The western Bohemian Massif is known for geodynamic phenomena such as earthquake swarms, CO2 dominated free gas emanations of upper-mantle origin, and Tertiary/ Quaternary volcanism. Among other explanations, a small-scale mantle plume has been suggested. We used data from the international passive seismic experiment BOHEMA (2001-2004) and of a previous seismic experiment to investigate the structure of the upper-mantle discontinuities at 410 km and 660 km depth (the ‘410’ and the ‘660’) beneath the Bohemian Massif with the P receiver function method. More than 4500 high-quality receiver function traces could be utilized.
Two stacking techniques were used: stacking by station (common station method, CSM) and stacking by piercing points in the mantle transition zone (common conversion point method, CCM). Since the station spacing is very close, rays from different stations have similar piercing points in the mantle transition zone. Therefore CCM is sensitive in the transition zone and CSM is sensitive to the uppermost structure of the mantle. The CSM shows delayed conversion times from the 410 km discontinuity beneath the western Bohemia earthquake region, which indicate a slow uppermost mantle. When stacking our data by CCM, we observe thickening of the transition zone towards the Alpine foreland, which agrees with tomographic results by Piromallo and Morelli. The thickness of the mantle transition zone beneath the western Bohemian Massif is normal, with a faint hint to thinning in the northern part.
Our conclusion is that a plume-like structure may exist in the upper mantle below the western Bohemia earthquake region, but with no or only weak imprint on the 410 km discontinuity.
The western Bohemian Massif in central Europe is situated at the junction of three Variscan structural units: the Saxothuringian in the north, the Teplá-Barrandian and Moldanubian units in the south (Fig. 1). The Palaeozoic suture between the Saxothuringian and Teplá-Barrandian/Moldanubian units has been reactivated since the Upper Cretaceous/Tertiary as part of the European Cenozoic Rift System. This led to the evolution of the 300 km long and 50 km wide ENE-WSW trending Eger (Ohre) Rift. The western part of the Eger Rift is known for geophysical and geological phenomena such as the occurrence of earthquake swarms, CO2 dominated free gas emanations of subcontinental lithospheric mantle signature in mineral springs and mofettes, Tertiary/Quaternary volcanism and neotectonic crustal movements (e.g. Špicák & Horálek 2001; Ulrych et al. 2003; Bräuer et al. 2005).
To explain the observed phenomena, several possible scenarios have been suggested: a small-scale mantle plume (Granet et al. 1995), lithospheric thinning beneath the Eger Rift (Babuška & Plomerová 2001; Heuer et al. 2006), and presently ongoing magmatic processes near the crust-mantle boundary, including magmatic underplating (Bräuer et al. 2005; Geissler et al. 2005).
Seismic tomography investigations of the French Massif Central and the Eifel volcanic region (Rhenish Massif) indicated plume-like structures in the upper mantle (Granet et al. 1995; Ritter et al. 2001). These regions are geologically comparable to the western Bohemian Massif as all of them are of Variscan origin and were overprinted by the European Cenozoic Rift system with its Tertiary and Quaternary volcanism and recent CO2 emanations. However, due to the geometry of the tomography experiments, the plume-like structures could only be resolved down to depths of 270 and 400 km for the French Massif Central and the Eifel region, respectively. Granet et al. (1995) suggested that the hot mantle finger might have its origin in the upper mantle near a thermal boundary layer, like the top of the mantle transition zone at 410 km depth, and predicted other diapiric mantle upwellings beneath the Rhenish Massif, the Bohemian Massif and the Pannonian Basin. Goes et al. (1999) observed low seismic velocities in the lower mantle beneath central and western Europe and suggested that they may represent a mantle plume and a common source for Cenozoic volcanism in central Europe. However, the relationship of this deep-seated low-velocity structure to diapiric upwellings in the upper mantle is not yet resolved.
Grunewald et al. (2001) and Budweg et al. (2006) carried out receiver function investigations in the Rhenish Massif of western Germany (Eifel region). They further support the idea of a plume beneath the Eifel by their findings of an approximately 200 km wide and 15 to 25 km deep apparent depression of the 410 km discontinuity in the upper mantle beneath the Eifel region.
