Geophys. J. Int. (2007) 169, 675–682 doi: 10.1111/j.1365-246X.2007.03361.x

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FA S T T R A C K PA P E R

Upper mantle beneath the Eger Rift (Central Europe): plumeor asthenosphere upwelling?

Jaroslava Plomerova,1 Ulrich Achauer,2 Vladislav Babuska,1 Ludek Vecsey1

and BOHEMA working group1Geophysical Institute, Academy of Sciences of the Czech Republic, Bocnı II/1401, 141 31 Prague 4, Czech Republic. E-mail: jpl@ig.cas.cz2EOST – Institut de Physique du Globe Strasbourg, R. Descartes 5, Strasbourg, France

Accepted 2007 January 10. Received 2007 January 2; in original form 2006 March 21

S U M M A R YWe present the first results of a high-resolution teleseismic traveltime tomography and seis-mic anisotropy study of the lithosphere–asthenosphere system beneath the western BohemianMassif. The initial high-resolution tomography down to a depth of 250 km did not imageany columnar low-velocity anomaly which could be interpreted as a mantle plume anticipatedbeneath the Eger Rift, similar to recent findings of small plumes beneath the French MassifCentral and the Eifel in Germany. Alternatively, we interpret the broad low-velocity anomalybeneath the Eger Rift by an upwelling of the lithosphere–asthenosphere transition. We alsomap lateral variations of seismic anisotropy of the mantle lithosphere from spatial variations ofP-wave delay times and the shear wave splitting. Three major domains characterised by differ-ent orientations of seismic anisotropy correspond to the major tectonic units—Saxothuringian,Moldanubian and the Tepla-Barrandian—and their fabrics fit to those found in our previousstudies of mantle anisotropy on large European scales.

Key words: Eger Rift, seismic anisotropy, tomography.

1 I N T RO D U C T I O N

The Tertiary Eger (Ohre) rift (ER), a 300-km long ENE–WSW strik-

ing structure in the western Bohemian Massif (BM), is a part of

the European Cenozoic Rift System (ECRIS, e.g. Prodehl et al.1995). The ER developed along an older Variscan suture between

the Saxothuringian (ST) and the Tepla-Barrandian unit (TBU), as

an eastern continuation of the ST and Moldanubian (MD) contact

(Fig. 1). Granet et al. (1995) suggested that the ECRIS may have

a common source of a ‘plume-like’ volcanism in the mantle, mani-

fested by the existence of ‘baby-plumes’ beneath the Variscan Mas-

sifs. Such relatively narrow plumes have been found beneath the

French Massif Central (Granet et al. 1995) and the Eifel (Ritter

et al. 2001).

Active tectonics of the western ER is primarily manifested by

Cenozoic volcanism (Ulrych et al. 2000), gas emanations contain-

ing high proportions of mantle derived CO2 and He (Weinlich et al.1999), as well as by the periodic occurrence of shallow earth-

quake swarms (Fischer & Horalek 2003) and neotectonic crustal

movements (Bankwitz et al. 2003). The question of a possible

source of ascending mantle fluids is important namely in relation

to the increase of 3He/4He ratios between 1993 and 2003, observed

by Brauer et al. (2005). The authors interpret the observation as

an evidence for ascending mantle-derived melt and as an indica-

tion of a possible future unrest in a presently inactive volcanic

region.

The concentration of various presently ongoing geodynamic phe-

nomena in a small area of the western ER calls for a study focused

on the interaction between near-surface tectonics and geodynamic

processes at depth. The BOHEMA project (BOhemian Massif HEt-

erogeneity and Anisotropy, Babuska et al. 2003) brought together

geoscientists from the Czech Republic, Germany and France for a

cooperative study of the structure and dynamics of the lithosphere

and asthenosphere of this region. Central to the BOHEMA project

was a passive seismic experiment in 2001–2003. In this short paper,

we present initial results of the seismic tomography and velocity

anisotropy study focusing on the characterization of the structure of

the upper mantle beneath the major tectonic units and on the pos-

sible existence of a ‘plume-like’ low-velocity structure below the

ER.

