The lithosphere in central Europe---seismological and petrological aspects

Tectonophysics, 207 (1992) 141-163

Elsevier Science Publishers B.V., Amsterdam

141

The lithosphere in central Europe-seismological european and petrological aspects *

V. BabuSka and J. Plomerov6

Geophysical Institute, Czechoslouak Academy of Sciences, BoEni II, 14131 Prague 4, Czechoslovakia

(Received March 13, 1990; revised version accepted September 23, 1991)

ABSTRACT

BabuSka, V. and Plomerovi, J., 1992. The lithosphere in central Europe-seismological and petrological aspects. In: R.

Freeman and St. Mueller (Editors), The European Geotraverse, Part 8. Tectonophysics, 207: 141-163.

The lithosphere thickness in the Variscan belt of central Europe varies between about 60 and 150 km with typical values

of loo-120 km. Our estimates, derived from directionally independent representative P-wave residuals, are in good agreemen; with magnetotelluric determinations of a layer with increased conductivity in the upper mantle. The large-scale

anisotropies of the subcrustal lithosphere beneath four seismological stations determined from spatial variations of relative

P residuals vary between 6.5 and 15.2% for P velocities; the S-wave anisotropies determined from SKS polarizations vary between 2.2 and 6.7%. These values are in reasonable agreement with the anisotropies of peridotites determined in

laboratory. Systematic spatial variations of the directional terms of relative residuals in dependence on azimuths and

incidence angles suggest the existence of large dipping anisotropic structures in the subcrustal lithosphere. The residual

patterns at most stations in the Saxothuringicum, Rhenohercynicum and in the Massif Central imply northwesterly

orientated dips of the anisotropic structures while stations in the Moldanubicum, the Alpine Foredeep and most of the Alps

north of the Insubric line, suggest southeasterly orientated dips. In our interpretation the dipping anisotropic structures may

represent paleosubductions which retain olivine preferred orientations originating from an ancient oceanic lithosphere. The

Variscides of central Europe may thus represent a collision zone characterized by two systems of paleosubductions divergent

relative to the suture between the Moldanubicum and the Saxothuringicum.

Introduction

To date, several attempts have been made to interpret the Variscan belt of central and western Europe from the point of view of plate tectonic development (e.g., Behr et al., 1984; Matte, 1986; Franke, 1986). However, full understanding of the deep continental structure in terms of plate tectonics is very difficult to achieve. Most information on Variscan structures comes from investigations of near-surface tectonics and seismic refraction and reflection surveys of the crust. As

* Publication no.130/90, Geophys. Inst., Czechosl.Acad.Sci,

Prague

Correspondence to: V. BabuSka, Geophysical Institute, Czech-

oslovak Academy of Sciences, BoEni II, 141 31 Prague 4,

Czechoslovakia.

to the subcrustal lithosphere, we have only a very limited amount of data. However, thanks to the olivine-preferred orientation and the corresponding seismic anisotropy we could obtain valuable information on the geodynamic processes which may have led to the present-day continental structure (BabuSka et al., 1984b).

This paper presents the results of tomographic research on the deep lithosphere structure in central Europe, namely the region of the Variscan belt. We transform variations of P-wave velocities into 3-D models of velocity perturbations, models of the lithosphere thickness and models of high- and low-velocity directions in the subcrustal lithosphere. These tomographic studies (e.g., BabuSka et al., 1988) are based on the investigation of relative P-wave residuals computed at networks of seismological stations. We use teleseismic P-

0040-1951/92/$05.00 0 1992 - Elsevier Science Publishers B.V. All rights reserved

142 V. HAHLJSKA AND J. PLOMEKOVA.

arrival times (from ISC bulletins and local net-

works of stations) of events at epicentral dis-

tances of 20-loo”, compute absolute residuals

relative to the Jeffreys-Bullen (JB) Earth model,

and introduce several corrections: for the elliptic-

ity, altitude of a station and crustal corrections.

For each seismic ray these crustal corrections

take into consideration the average velocity and

thickness of the crust beneath a station, and

thicknesses and average velocities of sediments if

they exist. These corrections compensate a delay

caused by a thicker or a thinner crust, or by

crustal velocities different from the JB model.

The crustal correction can be understood as a

reduction of the residuals to the M discontinuity

at a reference depth of 33 km. We tested the

effects of possible errors in determining the

crustal parameters and found them small, relative

to the observed residual variations (Babushka et

al., 1984a). We have shown elsewhere (Aric et al..

1989) that the relief of the lithosphere-astheno-

sphere transition cannot change significantly as a

result of incorrect crustal corrections beneath a

station.

The computation of relative residuals, usually

called normalization, deserves special attention in

each investigated region and each data set. The

normalization minimizes effects originating in

deep mantle paths and in source regions, as well

as mislocation effects. A reference basis of rela-

tive residuals is traditionally formed by an aver-

age residual of an event and its level can differ

for different data sets and different systems of

reference stations. It is also important to choose

reference stations in such a way that the direc-

tional pattern of the reference basis, which is

projected onto the pattern of the relative residu-

als at each station, does not distort the results.

When applied carefully, the normalization is a

powerful tool in computations of relative residu-

als and allows us to avoid event relocations, which

are otherwise necessary in tomographic studies

whose earthquake sources are located within the

investigated part of the Earth’s volume.

Relative P residuals at i-tb station

%,k = ‘/wk %, j pi j _Ri k/ 5 2 S

I

Nk . ! . . .nu&r of events in k-th source region Ri,j.. *relative residuals of j-tb event

I 3-D inversions of n-layer models; P-m PEtTDRMIONR in block wimtions of 3-D inhomgeneities of the upper mantle

at each station (i)

a directionally dependent ten ~~ATIVB RBsID(I~ Ri

azimuth-incidence dgle dependent terrs DIRBCI?O~ Tw)Gs Yii k I

fii = l/w CRi,k iii k = $ k - q for k-tb source region I I

N...nunber of source regions vith steeply in&ding waves from

Ii = I/?$ cRi,s NA,..nWr of azimuth intervals

each~a&utb interval I

Ri s = l/RR ~i,k , Ns...nueber Of source reqions in s-th

I azimuth interval HODEL OF Tm LI!rROspBwE TIi1CKNl3ss I

POLAR DIAGRAMS of pi );(AZ,I) S~OU directions of relatively hi d and lov velocities

9 HODRL OF ANISOTROPIC STRDCTURR

OF sDBauJsT& LmiOsPRRRR

Fig. I. Flow chart of the relative P-residual processing.

