots, the presence of reflected waves coming from depth and sampling a large volume of the lithosphere allows the production of a 2-D model based on the matching of reflected and diving wave observations. This provides a picture of a lower lithosphere including alternating high and low velocity layers. The presence of low velocity zones (LVZ) is necessary to fit observed travel times, apparent velocities and critical distance estimates. For shot B-2 (Fig. 51, a 29.5 km thick crust with an average velocity of 6.35 km * s-l, derived from the crustal modelling, is assumed. The same crustal average velocity has been used for the model derived from shot X-2 (Fig. 6) with a thickness of 32 km to fit the Pn travel time. The structure is modelled relative to the Moho. Strong Pn arrivals with an apparent velocity of 8.0 km. s-l appear in the B-2 record section at a reduced time of 6 s. In the X-2 record section the Pn arrivals are weaker but still clear. They can be correlated at a reduced time of 6.25 s to a distance of 250 km with an apparent velocity of 8.0 km. s-l. The crust under the first segment of the profile from X-2 to the Guadalquivir basin is governed by the structural complexity of the Betic Cordillera. Late energetic arrivals paralleling the mantle phases are probably multiples due to near source crustal effects. They are present throughout the record section. The next phase, Pi, appears at a critical distance of 250 km and about 1 s after Pn for the records from shot B-2. In an enlarged plot of the records from shot X-2 this phase, with more energy than Pn, appears at a critical distance of 200 km and delayed by 0.75 s. For a model with a constant velocity of 8.0 km. s-l between the Moho and the corresponding mantle interface, the critical distance is much too large when the depth of the sub-Moho reflector is adjusted to provide the correct times on both profiles. A low velocity zone with a velocity of 7.8 km . s-l and thickness of 7 km, must be introduced on top of this intra-mantle interface to fit the travel times, critical distances and also apparent velocities of these reflected phases. These arrivals are then

43

ascribed to a 7.8-8.2 km . s-’ velocity discontinuity at 48 km depth in the model derived from shot X-2 and 49 km from shot B-2. The main difference between the models derived from the two shots is caused by the interpretation of the next mantle phase, P,. This appears with a critical distance of about 300-320 km in both cases. The delay with respect to the previous phase is however significantly larger in the X-2 record section. Assuming a simple planar model this feature forces the introduction of an LVZ, with a velocity of 7.7 km . SC’ into the model. The arrivals are then ascribed to a 7.7-8.45 km * s- ’ velocity discontinuity at a depth of 70 km in the model derived from shot X-2. Were the travel times from shot X-2 to be fitted without an overlying LVZ by increasing the depth of the interface, neither the phase curvature nor the critical distance could then be fitted correctly. For the profile from the shot B-2 on other hand, the same considerations make the introduction of a LVZ unnecessary, and the reflector is found at 66 km depth. The extension of this LVZ from shotpoint X to the northwest along the profile is controlled in the 2-D modelling by the arrivals from shot X-2 which show a low apparent velocity at large distances. This enables the LVZ to extend in the model up to approximately 425 km from shotpoint X. A third mantle phase P3, with a critical distance around 430 km, occurs in the X-2 record section. A similar phase also exists in the B-2 record section with the same critical distance, but less well constrained. In the 1-D interpretation, an 8.0 km. s-' LVZ at around 80 km depth, primarily constrained by the X-2 data, must be included to fit the critical distances and arrival times correctly. These reflections can then be modelled as originating from an interface at about 87/88 km depth with an 8.6 km. s-* velocity layer below. For 2-D modelling, the travel times from shot B-2 can be fitted without this layer because the distant arrivals are already delayed by propagation through the 7.7 km. se1 layer above 70 km in the southern part of the line. In this case, as the critical distance is poorly defined, there is no need to retain the LVZ in the north and modify the reflector depth to fit the travel

44

ILIHA

times. Observations from shot X-2 require the LVZ to be preserved in the southern part of the profile. The resulting 2-D model includes an 8.0 km s- ’ LVZ at 82 km depth, vanishing to the north at a minimum distance of about 380 km from shotpoint X. The extent of the LVZ along the profile is limited only by the zone explored by the B-2 shot. It could extend up to 400 km from X. The finai velocity-depth models and reflectivity synthetic seismograms for the individual profiles are shown in Figures 5b and 6b. The 2-D ray-tracing model is presented in Figure 7.