In the western Bohemian Massif, a similar diapiric mantle upwelling was suggested by Granet et al. (1995) and might be plausible due to the similar geological and geodynamic situation as in the French Massif Central and the Rhenish Massif, for example, Tertiary/Quaternary volcanism and degassing of CO2. Teleseismic traveltime tomography studies show reduced seismic velocity at least down to 250 km depth beneath the western Bohemian Massif, but do not give any clear evidence for the existence of a diapiric low-velocity anomaly which could be interpreted as a mantle plume anticipated beneath the Eger Rift (Plomerová 2007; Koulakov et al. 2009). A regional receiver function study by Geissler et al. (2008) showed delayed arrival times of Ps phases from the 410 km and 660 km discontinuities in central Europe, while the differential traveltimes for the transition zone are as predicted by the Earth reference model IASP91 (Kennett 1991; Kennett & Engdahl 1991). This indicates that the delays are caused mainly by the seismic velocity structure (vp/vs) in the upper mantle.
In this study we show results of a P-receiver function analysis of the discontinuities at 410 and 660 km depth which was aimed at investigating possible influences of the suggested plume-like structure on the topography of the discontinuities. Our densely spaced network of permanent and temporary stations allows for a much more detailed investigation of the structure of the 410 km and 660 km discontinuity than in the studies by Geissler et al. (2008) and Grunewald et al. (2001). However, our station network has a smaller aperture than the networks in these previous studies.
The two main seismic discontinuities in the mantle at 410 and 660 km depth are generally attributed to pressure- and temperature-induced phase transformations in olivine and other minerals (e.g. Helffrich 2000; Lebedev et al. 2002). Olivine is transformed to a spinel structure at a depth of about 410 km and breaks down to perovskite and magnesiowüstite at about 660 km. These exothermic and endothermic phase transformations have Clapeyron slopes dP/dT of approximately +3 MPa K-1 at the ‘410’ and -2 MPa K-1 at the ‘660’, respectively (Helffrich 2000). Thus, if the 410- and 660-km discontinuities are entirely due to these phase transformations, regions of abnormally low temperature such as subduction zones should correspond to elevation of the ‘410’ to lower depths and depression of the ‘660’ to greater depths, so that the distance between the discontinuities is increased. Regions of increased temperature like plume conduits and their surroundings should correspond to depression of the ‘410’ and elevation of the ‘660’, so that the distance between the discontinuities is reduced. Applied to receiver function studies, this means that an increased temperature of 100-150 K in the mantle transition zone corresponds roughly to a depression of the ‘410’ by about 10 km, which delays the P-to-S converted phase by about 1 s compared to the IASP91 model.
As in previous studies, we use the differential delay times of P-to-S converted phases at the ‘410’ and ‘660’ as an estimate for the thickness of the mantle transition zone (thickness according to IASP91 is 24.0 s or 250 km).
The study was carried out with data from the international passive seismic experiment BOHEMA (BOhemian Massif HEterogeneity and Anisotropy 2001-2004) (Plomerová 2003), which includes regional permanent observatories, and data from a previous temporary experiment by Geissler et al. (2005). The distribution of the 117 stations is shown in Fig. 1, location and further parameters are given in Table 1. The selected teleseismic events have epicentral distances between 30° and 95°, magnitudes larger than 5.5 and a clear P onset with a signal-to-noise ratio larger than 3. The distribution of events is shown in Fig. 2. Altogether, more than 4500 individual receiver function traces were used for investigating the upper-mantle discontinuities.
The P receiver function method has been used successfully for quite a long time to investigate crustal and upper-mantle structure. The theoretical background of the technique was described, for example, by Vinnik (1977), Langston (1979), Kind & Vinnik (1988), Kosarev et al. (1993) and Zandt et al. (1995). This technique analyses waves that are converted from P to SV waves at a horizontal seismic interface, neglecting seismic anisotropy. The delay time between the converted wave and the direct P wave is used as a measure for the depth location of the seismic interface. The method basically comprises the following steps: restitution of the broad-band ground displacement to make signals of different sensors better comparable, component rotation from the ZNE coordinate system into the LQT coordinate system, deconvolution in the time domain as a source-equalization, and finally moveout correction, filtering, and summation of many traces to improve the signal-to-noise ratio of the converted phases. The moveout correction is carried out for a reference epicentral distance of 67° (slowness of 6.4 s deg-1). The time of the P onset is set to zero, so that the onset time of all following converted phases is referred to as ‘delay time’ relative to the P onset.
