Acta Geophysicavol. 59, no.3, 2011, pp. 597–617

DOI: 10.2478/s11600-011-0009-1

Seismic Model and Geological Interpretation of theBasement Beneath the Doupovské HoryVolcanic

Complex (NW Czech Republic)

Jan VALENTA1, Milan BROŽ1, Jiří MÁLEK1, Bedřich MLČOCH2,Vladislav RAPPRICH2, Zuzana SKÁCELOVÁ2

and the Doupov Working Group*

1Institute of Rock Structure and Mechanics,Academy of Sciences of the Czech Republic, Prague, Czech Republic

e-mail: valenta@irsm.cas.cz (corresponding author)2Czech Geological Survey, Prague, Czech Republic

*Members of the Doupov Working Group are listed in Appendix.

Abstract

The Doupovské Hory Volcanic Complex (DHVC) is thebest-preserved large volcanic suite of the Cenozoic intraplatevolcanism in the Bohemian Massif. However, many uncertain-ties remain in the geological setting of its basement. In summer2008, two seismic refraction profiles ran across this area to re-veal the depth of the volcanic rocks and the underlying geolo-gical structure.

A high-velocity body was detected beneath the central partof the DHVC at depths of about –500 m a.s.l., inclined to theNW at an angle of about 45°. This body is most likely formedby the basic and ultrabasic rocks (amphibolites, eclogites, ser-pentinites, metagabbros) of the Equivalent of the MariánskéLázně Complex (EMLC). Subvolcanic rocks of the DIC penet-rating these basic and ultrabasic metamorphic rocks have verysimilar physical properties. Therefore, intrusive body of theDIC can be hardly distinguished from rocks of the EMLC byseismic refraction investigation.

Key words: seismic refraction, seismic tomography, DoupovskéHory Volcanic Complex, Eger Graben, Mariánské Lázně Com-plex.

________________________________________________© 2011 Institute of Geophysics, Polish Academy of Sciences

Author-created version

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1 INTRODUCTIONThe Doupovské Hory Volcanic Complex (DHVC), located in the

western part of the Eger Graben (Fig. 1), is a dominant volcanic suite ofthe Cenozoic intra-continental magmatism of the Bohemian Massif. Eventhough it is important for understanding the Eger Graben formation andassociated anorogenic magmatism, it remained mostly unknown for along time. New geological mapping at scales of 1:50 000 and 1:25 000 wascarried out in this area in the last 20 years (e.g., maps by Hradecký (ed.),1998, 2005 unreferenced; see Notes section) and was followed by a subse-quent set of studies published in the recent few years (e.g., Rapprich andHolub 2008, Skácelová et al. 2009, Haloda et al. 2010, Holub et al. 2010,Mlčoch and Konopásek 2010). Although the thickness of the volcaniccomplex probably does not exceed 800 m, the structure of its deeperparts and the basement can be reconstructed on the basis of the outcropsfound by geological mapping, several boreholes, located mostly at thenorth-western margin (targeted on mineral waters), crustal xenoliths in

Fig. 1: Location of the study area within the European Cenozoic Rift System(after Dèzes et al. 2004): BF – Black Forest, DHVC – Doupovské Hory VolcanicComplex, EG – Eger Graben, FP – Franconian Platform, HG – HessianGrabens, LRG – Lower Rhine Graben, OW – Odenwald, VG – Vosges, URG –Upper Rhine Graben.

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volcanic rocks, and geophysical data. The understanding of the formationand setting of such a volcanic complex located at the contact of Variscancrystalline domains was crucial in the need of more data. The benefits ofsystematic deep drilling would not be adequate to the expenses of suchresearch. Therefore, the study of the DHVC basement needed a differenttechnique to penetrate the volcanic rocks at much lower cost. Our re-search, based on seismic experiment, was aimed at filling this gap andcontributing to the model of the internal structure of the DHVC and itsbasement using a seismic model.

2 GEOLOGICAL SETTINGSThe DHVC is situated within the Cenozoic Eger Graben (Ohře Rift),

which belongs to the European Cenozoic Rift System (ECRIS; Dèzes etal. 2004). The Eger Graben is interpreted as an incipient rift system thatunderwent two distinct phases of extension. The first phase, characterisedby a NNE-SSW to NS oriented horizontal extension between the end ofthe Eocene and early Miocene, was oblique to the rift axis and caused thedevelopment of a fault system characterised by the en-echelon pattern ofE-W (ENE-WSW) faults. These faults defined a number of small, shallow

Fig. 2: Simplified geological map of the DHVC (Doupovské Hory VolcanicComplex): KHF – Krušné Hory (Erzgebirge) Fault, LF – Liboc Fault, PF– PetrovFault, SF – Střezov Fault.