THE LITHOSPHERE IN CENTRAL EUROPE 143

To suppress the uneven distribution of earthquake foci as to their distances and azimuths around an investigated region, we compute average relative residuals Ri,k from relative residuals at the i-th station and foci in the k-th source region-a segment usually of 10” in the epicentral distance and 20” in the azimuth relative to the centre of the investigated region. This kind of averaging residuals over source segments also helps to eliminate the well-known reading errors. Only those relative residuals Ri,j (i stands for a station, j for an event) that do not differ from the average more than k 2 s (see Fig. 1) are used for computing Ri,k (k denotes k-th region). Thus the stations with almost no difference in the absolute residuals can exhibit different directional dependence of the relative residuals which are one order higher than mean errors of the relative residuals (BabuSka et al., 1990).

Figure 1 gives brief information about the fur- ther processing of the relative residuals R+ (for a full flow chart of the method and explanation see BabuSka at al., 1989 and 1984a, respectively). Traditional tomographic studies of the Earth con- vert P residuals of single events into 3-D velocity perturbation models which describe a priori determined rectangular blocks beneath the region down to an arbitrarily chosen depth. In contrast, we compute and present here a set of discrete values of residuals in the region. The distribution of these values depends only on the distribution of the stations. The residuals are divided at ‘each station into a directionally independent term (Ri), which is called a representative residual, and azimuth-incidence angle dependent terms (Ri,k) called directional terms (see Fig. 1). The former is used to model the lithosphere thickness of a region, the latter maps directions of relatively high and low velocities.

P-velocity inhomogeneities in the lithosphere-as-

thenosphere system

Three-dimensional inversion methods pio- neered by Aki et al. (1976) and extended by a number of other authors (e.g., Koch, 1985; Spak- man, 1988; Granet and Cara, 1988) explain variations of P-wave velocities in a region in terms of

velocity perturbations in rectangular blocks. We have computed and published several models of velocity perturbations in the mantle for several

regions of Europe-e.g., beneath central Europe (BabuSka et al., 1988), the Alps and their surroundings (BabuSka et al., 1990), and the eastern Alps (Aric et al., 1989)-using the original or modified ACH method (Aki et al., 1976; Granet and Cara, 1988). A common feature of all the inversions, however, is evident dependence of perturbations on block approximations of a region and on block sizes and their locations. The station coverage of the European region, especially in its eastern part, does not allow us to compute small enough blocks which would fit the scales of complex European tectonics. Neverthe- less, large-scale tectonic features known from surface geology are detected in the distribution of velocity perturbations of upper layers in our published models. The models show, e.g., the high- velocity perturbations beneath the Alps down to several hundred kilometres (BabuSka et al., 1988; Spakman, 1988). A special study on the deep structure of the Alps (BabuSka et al., 1990) con- firms the existence of two high-velocity inhomogeneities beneath this mountain range separated by a material with lower velocities in the central Alps. This is in good agreement with a model of European lithosphere (BabuSka et al., 1985; see also Fig. 2 in the next section) where we detected two lithospheric roots beneath the Alps-one in the western and the other in the eastern part. This means that we also observe a correspondence of velocity perturbations in upper layers with models of lithosphere thickness derived from lateral variations of the representative residuals (see Fig. 1). Good correlation is also found in regions of subsidence-e.g., the PO Plain, or the Pannonian Basin-where all 3-D inversions give low-velocity perturbations and where the lithosphere models detect shallow depth of the lithosphere-asthenosphere transition (around 60 km).

On the other hand, in regions where we observe strong directional dependence of the residuals (see section 5) the 3-D inversion techniques, which do not allow for anisotropic wave propagation, can map the extremal velocities in particular directions down to deeper layers. This can in-

144

crease the amplitude of perturbations at bottom layers and bias the final models (BabuSka et al., 1990). Until now, no appropriate 3-D inversion method, which would take into account directional variations of velocities in individual blocks and which could be applied to real data has been developed. Therefore, the traditional 3-D inversion results have to be considered as a first ap- proximation, especially in regions with complex anisotropic structures, such as we observe, e.g., in Europe. The limitation of 3-D results is reflected in values of the variance improvement of the inversions, which usually do not exceed 40%. This means that less than half of the observed residuals are explained by the isotropic models. Of course, the quality of observations affects this parameter, but we invert pre-processed average residuals R,,, (see Fig. 11, whose reading errors are mostly eliminated. The anisotropic effects are probably so pronounced that an isotropic inversion performed with non-bulletin high-quality data exhibits relatively low variance improve-

I-

V. BAHUSKA AND J. PLOMEROVA

ments of about 40% as well CM. Granet, pers. commun., 1990).

Model of lithosphere thickness

Relative residual-lithosphere thickness relation

To model the thickness of the lithosphere, we have used the representative residuals Rj computed at each station (see Fig. 1). Our modelling of the “seismic” lithosphere as a high-velocity layer overlying the low-velocity asthenosphere (Sacks and Snoke, 1984) is based on the assumption, similar to that of Poupinet (1979). We assume that changes of the lithosphere thickness are proportional to the observed variations of the representative residuals, which reflect the average velocity variations beneath the individual stations regardless of the directions of arriving waves. The regions of relatively thin lithosphere are characterized by positive relative residuals-a de- ficiency of high-velocity lithospheric material-

ORTJ-I dERMAN- PoL,Sk

PLATFhRM

EAST EUROPEAN

i

PLATFORM

I ‘/I nn I \

1

\

Fig. 2. Model of the lithosphere thickness (isothickness contours in km) derived from the representative residuals R, (see Figs. 1 and 3) computed at individual seismological stations (dots).

THE LITHOSPHERE IN CENTRAL EUROPE

and, on the other hand, negative relative residuals indicate regions of relatively thick lithosphere -abundance of high-velocity material.

To derive a relationship between the lithosphere thickness and the representative residuals (BabuSka et al., 1987), no a priori assumption about the velocity-depth distribution was made. We assume that effects of lateral changes of average velocities in the entire subcrustal lithosphere are small in comparison with the effects of lateral changes of the lithosphere thickness, as a consequence of the velocity contrast between the lithosphere and the asthenosphere. For the representative residual-lithosphere thickness relation, we adopted the lithosphere thickness according to Panza et al. (1980) for the most positive residuals in the Belgo-Dutch Platform ( + 0.8 s = 50 km) and the depth of the mountain root according to Baer (1980) and Mueller (1982) for the most negative residuals in the western Alps (-1.0 s = 220 km). By linear interpolation we derived for central Europe a gradient of 9.4 km/O.1 s (Plomerova and BabuSka, 1988).

To analyze possible effects of lateral variations of temperature in the crust and uppermost mantle on the model of lithosphere thickness derived from P residuals we estimated changes of mean P velocity in the subcrustal lithosphere beneath the Moldanubicum, as a cold stable core of the Bo- hemian Massif and beneath the Pannonian Basin, as a hyperthermal region (Cermak et al., 1991). The resulting 0.1 km/s change of the mean velocity as a consequence of lateral variation of temperature in the lithosphere would produce approximately 0.6 km thickening of the lithosphere beneath the Pannonian Basin in comparison with the model which does not take into account the temperature effects. As the accuracy of the lithosphere thickness model derived from P residuals is estimated at k 10 km beneath high-quality stations, the temperature effects are one order lower and can be neglected.