Before the summation of individual traces, their quality was visually checked. Apart from the criterion that the signal-to-noise ratio of the P wave was at least 3, further quality criteria were low noise in the receiver function traces before P onset, existence of a Moho signal and/or Moho multiple reverberations, no extreme noise-like unrealistically large amplitudes or ‘ringing’ in the later part (20 or more seconds after P onset) of the receiver function trace. Thus, we obtained 4508 individual receiver function traces.
We carried out filter tests to maximize the energy of the converted phases at about 44 s (discontinuity at 410 km) and 68 s (discontinuity at 660 km) and decided to use a bandpass filter from 3 s to 20 s.
The amplitudes of P-to-S conversions from the upper-mantle discontinuities (named P410s and P660s hereafter) are about four times weaker than the Moho conversion signal (PMs) and not always obvious in single receiver function traces. Hence, it is necessary to stack individual traces to reduce noise and enhance converted phases from the ‘410’ and ‘660’.
We used two different approaches to sort the receiver function traces: first, we read delay times of the conversions at 410-km and 660-km depth from common station stacks (Common station method, CSM) with at least 15 individual traces and good signal quality (approach of Geissler et al. 2008, see Fig. 3). In this approach, mainly effects below the stations (in the uppermost mantle) are emphasized.
Second, we stacked the data according to their piercing points at 410 km and 660 km depth, respectively (similar to Grunewald et al. 2001). In this approach (CCM), mainly effects at the ‘410’ and ‘660’ are emphasized. Effects from other depth regions superimpose destructively by trend.
Due to the ray geometry, the information on the ‘410’ and ‘660’ in the receiver functions originates from a region ca. 80-170 km (‘410’) and 150-310 km (‘660’) away from the receiver. Ps piercing points were calculated using the IASP91 Earth reference model (Kennett 1991). We divided the area into 43 geographical boxes of 1°× 1°, which corresponds to approximately 70 km (E-W) × 112 km (N-S) at the surface (see Fig. 4). We added a rim of 0.25° around each box so that they overlap, which slightly smoothes the results of neighbouring boxes. The size of the resulting 1.5°× 1.5° boxes corresponds approximately to the dimensions of the Fresnel zone and hence to the resolution of our data. The size of the Fresnel zone for data with a dominant period of 3-5 s is about 200 km at 410 km depth (Grunewald et al. 2001). It is therefore not possible to resolve structures significantly smaller than 150 km laterally at the depth of the mantle transition zone with the receiver function method.
As most events were recorded from northeastern directions, the area northeast of the western Bohemian Massif, between Berlin and Prague, is sampled best. More recordings from stations in southeastern Germany would be necessary to get a better resolution of the ‘410’ and ‘660’ beneath the western Bohemian Massif. Nevertheless, there is still fairly good data coverage beneath and to the west of the western Bohemian Massif. To the south, there is a data gap especially in the piercing points at 660 km depth.
To study the effect of lithospheric/asthenospheric structure on the delay times of P410s and P660s, we analysed synthetic seismograms in the same way as the real data (see the Appendix). The reflectivity method was used to compute the theoretical seismograms (Kind 1985).
To assess the robustness of our data and the influence of noise, we tested different distributions and sizes of the boxes, and stacked different subsets of the data within individual boxes. The observed variations represent mean topographic variations of the mantle discontinuities within the boxes. The error introduced by picking delay time values and by bandpass filtering is relatively small (±0.1 s). Together, this amounts to uncertainties of ±0.3 s for delay time values of both the ‘410’ and ‘660’ obtained with the CCM. For the transition zone thickness, the maximum uncertainty for delay time differences adds up to ±0.6 s.
As already mentioned before, we applied two stacking approaches to study the shallow and the deep upper mantle. First, we look at common station stacks, second at common conversion point stacks.
.1 Common station stacks
This is the usual approach in areas with wide station spacing. However, also in areas with dense station distribution this approach can give important information on the uppermost mantle structure. Individual traces from different azimuths and with piercing points relatively far away from the actual station are stacked, however, the shallow structure below the stations is sampled by rays from all events.