tion is good up to 570-~0 km, where the signal to noise ratio becomes very poor. This can be attributed to attenuation in the Duero sedimentary basin and to the high cultural noise in that area. The marine shot D-1 was poorly located. This results in the absolute time-distance information being uncertain. The critica point reduced time for Pn on record sections from shots D-2 (close to D-l), C-l and C-2 (about 150 km to the north; Fig. 1) as well as the same time from two previous DSS profiles in southwest Portugal and Spain (Caetano, 1983; Grupo de Trabajo de Perfiles sismicos profundos, 1983) are all at approximately 6 s with respect to an 8.0 km. se1 reduction velocity. For consistency the D-1 data were shifted by 6.0 km with respect to the nominal shotpoint coordinates. Using the same average crustal velocity of 6.35 km - S-I as for shots B-2 and X-2, the Moho depth was caiculated to be 32 km, similar to that obtained beneath other lines in the region. The apparent velocity for Pn is about 8.2 km + S -’ which is clearly higher than previous esti-

Profile D-A

Profile D-A (Fig. 8a>, recorded on a SSWNNE line from shotpoint D off-shore Faro in southwest Portugal, is unreversed. It intersects profile B-X 250 km from shotpoint D and 400 and 320 km from shotpoints B and X, respectively. It traverses the Hercynian domain, examining the Iberian lithosphere along a different azimuth from profile B-X (Fig. 1). Energy propaga-

i00r

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t

Profile D-A

Fig. 7. Profile B-X. 2-D ray-tracing model. The intersection point with profile D-A is indicated.

6""

j

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mates. Another significant difference with respect to the B-X profile is the first sub-Moho arrival. This has a large critical distance of 280 km, is very energetic, is delayed by 1.4 s with respect to Pn and has a high apparent velocity. Modelling shows that this phase should be equated to Pi, the second mantle reflection on the B-X profile. The first discontinuity beneath the B-X profile, marked by an increase to 8.2 km * s- ’ at 48 km depth, is not seen on the D-l records. As profile D-A intersects B-X, a first assumption is that reflection P, is produced by the same interface in both cases, The first model to be considered for D-l includes a single layer of velocity 8.2 km. s-l (derived from Pn> above this

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45

mantle interface. In order to fit both travel times and the critical distance, a LVZ (of velocity 7.7 km-s-l) must exist above the interface which is now at 61 km depth (Fig. 8b). Models assuming that the Pn apparent velocity of 8.2 km +s- ’ results from an updip phase from the layer beneath the Moho which has a true velocity of 8.0 km - se1 have also been tested. The critical distance can then be fitted without introducing a LVZ and the depth of the interface is calculated as 60 km. However, the travel time branch can not be fitted correctly. If it is insisted that the reflector has the same depth of 68 km as beneath profile B-X, a second type of models, including azimuthally varying ve-

DISTANCE(km) VRED = 8.0km/s Fig. 8. Profile D-A. Mantle arrivals from shot D-1. (a> Record section. (b) Synthetic seismograms and C(Z) model for the first type of model (see text). (c) Synthetic seismograms and u(z) model for the second type of model (see text).

46

locities, may be examined. For a sub-Moho velocity of 8.2 km. s-l, and including a thin LVZ above the mantle reflector as observed beneath profile B-X, the critical distance value of 300 km requires a sub-reflector velocity of 8.6 km . s-l (Fig. 8~). Another later mantle phase can be correlated on the D-l record section, but neither its apparent velocity nor its critical distance are well constrained. This would correspond to P,. The simplest model of the first type described above, including non-azimuthally varying velocities, would involve an 8.45 km. se1 overlying an 8.6 km . SC’ layer as obtained in the model derived from shot B-2. An 8.0 km . se1 LVZ, similar to that in the model from shot X-2, must be included above the boundary between the two layers. The depth to this boundary is constrained at approximately 80 km, being shallower than beneath the B-X line. Synthetic seismograms obtained using the reflectivity method (Fig. 8b) show a good agreement with observed data. The second type of model assuming the same depth of reflector beneath profiles B-X and D-A, but with an azimuthal variation in velocity, can also be evaluated for this deep reflector at about 88 km depth. The velocity beneath it has to be greater than the 8.6 km. s-l derived beneath the B-X profile, this being the velocity derived for the P, reflector from shotpoint D, at a depth of 68 km. Although the fit of the synthetic seismograms (Fig. 8c) is ‘poor, they do demonstrate that a velocity of 8.75 km * s-’ is the minimum value required beneath the deepest reflector even if there is an LVZ on top of it, as for the model derived from shot X-2. It is considered that, although there are notable differences in the resulting velocity-depth function, the record section from shot D-l shows a similar response of a layered subcrustal lithosphere, to those from shotpoints B and X. The reflectors derived for the first type of model from shotpoint D may be linked with those found beneath B-2-X-2 in a single model using 2-D ray-tracing, assuming they dip to match the depths at the intersection of the lines. It is found that the thickness of the LVZ on top of the 8.45 km. s- ’ reflector must be reduced from the

