Refraction-seismic investigations of the northern Massif Central (France)

TECTONQPHySICS Tectonophysics 275 (1997) 99-117

ELSEVIER

Refraction-seismic investigations of the northern Massif Central (France) Hermann Zeyen a,,, Olaf Novak a,l, Michael Landes a,2, Claus Prodehl a, Lynda Driad b, Alfred Him b a Geophysikalisches lnstitut, Universitiit Karlsruhe, Hertzstrasse 16, D-76187 Karlsruhe, Germany t, lnstitut de Physique du Globe, Universitg de Paris IV, 4, Place Jussieu, F-75230 Paris, Cedex 05, France

Received 30 November 1995: accepted 17 July 1996

Abstract New refraction-seismic data have been obtained in the northern French Massif Central. Seven shots were recorded simultaneously on two parallel, NNW-SSE (i.e., perpendicular to the surface structures) oriented profiles allowing for inline and fan interpretations. The results of a 2D-interpretation reveal a large-scale updoming of the Moho of about 2 km, compared to average western Europe, corresponding to the Cenozoic uplift of the Massif Central. Local crustal thinning of up to 20% is restricted to the Cenozoic grabens as seen at the surface. Beneath the volcanic area of Cantal, one of the largest Cenozoic European shield volcanoes, which was active in Miocene to recent times, significant crustal thickening of 2-3 km is observed. Velocities in the upper mantle are generally in the range of 7.9-8.1 km/s beneath the Limagne graben and surrounding undisturbed areas. Underneath the volcanic area of C6zallier, however, velocity is reduced to about 7.7 kin/s, corresponding to the area of crustal thickening. Therefore, the low velocity in this zone is explained by the remainders of cooling magma chambers and heating by ascending hot plume material. The Cenozoic tectonic evolution of the Massif Central has affected crust and upper mantle in different ways. Oligocene rifting corresponds well to local crustal thinning, but no correlation is detected to velocities of the upper mantle. In contrast, the volcanic event occurring since Miocene times is related to upper mantle velocity reduction and cmstal thickening due to underplating. Keywords: crust; velocity structure; underplating; cmstal thinning; Limagne

1. Introduction Since the 1960s the French Massif Central (Fig. 1) has been studied by numerous refraction-seismic ex* Corresponding author. Present address: Institutionen f6r Geofysik, Uppsala Universitet, Villav~igen 16, S-75236 Uppsala, Sweden. Tel.: +46 (18) 183-781; Fax: +46 (18) 501-110; E-mail: [email protected] 1Present address: Dublin Institute for Advanced Studies, 5, Marrion Square, Dublin, Ireland. 2 Present address: Dublin Institute for Advanced Studies, 5, Marrion Square, Dublin, Ireland.

periments, most of the lines running roughly in N - S direction (Fig. 2). A compilation of the results until 1973 is given by Perrier and Ruegg (1973). The reported crustal structure is dominated by a thinning of the crust from 30-32 km in the western, undisturbed part of the Massif Central to 24 km beneath the Limagne graben, a Cenozoic rift structure, cutting the northern part of the Massif Central in approximately N - S direction. This thinning is the culmination of a broad updoming of the crust-mantle boundary (Moho) to 27-28 km depth beneath the eastern part of

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H. Zeyen et al. / Tectonophysics 275 (1997) 9 9 - 1 / 7