2 T O M O G R A P H I C I M A G I N G O F T H E

U P P E R M A N T L E B E N E AT H T H E

B O H E M I A N M A S S I F

Deviations in seismic wave speed relative to a radial earth velocity

reference model result in isotropic images of the velocities in the

Earth interior. We present tomographic images of the upper mantle

C© 2007 The Authors 675Journal compilation C© 2007 RAS

676 J. Plomerova et al.

Figure 1. Simplified map of the Bohemian Massif showing the major tectonic units and faults. Two lines locate the cross-sections presented in Fig. 2(c).

beneath the BM down to 250 km, with the highest resolution in

its western part, beneath the BOHEMA array (Figs 2a and b). The

BOHEMA array of seismic stations covered a territory of about

270 × 150 km, with its long axis oriented perpendicularly to the

strike of major tectonic units and to the ER. The network con-

sisted of 61 permanent and 92 temporary stations operating be-

tween 2001 October and the end of 2003, with a core of recordings

in 2002 (Plomerova et al. 2003). About 2/3 of the network consisted

of three component broad-band stations with the remaining being

short-period stations. Spacing of stations was generally less than 30

km, while in the central part of the array the spacing was as dense as

10–15 km, hence allowing for a lateral resolution of approximately

20 km in the upper mantle.

For this preliminary 3-D seismic tomography about 80 per cent

of the total data were analysed, resulting in 13 506 P-wave arrival

times with high precision. Arrival times of the first coherent ex-

tremes (‘relative picks’) were measured manually across the array

along with arrival times of the first sharp breaks (‘absolute picks’)

on a small number of highest-quality recordings. Then the absolute

arrival times were calculated at the stations. A waveform correlation

with the ‘Seismic Handler’ software (Stammler 1993) was applied

in case of less clear waveforms. Based on the estimated accuracy of

the timing, the data were sorted into three quality classes (±0.05,

±0.1 and ±0.2 s) and weighted accordingly in the inversion and in

the P-residual analysis. A good coverage of azimuths and ray pa-

rameters was achieved with 230 teleseismic events (see Fig. 1). For

our P-wave model we use 151 stations with at least 10 observations

each. P-wave traveltime residuals are calculated by subtracting the

expected traveltime of the IASP91 model (Kennett & Engdahl 1991)

from the observed traveltime. To exclude traveltime effects due to

heterogeneities in the source region or event mislocation, we sub-

tract the mean residual of each earthquake. This results in relative

residuals which are mainly generated by seismic velocity perturba-

tions underneath the station network. These relative residuals vary

from −0.7 to +0.7 s at our stations. Since we used a number of

different acquisition systems and types of seismometers (short pe-

riod as well as broadband (Plomerova et al. 2003), all recordings

were normalized by deconvolving the seismometer characteristics

and convolving the signal with a WWSSN-seismometer type char-

acteristic before starting the picking in order to minimize traveltime

effects related to acquisition system differences.

We use a 3-D tomographic inversion procedure to map the seismic

velocity variations in the mantle, known as the ACH-method (Aki

et al. 1977). While our updated version of this method (Evans &

Achauer 1993) is not as sophisticated as some of the more recent

codes, it does have some benefits missing in newer codes. Not only

is the program very robust, but also it does allow for a thorough

resolution analysis since it uses a full matrix inversion approach,

hence allowing for the calculation of the posteriori covariance and

resolution values, and could therefore, be regarded as a first-step

tool before carrying on with more detailed analysis. In addition,

does it contain an average-and-offset scheme (Evans & Achauer

1993) which serves as a smoothing filter and diminishes possible

effects from the artificial block boundaries. Our model consists of

nine layers with 20 (crustal layers) 60 (deepest mantle layer) km large

horizontal blocks, enlarging with depth down to 380 km. The starting

model is based on a 1-D local crustal velocity model derived from

Deep Seismic Sounding (DSS) measurements (Beranek & Zatopek

1981; Mayerova et al. 1994; Hrubcova et al. 2005) and receiver

function methods (Geissler et al. 2005; Wilde-Piorko et al. 2005),

and the 1-D IASP91 velocity distribution (Kennett & Engdahl 1991)