As we work with the relative residuals, the reference level formed by the residuals averaged over the reference stations plays an important role. This is significant for the lithosphere thickness in central Europe displayed by the isothickness contours in Figure 2. This map is based on

47

45

43_

PYM -5 ASS0

7- :

0 2 4 6

-1 -5 GRC A -3 ALOR

145

Fig. 3. Representative residuals Ri (in tenths of a second,

without the crustal corrections) at stations in the Massif

Central and its southwestern surroundings used for the model

of the lithosphere thickness (Fig. 2).

the study of P residuals in central Europe (BabuSka et al., 1984, 1987) and complemented with the results of several regional studies-e.g., in the Alps (BabuSka et al., 1985; Aric et al., 1989) or beneath the GRF array (Plomerova and BabuSka, 1988). For all data sets we kept the European gradient of the residual-lithosphere relation, however, the relation could not be applied directly to the residuals due to different reference levels. The residuals of various data sets were homogenized by a shift, the magnitude of which was determined from a comparison of relative residuals of several stations situated in overlapping areas. This is also the case of the lithosphere thickness in the westernmost part of the region depicted in Figure 2. We present here a more detailed model estimated from P-arrival times recorded at French stations in the Massif Central and its southwestern surroundings in the period 1985-1986. Figure 3 shows the Ki distribution in this region, computed without the crustal corrections relative to average residuals ever all stations in the western part of the Euro- pean region (see Fig. 5). The accuracy of these

146 v. BABLJSKA AND J. PLOMEHOV~

residuals is kO.1 s with the exception of stations PLDF, PYM and SSB, which reported less data.

Crustal thicknesses vary between 27 and 30 km beneath the Massif Central and also changes of average crustal velocities are small (Mostaanpour, 1984). Therefore, the crustal effects could be

included as a constant term in the residual shift which was applied to correct the difference in the reference levels mentioned above. Eight high-quality stations of the region-LOR, LBF, SSF, AVF, TCF, SMF, MZF and CAF (Figs. 3 and 5) -were used to link up the new results with the

WI 1 CLL

11’ 12’

H

I

J as Cl

c4

SE

MM ZEV ZTT Saxothuringirn Moldanubian

Frankcnw~td Fichtelgcbirge Oberpfalzer Wald DEKORP 40 I

TWT

Fig. 4. (a) Tectonic sketch map of the region of the contact zone of the Moldanubicum and Saxothuringicum with seismological

stations (triangles) and the GRF array (dots). Squares denote magnetotelluric measurements, numbers ascribed to them arc depths

(in km) of a layer with increased electrical conductivity in the upper mantle; dot-dashed lines are isothickness contours of the

lithosphere model depicted in Figure 2. (b) Structural interpretation of the DECORP 4 profile (Vollbrecht et al., 1989):

I = Saxothuringian crust; 2 = Moldanubian crust; 3 = Erbendorf body; 4 = nappes. (c) Schematic cross section through the model

of the lithosphere thickness along profile KLM (see Fig. 4a) beneath the GRF array (Plomerova and Babushka, 1988).


previous model (BabuSka et al., 1988). The representative residuals differ by - 0.2 + 0.04 s on the average and, therefore, we shifted all the residuals by this value and transformed them into a model of the lithosphere thickness of the Massif Central and its southwestern surroundings (Fig.

2).

Lithosphere thickness beneath Variscan Europe

This study focuses on the central and western European Variscides. A detailed description of the lithosphere-asthenosphere transition beneath other parts of central Europe was published elsewhere (e.g., BabuSka et al., 1987, 1988). Typical thicknesses of the Variscan lithosphere are around loo-120 km. The lithosphere be- comes thicker along the southern rim of the belt -beneath the southern part of the Bohemian Massif and the Alpine foredeep. The thickening of the lithosphere continues to the south beneath the Alps where roots as deep as 200 km are produced by the collision of the European and Adriatic plates. The lithosphere-asthenosphere transition sinks down to depths of over 140 km also beneath the northern part of the Massif Central whereas its distinct upwelling to depths of about 60 km is observed beneath the northern part of the Rhine Graben and the southern part of the Rhenish Massif. A similar regional thinning appears at the centre of the Massif Central and other regions of active Tertiary volcanism. In general, the characteristic SW-NE trend of the Variscan belt in central Europe is reflected in the isothickness contours beneath the Bohemian Massif and the tectonics units of southern Ger- many.

The representative residuals are usually determined with the accuracy of kO.1 s in good-quality stations. The corresponding accuracy of lithosphere thickness variations is about & 10 km. However, the exact level of the lithosphere- asthenosphere transition remains an open question. An independent possibility of checking the thickness of the “seismic” lithosphere is a comparison with the results of magnetotelluric investigations. This can be done due to the fact that a decrease of the seismic velocities at the top of the asthenosphere is probably caused by partial melt-

TABLE 1

Thicknesses of “seismic” and “magnetotelluric” lithosphere

Region Lithosphere thickness (km)

“Seismic” “Magnetotelluric”

Belgo-Dutch Platform 60- 80 55- 70

Rhenish Massif 60- 80 40- 90

Southern Rhine Graben 80-100 80- 95 Bohemian Massif 100-140 85-150

Western Carpathians 100-140 85-130 Pannonian Basin 60- 80 40-100

ing and is thus accompanied by an increase of electrical conductivity (Fournier et al., 1963). In a special study of this topic, Praus et a1.(1990) found good agreement between depths of the lithosphere-asthenosphere transition determined from seismic and magnetotelluric measurements. The agreement is observed both in the scale of large tectonic units (see examples in Table 1) and also in details-measurements on profiles (see Fig. 4). A study of the lithosphere structure beneath the contact zone of two Variscan units, the Saxothuringicum and Moldanubicum, beneath the GRF array (Plomerova and BabuSka, 1988) showed that the lithosphere-asthenosphere transition dipped steeply to the south-southeast and isothickness contours of this zone followed a prominent SW-NE direction of the Variscan fold belt in central Europe. The magnetotelluric measurements along two crossing profiles, one paral- lel to the GRF array (Fig. 4) and the other perpendicular to it, confirm both the depths of the transition and its topography. Figure 4 also shows a structural interpretation of the reflection seismic profile Decorp 4 (DEKORP Research Group, 1988) and a cross section through the model of the lithosphere thickness beneath the GRF array. The steep dip of the lithosphere- asthenosphere transition to the south-southeast, which is also indicated by a dipping highly con- ductive layer in the mantle (Praus et al., 19901, has its equivalent in the structure of the crust. A highly reflective horizon dipping. steeply to the southeast is interpreted as a deep-reaching decollement along which the Moldanubicum was overthrust onto the Saxothuringicum (Vollbrecht et al., 1989). The lithosphere thickness of the