the Massif Central. The area of updoming, which comprises the volcanic areas and the grabens of the Massif Central has been correlated with very low velocities in the uppermost mantle of only 7.2-7.4 km/s which increase to 8.4 km/s at a depth of 45 km. This sub-Moho low velocity has been explained by heating due to lithospheric thinning to about 50 km (Perrier and Ruegg, 1973) and by the presence of partial melt at Moho level (Lucazeau et al., 1984). A new refraction-seismic project has been carfled out jointly by the Geophysical Institute of the University of Karlsruhe and the IPG de Paris in fall 1992. The project formed part of a program integrating petrological xenolith studies (Wilson and Downes, 1992; Werling, 1994), teleseismic tomography (Granet et al., 1995) and the refraction-seismic experiment with the aim to study the structure of the crust and upper mantle under the Limagne graben and the surrounding young volcanic areas. In this paper we describe the P-wave structure of the crust and uppermost mantle obtained from the refraction-seismic experiment below the central Limagne graben and south of it under the recent volcanic areas of Cantal and Dev~s/Velay where hitherto no profiles existed perpendicular to the strike of the structures. We will show that the crustal structure published by Perrier and Ruegg (1973) has been largely confirmed, but that the area and amplitude of the low velocity anomaly in the uppermost mantle has to be strongly reduced in order to fit the new seismic-refraction data, which is in accordance with teleseismic and petrological data. 2. Tectonic evolution

The Massif Central forms the central part of the European Hercynian belt. During the Hercynian orogeny the area was folded and cut into blocks by strike-slip faults like the Sillon-Houllier fault with an associated displacement of more than 50 km (Fig. 1). Widespread metamorphism and plutonism are responsible for the actual composition of the basement, formed mainly by gneisses and granitic batholiths of Hercynian age (Burg and Matte, 1978; Ledru et al., 1994). After erosion to a peneplain and a marine transgression during Jurassic and Cretaceous times the area was reactivated during the Alpine orogeny.

In the Oligocene, parallel to the formation of the European Cenozoic Rift System (Ziegler, 1992), a system of halfgrabens developed in the northern part of the Massif Central, the most important one being the Limagne graben (Fig. 1). Its Cenozoic sedimentary thickness reaches a maximum of about 2.5 km on its western side north of Clermont-Ferrand (Morange et al., 1971). Other Cenozoic grabens are the basins of Roanne, Montbrison, Ambert and Le Puy east of the Limagne graben (Fig. 1). The border faults of these basins are generally striking in N to NW direction, parallel to the maximum horizontal stress from Oligocene to recent (Carbon, 1992). Parallel to the graben formation in the northern Massif Central the whole area started to rise in the Oligocene (Carbon, 1992) leading to the actual average elevation of approximately 1400 m above sea level and denudation of the basement. Little volcanism accompanied this updoming in its first stage. Since Miocene times, however, strong volcanic activity in the west and southeast of the Limagne graben has made this area the main volcanic province of the European Cenozoic Rift System. The principal phases of eruption correspond to Middle and Upper Miocene (8-4 Ma) and to almost recent times (3~0.3 Ma). After that time volcanism continued in the northern (Cha~ne des Puys) and southeastern (Velay) provinces (Fig. 1). The volcanism is concentrated mainly in two areas. Along the western edge of the Limagne graben a chain of volcanoes runs approximately in N-S direction, comprising, from north to south, the Chaine des Puys, the large strato volcano of Mont Dore, the basaltic plateau of the Crzailter, the central volcano of the Cantal and the Aubrac Mountains. Southeast of the Limagne graben, surrounding the Le Puy basin, the volcanic areas of Velay, Dev~s and Vivarais contain some of the youngest volcanoes. The composition of the extrusives varies widely from andesitic to basaltic and phonolitic. Many lower crustal and upper mantle xenoliths are found, brought to the surface by the volcanic activity. Lower crustal xenoliths contain mafic as well as acid granulitic rocks of Hercynian formation age. Compared to other areas in Europe where xenoliths from the lower crust have been found, acid rocks, with an occurrence of approximately 40%, are fairly abundant in comparison to mafic ones (Downes, 1993). Man-

14. Zeyen et aL /Tectonophysics 275 (1997) 99-• 17

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tie xenoliths give a mainly therzolitic to harzburgitic composition of the mantle above about 80 km depth (Werling, t994). Their study resulted in different models of the generation of the graben and volcanism

ranging from the idea of a triple junction (Brousse, 1974), an uprising plume (e.g., Lucazeau and Bayer, 1982; Granet et a l . 1995) and the formation of severat small-scale diapirs (Nicolas et al., 1987).