Fig. 3 shows the complete sum traces of all 117 station sites analysed. The most prominent signal is the conversion from the Moho (PMs), followed by crustal reverberations (mm). The conversions from the discontinuities at 410 km and 660 km depth can be seen at about 44 s and 68 s, respectively. Unfortunately, not all stations have enough single traces or show reliable Ps phases from the mantle transition zone. Therefore, only the ‘good’ stations with at least 15 single traces are shown in Fig. 5. For 19 stations we observe clear arrivals from the ‘660’, but not from the ‘410’. This might be due to higher order reverberations from shallow interfaces or complicated structure at the ‘410’. The measured delay time values are marked by red and green circles and listed in Table 1. Table 1 also gives the delay time difference between the two discontinuities as a measure of the thickness of the mantle transition zone.
The P410s delay times vary between 43.8 s (BG29) and 46.7 s (BG19) (Table 1). In the corresponding map (Fig. 6a), this is shown by green and reddish colours meaning normal and slightly increased delay times compared to the Earth reference model IASP91. However, some structure is visible: ‘normal’ values of 44.1 s ± 0.5 s dominate in the N and W of the investigated area, while red shades show increased delay times in the centre and SE (in the Bohemian Massif). The grey stations are stations where we could not determine a delay time value in the sum trace, mostly because of a lack of individual traces, or sometimes because of a missing clear coherent conversion signal.
The P660s delay times vary between 67.1 s (CLL) and 69.5 s (LAC2). Similar to the ‘410’, normal values (68.1 s ± 0.5 s according to IASP91 reference model) can be observed in the N and W, while slightly increased delay times dominate in the centre and SE (Fig. 6b). The delay time difference between the two discontinuities varies between 21.6 s (BG19) and 24.7 s (GRC1). As visible in Fig. 6(c), ‘normal’ values around 24.0 s dominate the whole investigated area. However, in the centre and north they are mixed with red shades corresponding to shorter time differences.
Since most of the analysed events arrive from northern to eastern backazimuths, the sampled region is slightly shifted to the NE of the individual stations.
.2 Common conversion point stacks
For this stacking approach, we divided the investigated area in 43 geographical boxes as described in Section 3 and shown in Fig. 4. This is possible because we have a dense station spacing and thus a dense coverage of piercing points of the rays at 410 km and 660-km depth. For the analyses we considered only boxes containing at least 20 individual traces. The number of traces within a certain box is usually different for piercing points at 410 km and 660 km depth. Figs 7(a) and (b) show examples of two boxes, displaying all individual traces as well as final common conversion point stacks (sum trace on top).
Fig. 8(a) shows the common conversion point stacks for the ‘410’. The measured delay times vary between 43.8 s (Box 2) and 45.2 s (Boxes 20, 31 and 38) (Table 2) with an average of 44.7 s. In the top traces in Fig. 8, only a poor signal of the ‘410’ is recorded. This might again be due to higher order reverberations from shallow interfaces or complicated structure at the ‘410’. The values are shown colour-coded in Fig. 9(a). While the boxes in the north and northeast show ‘normal’ values in regard to the reference model IASP91 (i.e. 44.1 ± 0.5 s, green colour), most boxes in the centre, west and southeast of the investigated area show delays of conversion signals from the ‘410’. Boxes 27 and 29 seem to be anomalous. However, Table 2 shows that the delay time values in both boxes (44.5 s) is very close to ‘pink values’ starting at 44.6 s.
A similar image is shown in Fig. 8(b) for common conversion point stacks for the ‘660’. Here the delay times of individual boxes vary between 67.9 s (Boxes 14 and 18) and 70.5 s (Box 43, Table 2) with an average of 68.7 s. The values are shown colour-coded in Fig. 9(b). There is a clear structure visible: in the north and partly the centre of the investigated area, delay times of the ‘660’ show normal values of 68.1 ± 0.5 s (green colour), according to IASP91. Southwards, the signals are clearly and increasingly delayed (blue shades) by up to more than 2 s.
The time difference between P660s and P410s is also given in Table 2 and plotted into Fig. 9(c). Values vary between 22.9 s (Box 7) and 25.7 (Box 39) with an average of 24.0 s. A clear majority of the boxes shows normal values of 24.0 ± 0.5 s according to IASP91. However, there are also a few boxes showing slightly increased (blue) or decreased (pink) values. Table 2 shows that the blue boxes 2, 8 and 29 have values not far away from normal (green) values. However, between neighbouring boxes 7 and 8 as well as 31 and 40 there remains quite a large step.