148 V. BABUkA AND J. PlDMEROVli

region depicted in Figure 2 can be compared with the results of average elastic properties of the upper 200-250 km by Panza (1984). However, that author pointed out that only in the areas where there was a distinct decrease of velocities within the upper and lower layers of his model, could the isolines be considered representative of the depth to the lithosphere-asthenosphere transition. This has to be taken into consideration in comparison with other estimates of lithosphere thicknesses. Due to a long-wavelength and an integral sampling of the lithosphere by the surface waves, the results describe mainly “the gross features of the lithosphere-asthenosphere system” (Panza et al., 1980). Both maps show lithospheric thickening beneath the Alps, however, the surface waves detect only one mountain root. The prominent SW-NE direction of Variscan belt in central Europe detected in the map of the lithosphere thickness constructed from P-wave residuals (see Fig. 2) differs from the approximately N-S orientation of “lithosphere thickness contours” in the map by Panza et al. (1980). A recent study by Yanovskaya et al. (1990) estab- lished an azimuthal anisotropy from the surface waves beneath central Europe which could affect the previous estimates of elastic properties of the lithosphere-asthenosphere system.

Another estimate of the lithosphere thickness based on the dispersion of surface waves was made by Chang and Marillier (1987) in the western part of the Variscan Europe, where they estimate the lithosphere thickness to be at ieast 75 km, but more likely, more than 100 km. The latter estimate is close to our loo-120 km typical thickness (see Fig. 2).

The marked lateral variations of the lithosphere-asthenosphere transition with depth should be reflected in the convection flows in the viscous mantle beneath the rigid lithosphere and thus affect the stress field. Pick and Charvatova- Jakubcova (1988) used Runcorn’s equations to compute the stresses caused by convection flows from the outer gravity field. While the pattern of stresses beneath the Variscan belt in central Eu- rope is relatively smooth and without extremes, the stress field due to the flow beneath the Alps is affected by the lithosphere roots. The orienta-

tions of stresses computed from the gravity anomalies are in accord with the dip directions of the lithosphere thickening and thus seem to support the topography image of the lithosphere- asthenosphere transition derived from the representative P residuals.

Petrology and elastic-wave velocities of materials of the deep lithosphere

The large variations of the observed P residuals at single stations, both as to their average values (representative Ri residuals) and as to their directional dependence (azimuth-incidence angle dependent terms g;i,k), can be only partly explained by crustal effects, and their main cause has to be sought in the upper mantle. In this section we want to review briefly the laboratory data on P and S velocities in the main upper mantle rocks and their velocity anisotropies at high pressure conditions, which simulate the envi- ronment of seismic wave propagation in the deep continental lithosphere.

Peridotites, pyroxenites and eclogites are the most important representatives of the upper mantle materials according to their physical properties, namely densities and elastic wave velocities. Direct evidence of the petroiogy of the upper mantle is derived from xenoliths and from ophiolites. Another source of the information is in ultramafic massifs, which were uplifted by tectonic processes from the mantle along faults, especiahy in collision zones.

A recent analysis of heavy mineral concen- trates from upper mantle samples made by Schulze (1989) confirmed the previous estimates of an average composition of upper-mantle xenoliths which led to the conclusion that the upper mantle was predominantly peridotite. The author found that the amount of eclogite in the upper 200 km of the subcontinental mantle was perhaps less than 1% by volume overall. Therefore, it is reasonable to base any seismological model of the subcrustal continental lithosphere on the elastic properties of peridotite specimens.

Peridotite, also called oiivinite, is composed of olivine as the dominant mineral, two aluminous pyroxenes and sometimes garnet and chromifer-


TABLE 2

Compressional and shear-wave velocities and their anisotropy coefficients (k) in minerals relevant to peridotite composition of deep lithosphere

P velocities &m/s)

” maX “min “mean

S velocities (km/s) Source of data

k (%) ” max L, min

0 mean k (%I

Olivine 9.89 1.12 8.42 24.6 5.53 4.42 4.89 22.3 Kumazawa and Anderson (1969) Bronzite 8.30 7.04 7.78 16.4 4.99 4.27 4.12 15.6 Kumazawa (1969)

Diopside 8.60 6.94 7.7 21.4 4.83 3.94 4.38 20.3 Aleksandrov et al. (1963), Volarovich et al. (1975)

Garnet(pyrope) 8.90 8.85 8.88 0.6 4.96 4.91 4.94 1 Babubka et al. (1978) Spine1 10.31 9.10 9.71 12.5 6.61 4.50 5.56 37.9 Verma (1960)

ous spine1 in a minor quantity. Depending on their mineral content, special varieties of peridotites are distinguished-dunite (almost exclu- sively composed of olivine), lherzolite (olivine, orthopyroxene and clinopyroxene), wehrlite (olivine and clinopyroxene) and harzburgite (olivine and orthopyroxene).

Oehm et al. (1983) analyzed upper-mantle xenoliths from basalts of the Northern Hessian Depression and came to the conclusion that the average modal composition of 73 vol.% olivine, 18% orthopyroxene, 7% clinopyroxene and 1.3% spine1 was close to xenoliths from the Eifel, Mas- sif Central and Spain, and from worldwide sampling (Wedepohl, 1987). It is thus reasonable to take such a composition as a basis for a model of the subcrustal lithosphere beneath central Eu- rope.

It is interesting to note that an average composition of the oceanic uppermost mantle is very

similar to analyses made for continents. Dick (1987) made modal analyses of 266 peridotite samples dredged at six ocean ridges and gave an average composition of 74.8 f. 5.3% olivine, 20.6 + 3.7% enstatite, 3.6 k 2% diopside, 0.5 f 0.2% Cr-spinel, and 0.5 k 1.2% plagioclase.