102

H. Zeyen et al. / Tectonophysics 275 (1997) 99-117

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3. Experimental setting In the context of Cenozoic graben formation associated with subsequent relatively strong volcanism and due to the availability of xenoliths which reflect the recent thermal state of the lithosphere beneath the Massif Central, a combined petrologic, teleseismic and refraction-seismic project was carried out

in cooperation with different German and French institutions (Werling, 1994; Granet et al., 1995). The aim of the refraction-seismic experiment was to investigate the effect of the graben formation and the volcanism on the structure of the Hercynian crust and to obtain absolute velocities in the crust and uppermost mantle which were to be compared to velocity variations obtained by the teleseismic study

103

H. Zeyen et al./ Tectonophysics 275 (1997) 99-117 Table 1 Shotpoints of the refraction seismic project LIMA 92 Shotpoint

Date

Latitude (N)

Longitude (E)

Topography (m)

Charge (kg)

PIO MON GEO EYM LAN PAU CHE

10/17/92 10/17/92 10/21/92 10/19/92 10/20/92 10/20/92 10/17/92

46°04~27.22" 46°00'28.18" 45o42, 15.77" 45040'45.57" 45°26'54.83" 45o09'47.50" 44°55~45.47"

2°44'10.59" 3003'42.95 " 3054, 15.02" 1°47'03.15" 2o36' 15.62" 3o38'57.30" 4°31' 14.71"

800 510 990 650 820 1140 810

400 400 400 800 400 400 800

and to petrologically expected velocities for Moho P - T conditions and compositions as revealed by

xenolith analyses. In order to accomplish both objectives at the same time the experiment was given a special design (Fig. 2). Seven shots (Table 1) were simultaneously recorded on two profiles by 150 digital, one-component SGR stations with 2 Hz geophones from PASSCAL, U.S.A. (Program for Array Seismic Studies of the ContinentAl Lithosphere). 105 stations were deployed along 230 km on the southern profile which runs parallel to the northern profile at a distance of about 70 km. This profile crosses two of the young volcanic fields, the Velay/Dev~s in the east and the C6zallier in the west and is centred at the southern end of the Limagne graben. The northern profile with the remaining 45 stations crosses the central Limagne graben in WNW-ESE direction 30 km north of Clermont-Ferrand with a total length of 150 km. Because of unfavourable noise conditions, in the graben itself no stations were deployed. The station spacing is generally between 2 and 2.5 km. The northern profile was planned in order to get information on the crustal structure below the Limagne graben mainly by inner crustal and PMP (Moho reflection) phases. The shot points Pionsat (PIO) and St. Georges (GEO) were placed such that their distance to the opposite graben border corresponded to the expected critical distance for PMP reflections (approximately 70 km). The central shot, Montcel (MON), was intended mainly to image the transition from stable Hercynian areas to the volcanic zone and the graben by wide-angle reflections recorded on the near-fan profile discussed below. The southern profile was intended to image the influence of the young volcanic areas on crustal structures

and to give information on the P-wave velocity in the uppermost mantle undemeath these volcanic areas. With an overall length of 230 km the profile is long enough to allow for reversed Pn observations from the extremal shot points Eymoutiers (EYM) and Le Cheylard (CHE). Details of the crustal structure were expected to be imaged by the intermediate shots Lanobre (LAN) and Paulhaguet (PAU). Due to the simultaneous recording on both profiles reversed fan-type data should give more detailed information on the crustal structure, and in particular reveal details on the form and extent of the Moho updoming under the Limagne graben. Therefore the distance between both profiles was fixed to 70 km, the expected critical distance of PMP reflections. 4. Overview of the data