.1 An anomaly within the upper mantle?
To answer this question, we look at the observations of the common station method (CSM) as they are more sensitive to shallow structure. We observed roughly similar patterns for conversions at the ‘410’ and ‘660’, that is, ‘normal’ delay time values in the north and west and increased delay times in the centre and southeast (Figs 6a and b).
Comparison with the synthetic data (Fig. A1, Table A1 in the Appendix) shows that the observed delays might be either caused by reduced velocities (increased vp/vs) in the upper mantle (e.g. asthenosphere material in models BOH-1D and -1E), or by increased temperature of about 200 K at the depth of the ‘410’ (models BOH-3D and BOH-3I). Obviously, an increased temperature of the whole mantle transition zone does not fit with our data, as this would ‘decrease’ delay times of the ‘660’ compared to IASP91 (models BOH-3G, -3C and 3D in the Appendix). An increased temperature at the ‘410’ and decreased temperature at the ‘660’ might fit the data but seems unrealistic. Therefore, the most plausible explanation for the observed common station delays is a limited volume of lower velocities in the upper mantle above the ‘410’, which may reach down to the ‘410’ and depress it. This is in accordance with the results of traveltime tomography (Plomerová 2007; Koulakov et al. 2009). The thickness of the mantle transition zone (Fig. 6c) shows values as predicted by IASP91. There are several stations that show decreased differential delay times (reddish colours in Fig. 6c) in the central and northeastern part of the area. This could indicate that a possible temperature anomaly in the upper mantle is reaching down to the 410 km discontinuity, but keeping in mind the uncertainty of ±0.6 s and the size of the corresponding Fresnel zones this interpretation is not supported by the small aperture data.
.2 An anomaly in the mantle transition zone?
In contrast to the CSM, considering the data by common conversion point (CCM) takes the real position of the P-to-S conversion into account, allowing a detailed study at depths of the mantle transition zone. In this approach, traces recorded by different stations are stacked, which means that existing near-receiver structure is smeared over a wide area.
For the discontinuities at 410 and 660 km depth, the CCM (Figs 9a and b) generally shows a similar structure as the CSM (Figs 6a and b). Additionally, a deepening of the ‘660’ can be observed in the very south.
The differential delay times for the ‘660’ and ‘410’ (Fig. 9c) generally concur with the value predicted by IASP91 reference model (24.0 s). Taking the size of corresponding Fresnel zones and uncertainty of ±0.6 s into account, this indicates normal thickness and temperature of the mantle transition zone. Boxes 18 and 19 might be interpreted as indicative of a thinning of the mantle transition zone. Comparing with the synthetic models, an increased temperature in the upper mantle above (pronounced asthenosphere or plume-like structure) and at the ‘410’ is indicated, possibly combined with a slight depression of the 410 km discontinuity.
This means that with the CCM we clearly see a velocity reduction in the upper mantle. However, a shift to north or northeast (our main azimuthal directions) and the depth extend down to the 410 km discontinuity cannot be determined with confidence. Better azimuthal ray coverage is therefore essential for this approach. However, the reduced thickness shown in boxes 18 and 19 (Fig. 9c) can serve as an upper limit for estimating a possible temperature increase at the 410 km discontinuity by about 100-150 K.
.3 Implications for the structure of the upper mantle
Our findings of a low-velocity anomaly in the upper mantle beneath the northwestern Bohemian Massif and surroundings are based on delays of the Ps conversion signals from the ‘410’ and ‘660’ of up to more than 2 s, compared to IASP91 reference model. We also have some weak signs that this anomaly may extend even down to the discontinuity at 410 km depth. Our results support the results of previous studies. (2005, 2008) also found delays of up to 2 s at some stations in the same area. However, they used much less data as in this study. In their P-wave tomography study, Piromallo & Morelli (2003) found a pronounced low-velocity anomaly of about -2 per cent in the 50-200 km depth range beneath the Bohemian Massif as well as beneath the Rhenish Massif and Massif Central. However, it disappears towards the transition zone. Plomerová (2007) found a broad low-velocity anomaly beneath the Eger Rift rather than a plume-like, columnar upwelling, which they interpret as an upwarping of the lithosphere-asthenosphere transition. Koulakov et al. (2009) present a model of P- and S-wave velocity anomalies beneath Europe from tomographic inversion. Both P and S models show a low-velocity zone of -2 per cent only in the uppermost mantle between ca. 80 and 250 km depth, yet they deliberately leave the question about the presence of a hypothetical plume at greater depths open.