Petrophysical data important for interpreting seismological observations are given in Tables 2-4. Table 2 shows P- and S-wave velocities and their anisotropy in the main minerals of the subcrustal lithosphere. We can use average compressional velocities in single crystals for simple computations of velocities in isotropic aggregates. The average modal composition of upper-mantle xenoliths published by Oehm et al. (1983) yields an average velocity of 8.27 km/s. Admixtures of 50% bronzite or 50% diopside in an olivine aggregate reduced the P velocity to 8.1 and 8.06 km/s, respectively. On the other hand, if we consider extremal cases of admixtures of 20%

TABLE 3

Compressional wave velocities in upper mantle rocks and their anisotropy coefficients (k) at high hydrostatic pressure

Number Range of Average Range of Average of Pressure Source of data

Dunite

of mean u of mean anisotropy samples (km/s) u (km/s) k (%)

8 8.00-8.66 8.32 7.2-15.0

Lherzolite 6 7.82-8.37 8.06 3.8-11.1 Olivinite 4 7.45-8.6 7.95 3.1-13.2 Harzburgite 3 7.58-7.71 7.64 9.2-12.3 Pyroxenite 3 8.01-8.29 8.18 0.5- 5.4 Eclogite 11 7.9 -8.61 8.24 0.7- 3.9

anisotropy (GPa)

k (%I

10.6 1

1.5 1 10.2 1.5 10.8 1 3.5 1 2.3 1

Birch (1960), Christensen

(19661, Babushka (1972) BabuSka (1976) Levitova (1979) Kroenke et al. (1976) Birch (1960), Kroenke et al. (1976) Manghnani et al. (1974)

150 V HABU~KA AND J. YLOMEKOVA

garnet or 10% spine1 in an ohvine aggregate, we get a veIocity increasing to 8.51 and 8.55 km/s, respectively. These simple calculations demonstrate that the P velocities in the subcrustal lithosphere as large as 8.7-8.9 km/s are practically impossible to explain in terms of isotropic aggregates of realistic compositions. Such velocities were, therefore, sometimes regarded as indirect evidence of velocity anisotropy (Fuchs, 1977; An- sorge et al., 1979).

Olivine, orthopyroxene (enstatite or bronzite, both with simifar elastic properties) and diopside, have single-crystal anisotropies which could con- tribute to the overall anisotropy of their aggregates. However, Christensen (1984) showed that an admixture of pyroxene decreased the anisotropy of olivinic rocks. This is confirmed by laboratory measurements of other authors (Table 3). The average P-velocity anisotropies in ohvinic rocks vary between 7.5 and 10.8%, whereas in pyroxenite and eclogite they attain only 3.5 and 2.3%, respectively. It has to be emphasized that both the range of velocities and the anisotropy resulting from laboratory measurements depend on several factors, such as the size and quality of samples, number and orientations of directions of velocity measurements, as we11 as their accuracy.

Shear-wave velocities and their anisotropy (Ta- ble 4) have to be evaluated with particular cau- tion because their values in anisotropic media depend both on directions of wave propagation and polarizations of shear waves. In general, the larger number of directions and polarizations is used for investigating a massive rock, the larger variations of velocities and their anisotropy are observed. This may be one of the reasons why the anisotropy coefficients in the same rock type, like in dunite from Twin Sisters (Table 41, significantly vary.

A special case of velocity anisotropy measurements are studies of shear-wave splitting in a single direction of wave propagation. This applies to, e.g., studies of large-scale anisotropy by means of SKS polarization analysis, which, however, represent only subvertical directions of wave propagations. Since in anisotropic media the velocities vary in dependence on directions of wave propagation with respect to an orientation of symmetry elements, the estimates of the magnitude of anisotropy determined by this method (e.g., Silver and Chan, 1988; Vinnik et al., 1989) may not be representative as the measurements are hardly orientated in the directions of extremal velocities. As an example we can imagine the shear-wave

TABLE 4

Shear-wave velocities and anisotropy coefficients (k) in olivinic ultramafites

I,mex t,min

(km/s) (km/s)

k f%f Source of data

Dunite (Webster) 4.44 4.38 4.40 1.4

Dunite (Mt. Dun) 4.64 4.43 4.54 4.6

Dunite (T. Sisters) 4.YO 4.70 4.83 4.1

Dunite (Transvaal) 3.98 3.75 3.90 5.9

Dun&e A (T. Sisters) 4.696 4.362 4.529 7.4

Dunite B CT. Sisters) 4.980 4.747 4.864 4.8

Peridotite 1 (Hawaii) 4.62 4.29 4.49 7.3

Peridotite 2 (Hawaii) 4.75 4.62 4.68 2.8

Dunite (T. Sisters) 4.89 4.51 4.74 8.0

Dunite CT. Sisters) 5.01 4.57 4.79 9.2

Lherzohte KHZ (N.Mex.1 4.85 4.42 4.63 9.2

Lherzolite SBl (Calif.) 4.R9 4.6 I 4.75 5.9

Lherzolite Ba2 (Calif.) 4.86 4.53 4.69 7.0

Simmons (1964)

Christensen and

Ramananantoandro (1971)

Christensen (1966)

BabuSka ( 1972)

BabuSka (19761


splitting for wave propagations and wave dis- placements along the crystallographic axes of a single crystal of olivine (Kumazawa and Ander- son, 1969). The anisotropy of S waves propagat- ing along b and c axes attains 10% whereas along axis a it is practically zero.

Large-scale anisotropy

While in laboratory investigations of elastic anisotropy the directions of wave propagation can be chosen with respect to a rock structure and the symmetry of the anisotropic medium, the possibilities in seismological observations of large-scale anisotropy are limited by the distribution of the sources of waves and receiving stations. As the anisotropy of seismic velocities is a

3-D phenomenon, specially designed experiments will be needed in the future to resolve regional anisotropies as to their symmetry and orientation of symmetry elements. At present, the available observations allow us, with some uncertainty, to detect seismic anisotropy in some regions and to speculate on simple models, which are probably still far from reality.

In this section we characterize the deep lithosphere beneath seismological stations (Fig. 5) according to the directional dependence of P-wave velocities derived from relative residuals and compare the P-wave anisotropies with the anisotropy determined from the SKS polarizations.

Variations of the directional terms ki,k of the residuals relative to the directional average Ri at

TABLE 5

Directional averages k, at stations in the western part of the region

Station

AGO

ARV

AVF

BDI

BGF

BHG

BNS

BSF

CAF

CDF

CDR

CRE

cm

DBN

DIX

DOU

EMS

ENN

EPF

FUR

FVI

GRC

GRF

HAM

HOF

LBF

LFF

LLS

di (s)

0.34

-0.16

- 0.46

- 0.27

- 0.29

- 0.08

0.59

- 0.57

-0.12

- 0.37

0.07

0.04

-0.31

0.64

0.34

0.19

0.12

0.19

- 0.25

0.33

- 0.03

- 0.28

0.11

0.79

- 0.04

-0.58

0.05

- 0.01

Mean

error

+0.11

0.17

0.08

0.18

0.09

0.15

0.16

0.04

0.10

0.12

0.07

0.19

0.13

0.17

0.12

0.09

0.11

0.14

0.13

0.11

0.13

0.10

0.12

0.25

0.10

0.06

0.07

0.12

NAZ

15

17

17

13

16

16

14

17

17

16

16

15

16

15

17

16

16

16

17

16

16

14

17

10

15

17

17

17

Station I& (s)