ILIHADSSGROUP

intersection of both profiles towards shotpoint D. The observed and calculated travel times may be equated but the fit of critical distance values becomes worse. As the zones of the reflector explored beneath the profiles are not strictly coincident, the model is not well constrained. For the deeper, 8.6 km * s-l reflector the difference between the depths derived from the two profiles is similar to that for the previous reflector. To fit the travel time data it is again necessary for the overlying LVZ to thin towards shotpoint D. In summary, two types of model can be invoked to explain the mantle seismic data. In the first type, the same layer velocity values (except immediately beneath Moho), are assumed to occur beneath lines D-A and B-X. The two mantle reflectors are then identified at 61 and 80 km depths beneath line D-A, 7 and 6 km shallower than beneath line B-X. For the second type of model it is assumed that the reflectors are approximately horizontal beneath the intersection of lines B-X and D-A with the depths derived from the shot D-l data being the same as those derived from shots B and X. In this case the layer velocities must vary with azimuth. Thus, either lateral heterogeneity or azimuthal anisotropy is present at certain depths in the subcrustal lithosphere beneath the two profiles. Discussion anisotropy

of lithospheric

heterogeneity

and

The Hercynian orogeny has been the main crust- and lithosphere-forming event in the Iberian Peninsula, except in the Alpine Betic Cordillera. This Hercynian continent has subsequently behaved as a single entity for large scale motions, for example the Cretaceous sinistral displacement with respect to Europe creating the Gulf of Biscay. It may have been partly reworked at its margins, by Eocene compression in the Pyrenees and later extension across its Mediterranean-Balearic boundary. Internal deformation may have been related to the general distensive episode of the opening of the Atlantic and to localized shear during the pre-Pyrenean rotation (Vegas, 1988). The seismic data reported here were collected

A DEEP

SEISMIC

SOUNDING

INVESTIGATION

OF THE

LITHOSPHERE

over the southwest part of the Peninsula. Apart from that obtained from the first 160 km from shot X-2 through the Betic Cordillera, all data pertain to lithosphere consolidated during Hercynian times. However, lithospheric segments sampled by profiles from shotpoints B, X, C and D are different as to the nature of the basement and their place in the Hercynian orogeny (Quesada, 1991). The northern part of profile B-X lies east of the Porto-Tomar fault, inside and along the strike of the southwest limit of the pre-Hercynian Central Iberian autochthon. The crust is of Precambrian origin, subvertically faulted and massively intruded by Palaeozoic granitic plutons. Here and in the Central System to the east, which is traversed by crusta profiles from land-shotpoint G and reached by mantle profiles from shotpoints D and C, the strike of these structures is that of the general IberoArmorican arc, being NNW-SSE at the western edge and oriented almost E-W in the centre. During the collisional phase of the Hercynian orogeny, compression occurred normal to this strike. Subsequently, during the peneplanation and root erasing stage, extension occurred exposing the huge granitic intrusions as well as deep metamorphics along normal faults, in the form of core complexes (e.g., Casquet et al., 1988). To the south, along profile B-X, the importance of Hercynian magmatism decreases and the line crosses to the other Precambrian unit, the Ossa Morena zone, through the BadajozCdrdoba shear zone (Burg et al., 19801, possibly a cryptic Cadomian suture reactivated as a Hercynian wrench fault zone. The Qssa Morena zone, the southwest imbricated margin of the preHercynian Iberian autochton, is bordered, along the South Iberian shear zone ~Crespo-~lanc and Orozco, 1988) by the Beja-Acebuches ophiolites and oceanic sedimentary sequences of the Pulo do Lobo terranes accreted during the Hercynian orogeny (Munha et al., 1989). It contains, however, only very few subduction related igneous rocks. While the amount of convergence, subduction, and compression at the southwest Iberian ophiolite belt may have been modest, the corresponding South Iberian shear zone as well as the Badajoz-C~rdoba shear zone have accommo-