Beneath the geologically comparable Eifel volcanic region, Grunewald et al. (2001) found an apparent depression of the 410 km discontinuity by about 20 km, which points to a zone of increased (by about 200-300 K) temperature and/or reduced seismic velocities in the mantle above 400-km depth which they interpreted to support the idea of a plume-like structure in the region. Beneath the western Bohemian Massif, their data show normal to slightly increased depth of the ‘410’. However, the results regarding the Bohemian Massif are based on very few stations.
No indication for an anomaly of the ‘660’ was found under the Eifel, which argues against a strong thermal anomaly at the bottom of the mantle transition zone (Grunewald et al. 2001; Budweg et al. 2006). The same is the case for the ‘660’ beneath the western Bohemian Massif investigated in our study. However, we need to point out that a structure significantly smaller than 150 km laterally cannot be resolved by the receiver function method. The velocity anomaly in the upper mantle observed by Koulakov et al. (2009) is of the dimension of 150 km. The 1 s isochrones at 410 km depth have a diameter of comparable size (200 km, Grunewald , 2001). P-wave tomography by Piromallo & Morelli (2003) does also not show any significant velocity anomaly in the depth range between 600 and 700 km. The tomographic inversion by Koulakov et al. (2009) supports these findings.
In the southern part of the investigated area, that is, at the transition between the southern Bohemian Massif and the Eastern Alps, we map a clear deepening of the ‘660’ in the CCM. The delay of arrival times of up to more than 2 s corresponds to a deepening of the ‘660’ of about 20 km or a decreased temperature of about 150-200 K. However, this southernmost area is in the geotectonic context of the Alpine foreland and beyond the scope of our study
The thickness of the mantle transition zone is normal in most parts of our investigated area. There is a hint to a slightly thinned mantle transition zone in boxes 18 and 19. We consider the increased thickness of the mantle transition zone in boxes 39 and 40 as significant (Fig. 9c), even though it is at the rim of the investigated area. It is in accordance with the results of Geissler et al. (2008) and Piromallo & Morelli (2003). Koulakov et al. (2009) show increased P velocities at a depth of 500 km in the very south of Germany, while at 700 km depth the velocity structure is not anomalous.
If the manifold processes that point to magmatic activity beneath the western Bohemian Massif should have their origin in a diapiric mantle upwelling as predicted by Granet et al. (1995), the absence of a distinct topography of the upper-mantle discontinuities leads to the conclusion that this predicted mantle finger shows only a weak imprint on the 410 km discontinuity and no effect on the ‘660’. If it exists, it either needs to have its origin above the mantle transition zone, or its diameter is significantly smaller than 150 km at the depth of the mantle transition zone and thus too narrow to be resolved by teleseismic receiver functions. This is in accordance with the tomographic study of Koulakov et al. 2009.
As already stated in Geissler et al. (2008) and references therein, the synthetic analysis shows that most of the apparent deepening/upwelling of the upper-mantle discontinuities can be explained by reduced/increased seismic velocities in the crust and uppermost mantle. This was already proposed by Kind & Vinnik (1988). The necessary higher vp/vs ratios can be partly caused by different chemical composition, as well as by anisotropy.
Consistent delays of the Ps converted waves from the ‘410’ and ‘660’ in common station stacks clearly indicate a low-velocity anomaly (high vp/vs anomaly) in the upper mantle beneath the northwestern Bohemian Massif and surroundings, in accordance with previous studies. This is supported by delays of the Ps converted waves from the ‘410’ and ‘660’ in common conversion point stacks. There are slight indications that this anomaly may extend down to the discontinuity at 410-km depth, because decreased differential delay times between conversions from the ‘410’ and ‘660’ were observed by common conversion point stacking. As the IASP91 global reference model does not contain an asthenosphere, an increased vp/vs ratio within an existing asthenosphere beneath the investigation area might explain at least some part of the observed time delay of the discontinuities of the mantle transition zone. Similarly, a plume-like structure in the upper mantle above 410 km may explain the observed data. A deepening of the ‘660’ was observed beneath the southern part of the investigated area by the CCM. The thickness of the mantle transition zone is normal to slightly thinned in the northern and central part of the investigated area, while it increases in the very south.