Figs. 3-5 show the record sections for P-waves with the correlations computed from the models presented in the next section plotted with a reduction velocity of 6 km/s. The data quality is very good and the high dominant frequency of about 15 Hz allows good structural resolution. The record sections are all filtered with a 20 Hz low-pass filter in order to enhance the signal-to-noise ratio. Four phases can be recognised in all inline sections (Figs. 3 and 4): a direct wave (Pg) with a velocity of about 5.9 km/s, two inner crustal reflections (PI1P and PI2P) and the PMP reflection from the Moho. First arrivals from the uppermost mantle (Pn) can only be identified in five out of eleven record sections at distances larger than 130 km (Fig. 5). Furthermore, data from shot CHE, recorded on the northern line, show clear secondary arrivals between Pn and PMP phases which are interpreted as reflections from an interface within the subcrustal lithosphere (Fig. 4e). In

104

H. Zeyen et al. / Tectonophysics 275 (1997) 99-117

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H. Zeyen et al./Tectonophysics 275 (1997) 99-117

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DISTANCE (kTn) Fig. 3 (continued).

75

50

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0

(d)

H. Zeyen et al. / Tectonophysics 2 75 (1997) 99-117

106

general, the PMP reflection is the most prominent phase in the record sections at distances beyond 70 km. Pn arrivals are mostly weak except for the inline data of shot GEO and the fan data of shot PAU (Fig. 4c,e). Comparing the record sections of the two western shots on the southern line (Fig. 3a,b)

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H. Zeyen et al./Tectonophysics 275 (1997) 99-117

107

in the area of the Limagne graben and in one of the sections (Fig. 6c) also at about 40 km east of the graben.

5. Interpretation 4

5.1. Southern line

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as reverberations originating in a layered lower crust, covers this reflection. The fan-type data are represented using a constant velocity normal moveout (NMO) technique. Seismograms are projected to their midpoints (reflector points) (Fig. 6). In this way one can immediately recognise variations in the crustal structure by delayed or early arrivals and attribute them to the approximate geographical area from where they are reflected. In this case the NMO has been calculated for the PMP reflection with a one-layer model, having an average velocity of 6.17 km/s and a thickness of 29.5 km, corresponding to an average crustal model for the southern profile. PuP arrivals earlier than 0 s show crustal thinning or higher average crustal velocity whereas arrivals later than 0 s indicate thickening of the crust or a smaller average crustal velocity. With only one fan shot it is not possible to distinguish between variations in Moho depth or mean crustal velocity. But since several overlapping fans have been recorded which illuminate the interesting area at different angles this distinction can be made. Especially in the sections of the northern shots (Fig. 6a-c) early PuP arrivals are clearly seen

The different inline profiles and long-distance fan sections have been interpreted by 2D-modelling using the ray-tracing program MACRAY (Luetgert, 1992). On the southern profile (Fig. 7) an upper crust with a mean velocity of 5.9 km/s and a thickness of 11 to 12 km is underlain by a middle crust with velocities of 6.3-6.4 km/s and a lower crust with a velocity increasing from about 6.6 km/s at a depth of 23 km to 6.7 km/s at the crust-mantle boundary. Due to the large critical distances of the two inner crustal reflections (90-100 kin) lateral resolution of the depth of the corresponding reflectors is relatively poor. Therefore, depth variations of up to 2 km cannot be excluded. The traveltimes of all phases are clearly delayed in the area of the volcanic zone of Crzallier, indicating a low velocity of the near-surface volcanic edifices which are mainly composed of pyroclastic materials. The Moho depth is interpreted beneath most of the profile between 28 and 29 kin. Beneath the volcanic area of Crzallier, however, a crustal thickening of 2-3 km to approximately 31 km is modelled in order to explain a delay of up to 0.3 s of the PMP reflections from shot EYM between 140 and 160 km distance and some strong, delayed reflections on several traces of shots LAN (60-80 km) and PAU (100-130 km). This thickening cannot be explained by lower crustal velocities in the western part of the Massif Central since the velocities are constrained by other phases, including the Pn arrivals from shot CHE. On profiles LAN and PAU, however, strong additional, earlier arrivals are observed in the same distance range which can be fitted by a reflection from a depth of 28-29 km. From these observations we infer that beneath the volcanic area of Crzailier an additional fourth layer separates lower crust and mantle. Missing correlation between adjacent traces and high amplitudes indicate a complicated structure, possibly a fine-scale interlayering of peridotitic and basaltic material.