just underneath. The smooth crustal model assumes average crustal

velocity, depth of Moho, as well as thickness and average velocity

of unconsolidated sediments.

Tests were made for inversions with and without station term

corrections and with and without upper crustal corrections. For the

C© 2007 The Authors, GJI, 169, 675–682

Journal compilation C© 2007 RAS

Upper mantle beneath the Eger Rift 677

Layer 2: 15 - 35 km

Layer 3: 35 - 70 km

Layer 4: 70 - 105 k m

Layer 5: 105 - 141 km

Layer 6: 141 - 177 km

Layer 7: 177 - 251 km

P velocity perturbations

-5 %0

Figure 2. Plane views (a) of tomographic images of smoothed velocity perturbations (preferred model, including station term and upper crustal velocity

corrections—see text for details), along with the seismic stations (triangles); (b) Diagonal elements of resolution matrix; unresolved blocks are shown in light

grey, well resolved blocks are in dark grey; (c) Vertical cross-sections through the velocity perturbation model along and across the Eger Rift (see Fig. 1 for

their locations). Higher (lower) than normal velocities wrt the IASP91 reference model are shown in blue (red) colours.

corrections a VELEST-type (Kissling et al. 1994) 1-D inversion for

a single layer (5 km thick) and/or a two-layer case (15 km thick) was

carried out to search for the maximum residual reduction due to the

heterogeneities directly beneath the stations. See Evans & Achauer

(1993) for details of this approach, called ‘layer-stripping’. The

velocity heterogeneities found by this approach were included as

a priori information (‘station terms’) in the ACH inversion

procedure.

Our chosen 3-D velocity model explains 70.6 per cent of the mea-

sured traveltime residuals with the diagonal elements of the resolu-

tion matrix being larger than 0.7 for most parts of each layer, but

the edges and the southeastern part of the model (see the marked re-

gions with resolution values larger than 0.7 in Figs 2a and b). It also

contains corrections for station terms and upper crustal velocity het-

erogeneities. The corrections influence mainly the two crustal and

the two uppermost mantle layers (smearing between crust and man-

tle). Anomalous signal in the uppermost mantle gets smaller, while

there’s essentially no influence on the deeper layers. In Figs 2(a) and

(c) we present perturbations of our preferred model of the compres-

sional wave speed relative to the IASP91 model as depth slices and

on two vertical cross-sections.

In order to get a better handle on the resolution limits of our

model and possible errors due to vertical smearing between layers,

we did a number of synthetic simulations, including checkerboard

tests and tests for models resembling structures of ascending plumes

like published for the Massif Central and the Eifel region (Granet

et al. 1995; Ritter et al. 2001). In addition, we did a number of

simulations in which we changed the depth-range and the depth of

the top of the synthetic anomaly to evaluate the amount of vertical

smearing. For all our synthetic modelling cases we used a 3-D ray

tracer (Koch 1985) and the same data distribution as in the real

data set. Realistic errors were added to the data using a Gaussian

error routine. Our synthetic models confirmed that with the available

station-data combination we are able to distinguish between cases

C© 2007 The Authors, GJI, 169, 675–682

Journal compilation C© 2007 RAS

678 J. Plomerova et al.

Layer 2: 15 - 35 km

Layer 3: 35 - 70 km

Layer 4: 70 - 105 k m

Layer 5: 105 - 141 km

Layer 6: 141 - 177 km

Layer 7: 177 - 251 km

0.0 0.3 0.5 0.7 0.9 1.0

Diagonal elements of resolution matrix

Figure 2. (Continued.)