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initial depocentres, characterised by very small subsidence rates that grad-ually merged during the growth and linkage of the normal fault segments(Rajchl et al. 2009). The origin of the paleo-stress field of the oblique, ex-tensional phase remains controversial and can be attributed either to theeffects of the Alpine lithospheric root (Michon et al. 2003, Michon andMerle 2005) or (perhaps more likely, because of the dominant volcanismat the onset of the Eger Graben formation) to the doming due to thermalperturbation of the lithosphere (Dèzes et al. 2004). The second, orthogo-nal extensional phase is explained by stretching along the crest of a grow-ing regional-scale anticlinal feature, which supports the recent hypothesisof lithospheric folding in the Alpine-Carpathian foreland (Dèzes et al.2004, Bourgeois et al. 2007, Rajchl et al. 2009).

The DHVC represents a moderately eroded multiphase volcanic com-plex. The activity initiated in the earliest Oligocene – mammal zone MP-21 (Fejfar and Kaiser 2005) with predominantly explosive activity(Hradecký 1997). The initial explosive activity produced a sequence of py-roclastic and resedimented pyroclastic deposits up to 80 m thick. The ac-

Fig. 3: Geology of the crystalline basement surface beneath the post-Variscansedimentary and volcanic rocks (adapted after Mlčoch 2001, 2003): DHVC –Doupovské Hory Volcanic Complex, EMLC – Equivalent of the Mariánské LázněComplex, KHF – Krušné Hory (Erzgebirge) Fault, LF – Liboc Fault, MLC –Mariánské Lázně Complex, PF– Petrov Fault, SF – Střezov Fault.

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tivity subsequently became predominantly effusive in character (Rapprichand Holub 2008) associated with weak Strombolian and possibly alsoHawaiian eruptions. The evolution of the DHVC was interrupted severaltimes by episodes of volcanic edifice decay, including sector collapses(Rapprich et al. 2007, Sakala et al. 2010). The volcanic activity lasted untilthe Early Miocene (about 20 Ma) (Sakala et al. 2010), when a group ofmonogenic cones has formed on the northern periphery of the DHVC.The total preserved thickness of the DHVC reaches 600 m.

The composition of lavas erupted within the DHVC is dominated byalkaline basaltic rocks including foidites, picrobasalts, basanites,tephrites, alkali (olivine) basalts and trachybasalts (Shrbený 1982, Rap-prich and Holub 2008). Differentiated rocks (trachyandesites, trachytesand phonolites) occur scarcely in the form of dikes, mostly trending NW-SE (Skácelová et al. 2009).

The alkaline intrusive rocks cropping out in the central part of theDHVC near the abandoned town of Doupov, the so-called Flurbühl intru-sion or Doupov Intrusive Complex (DIC) (Fig. 2), have been known forover 100 years (Wiesbaur 1901). The rock was originally described astheralite, and several geologists started to collect thin sections from thislocality, unfortunately without publishing their descriptions or interpreta-tions. Shrbenný (1982) published the first silicate analyses of five rocksamples from the subvolcanic complex of Flurbühl, but again with no in-terpretation added. This locality has become a focus of modern petrologi-cal research (Haloda et al. 2010, Holub et al. 2010). The first results haveshown a broad variety of intrusive rocks present within this intrusivebody cropping out on a relatively small area. The DIC comprises olivineijolite, ijolite, urtite, essexite, foid-monzonite and foid-syenite, and it wasemplaced into the early DHVC edifice about 29-30 Ma (Holub et al.2010).