BENEATH

THE

IBERIAN

PENINSULA

47

dated large scale oblique Hercynian accretion of the South Portuguese terrane. This constitutes a relatively narrow localized domain of important sinistral shear in the Iberian Peninsula (Burg et al., 1980). Crustal phases from shotpoints C and LI traverse the South Portuguese terrane, accreted by the Hercynian orogeny to the pre-Cambrian core by closure of the ocean and abduction of the ophiolites. Transtensional suites of volcanic and plutonic rocks and sedimentation in pull-apart basins in the pyrite belt of the South Portuguese terrane bordering the Pulo do Lobo zone, indicate that strike-slip along the South Iberian shear zone was a dominant influence on its accretion to the Ossa Morena autochthon. Further to the south and west, a southwest verging thrust and fold belt developed in the erogenic culm sediments, resting above an exotic Precambrian crust not otherwise involved in Hercynian crustal evolution. In summary, the Hercynian lithosphere sampled by the ILIHA experiment has an extremely heterogeneous origin. It has: (1) come from different and previously differentiated crusts from opposite sides of the suture along the South Iberian shear zone; this could provide ample reason for contrasts in the wave fields propagated from shotpoints C and L> with respect to the others, (2) incorporated oceanic lithosphere, probably in northeastward directed subduction along this suture; this could have caused, locally, a dipping anisotropy with this azimuth inherited from the oceanic lithosphere, as for models inferred elsewhere from spatial variations of P-wave travel time residuals (BabuSka et al., 1984; BabuSka and Plomerova, 1990), (31 included zones of extreme Hercynian shear deformation, along not only the South Portuguese suture but also the neighbouring Badajoz-Cordoba zone; this could have induced a local orientation into the fabric of the lithosphere along a narrow band, resulting in anisotropic heterogeneity in the sense suggested by Vauchez and Nicolas (19911, (41 involved a gradation in the amount of crust and mantle differentiation and intrusion both

ILIHA

48

from the southwestern and from the northern external zones to the internal domain, (5) been submitted successively to compression and thickening and then to extension, peneplanation and crustal normal faulting in a constant direction normal to the main structural strike, to a degree and orientation dependant on the locality. This could have resulted, if compression was the dominant stage, in lithospheric strain consistent with the measured E-W orientation of fast split SKS waves (Silver and Chan, 1988, 1991; Vinnik et al., 1989), a direction which would be expected to change along the Ibero-Armorican arc. The seismic structure of the lithosphere derived from the present study is more uniform than expected from this domain of highly heterogeneous origin. Significant lateral heterogeneities are not recognised within the subcrustal lithosphere as would be expected if the frozen-in Hercynian structure was evident at these depths. The several mantle phases distinguished in the ILIHA long range observations indicate the heterogeneity of the lower lithosphere. However, it only exists in the vertical plane, at a scale of some ten kilometers. The systematic phase patterns, seen over quite large distances, demonstrate that the corresponding layering must be nearly horizontal and quite extensive, as observed elsewhere

in Europe and discussed above in the section “Mantle structure”. Severe lateral heterogeneity would not permit the development of such persistent phases. Their generation might be supposed to depend only on details of the structure encountered by the ray bundle on entrance into the mantle around 50 km from the shotpoint. The recorded wavefields might thus differ owing to the shotpoint sites being quite different. Shot X-2 is in the Alboran sea south of the Betic Cordilleras, D-1 is in the Atlantic off the South Portugese terrane accreted to Iberia in Hercynian times, and B-2 lies off the Central Iberian autochthon, massively intruded and reworked in the Hercynian. Despite this, the mantle wave patterns have important similarities, their generation seeming to be robust with respect to local crustal structure. This is an important observation in relation to the significance of the resultant lithospheric models. Two or three high-velocity reflectors are documented in the upper 60 km of the mantle lithosphere. Reversed observations on the B-X profile can be explained with continuous, sub-horizontal layers in the mantle over a distance of 690 km. Beneath the southern part of this line, low-velocity zones above the two deeper boundaries at 68 and 86 km are well constrained. The derived model may be compared with the lower resolution, 1-D layered models obtained

b)

70,

I

Fig. 9. (a) Profile D-l, u(z) models.