H. Zeyen et al./Tectonophysics 275 (1997) 99-117

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H. Zeyen et al./ Tectonophysics 275 (1997) 99-117

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The P-wave velocity in the uppermost mantle can only be indirectly inferred. Only non-reversed Pn-arrivals are observed from shot CHE (Fig. 5a). For most of the profile the Pn-velocity is inferred to be 7.95 + 0.1 km/s. However, a delay of 0.2 s of all Pn-arrivals at distances larger than 200 km from the shot point indicates a reduction of uppermost mantle velocities between km 75 and 100 of the profile, corresponding again to the volcanic area of Crzallier. The apparent velocity here is 7.5-7.6 km/s which, due to the thickening of the crust in the same area, corresponds to a true velocity of about 7.7 km/s. Due to the good crustal control based on the intermediate shot points we believe that, in spite of not being reversed, a velocity of 7.95 i 0.1 km/s is representative for the uppermost mantle near the southern end of the Limagne graben. Only beneath the recently active volcanic areas a velocity reduction is indicated.

5.2. Long-distance fan shots The shots EYM and CHE, recorded on the northern line (Fig. 4d,e), can also be considered as quasi inline shots at least for distances above 100 km. Both record sections show Pg, PIP and PMP arrivals as well as weak but clear Pn arrivals (Fig. 5b,c). While the reflected waves and the Pg come from different areas, the Pn from both shots image the uppermost mantle below

the Limagne graben, continuing to both sides (Fig. 8). The small intercept time for Pn arrivals at short distances (160-190 km from shot CHE and 140-170 km from shot EYM at about 5.8-5.9 s, Fig. 5b,c) indicates a Moho at about 26 km depth. Beyond that distance Pn is delayed in both profiles. For shot CHE the delay occurs along a short range between 190 and 200 km distance indicating a rather steep depth increase of Moho towards the west to about 29 km. Pn data from shot EYM show a similar increase of crustal thickness on the eastern side of the graben. The quasi reversed Pn-velocity below the graben is 8.0 to 8.1 km/s. The record section of shot CHE shows a further remarkable feature: strong correlatable arrivals in the time range between the Pn and PMP arrivals. This energy may originate at approximately 45 km depth where the velocity increases to about 8.3-8.4 km/s as inferred from the critical distance. Since this is the only record section where this energy can be seen this interpretation is only confirmed by the interpretation of the older, N-S-running profiles, where Perrier and Ruegg (1973) have reported a similar reflector in the upper mantle.

5.3. Northern line The northern line was observed mainly in order to obtain information on the crustal structure below

112

H. Zeyen et al./Tectonophysics 275 (1997) 99-• 17

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600

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Fig. 8. Map with indication of source areas of different wave types under the assumption of horizontal reflectors. UTM coordinates are given. the Limagne graben. Nevertheless, a Pn wave with strong amplitudes is observed on one of the three shots (GEO, Fig. 4c), yielding a very low apparent velocity of about 7.5 km/s. From a 1D-interpretation, a Moho depth of 25 km with an underlying low-velocity mantle might be inferred. In this case the high amplitudes should be due to a high vertical velocity gradient in the uppermost mantle as was interpreted earlier for a profile following the Limagne graben along its axis (Perrier and Ruegg, 1973). Since all other Pn observations yield velocities of about 8 km/s under undisturbed areas as well as

under the graben and since the source area of the P. waves of shot CHE recorded along the northern profile is at only about 10 km distance from the one of shot GEO it seems unlikely that such a low velocity should exist under an area which, in case of a horizontal Moho, would lie outside the graben area and north of the volcanic area. Therefore we propose a 2D-model based on two assumptions: (1) the true velocity under the graben is about 8 km/s, and the observed low velocity is an apparent velocity due to an increase of crustal thickness along the profile; and (2) the strong amplitudes are due to

H. Zeyen et al. / Tectonophysics 275 (1997) 99-117

113

LIMA 92: Northern Profile Limagne gmben PIO

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Moho depth derived from fan and inllne registrations J

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St. Georges

z~~

Lanobre

~

(fan)

(fan)

P a u l h a g u e t (fan) N o r t h e r n profile

-26

S o u t h e m profile t-

15.