The Eger Graben (Ohře Rift) follows the older Variscan suture be-tween the Saxothuringian Zone and the Teplá-Barrandian Unit (Jindřich1971, Kopecký 1986, Mlčoch 2003, Babuška et al. 2010, Mlčoch andKonopásek 2010) in the north-western part of the Bohemian Massif. Sev-eral geological units are combined in the basement of the DHVC in theform of a heterogeneous crustal mosaic. The Saxothuringian Zone ex-tends from the north as far as the Střezov and Petrov Faults in the base-ment of the DHVC (Fig. 2). The Variscan Nejdek–Eibenstock Plutonplunges beneath the DHVC from the west, and the rocks of the Teplá-Barrandian Unit form the base to the south. The basic and ultrabasicrocks of the Equivalent of the Mariánské Lázně Complex (EMLC) are sit-

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uated beneath the centralpart of the DHVC(Fig. 3). The EMLC con-sists of amphibolites,eclogites, serpentinitesand possibly alsometagabbros (Mlčoch2006). On the southeast,the DHVC covers Permo-Carboniferous continen-tal sediments of the ŽatecBasin limited by the Stře-zov (on the north) andLiboc (on the west) faults(Mlčoch 2003, Fig. 3).Sediments do not exceed800 m thickness and con-sist of alternating layersof claystones, siltstones,sandstones and conglom-erates (Mlčoch andMartínek 2002).

3 GEOPHYSICALSETTINGS

While the surface and near-surface geological structure of the DHVCis becoming well investigated, the deeper structure needs more data. Theonly geophysical research in the vicinity of the DHVC aimed to investi-gate deep structures was the SUDETES 2003 seismic refraction experi-ment (Guterch et al. 2003, Grad et al. 2008). Even though this seismicprofile missed the DHVC, it identified a large scale high velocity body atdepth greater than 2 km which most likely reaches also the DHVC area(Grad et al. 2008).

In order to contribute to the knowledge of the DHVC basement, theseismic refraction experiment was carried out at the end of June 2008.Five quarry blasts were used as sources of seismic energy (Fig. 2). In or-der to minimise the industrial noise, the measurement was carried out inthe evening, between 6:00 and 7:00 PM. The five quarries in which theblasts were affected are active quarries located at the margins of theDHVC, or close to this volcanic complex. One blast from each quarry was

Fig. 4: Sample of seismograms recorded atvarious offsets from the quarry blast at the Tašovicequarry. The data are taken from profile 2, thetraces are normalised. All the P-wave onsets arevery clear.

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employed. The quarries are: Horní Tašovice, Mokrá, Stráž nad Ohří,Úhošťany and Krásno (Fig. 2). The charges of the individual blasts werebetween 200 to 400 kg of explosives. The origin times were measured us-ing the BR3 seismic stations (Brož 2000).

The seismic waves were recorded using 30 mobile seismic stations de-ployed throughout the area along two profiles between the quarriesMokrá–Stráž nad Ohří (profile 1) and Horní Tašovice–Úhošťany (pro-file 2). The average distance between the individual stations was 3 km.Additional stations were deployed in the central area of the DHVCaround the profile crossing (Fig. 2). The seismic stations were of twokinds: the Gaia stations with LE3D sensors (Institute of Geophysics,Academy of Sciences of the Czech Republic); and the BR3 stations withSM6-3D sensors (Institute of Rock Structure and Mechanics, Academy ofSciences of the Czech Republic). The geophones were placed in shallowholes to ensure good ground-receiver coupling.

The seismograms were recorded with a 250 Hz sampling frequency;the time at the individual stations was synchronized using the GPS signal.The records were continuous; recording started with the deployment ofthe stations and stopped after the last blast was fired (Fig. 4).

4 DATA PROCESSING

4.1 Raw data processing

The processing of recorded data was carried out in the simplest possi-ble manner in order not to lose any information obtained in data (e.g., byfrequency filtering, etc.). It consists of synchronisation of the instrumenttime with the GPS time in cases when needed, cutting of time windowswith seismic signal and conversion to the SEG-Y format.

Then the P-wave onsets were picked. The S-wave onsets were not clearenough to enable reliable readings and therefore the S-waves were not in-cluded in further processing. It was possible to obtain 123 travel-timereadings in total. The resulting travel-time curve does not show any sub-stantial spatial changes in velocities (Fig. 5). The dataset can be dividedinto two groups based on the apparent velocity (travel-times for offsetslower than 5.5 km have lower apparent velocities, 5.2 km/s on average,than the rest of them). Hence in the first approximation the data can beinterpreted using a two-layer model. To obtain a reasonable two-layermodel, the P-wave onsets were processed using the 2D plus-minusmethod.

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4.2 The 2D data processing

The first arrivals at the stations located along individual profiles werepicked and processed using the plus-minus method (Hagedoorn 1959).The data were processed using a two-layer model. The velocities in theupper layer were determined using the data from the vicinity of the indi-vidual quarries and then interpolated along the profiles. The depths andvelocities of the lower layer were then computed. The resulting modelsare plotted in Figs. 6 and 7.