with and without an ascending baby plume in the area of interest

and dense station spacing. The tests also showed that there is quite a

significant amount of smearing along ray paths into the uppermost

layers (layers 2 and 3) with the anomalies situated in layer 4 (and

less so in layer 5), due to the subvertical incident angles of the

teleseismic rays, getting larger to the edges of the model and that,

therefore, one has not a good control on crustal structures. In these

cases the smearing might arrive at almost 50 per cent of the original

amplitude. For cases where the top of a possible anomaly is at 150 km

or deeper, the smearing is essentially restricted to adjacent layers

with the amplitudes there being of the order of 20 per cent of the

original synthetic amplitude.

Velocity perturbations in the tomographic images, corrected for

upper crust and station terms (Fig. 2a), include the entire BM and

exhibit some features which correlate with main tectonic structures.

The perturbations show a difference in crustal thickness of the ST

and MD, the latter having the Moho deeper (layers 2 and 3), and

also a southward thickening of the mantle lithosphere beneath the

MD (layers 4, 5 and 6). In general, the velocity perturbations are

small (exceptionally exceed 2 per cent) and do not cluster to image

large-scale velocity heterogeneities in the upper mantle.

In particular, we were searching for a narrow and vertical low-

velocity anomaly beneath the ER, which would fit the ‘baby-plume

concept’ below the Variscan Massifs (Granet et al. 1995; Ritter et al.2001). However, the P-velocity tomography of the upper mantle

below the BM does not image any distinct heterogeneity down to

250 km which could be unambiguously interpreted as a magma

chamber or a mantle plume. There are weak negative perturbations in

the western part of the ER, which can be traced down to the bottom of

the model. However, they do not form a tube-like low-velocity body,

which could be associated with an uprising plume below the rift.

Moreover, there are comparable well resolved negative perturbations

in other parts of the model.

To show velocity variations related to the ER, we construct two

cross-sections through the velocity perturbation model—along and

across the rift (Fig. 2c). The perturbations were smoothed in a band

C© 2007 The Authors, GJI, 169, 675–682

Journal compilation C© 2007 RAS

Upper mantle beneath the Eger Rift 679

% % %

P-velocity perturbations

Along the Eger Rift

Across the Eger Rift

’

’

Figure 2. (Continued.)

30 km wide along each profile located in Fig. 1. The width of the

bands corresponds to the size of the mantle blocks in the tomography.

The low-velocity anomaly forms a broader ridge which parallels the

rift axis and not a dome structure, which would reflect an interaction

of a deep seated mantle plume with the lithosphere. Alternatively, we

can interpret the laterally elongated low-velocity region below the

ER as an asthenosphere updoming (a lithosphere thinning), mapped

also in a larger-scale model of the lithosphere thickness of central

and western Europe based on data of permanent seismic observato-

ries (e.g. Babuska et al. 1987; Babuska & Plomerova 1992, 2006),

or in a detailed seismic study of the ST-MD mantle lithosphere con-

tact (Plomerova et al. 1998; Babuska & Plomerova 2001). At our

present stage of knowledge, we leave the question of the existence

of a hypothetical plume in the upper mantle deeper than 250 km

open.

3 F RO Z E N A N I S O T RO P Y I N B L O C K S

O F M A N T L E L I T H O S P H E R E

Besides the 3-D isotropic velocity tomography, body-wave velocity

anisotropy of the western part of the BM was studied. We have ap-

plied the method by Babuska and Plomerova (1992; see also Sıleny

& Plomerova 1996), which has already been used successfully in

other data sets and different regions (Babuska & Plomerova 2006,

for a review). To study the spatial variations of P velocities we

construct P residual spheres for individual stations. The spheres

show that part of the relative residuals (directional terms) which

depends on the direction of propagation through the lithosphere.