-28

o tO

-30

-32 450 W

500

550

UTM easting [km]

600 E

Fig. 10. Interpretation of the PMP arrivals of the fan shots in terms of Moho depth. The corresponding model has one crustal layer with an average velocity of 6.19 krn/s. Results from northern profile and fan data, which both give information from beneath the Limagne graben, are similar, whereas the southern profile, located only a few kilometres south of the Limagne graben, shows a very different structure. The coordinates of the horizontal axis correspond to UTM coordinates to enable comparison with Fig. 8.

114

tt. Zo,en et al,/Tectonophysics 275 (t997) 99-t 17

O°

t °

2°

3°

4°

5"

6"

43"

43" O"

1"

2"

3°

4"

5"

6~

Fig, 11, New Moho map for the area of the Massif Central (courtesy of C, Bauer). White circles indicate sites of published Moho depths. lsolines every kilometre. Light grey, shallow Moho; dark grey, deep Moho.

structural effects which focus the rays in the area of observed strong Pn arrivals. Fig. 9 shows the model of crustal structure which results in this case, The crust is asymmetric being thinnest (26 kin) near the

western border of the Limagne graben. Towards the west the coast thickens abruptly by 4. km (to about 30 kin) whereas to the east it thickens only gradually reaching 28 km at the eastern end of the line. Within

H. Zeyen et al./ Tectonophysics 275 (1997) 99-117

the crust two reflectors are distinguished. Between 11 and 12 km depth velocity increases slightly from 6.0-6.1 to about 6.1-6.2 km/s. At about 20 km depth the lower crust starts having velocities between 6.6 and 6.8 km/s. Both interfaces seem to be essentially horizontal with a depth variation of not more than 1-1.5 km. The corresponding ray theoretical amplitudes fit well the observed amplitude ratios of P,, PuP and PTP waves. 5.4. Fan shots

The position of the two lines was selected such that quasi fan shots could be observed at the expected critical distance range for PMP reflections. This is why in the record sections (Fig. 6) the main amplitudes are PMP arrivals. The NMO correction for all profiles was performed assuming a single-layer crustal model with a mean velocity of 6.17 km/s and a crustal thickness of 29.5 km, corresponding to the average model of the southern line. Especially the northern shots recorded on the southern profile yield good data and show clearly early arrivals of up to 0.6 s between about 500 and 530 km (UTM coordinates in km) which corresponds to the area just below the Limagne graben. A second area of early PMP arrivals (up to 0.5 s) is observed around 550 km (UTM) in the records of shots GEO and PAU. Qualitatively, the traveltime curves of the western fan anomaly show the same behaviour as the Moho structure of the model of the northern line. The distance between minimum and maximum arrival times is about 30 km and the area of minimum arrival times is situated below the Limagne graben near its western border. It is interesting to note that on the fan records of shot PIO between km 485 and 495 (UTM), just below the volcanic area of the ChaSne des Puys, an arrival time delay of 0.2 s is observed. It cannot be decided, however, whether it has to be interpreted by a crustal thickening, as it is interpreted beneath the volcanic area of Ctzallier on the southern profile, or by a reduction of mean crustal velocity. Fig. 10 shows a quantitative interpretation of the delay times. This interpretation was performed with a one-layer model, maintaining a laterally constant mean velocity and varying the crustal thickness. Implicitly, with this approach one supposes that along

115

the ray paths the thickness of all layers varies more or less in the same way. Fig. 9 indicates that at least east of the graben this might be true. The mean velocity was then varied in order to obtain a good fit of the models for the anomaly around 510 km (UTM). The mean crustal velocity results to 6.19 krn/s. The model of the western anomaly confirms the inline model of the northern line fairly well, giving a minimum Moho depth below the western half of the Limagne graben of 25-26 km. The eastern anomaly is explained by a Moho updoming of about 3 km. Fig. 11 shows an updated contour map of the depth to Moho for the area of the Massif Central using all available data until present (Bauer, 1995).