B-X,

u(z) models.

Solid line corresponds

Solid line corresponds

to shot X-2, dashed

to the first type of model. Dashed

Velocity values have been corrected

DSS GROUP

to accommodate

--

1

line corresponds

line corresponds the curvature

to shot B-2. (b) Profile

to the second of the Earth.

D-A,

shot

type of model (see text).

A DEEP SEISMIC

SOUNDING

INVESTIGATION

OF THE LITHOSPHERE

from surface wave dispersion studies between station pairs from Porto, Toledo and Malaga (Badal et al., 1990). These will be refined following the temporary deployment, during the ILIHA project, of the NARS broadband array (Paulssen, 1990). These models consistently indicate the presence of a shear wave tvz within the lithospheric lid of the mantle. Its base at a depth of about 60-70 km, coincides with the P2 phase reflector, observed from all the three shotpoints B, X and D. Although not strictly geographically coincident, this agreement lends support to the idea that the low- to high-velocity interface at a depth of about 66 km is a widespread phenomenon beneath Iberia. The velocities obtained for the seismic models must be corrected to account for curvature of the Earth. The necessary correction results in rock velocities of about 8.15 km-s-l at 50 km, 8.35 km * s-’ at 65 km, and 8.5 km-s-’ at 90 km depth beneath profile B-X (Fig. 9a>. These values may be discussed with respect to pressure and temperature derivatives of candidate mantle rocks (Fuchs, 1979, 1983). The derived values are higher than expected for laterally isotropic conditions, providing indirect evidence of anisotropy. Similar values have already been reported in other continental areas (Fuchs, 1979; Bean and Jacob, 1990) and justified using anisotropic models. These flatlying zones of limited thickness but with significant latera continuity call for regional or even large scale mechanisms producing homogeneously elastic parameters in the lower lithosphere. Profile D-1, which intersects B-X, seems to provide evidence of lateral heterogeneity in the form of either slight interface dips or azimuthal anisotropy at specific depths in the mantle beneath the Iberian Peninsula. The Pn velocity from shotpoint D is higher than determined from shotpoints B or X. The first mantIe reflector, modelled beneath profile B-X at a depth of 48 km is not observed from shotpoint D either because of the lower velocity contrast across this reflector at this azimuth or because it is not present in the explored zone. A significant difference is also observed for the next two mantle reflectors with either an 8% difference in depth or a 3% differ-

BENEATH

THE IBERIAN

PENINSULA

49

ence in velocity of the layer below the reflector between the intersecting lines (Fig. 9b). Azimuthal anisotropy from Pn observations immediately beneath the Moho (Barnford, 1977) has been attributed by Fuchs (1983) to flow orientation along the m~mum shear stress direction derived from surface observations and focal mechanisms in southwest Germany. It may thus be attributed to present-day crustal stress pervading the uppermost mantle. Large scale anisotropy at the base of the lithosphere (Him, 1977) could be related to flow at the transition between the lithospheric plate and the deeper mantle locked to the present-day hotspot reference frame. There is an incompatibili~ between shear wave velocities derived beneath the Iberian Peninsula from Rayleigh and from Love waves (Maupin and Cara, 19921, suggesting an anisotropy at depths greater than the 100 km penetrated by the ILIHA DSS experiment. The surface wave data are not adequate to resolve a possible azimuthal term but it is suggested that with such a term the models could account for observations of SKS splitting beneath Iberia (Silver and Chan, 1988; Vinnik et al., 1989). Anisotropy above the present-day base of the lithosphere, in the depth range sampled by the ILIHA DSS experiment, could be related to recent motions. It could also be inherited from a particularly vigorous previous flow episode enabling lattice preferred orientation mechanisms in olivine (Nicolas and Christensen, 1987) in deep, subhorizontal, heated shear zones. This accommodates differential motion in response to vertically inhomogeneous applied stress as proposed by Bean and Jacob (1990) beneath Britain and Ireland. Evaluation of the whole ILIHA DSS data set should further constrain the picture in due course. Acknowkdgements The ILIHA project has been planned as part of the European Geotraverse (EGT) coordinated by the European Science Foundation. The project was funded by the Stimulation programme of the Commission of the European Communities, Directorate General XII, to support the participating institutions. It was also supported by the