These terms are obtained by subtracting a directional mean at each

station computed as an average relative residual filtered in the az-

imuth incidence angle space. The directional mean represents an

isotropic velocity in the volume beneath a station. As the direc-

tional mean forms a reference level in each sphere, the spheres can

be compared regardless of differences in the lithosphere thickness.

The directional terms are smoothed and plotted in polar projection

of lower hemisphere, in which the ray azimuth and angle of prop-

agation within the mantle lithosphere are the parameters. Often we

observe so called ‘bipolar’ pattern of the P spheres, which means

that negative directional terms consistently prevail in one half of the

hemisphere, while positive ones dominate in the other. We primar-

ily assume that only long wavelength variations in relative residuals,

normalized by a station directional mean, can be related to mantle

lithosphere anisotropy. The negative and positive residuals thus can

be related with relatively high- and low-velocity directions within

the lithosphere beneath individual stations relative to the average

velocity beneath each station, linked with the directional means.

The method of relative delay times analyzed from point of view of

seismic anisotropy integrates into a family of methods used to im-

age continental fabric, none of which alone has the ability to fully

characterize seismic anisotropy (Fouch & Rondenay 2006).

Of course, there is a trade-off between anisotropy and velocity

heterogeneities for body-wave traveltimes. In general, the trade-off

makes it difficult to separate out the respective contributions and we

can rely only on indirect indications. Effects of non-compensated

small-size or distant heterogeneities appear as localized perturba-

tions in otherwise simple bipolar or ‘no’ pattern of the spheres (see

also Fig. 3 and its caption). The former pattern extends laterally

up to several hundreds kilometres, while the latter is often linked

with prominent tectonic sutures or dominant faults cutting the whole

lithosphere (Babuska & Plomerova 2006). Therefore, it would be

difficult to explain sudden changes of the P-sphere pattern, associ-

ated with prominent tectonic features, by velocity heterogeneities

in the deep mantle. If it were acceptable for a small-size region,

with the lateral extend of few times of the lithosphere thickness,

the explanation by velocity heterogeneities in the sublithosphere

mantle would be impossible in case of large regions several times

exceeding the lithosphere thickness. Such characteristics of the Pspheres, described above, along with geographical and backazimuth

variations of the shear wave splitting allow us to associate the di-

rectional velocity variations with seismic anisotropy of the mantle

lithosphere.

Three different types of residual patterns were recognized in the

western BM (Fig. 3). Regions with similar patterns delimit the man-

tle lithosphere domains defined by the P-wave anisotropy. The ST

pattern exhibits the negative residuals for northern propagation di-

rections, the MD pattern is characterized by negative residuals (rel-

atively high-velocity directions) for southern propagation, and the

TBU pattern by negative residuals for waves arriving from NE–SE.

Divergently dipping fabric of the ST and MD mantle lithospheres

was also modelled in the western rim of the BM (Plomerova et al.1998). A pilot study of the mantle lithosphere anisotropy of the BM

(Plomerova et al. 2005) indicated that a fabric of lithosphere root

below the TBU, with the high velocities plunging dominantly to

the E, differs from those of the ST and of the MD, where the high

velocities plunge to the NE–NW and to the S, respectively. The

densely spaced observables of P-wave anisotropy extracted from

the BOHEMA array confirmed the existence of three differently ori-

ented anisotropic structures in the upper mantle and allowed us to

map their approximate boundaries. Stations with a similar P-pattern

C© 2007 The Authors, GJI, 169, 675–682

Journal compilation C© 2007 RAS

680 J. Plomerova et al.

Figure 3. Examples of smoothed P-residual spheres, reflecting directional dependence of relative P velocities in the mantle lithosphere beneath stations (relative

means at stations subtracted), are shown for three regions with different orientation of fabric in the mantle lithosphere. Blue triangles and red circles stand

for negative and positive relative residuals, respectively, crosses mark residuals close to zero. Colour dots in the map stand for stations with one of the three