6. Conclusions The model presented here shows the following main features: (1) A three-layer crust with velocities of 5.9-6.0 km/s in the uppermost 10-12 km, 6.3-6.5 krn/s in the middle crust and 6.6--6.8 km/s in the lower crust below 18-20 km. The reflected phase corresponding to the boundary between middle and lower crust is not seen well on all profiles, on other profiles it has strong amplitudes already at precritical distances but it is not a continuous phase. Therefore we interpret the velocity increase in the lowermost crust as well as the observed reflections by layers of basic intrusions into a generally felsic middle and lower crust. In this case the lower crust seems to consist of essentially two kinds of material, the more felsic one with velocities smaller than 6.6-6.8 km/s and the basic one with higher velocities (6.9-7.0 km/s). This view is supported by lower crustal xenoliths which contain an exceptionally large amount of felsic material (Downes, 1993). The bimodal velocity distribution could explain the strong, apparently precritical, arrivals in the seismic data (see e.g., shot EYM, Fig. 3a). (2) Mean crustal thickness is 28-29 km. Thinning of up to 4 km has been observed, localised in narrow zones just beneath the Cenozoic rift basins. Beneath the volcanic area of Ctzallier of Miocene to recent age two distinct reflections in the PMP range indicate a crustal thickening of 2-3 km due to a layer which can be explained by underplated material. Similar observations were made by Perrier and Ruegg

I 16

H. Zeyen et al. / Tectonophysics 275 (1997) 99-117

(1973) who have reported crustal thickening of about 3 km beneath the Crzailler/Cantal area, which is mainly due to an additional lowermost-crustal layer of unknown velocity. (3) Pn velocities are generally 7.9-8.1 km/s. Only beneath the volcanic area of C6zallier lower velocities of approximately 7.7 km/s are found. These results show that the rifting, occurring mainly in the Oligocene, affected the crust by forming several subparallel basins with up to 2.5 km of sediment infill and by thinning the crust by up to 4 km, mainly at the expense of the lower crust. The total thinning of the crystalline crust in the basin areas amounts to 15-20%. This thinning is, however, restricted to the areas where basins are present. Comparing the overall crustal thickness of 28-30 km with other Hercynian areas and surrounding profiles (Perrier and Ruegg, 1973; Prodehl et al., 1996), one observes a Moho updoming of about 2 km. This updoming corresponds well to the amount of Cenozoic uplift of the Massif Central. Therefore, this rather shallow position of Moho should not be considered as rift-related crustal thinning but as due to thermal uplift of the whole crust in connection with the rise of a plume (Granet et al., 1995). Velocities in the uppermost mantle, however, have not been affected by rifting. In contrast to earlier results (Him and Perrier, 1974) Pn-velocities of 8.0 4- 0.1 km/s in most of the area are normal velocities and correspond to crust of Hercynian age. Very low velocities of 7.2-7.4 km/s beneath the Limagne graben could not be confirmed. They may have been the result of essentially 1D-interpretations of low apparent velocities near the southern border of the Limagne graben. However, low velocities in the uppermost mantle beneath the volcanic area of Crzallier, reported from refraction seismics (Perrier and Ruegg, 1973) and teleseismic models (Granet et al., 1995) are confirmed and may be explained by heating due to the uprise of hot plume material (Sobolev et al., 1996). Neither partial melt nor anisotropy are needed to explain velocity reduction. Therefore we conclude, that rifting affected mainly crustal thickness whereas seismic velocities in the mantle are mainly modified by the uprise of a plume and the related volcanic activity. An exception from this is the crustal thickening beneath the Crzallier which is interpreted as due to underplating.