The depth cross-sections cannot be easily interpreted in geologicalterms. The cover-bedrock boundary is more or less flat, located at depthsof about 200 m below sea level, which is approximately 1 km below thecurrent surface. This does not reflect the geological situation known atthe margins of the DHVC (e.g., in the erosional gorges of the Ohře River– about 300-400 m a.s.l.). However, in the centre of the area it could rep-resent the surface of the underlying EMLC.

In contrast, the velocity sections seem to reveal several interesting as-pects (Fig. 7). The overall refracting velocities along profile 1 (Mokrá–Stráž nad Ohří) are lower than those along profile 2. This difference,which is about 4-5 % on average, can be the result of seismic anisotropy.The profile perpendicular to the direction of the Eger Graben (profile 1)has lower velocities than the profile parallel to it (profile 2). A similar dif-

Fig. 5: Reduced travel-time curve for the P-waves. Reduction velocity is 5.5km/s. The data set consists of 123 travel-time readings. Note that the dataset canbe divided into two parts based on the apparent velocity. The travel-times foroffsets shorter than 5.5 km have lower apparent velocity (5.2 km/s on average)than the rest of them, suggesting a two-layer velocity model for interpretation.

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ference (in terms of direction and magnitude) was already observed dur-ing previous seismic measurements in the area (Málek et al. 2001).

Next, the values of the velocities are almost constant along the indi-vidual profiles. Larger variations of velocities are encountered along pro-file 2, where the variations are within the range of ±200 m/s. Such smallvelocity changes may be considered insignificant; however, with the helpof geological data certain interpretation could be done.

The highest velocities were registered in the southwestern and centralparts of profile 2 (Fig. 7). The profile starts in the granites of the Nejdek–Eibenstock Pluton. Between the kilometres 8-14 the velocity decreasesand consequently again increases on the contact with the EMLC. Thearea with higher velocities (14-21.5 km in Fig. 7) coincides with basic andultrabasic rocks of the EMLC. The P-wave velocities ranging from 6.18 upto 6.26 km/s correspond with the laboratory measurements (Hrouda and

Fig. 6: 2D data processing: depths of the refracting boundary. The results showthat the surface of the bedrock is more or less flat, at depths of about 1 km belowthe surface. No distinct anomalies can be identified.

Fig. 7: 2D data processing: velocities along the refracting interface (top of thebedrock). It is obvious that the velocities along profile 1 are lower than those alongprofile 2. This can be caused by the anisotropy of seismic velocities. The averagedifference in velocities is about 4%. The velocity of the overlying layer is 5.2 km/son average.

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Chlupáčová 1993 unreferenced; see Notes section) of the amphibolites(6.2 km/s) and serpentinites (6.0 km/s). The high velocities can be alsocaused by presence of massive dikes (non-vesiculated and hence withhigher seismic velocities) and subvolcanic rocks (Flurbühl intrusion withvelocity about 6.3 km/s). Area in the central part between the Liboc andStřezov faults (19-21.5 km) with the highest velocities may represent abody of basic and ultrabasic rocks of the EMLC (eclogites and peridotiteswith P-wave velocities up to 7.2 km/s after laboratory measurements;Hrouda and Chlupáčová 1993 unreferenced; see Notes section). Thishigh-velocity area ends with a sharp decrease of velocities and a certain de-pression of the refracting boundary (21.5-23 km in Figs. 6 and 7). Al-though the refracting boundary lies beneath the Permo-Carboniferoussediments, the decrease of velocities is connected with the Střezov Fault(Fig. 2) delineating the sediments from the north (Figs. 2 and 3). The rel-atively low velocities on profile 2 between 23-30 km coincide with thecontact of rocks of the Saxothuringian Zone. The velocities increase fromphyllites over mica-schist and orthogneisses to granulites.

The geological interpretation along profile 1 is more difficult due to alow contrast of computed refracting velocities. Nevertheless, the contactof the Teplá-Barrandian Unit and EMLC dips to NW (15-17 km) and Stře-zov Fault (23 km) is indicated.

4.3 The 3D data processing

After completing the 2D profile processing stage, the next step is the 3Dtomography inversion. The PStomo_eq tomography code by Ari Tryggv-ason (Tryggvason et al. 2002, Tryggvason and Bergman 2006, Tryggvasonand Linde 2006) was used for this purpose.