P-sphere patterns (red—Saxothuringian, dark blue—Moldanubian, light blue—Tepla-Barrandian), open circles mark stations with less clear pattern. Stations

with no pattern or with very small values of P residuals in the P spheres are marked by open black circles. The stations with a similar P-pattern cluster into

domains closely related to the three tectonic units.

cluster into domains closely related to the three tectonic units. The

southwestern boundary of the fabric of the mantle domain below

the TBU is projected on the surface near the Marianske Lazne Fault

(MLF) and to the northern part of the West-Bohemian Shear Zone

(WBSZ, see Fig. 1). The northwestern and southeastern boundaries

of the TBU domain parallel the ER and the CBSZ with offsets of

about 10–20 and 40 km, respectively.

Teleseismic shear wave splitting, the analogy of the light birefrin-

gence, is generally considered as a proof of upper mantle anisotropy,

and represents the most common method to study the seismic

anisotropy (Silver 1996; Savage 1999, for reviews). Elliptical po-

larization of the core–mantle refracted shear waves (SKS, SKKS),

exhibiting a non-zero content of energy on the T component, re-

sults from an interference of the split fast- and slow-, or SV and

SH (vertically and transversely polarized) shear waves generated

in an anisotropic mantle. To evaluate the splitting parameters—

polarization of the fast shear wave ψ defined by two Euler angles

(azimuth φ and inclination θ—measured upward from the positive

axis z oriented downward), and the time delay δt of the slow shear

wave, we use a method which minimizes the energy on the trans-

verse component (Sıleny & Plomerova 1996), originally developed

by Silver & Chan (1991) for evaluation of azimuthal anisotropy of

shear waves. We search for the fast polarization direction ψ in the

plane (Q-T) perpendicular to the ray path plane (L-Q) of the shear

phase. A rotation of the coordinate system in the plane (Q-T) by

angle ψ , and time shift δt imposed on the shear wave components,

yield the splitting parameters determined in 3-D. The rotation and

the time shift transform the elliptical particle motion in the L-Q-T

coordinate system into a linear particle motion.

Fig. 4 illustrates lateral and azimuthal variations of the fast shear

wave polarizations and time delays of the slow shear wave. We show

here only examples of the preliminary measurements of the splitting.

Azimuths of the polarized fast shear waves are coherent within the

SKS SKKS

SKS021012_2009:

011012_1502:

Shear-wave splitting

Backazimuth

Backaz

imuth

Saxothuringian

Moldanubian

Teplá-Barrandian

Figure 4. Azimuths of the fast shear wave polarizations and split delay

times of SKS and SKKS phases from two events with almost opposite back-

azimuths. The anisotropic parameters are shown as vectors as they are eval-

uated in the LQT ray parameter coordinate system with the use of all three

components.

tectonic units, but differ for the two events with foci at approximately

opposite azimuths relative to the region. Moreover, we have found

below the TBU and MD different polarizations for SKS and SKKS

phases from the same event, that is, for waves propagating through

C© 2007 The Authors, GJI, 169, 675–682

Journal compilation C© 2007 RAS

Upper mantle beneath the Eger Rift 681

the upper mantle with the same azimuth, but at slightly different

incident angles. Distinct dependence of the splitting parameters on

direction of propagation through the upper mantle shows that the

mantle fabric can hardly be approximated by a single layer with

horizontal ‘fast’ symmetry axis (e.g. Plenefisch et al. 2001). On the

other hand, joint interpretation of variations of the splitting with

backazimuth along with the P spheres leads to 3-D self-consistent

anisotropic models compatible for both types of independent obser-

vations of velocity anisotropy (Plomerova et al. 1998, 2005).