50

Swiss National Research Funds. The assistance of other national research councils providing equipment and personnel is also acknowledged. The Portuguese and Spanish navies also supported the project and carried out the sea-shooting. References Babushka, V. and Plomerova, J., 1990. F-residual study in the Iberian Peninsula. XXII Gen. Assem. Eur. Seismol. Comm., Barcelona, Abstr.: 137. Babushka, V., Plomerov& J. and Silenj;, J., 1984. Spatial variations of P residuals and deep structure of the European lithosphere. Geophys. J.R. Astron. Sot., 79: 363-383. Badal, J., Corchete, V., Payo, G., Canas, J.A., Pujades, L. and Set&, F.J., 1990. Processing and inversion of long-period surface-wave data collected in the Iberian Peninsula. Geophys. J. Int., 100: 193-202. Bamford, D., 1977. Pn velocity anisotropy in a continental upper mantle. Geophys. J.R. astr. Sot., 49: 29-48. Bamford D., Faber, S., Jacob, A.W.B., Kaminski, W., Nunn, K., Prodehl, C., Fuchs, K., King, K. and Willmore, P., 1976. A lithospheric seismic profile in Britain I. Preliminary results. Geophys. J.R. Astron. Sot., 44: 145-150. Banda, E., Suriiiach, E., Aparicio, A., Sierra, J. and Ruiz de la Parte, E., 1981. Crustal and upper mantle structure of the central Iberian Meseta. Geophys. J.R. Astron. Sot., 87: 779-789. Bean, C.J. and Jacob, A.W.B., 1990. P-wave anisotropy in the lower lithosphere. Earth Planet. Sci. L&t., 99: 58-65. Burg, J.P., Iglesias, M., Laurent, Ph., Matte, Ph., and Ribeiro, A., 1980. Variscan intracontinental deformation: the Coimbra-Cordoba shear zone (SW Iberian Peninsula). Tectonophysics, 78: 161-178. Caetano, H., 1983. Structure crustale de la zone Sud Portugaise et de la zone Ossa-Morena d’aprbs les etudes de sismologie experimentale. These 3e cycle, Univ. Paris VI, Paris, I44 pp. Casquet, C., Fuster, J.M., Gonzalez-Casado, J.M., Peinado, M. and Villaseca, C., 1988. Extensional tectonics and granite emplacement in the Spanish Central System. A discussion. In: E. Banda and L.A. Mendes Victor (Editors), Proc. 5th Workshop on the European Geotraverse (EGT), the Iberian Peninsula. European Science Foundation, Strasbourg, pp. 65-76. Cordoba, D., Banda, E. and Ansorge, J., 1987. The Hercynian crust in NW Spain: a seismic survey. Tectonophysics, 132: 321-333. Cordoba, D., Banda, E. and Ansorge, J., 1988. P-wave velocity-depth distribution in the Hercynian crust of northwest Spain. Phys. Earth Planet. Int., 51: 236-248. Crespo-Blanc, A. and Orozco, M., 1988. The southern Iberian Shear Zone: a major boundary in the Hercynian fold belt. Tectonophysics, 148: 221-227. Faber, S. and Bamford, D., 1979. Lithospheric structural