The PStomo_eq program is a 3D first-arrival tomography code. Thevelocity model consists of rectangular constant-velocity cells. The forwardmodelling is carried out using the first-order finite-difference solver of theeikonal equation (Podvin and Lecomte 1991, Tryggvason and Bergman2006). Hence only the first arrivals can enter the inversion. The benefit is,that the time field is always resolved, even in complex models. The inver-sion step is done using the conjugate gradient solver LSQR (Paige andSaunders 1982).

As stated before, receivers were deployed along the two profiles withadditional receivers in the centre of the area. Five shots (sources) wereused (Figs. 2 and 3). This configuration provides relatively good ray cov-erage in the centre of the area while it essentially decreases towards themargins (Fig. 8).

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First of all, the chessboard test for the current configuration ofsources and receivers was carried out. Next, the tomography of the regionof interest was computed.

4.4 Chessboard test

The chessboard test is a commonly used tool for evaluating tomographyresults. Since the number of sources and receivers in this study was relat-ively low and hence the dataset was sparse (Fig. 8), it was used to estimatethe resolving power of the current configuration.

The chessboard-model consisted of cells, size of 6 × 6 × 6 km, with ve-locities of 3 and 6 km/s. The 3 km/s cells could mimic the Permo-Car-boniferous and Tertiary sediments within the area while the 6 km/s cellsrepresent coherent volcanic and metamorphic rocks.

Synthetic travel times were then inverted using the same parametersas during the real-data inversion (inversion cell size was 1 × 1 ×1 km). TheRMS of the final model was 0.043 s. Two slices in different depthsthrough the inverted dataset are presented in Figs. 9 and 10. The slice inFig. 9 is for a depth 600 m while the second one, in Fig. 10, is for a depthof 1600 m below the surface. The resolved pattern is reasonable along thetwo aforementioned profiles as the ray coverage is the highest there. The

Fig. 8: Ray coverage of the investigated area.

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resolution is also essentially better in the shallow depth and is decreasingdownwards where the pattern is inevitably smeared.

It can be concluded that the tomographic inversion of the real data islikely to be reasonable along the two profiles, whereas the results else-where should be taken with care. Therefore, in contrast to many 3D stud-ies, we will limit our interpretation of the tomography results to the indi-vidual profiles only, leaving out the horizontal slices.

4.5 Real data processing

The complete dataset consists of 123 travel times. Considering this fact andthe spacing between the receivers, it is clear that the size of the (forward)cells (hence the smallest detectable size of possible anomalies) used for thechessboard tests (6 km) will be fairly close to the resolving power of thetomography.

Fig. 9: Results of the chessboard test for a depth of 600 m below the surface.The chessboard model consists of cells, size 6 × 6 × 6 km. The velocities of theindividual cells are 3 km/s (white boxes) and 6 km/s (the black boxes). Invertedvelocities are depicted as contour lines. Axes are labeled in km.

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For tomography computations, several different inversion (constant-velocity) cell sizes were tested. The cell size of 1 × 1 × 1 km, presentedhereinafter, seemed to give the best results. Smaller cell sizes resulted ininferior models (due to the poor ray coverage) while the larger ones pro-duced models which were unnecessarily smooth.

The starting model for the inversion was a simple 1D velocity modelwith a gentle linear increase of velocity. The velocity at the surface was setto 5 km/s and then it increased up to 7 km/s at a depth of 15 km. To en-sure that the velocity model settles with the data, the inversion was firstrun with a high level of smoothing (smoother 400), which gradually de-creased to the smoother 50 over four iterations. The inversion process wasstopped after the 20th iteration; however, the model remained almost un-changed from the iteration 10.

The final model (iteration 20) is depicted in Figs. 11 and 12. The RMSof the final model was 0.053 s and the residuals are plotted in Fig. 13.

Fig. 10: Results of the chessboard test for a depth of 1600 m below the surface.The model and parameters are the same as in Fig. 9.

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The resulting velocity model is not very detailed as could be expectedfrom the sparse shot/receiver configuration. However, it still provides cer-tain facts about the deep geological structure of the area.