LMR -0.07

LOR -0.56

LPG - 0.03

LPO 0.02

LSF - 0.27

MAF -0.10

MEM 0.42

MOX 0.09

MZF - 0.03

oss 0.17

PLDF 0.16

PYM 0.27

RJF -0.13

ROB - 0.53

SAL 0.46

SAX - 0.09

SMF - 0.48

SSB - 0.29

SSF - 0.46

STV - 0.57

TCF -0.18

TMA - 0.27

TNS 0.18

ucc 0.45

WET -0.11

WIT 1.27

WTS 0.42

Mean

error

kO.14

0.05

0.15

0.08

0.09

0.12

0.13

0.13

0.10

0.10

0.13

0.10

0.09

0.21

0.17

0.09

0.08

0.24

0.06

0.30

0.08

0.14

0.13

0.10

0.14

0.09

0.14

NAZ

15

17

16

17

17

17

16

17

16

17

15

15

17

10

15

16

17

12

17

13

17

17

16

16

16

16

16

, “U

T u

----

L

SF “

chFA

A

A

A

A

*G

OPL

DF

PYA

A A

AJF

W’

A C,

F SS

B

15--a

, ,

A

A

CT

I A

I PC

N A 15

IF

I_.,.

..#

.. RV

T

9 ,

11

13

Fig.

5.

Sei

smol

ogic

al

stat

ions

w

hose

IS

C

data

fr

om

the

1985

-198

6 pe

riod

ar

e used f

or

stud

ying

sp

atia

l va

riat

ions

of

P r

esid

uals

,

THE LITHOSPHERE IN CENTRAL EUROPE

a station (see Fig. 1) are expressed in polar diagrams. In these diagrams the directional terms are plotted in dependence on the azimuths and incidence angles of arriving waves. As the directional averages fii (Tables 5 and 6) form the zero level in each diagram, the directional depen- dences at individual stations can be compared. The diagrams constructed for stations in the Variscan belt are shown in Figures 6-8. The investigated region framed by latitudes 43-54”N and by longitudes O-22”E was divided into two subregions 13” wide in longitude with a 4” overlap (Fig. 5). We examined 13’7 stations, the data of which were extracted from the ISC bulletins of

153

1985 and 1986. In both subregions, as a reference base of the normalization we used average residuals computed over all stations with the condition that at least 30% of the stations recorded an event. We checked the diagrams of 24 stations which are in the overlapping parts of both subregions and found them almost identical (Figs. 7 and 8) in spite of the fact that the reference bases differ and the event set was not the same. The normalizing diagrams (Figs. 7 and 8) show the average absolute residuals in individual source regions and demonstrate the pattern which is plotted in reverse on the diagrams of the relative residuals. Therefore, the pattern of the normaliz-

4 6 Fig. 6. Diagrams of the directional terms Ri,k (see Fig. 1) of the relative P residuals at stations in the Massif Central. The perimeter of the diagrams corresponds to an incidence angle of SO”, steeply incident teleseismic waves are projected close to the centre-the inner circle corresponds to an incidence angle of lo”. Minus signs and crosses stand for negative and positive residuals, respectively,

and their size is proportional to the absolute value of the residuals.

ing diagrams as random as possible is the best sponding incidence angles. As crustal corrections

solution because such normaiization cannot bias at indiv~dua1 stations behave as constant terms,

the individual patterns of relative residuals. their omission cannot bias the spatial variations

Computations of the diagrams (Figs. 6-9) do of the directional terms displayed in the dia-

not include the crustal corrections and they are grams. It would be difficult to show all the dia-

constructed from I?i.k at the sea level with corre- grams and therefore, we only present representa-

Fig. 7. Diagrams of the directional terms of the relative P residuals in the western part of the region: RM = Rhenish Massif;

RG = Rhine Graben; MC = Massif Central: S = Saxothuringicum; M = Moldanubicum: A = Alps: P = PO Plain. The diagram in

the upper left represents the pattern of the average absolute residuals used for the normalization. For more detailed explanation see Figure 6.

TWE LITHOSPHERE IN CENTRAL EUROPE 155

tives of groups with a similar residual pattern. Figure 6 shows an example of individual diagrams for stations in the Massif Central. The pattern is not simple but very consistent regardless of the directional averages Ri (Table 5). In the western part of the Massif Central the fast directions point predominantly to the northwest in all stations. The eastern part of the Massif Central

seems to be slightly different, however, the negative residuals facing northwest still prevail. The consistency of the diagrams persists both in the northern part with a thick lithosphere and in the region of the lithosphere thinning in the central part (PYM, AGO, PLDF).

Many diagrams (Figs. 7 and 8) display highly organized patterns of groups of positive and neg-

IPP I (

Fig. 8. Diagrams of the directional terms of the relative P residuals in the eastern part of the region: S = Saxothuringicum;

M = Moldanubicum; A = Alps; P = PO Plain; PP = Polish Platform. The diagram in the upper right represents the pattern of the

average absolute residuals used for the normalization. For more detailed explanation see the caption of Figure 6.

1% V. BABUSKA AND J. I’LOMbKOVi


ative residuals, with maximum differences between 2 and 3 s at the majority of stations. Stations with similar patterns form groups which, in some regions, extend over several hundred kilometres. Often the diagrams have a consistent pattern of negative residuals for waves arriving at a station from one side and of positive residuals from the opposite side. This can be observed, e.g., at the stations in the Moldanubicum and in the northern part of the central and eastern Alps, where the waves from the south, southeast and southwest show mostly negative residuals (e.g., KHC, WET, FUR, BHG, MOA, Fig. 8). The residual patern changes to the north of the Eger Graben. Stations in the Saxothuringicum and farther to the north and northwest (BRG, MOX, CLL,) recorded fast arrivals for waves approaching stations mostly from the north and northwest. This pattern is observed also at stations in the Rhenish Massif and its northwestern surroundings (Fig. 7, e.g., BNS, MEM, ENN, DOU, SNF, UCC, WTS). Another region with consistent orientations of high velocities pointing to the northwest is the Massif Central (e.g., LPO, LFF, RJF, CAF, LSF, MAF, TCF, see also Fig. 6). It seems

that the pattern changes at station SSB situated at the eastern margin of this tectonic unit, but this station has scanty data in the period 1985 1986. Nevertheless, the residuals display an orientation pattern which is similar rather to that of the stations in the southern part of the western Alps (e.g., STV, ROB, CDR and LMR, Fig. 7) with high P velocities pointing to the northeast.