The ‘baby plume’ beneath the western Bohemian Massif predicted by Granet et al. (1995) would thus have no or only weak impact on the topography of the 410- and 660-km discontinuity. Our conclusion is that also with a considerably increased data set no indication of a plume in the mantle transition zone is detected, however from the receiver function point of view a plume-like structure may exist in the upper mantle below western Bohemia.
We would like to thank the Geophysical Instrument Pool Potsdam of GFZ Potsdam for providing mobile seismic stations for the BOHEMA experiment. W. Hanka helped with archiving the data in the GEOFON data centre. We are also grateful to all institutes providing data from the permanent observatories. For data processing we used K. Stammler's program SeismicHandler. Most figures were generated using the Generic Mapping Tools (GMT). We would like to thank the reviewers for their helpful and constructive comments. We thank the Deutsche Forschungsgemeinschaft (DFG) for the funding of the BOHEMA experiment, the EU for support within the NERIES project and GFZ Potsdam for further funding of B.H.
Appendix: Synthetic receiver functions
To study effects of upper mantle and mantle transition zone velocity structure on the delay times of MTZ conversions, we computed receiver functions from synthetic seismograms (Fig. A1) using the reflectivity method (Kind 1985). Starting with the global reference model IASP91, we changed crustal and uppermost mantle structure according to results from previous investigations of the Bohemian Massif (e.g. ., 2006, 2007; Plomerová. 2007), see Table A1. However, it is not trivial to construct an absolute velocity model from a regional traveltime tomography model (Plomerová. 2007), since in tomography the velocity variations are only valid in specific layers and do not allow for derivation of absolute velocity variations in 3-D. Even if Plomerova et al. chose IASP91 as the background model, it does not have to fit the actual velocity distribution in the upper mantle, since only residual residuals are inverted. The average variation of P-wave velocity in the layers of the regional model of Plomerová. (2007) is in the order of 2 per cent. As there is no cratonic lithosphere in the investigated area, we assume that the highest velocities correspond to IASP91 reference model. The IASP91 reference model contains no low-velocity zone representing the asthenosphere. An asthenospheric layer with reduced vp and vs and a vp/vs ratio of 1.85 was introduced in models BOH-1D and BOH-1E. The top of the asthenosphere is located between about 80 km (BOH-1D) and 130 km (BOH-1E). At the bottom, the asthenosphere converges to the IASP91 reference model between 271 and 371 km depth. As discussed by (2008), the delay times of the conversions from the mantle transition zone depend more on the vp/vs ratio within this layer than on the absolute velocity variations in the upper mantle.
In models BOH-3, we left the upper-mantle structure as in IASP91 reference model, that is, without an asthenosphere. We only changed the depths of the discontinuities of the mantle transition zone to study the temperature effects on these discontinuities. The depth (pressure) of mineral phase transitions which cause the seismic discontinuities is strongly related to temperature (e.g. Bina & Helffrich 1994; Lebedev et al. 2002).
As can be seen in Fig. A1, the first 20-30 s of the synthetic receiver functions are dominated by crustal primary conversions and reverberations. The boundaries of the asthenospheric layer in the models are 10-km-broad gradient zones, so that no distinct converted arrivals can be seen in the synthetic receiver functions. Clear arrivals can be observed for both the 410- and 660-km discontinuities. The delay times of these phases strongly depend on the chosen model:
Models BOH-1D and -1E clearly show a delay of the ‘410’ and ‘660’ conversions in relation to the IASP91 reference earth model. This is mainly the effect of the asthenospheric layer in the upper mantle. Models BOH-3b to -3I study the effect of depth (temperature) variations at 410 and 660 km depth. There is a clear trend to smaller differential delay times between ‘660’ and ‘410’ conversions with increasing temperature (thinner MTZ) (Table A1). Furthermore, there is an effect on the absolute delay times of both the ‘410’ and ‘660’ conversions: with increasing temperature, the depth to the 410 km discontinuity also increases about 4 km per 50 K, according to Bina & Helffrich 1994.