The most prominent feature of the resulting model is the high velocityanomaly (more than 5.9-6.0 km/s) in the centre of both profiles. It seemsto be elongated in the W-E direction, perpendicular to profile 1 (Mokrá–Stráž nad Ohří). The width of the body (the shorter axis, NW-SE direc-tion) is about 10 km, whereas its length (the W-E direction) exceeds therange of this survey. In addition, it seems to shift to the NW with depth,which is most likely caused by inclination of the high-velocity body to theNW (approximately under an angle of 45°). Although the inclination can-not be reliably derived solely from our dataset, it is supported by previousstudies (Novotný et al. 2009, Mlčoch and Konopásek 2010).

5 DISCUSSIONThe highest peak of the DHVC (Hradiště Hill) is 934 m a.s.l. and the

base of the volcanic complex is, according to the geological situation atits borders (e.g., in the erosional gorges of the Ohře River), about 400 ma.s.l. Therefore, only the topmost subset of the data represents the rocksof the DHVC.

Fig. 11: 3D tomography: distribution of P-wave velocities along the individualprofiles. The high-velocity area (velocities higher than 6.2 km/s) most likelyrepresents a body of basic rock – probably the EMLC.

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The high-velocity body (more than 5.9-6.0 km/s) in the centre of bothprofiles (Figs. 11 and 12) seems to reflect presence of coherent basic andultrabasic rocks – most likely amphibolites of the Equivalent of the Mar-iánské Lázně Complex (EMLC). The surface of the EMLC in the centralpart is probably just beneath the volcanic rocks of the DHVC at about –200 m a.s.l. The extent of the high-velocity area shifts northwards withdepth. This high velocity anomaly most likely corresponds to the high ve-

Fig. 12: 3D tomography results in perspective view. This view suggests that thehigh-velocity body is elongated in the SW-NE direction where it exceeds the rangeof this survey. The Mokrá–Stráž nad Ohří profile is roughly perpendicular to thebody and hence the width of the body can be estimated at about 12 km.

Fig. 13: Travel-time residuals for the 3D tomographic model.

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locity body in the area of the Eger Rift discovered within the SUDETES2003 experiment (Grad et al. 2003).

Based on the seimic model along profile 1, we suppose that the con-tact of basic rocks of the EMLC with the gneisses of the Teplá-BarrandianUnit (in the south and southeast) is inclined and dips to the north ornorthwest. It is in accordance with previous geological mapping andgravity survey (Mlčoch and Konopásek 2010). The northern boundary ofthe EMLC is probably steeper and follows the Střezov Fault in the DHVCbasement.

On the other hand, it is hard to interpret the high-velocity anomalyalong profile 2, since it is present along most of the profile. This is inagreement with the E-W elongation of the EMLC. Hence, the velocitygradient is pushed to edges of the profile (usually considered unreliable).Nevertheless, the borders seem to be steep, which well fits the estimatedboundary between the high-velocity basic and ultrabasic rocks and low-velocity granitic rocks (Mlčoch and Konopásek 2010). Our observationsupports the interpretations of depth of the Nejdek–Eibenstock Plutonexceeding 3 km (Dudek et al. 1991, Blecha et al. 2009) and the steeply dip-ping contact with the EMLC.

The high-velocity zone in the central part of the DHVC (x-coordinateabout 20 km along profile 2 and velocities higher than 6.3 km/s in Fig. 7)

Fig. 14: Geological interpretation of profiles 1 and 2.

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represents probably the basic and ultrabasic rocks of the EMLC in thecrystalline basement of the DHVC (Fig. 14). This is localised between theLiboc Fault to the west and the Střezov Fault to the north (Mlčoch 2003)under the Permo-Carboniferous sediments of the Žatec Basin. The thick-ness of the sedimentary sequence does not exceed 800 m. Profile 1 did nottouch this sedimentary basin and profile 2 reached only the very edge ofthe sedimentary infill. The seismic data, therefore, did not clearly detectthe presence of sedimentary low-velocity rocks.

The presence of an accumulation of high-density rocks in central areaof the DHVC is also supported by the positive gravity anomaly – anomalyNo. 8 in Sedlák et al. (2009) – exceeding 20 mGal. The most basic (andhence heaviest) rocks of the Flurbühl intrusion are melteigites (olivineijolites) with a density of about 3.1 g/cm3; the amphibolites of the EMLChave density of 2.95 g/cm3, metagabbros 3.0 g/cm3 and eclogites 3.2g/cm3. Such rock association in the central part of the DHVC would re-sult in gravity anomaly of about 8 mGal. The gravity data thus indicatesthe presence of high-density rocks – most probably of cumulitic mag-matic origin. Due to the lack of geological evidence, we prefer the modelcombining both, basic and ultrabasic metamorphic rocks in deeper struc-ture of the crystalline basement of the DHVC and subvolcanic alkaline in-trusion including heavy cumulates as the local gravity maximum(Fig. 14).