In Figure 9 we marked the regions with similar orientations of residual patterns. The Saxothu- ringicum, Rhenohercynicum and the Massif Cen- tral can be characterized by predominantly NNW-facing high velocities. The Moldanubicum as well as the deep lithosphere beneath the Alpine Foredeep and most of the Alps exhibit an opposite pattern-high velocities point to the southeast and south. This could be explained by a similar general orientation of large-scale seismic anisotropy in the deep lithosphere. The bipolar pattern of P residuals at groups of stations in particular tectonic units of central Europe was interpreted by BabuSka et al. (1984b) by systems of dipping anisotropic structures with olivine crystals orientated predominantly with a axes along the dip of structures. From the residual differ-

TABLE 6

Directional averages ki at stations in the eastern part of the region

Station

ARV

r?, w

-0.14

Mean

error

kO.18

N&

16

Station

LJU

R, (s)

-0.13

Mean

error

f 0.08

NAZ

16

ASS

BDI

BE0

BHG

BRG

BRN

BUD

CLL

CRE

CTI

FUR

GRF

HAM

HOF

KBA

KHC

KRA

KSP

0.28

- 0.37

0.10

- 0.01

- 0.23

0.85

- 0.03

- 0.37

0.21

- 0.27

0.27

0.03

0.63

- 0.09

- 0.47

- 0.24

- 0.22

- 0.23

0.28 13 LLS 0.22 15 MD1

0.10 16 MOA

0.15 16 MOX 0.10 16 NIE 0.12 14 OGA 0.13 14 oss 0.09 16 PRU 0.26 14 RAC 0.11 16 SAL 0.14 16 SAX

0.14 16 SOP 0.23 9 TRI 0.11 16 VKA 0.13 16 WAR 0.11 16 WET 0.09 16 ZAG 0.07 16 ZST

- 0.06 0.13 16

- 0.29 0.17 15

- 0.44 0.14 16

0.09 0.16 14

0.43 0.10 16

- 0.07 0.14 15

0.12 0.11 16

- 0.26 0.11 16

0.60 0.23 13

0.27 0.14 13

- 0.09 0.10 16

-0.14 0.11 16

- 0.21 0.12 16

- 0.05 0.10 16

0.36 0.26 11

- 0.09 0.15 16

0.32 0.08 16

- 0.07 0.13 16

158 V BABUSKA AND .I PLOMhKOVbl

ences the authors estimated P-wave anisotropies

between 6 and 11% which is in accord with the

laboratory determined anisotropies in olivinic

rocks (Table 3).

Let us have a closer look at those seismologi-

cal stations where data for both the shear- and

compressional-wave anisotropies are available.

Besides the diagrams of the directional terms of

the relative P-wave residuals, Figure 9 shows also

the directions of polarizations of the fast SKS

waves (azimuths are in Table 7) determined from

polarization analyses at seismological stations

STU (Kind et al., 1985), GRF (Silver and Chan,

1988), KHC and KSP (Makeeva et al., 1990), SSB

and NE15 (NARS, Vinnik et al., 1989). For the

last station we did not have any P-wave data,

therefore we used the P residuals of nearby sta-

tion ENN to compare estimates of P- and S-wave

anisotropies which are summarized for five sta-

tions in Table 7. When computing the P-ani-

sotropy beneath SSB we mainly considered the

diagrams constructed from an older set of data

(BabuSka et a1.,1984b), because, as mentioned

above, this station has scanty data in the period

1985-1986.

grams of P residuals (Figs. 6-9) can be inter-

preted by dipping anisotropic structures with high

and low velocities inclined in opposite directions

(BabuSka et al., 1984b) with the dip direction

approximately perpendicular to the azimuth which

divides the positive and negative residuals in the

bipolar diagrams and points to a side of negative

residuals. To compute the P-wave anisotropy

which corresponds to such a model, we evaluated

average negative and average positive residuals

and their differences T (see Table 7) in the oppo-

site parts of the diagrams of stations GRF, SSB,

KHC and ENN. Station KSP was not included as

its pattern is not of a bipolar character (see Fig.

8). The anisotropy coefficients k were computed

from the differences between these averages. To

determine the lengths of the ray paths, we used

the model of the lithosphere thickness (see Fig.

2). For simplicity we assumed that velocity ex-

tremes dipped at 45” and took rl,(min) = 7.9

km/s. An increase or a decrease of this value has

only a small influence on the resulting k as we

work with the residual differences.

The anisotropy is calculated for the wave prop-

agation in the subcrustal lithosphere of the thick-

nesses given in the previous section. The dia-

grams of SSB, GRF, KHC and ENN can be

characterized as bipolar with prevailingly early

and late arrivals from opposite directions. The

patterns are nearly symmetric to a subvertical

plane, whose orientation is close to the direction

of the particle motion of the fast S waves deter-

mined from SKS polarizations. As mentioned

above, the observed bipolar pattern in the dia-

The range of P-velocity anisotropies deter-

mined at four stations (Table 7) is between 6.5

and 15.2% which is comparable with the range of

anisotropies determined by laboratory measure-

ments on samples of dunite (7.2-lS%), lherzolite

(3.8-11.1%) and olivinite (3.1-13.2%, Table 3).

Also the S-wave anisotropies (2.2-6.7%) deter-

mined from SKS polarization analyses mentioned

above (c in Table 7 represents a delay of the

slow SKS relative to the fast SKS wave) have

similar variations to those determined by differ-

ent authors in samples of olivinic ultramafites

(1.4-9.2%, Table 4). The average P-wave

TABLE 7

P- and S-wave anisotropy coefficients (k) in subcrustal lithosphere

Station

GRF

KHC

SSB

ENN

NE15

Thickness (km)

crust Subcrustal

lithosphere

28 70

34 106

30 90

25 55

2.5 55

P waves

7 (s)

1.4

1.2

1.2

I .4

k(S)

11.x

6.5

7.x

IS.2

S waves

Azimuth

79”

90

140

70”

6 (s)

1

I.7

I

0.25

k (7;)

6.1

4.9

5.3

2.2


anisotropy determined beneath four seismological stations is 10.3% which is remarkably close to the mean value 9.8% of the average anisotropies computed from olivinic upper-mantle rocks- dunite, Iherzolite, olivinite and harzburgite (Ta- ble 3). Therefore, it can be concluded that the P-velocity anisotropy around 10% is a typical value of the subcrustal lithosphere, as derived both from the banch-scale tests of the rock samples and from the large-scale analysis based on the spatial variations of P residuals.

The average S-wave anisotropy of 4.8% determined from the SKS polarizations (Table ‘7) is lower than the average anisotropy derived from laboratory data (6%, Table 3). This is probably because the SKS analyses represent only subvertical wave propagation and only exceptionally can they coincide with the directions of velocity ex- trema whereas the laboratory measurements are made in several directions of wave propagation and for different polarizations.