ILIHA

DSS GROUP

contrasts across the Caledonides of Northern Britain. Tectonophysics, 56: 17-30. Fuchs, K., 1979. Structure, physical properties and lateral heterogeneities of the subcrustal lithosphere from longrange deep seismic sounding observations on continents. Tectonophysics, 56: l-15. Fuchs, K., 1983. Recently formed elastic anisotropy and petrological models for the continental subcrustal lithosphere in southern Germany. Phys. Earth Planet. Inter., 31: 93-118. Fuchs, K. and Mueller, G., 1971. Computation of synthetic seismograms with the reflectivity method and comparison with observations. Geophys. J.R. Astron. Sot., 23: 417-433. Grupo de Trabajo de Perfiles sismicos profundos, 1983. Perfiles sismicos profundos en Espana, 1981. Bol. Geol. Min., 44(4): 339-347. Guggisberg, B., Ansorge, J. and Mueller, St., 1984. Structure of the upper mantle under southern Scandinavia from Fennolora data. In: D.A. Galson and St. Mueller (Editors), Proc. 1st Workshop on the European Geotraverse (EGT), The Northern Segment. European Science Foundation, Strasbourg, pp. 49-52. Hauser, F. and Stangl, R., 1990. The structure of the crust and the lithosphere in Fennoscandia derived from a joint interpretation of P and S wave data of the FENNOLORA refraction seismic profile. In: R. Freeman and St. Mueller (Editors), Proc. 6th Workshop on the European Geotraverse (EGT), Data compilations and synoptic interpretation. European Science Foundation, Strasbourg, pp. 71-92. Hirn, A., 1977. Anisotropy in the continental upper mantle: possible evidence from explosion seismology. Geophys. J.R. Astron. Sot., 49: 49-58. Hirn, A., Steinmetz, L., Kind, R. and Fuchs, K., 1973. Longrange profiles in western Europe: II. Fine structure of the lower lithosphere in France (southern Bretagne). Z. Geophys., 39: 363-381. Hirn, A., Prodehl, C. and Steinmetz, L., 1975. An experimental test of models of the lower lithosphere in Bretagne (France) (1) (2). Ann. Geophys., 31: 517-530. Jacob, A.W.B., 1975. Dispersed shots at optimum depth-an efficient seismic source for lithosphere studies. J. Geophys., 41: 63-70. Kind, R., 1975. Propagation of seismic energy in the lower lithosphere. 2. Geophys., 40: 188-202. Maupin, V. and Cara, M., 1992. Love-Rayleigh wave incompatibili~ and possible deep upper mantle anisotropy in the Iberian Peninsula. Pure Appt. Geophys,, 138: 429-444. Mendes-Victor, L.A., Hirn, A. and Veinante, J.L., 1980. A seismic section across the Tagus valley, Portugal: possible evolution of the crust. Ann. Geophys., 36: 469-476. Mueller, St., Prodchl, C., Mendes, A.S. and Sousa Moreira, V., 1973. Crustal structure in the southern part of the Iberian Peninsula. Tectonophysics, 20: 307-318. Munha, J. Oliveira. J.T., Ribeiro, A., Quesada, C., Fonseca, P. and Castro, P., 1989. Accreted terranes in southern Iberia: the Beja-Acebuches ophiolite and related oceanic sequences. 28th Int. Geol. Congr., Washington, DC, Abstr., 2: 312-314.

A DEEP SEISMIC SOUNDING INVESTIGATION

OF THE LITHOSPHERE BENEATH THE IBERIAN PENINSULA

Nicotas, A. and Christensen, N.I., 1987. Formation of anisotropy in upper mantle peridotites. A review. In: K. Fuchs and C. Froidevaux (Editors), Composition, Structure and Dynamics of the Lithosphere-Asthenosphere System. Am. Geophys. Union, Washington, DC, pp. 137154. Paulssen, H., 1990. The Iberian Peninsula and the ILIHA Project. Terra Nova, 2: 429-435. Perrier, G. and Ruegg, J.C., 1973. Structure profonde du Massif Central franqais. Ann. Geophys., 29: 435-502. Quesada, C.. 1991. Geological constraints on the Palaeozoic tectonic evolution of tectonostratigraphic terranes in the Iberian Massif. Tectonophysi~s, 185: 225-245. Sapin, M. and Prodehl, C., 1973. Long-range profiles in western Europe. I: Crustal structure between the Bretagne and the Central Massif of France. Ann. Geophys., 29: 127-145. Silver, P.G. and Chan, W.W., 1988. Implications for continental structure and evolution from seismic anisotropy. Nature, 355: 34-39.

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Silver, P.G. and Chan, W.W., 1991. Shear wave splitting and subcontinental mantle deformation. J. Geophys. Res., 96: 16,429-16,454. Suriiiach, E. and Vegas, R., 1988. Lateral inhomogeneities of the Hercynian crust in central Spain. Phys. Earth Planet. Int., 51: 226-234. Vauchez, A. and Nicolas, A., 1991. Mountain building: strike-parallel motion and mantle anisotropy. Tectonophysics, 185: 183-201. Vegas, R., 1988. Alpine and recent geodynamic evolution of Iberia: crustal implications. In: E. Banda and L.A. Mendes Victor (Editors), Proc. 5th Workshop on the European Geotraverse (EGT), The Iberian Peninsula. European Science Foundation, Strasbourg, pp. 77-90. Vinnik, L.P., Farra, V. and Romanowicz, B., 1989. Azimuthal anisotropy in the Earth from observations at Geoscope and Nars broadband stations. Bull. Seismol. Sot. Am., 79: 1542-1558.