Babuska et al. (2002) and Plomerova et al. (2001) showed that

shear wave splitting diminishes around boundaries of domains with

different fabrics. The anisotropic signal is weak, unstable or ap-

pears as isotropic signal (null splits), if the broad-band shear wave

samples different fabrics on both sides of a boundary of upper man-

tle domains. Such a tectonic setting may explain the splitting at

station NKC, located close to the ST-TBU contact (ER). This sta-

tion shows nulls (Fig. 4; green and blue circles at 50.2N 12.4E),

or very weak splits (Plomerova et al. 2005) for the event from the

NE. On the other hand, there is also a broad zone of weak or null

splits paralleling the NE trend of the TBU-MD southeastern limit at

depth (see also Fig. 3). The splitting delay times are much smaller

there for the event from the west than the delays observed at sta-

tions in the ST, or in the same region, but for an event from the NE

(Fig. 4).

A large difference in split times of shear waves arriving from

opposite directions to stations in the SE part of the TBU and its

transition to the MD, as well as the different orientations of the

fast shear wave polarizations of SKS and SKKS phases (Fig. 4),

indicate a complicated 3-D structure of the mantle lithosphere in

this region. The observations also suggest that a major source of

the varying seismic anisotropy is a frozen fabric within the mantle

lithosphere. An anisotropy caused by a flow in the asthenosphere,

coined for interpreting the splitting observations in central Europe

for example by Bormann et al. (1996) and Plenefisch et al. (2001),

probably also plays a role in the overall mantle anisotropy of central

Europe.

To retrieve a 3D self-consistent model of the mantle lithosphere,

we need to apply the joint inversion of both the spatial variations

of P velocities (P spheres) and of a larger set of the shear wave

splitting parameters and that will be the subject of a future study

of anisotropy of the BM. Nevertheless, the two independent types

of observations presented in this initial study on velocity hetero-

geneities and anisotropy of the upper mantle beneath the ER are

complemetary (for a review of integration of different data sets and

methods see Fouch & Rondenay 2006) and show laterally variable

fabrics of mantle lithosphere with generally inclined symmetry axes

of anisotropy.

4 C O N C L U S I O N S

The high-resolution teleseismic P-velocity tomography in the

Bohemian Massif does not image any distinct magma chamber or

a mantle plume down to 250 km beneath the Eger Rift, similar to

small plumes found by other authors beneath the French Massif

Central and the Rhenish Massif. However, the broad and weak low-

velocity anomaly beneath the western Bohemian Massif could be

a result of a deep seated mantle plume interacting with the litho-

sphere. However, alternatively, we can interpret the anomaly by

an upwelling of the lithosphere–asthenosphere boundary reflect-

ing a zone of a lithosphere weakness at the margins of palaeo-

plates. The elongated upwelling beneath the Eger Rift fits general

trends of the relief of the bottom of the European lithosphere. We

also interpret the varying anisotropic signal, reflected both in the

shear wave splitting and in spatial distribution of relative P residu-

als, by an olivine fabric frozen in different domains of the mantle

lithosphere. Three domains with different orientations of seismic

anisotropy in the mantle lithosphere of the Bohemian Massif cor-

respond to the Saxothuringian, Moldanubian and Tepla-Barrandian

unit.

A C K N O W L E D G M E N T S

We thank the members of the BOHEMA working group J. Zednık,

P. Jedlicka, J. Horalek, T. Fischer, A. Bouskova, B. Ruzek, P. Kolar,

V. Vavrycuk, M. Broz, J. Malek, V. Nehybka, O. Novotny, M.

Granet, U. Achauer, T. Piquet, M. Korn, S. Wendt, S. Funke, M.

Brunner, D. Rossler, R. Kind, H. Kampf, W. Geissler, B. Heuer,

K. Klinge, T. Plenefisch, K. Stammler, M. Lindemann, K. Brauer

and P. Malischewsky for cooperation during the project. We are

grateful to both anonymous reviewers for their constructive reviews

and helpful comments on the original manuscript. Financial support

by the Czech Grant Agency, project No. 205/04/0748, is greatly

acknowledged.

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