The polar diagram of station KSP (Fig. 8) is difficult to interpret in terms of the bipolarity observed in the diagrams of other stations shown in Table 7 and, therefore, the P-wave anisotropy was not determined. Nevertheless, the SKS analysis proved large-scale anisotropy to exist beneath this station (Makeeva et al., 1990).

Discussion

In the section “Model of lithosphere thickness” we showed that laboratory data on the upper- mantle rocks demonstrated that velocity variations due to anisotropy were greater than those due to changes in composition. It can be expected that due to systematic mineral orientations (Christensen, 1984) large-scale anisotropic structures should play an important role in velocity variations derived from seismological observations. Seismic anisotropy in the upper mantle of central and western Europe was determined from independent observations of many authors: by Bamford (1973) and Fuchs (1977) from azimuthal variations of P, velocities; by Kind et al. (1985), Silver and Chan (1988), Vinnik et al. (1989) and Makeeva et al. (1990) from SKS polarization analysis; by BabuSka et al. (1984b) from spatial

variations of relative P residuals; by Wielandt et al. (1987) and Yanovskaya et al. (1990) from the dispersion of surface waves.

While the phenomena related to shear-wave splitting are generally accepted as a direct proof of anisotropy, a possibility of determining a large-scale anisotropy only from a systematic dependence of P-velocity variations is difficult to prove unambiguously. As an example, the system- atically early arrivals at seismological stations in the northern part of the Bohemian Massif and farther to the north for waves approaching stations from the north were interpreted by BabuSka et al. (1984b) as inclined anisotropic structures and by Wylegalla et al. (1988) as an inclined 400&m discontinuity dipping to the south. In combination with other independent anisotropy observations the explanation by inclined anisotropies seems to be more probable.

A preferred orientation of olivine crystals is generally accepted as the main cause of seismic anisotropy in the uppermost mantle both of oceans and continents. Fuchs (1983) related the anisotropy of P,, waves in southern Germany to the present stress field. On the other hand, BabuSka et al. (1984b) explained the anisotropic propagation of P waves in the deep lithosphere of central Europe by “frozen-in” olivine orientations inherited from an ancient oceanic lithosphere. The P-residual variations presented in this paper do not correlate with the stress field in the crust of central Europe (Mastin and Miiller, 1989) and therefore, we prefer an interpretation by systems of inclined anisotropic structures in the deep lithosphere which may represent paleo- subduction zones with inherited olivine orientations (BabuSka and Plomerova, 1989).

Au important question is the depth range of seismic anisotropy. As was mentioned above, the main cause of the large residual variations has to be sought in the mantle. The anisotropies calculated for the subcrustal lithosphere (section 4) with thicknesses derived from the representative residuals are in accord with the laboratory-determined anisotropies of olivinic ultramafic rocks (Table 4). The change in orientation pattern of P residuals at important tectonic boundaries, e.g., between the Moldanubicum and Saxothuring-

160 L’ HABLJSKA AND .I. 1’LOMtROV.i

icum, indicates that a transition between different orientations of anisotropic structures should not be very deep in the mantle. We can expect that due to different rheologies of the lithosphere and the asthenosphere, inherited anisotropy caused by “frozen-in” olivine orientations can be retained probably onIy in the rigid subcrustal lithosphere. Deeper in the mantle the large seismic anisotropy should mainly reflect a present convection flow.

In Figure 9 we distinguished two general patterns of the directional dependendence of P residuals. The stations in the Saxothuringi~um, the Rhenish Massif. the Belgo-Dutch Platform and the western part of the Massif Central belong to the “northern” pattern which is characterized by northwesterly oriented high-velocity directions. The “southern” pattern is oriented in the opposite direction and is characteristic for stations of the Moldanubicum, the Alpine Foredeep, the central and eastern Alps. This suggests that the subcrustal Iithosphere beneath a great part of the Alps is similar to the orientation of anisotropic structures of the deep lithosphere of the Moldanubicum, which we interpret by systems of paleosubductions. As a number of Archean com- ponents were found in the Moldanubian crust (Gebauer et al., 19891, some of the assumed paleosubductions may have taken place already in the Archean. The whole region of similar orientations of anisotropies may have continued in its development with a series of later subductions and accretion events (BabuSka and Plomerova, 1989).

The similarity of the residual patterns at stations in the Moldanubicum with the Alpine stations north of the Insubric line supports the idea of Matte (1986) that most of the basement of the Alps in the Tauern, the French-Swiss Alps and the western part of the Carnic Alps, and in our opinion most of the whole lithosphere, correspond to internal parts of the Variscan belt.

The orientation pattern of seismological stations in the southern part of the western Alps and their surroundings differs from the “MoIda- nubian” pattern which is characteristic for most of the Alps (Figs. 7 and 8). The intensive bending of the western part of this mountain chain corre-

sponds to the change of the residual patterns- the high velocities in the Swiss Alps point to the southeast whereas in the Italian-French Alps the high velocities are orientated toward the northeast.

If the anisotropy in dipping structures is due to the olivine orientation created in an ancient oceanic lithosphere, then each of the orientation patterns represents a region with a similar orientation of paleosubductions. Behr et al. (1984) and Matte (19861 proposed subduction zones with convergent polarities in the Variscides of central Europe whereas we derived from the directional terms of P residuals divergent systems of paleosubductions relative to the suture between the Saxothuringicum and Moldanubicum. The opposite dip of the assumed paleosubductions compared with our model is preserved also in the Rheno-Hercynicum. These authors assume a southward orientated subduction in the Rheno- hercynian basin whereas our diagrams for the stations in the Rhenish Massif show northwesterly orientated pale~)subductions (Fig. 71.

The bounda~ of the opposite orientation patterns (Figs. 4 and 9) represent an important suture between the Saxothuringicum and Moldanu- bicum, which was also detected by the Dekorp Research Group (1985). Outcrops of mafic rocks and serpentinized harzburgites along the so-called Ebendorf line (Behr et al., 1984) support the interpretation of this boundary as an important oceanic suture (Matte, 1986). The polarizations of the structures in the deep lithosphere of the Saxothuringi~um and the Moldanubicum~ as determined by the high P-velocity directions (Fig. 9) are similar to those of near-surface crustal structures and nappes which have a north to northwest vergence in the Saxothuringicum and a south to southeast vergence in the Moldanubicum (Franke, 1986).

The existence of large-scale anisotropy in the deep lithosphere has been demonstrated by independent seismological observations not only in central Europe. It is thus a great challenge for a near future to search for anisotropic models which would fit different observations of seismic anisotropy and to use such models in geotectonic interpretations.


Acknowledgements

We are grateful to two anonymous reviewers for their comments and to R. Freeman for critical reading of the manuscript and his careful edito- rial work.

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The lithosphere in central Europe---seismological and petrological aspects

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