Finally, we want to emphasise the importance of the configuration of the experiment. It was the first refraction-seismic experiment crossing the area roughly perpendicular to the strike of structures, allowing for detection of lateral variations of velocities and deep structures. Of utmost importance was the simultaneous recording of all shots on both profiles, yielding additional Pn arrivals and, above all, fantype profiles. Only these off-line recordings allowed to prove the lateral variation of crustal thickness beneath the Limagne graben and therefore the distinction between real low velocities and apparent ones due to downdip shooting.

Acknowledgements This work has been carried out in the frame of the Special Research Centre 108 of the Deutsche Forschungsgemeinschaft. Part of the funding came from INSU (Institut Nacional de la Science de l'Univers, France). Apart from the authors the following persons participated in the field experiment: U. Enderle, W. Kaminski, U. Kastner, J. Mechie, T. Nadolny, M. Tittgemeyer, V. Wehrle (Karlsruhe), J. Gagnepain, J. R6my, A. Simonin, M. Vadell, J. Verhille (Paris), S. Balzuhn, C. Haber, A. Rudloff (Berlin), C. Montana, B. Durrani (El Paso), C. Horan, T. Blake, P. Readman (Dublin), J.P. Canales, J, Dfaz, N. Vidal (Barcelona). We appreciate very much the help of PASSCAL, U.S.A., providing the field equipment and personnel for data acquisition (M. Alvarez and S. Michnick). We are grateful to J. Dorel (Universit6 de ClermontFerrand) who provided rooms for the head quarter during the experiment. The maps were drawn using the General Mapping Tools (GMT; Wessel and Smith, 1995). This is publication no. 564 of the Special Research Centre 108 and no. 723 of the Geophysical Institute of the University of Karlsruhe. J. Luetgert and E. Fltih helped to improve the manuscript with their comments on an earlier version of the manuscript.

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sium Jung, G~ologie, gromorphologie et structure profonde du Massif Central franqals. Plein Air Service, Clermont-Ferrand, pp. 295-309. Nicolas, A., Lucazeau, E and Bayer, R., 1987. Peridotite xenoliths in Massif Central basalts, France: textural and geophysical evidence for asthenospheric diapirism, In: H.P. Nixon (Editor), Mantle Xenoliths. John Wiley and Sons Ltd., Chichester, pp. 563-574. Perrier, G. and Ruegg, J.-C., 1973. Structure profonde du Massif Central francais. Ann. Grophys., 29: 435-502. Prodehl, C., Mueller, St. and Haak, V., 1996. The European Cenozoic rift system. In: K.H. Olsen (Editor), Continental Rifts: Evolution, Structure, Tectonics (CREST). Elsevier, Amsterdam, pp. 133-212. Sobolev, S., Zeyen, H., Stoll, G., Werling, E, Altherr, R. and Fuchs, K., 1996. Upper mantle temperatures from teleseismic tomography of French Massif Central including effects of composition, mineral reactions, anharmonicity, anelasticity and partial melt. Earth Planet. Sci. Lett., 139: 147-163. Weding, F., 1994. Die thermische Entwicklung des lithosph~irischen Mantels im Bereich des franzOsischen Riftsystems abgeleitet aus der Mineralchemie von Mantelxenolithen. PhD thesis, University of Karlsruhe, 291 pp. Wessel, P. and Smith, W.H.E, 1995. New version of the Generic Mapping Tools released. EOS Trans., AGU, 76: 329. Wilson, M. and Downes, H., 1992. Mafic alkaline magmatism associated with the European Cenozoic rift system. In: P.A. Ziegler (Editor), Geodynamics of Rifting, Volume I. Case History Studies on Rifts: Europe and Asia. Tectonophysics, 208: 173-182. Ziegler, P.A., 1992. European Cenozoic Rift System. In: P,A. Ziegler (Editor), Geodynamics of Rifting, Volume I. Case History Studies on Rifts: Europe and Asia. Tectonophysics, 208:91-111. -

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