The refracting boundary from the plus-minus method, unfortunately,does not reflect a single geological boundary. The geological surround-ings at this depth are complicated and do not represent a 2D layeredmedium, and it is, therefore, hardly possible to interpret these results inclearly geological terms. However, the seismic anisotropy (difference ofthe refraction velocities on profiles 1 and 2) suggests that the foliation ofthe metamorphic rocks of the basement is oriented in the same direction(i.e., SW-NE to W-E) as the profile 2 and caused the distinctively highervelocities than those on profile 1.

6 CONCLUSIONS• The central part of the Doupovské Hory Volcanic Complex

(DHVC) and its basement consists of high-velocity rocks. Accord-ing to the known geology, the basic and ultrabasic metamorphicrocks (amphibolites, eclogites, serpentinites, metagabbros) of theEquivalent of the Mariánské Lázně Complex (EMLC) are respon-sible for the seismic anomaly at depths.

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• Therefore, intrusive body of the DIC can be hardly distinguishedfrom rocks of the EMLC by seismic refraction investigation.Hence, we may not exclude the possible continuation of the Flur-bühl intrusion to deeper parts.

• According to the velocity model, the contact between the EMLCand the Teplá-Barrandian Unit dips to the north or northwest un-der an angle of about 45°.

• The geometry of the boundary between the high-velocity basicrocks and the low-velocity granitic rocks seem to be steep and be-low 3 km.

• The refracting boundary could represent the surface of the EMLCin the centre of the area (about –200 m a.s.l.).

• The anisotropy of seismic velocities might be associated with thegeneral geological structure (foliation of metamorphic units) ofthe area.

Acknowledgements. The research was supported by ProjectIAA300460602 of the Grant Agency of the Czech Academy of Sciences.The study was completed within the framework of the Research Plans ofthe Institute of Rock Structure and Mechanics (AVOZ30460519), theCzech Geological Survey (MZP0002579801) and the Geophysical Insti-tute (AVOZ30120515).

Special thanks are due to the Czech Ministry of Defence for allowingthe seismic experiment to be carried out inside a military training area,usually not accessible to civilians. It is the authors’ pleasure to thank AriTryggvason for providing them with his excellent PStomo_eq tomogra-phy code. Last but not least, we thank to Stanislaw Mazur and an anony-mous reviewer for constructive comments and to Marek Grad for carefulediting of the manuscript.

APPENDIX

Doupov Working Group comprises: A. Boušková, P. Hrubcová,P. Jedlička, P. Kolář, J. Kotek, M. Novotný, B. Růžek, and J. Zedník(Geophysical Institute, Academy of Sciences of the Czech Republic,Prague), L. Čermák, R. Gaždová, P. Hanzal, P. Kolínský, O. Novotný,V. Stejskal, J. Štrunc, and L. Žanda (Institute of Rock Structure and Me-chanics, Academy of Sciences of the Czech Republic, Prague).

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Seismic Model Of The Basement Beneath The Doupovské Hory

NOTES

The data used in this study were presented/described in following un-published materials from the Czech Geological Survey, Prague: (i) Basicgeological map 1:50 000, Kadaň with explanations; and (ii) Basic geologi-cal map 1:25 000, Kyselka with explanations. Both edited by P. Hradecký.(iii) Geological model of the western part of the Bohemian Massif in rela-tion to Deepborehole KTB-1 (in Czech) by Hrouda, F., and M.Chlupáčová. All documents are available in the archive of the CzechGeological Survey in Prague.

Hradecký, P. (ed.) (1998), Basic geological map 1:50 000, Kadaň with explana-tions, Czech Geological Survey, 11-22 (unpublished).

Hradecký, P. (ed.) (2005), Basic geological map 1:25 000, Kyselka with explana-tions, Czech Geological Survey, 11-223 (unpublished).

Hrouda, F., and M. Chlupáčová (1993), Geological model of the western part ofthe Bohemian Massif in relation to Deepborehole KTB-1, Czech Geolo-gical Survey, Prague (unpublished, in Czech).

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Received 27 May 2010Received in revised form 12 January 2011

Accepted 19 January 2011

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