Crustal structure of the Eastern Alps along the TRANSALP profile from wide-angle seismic tomography

Tectonophysics 414 (2006) 51 – 69 www.elsevier.com/locate/tecto

Crustal structure of the Eastern Alps along the TRANSALP profile from wide-angle seismic tomography Florian Bleibinhaus *, Helmut Gebrande Department of Earth and Environmental Sciences, Geophysics Section, University of Munich, Germany Received 10 September 2004; accepted 4 October 2005 Available online 10 January 2006

Abstract The objective of the TRANSALP project is an investigation of the Eastern Alps with regard to their deep structure and dynamic evolution. The core of the project is a 340-km-long seismic profile at 128E between Munich and Venice. This paper deals with the Pwave velocity distribution as derived from active source travel time tomography. Our database consists of Vibroseis and explosion seismic travel times recorded at up to 100 seismological stations distributed in a 30-km-wide corridor along the profile. In order to derive a velocity and reflector model, we simultaneously inverted refractions and reflections using a derivative of a damped least squares approach for local earthquake tomography. 8000 travel time picks from dense Vibroseis recordings provide the basis for high resolution in the upper crust. Explosion seismic wide-angle reflection travel times constrain both deeper crustal velocities and structure of the crust–mantle boundary with low resolution. In the resulting model, the Adriatic crust shows significantly higher Pwave velocities than the European crust. The European Moho is dipping south at an angle of 78. The Adriatic Moho dips north with a gentle inclination at shallower depths. This geometry suggests S-directed subduction. Azimuthal variations of the first-break velocities as well as observations of shear wave splitting reveal strong anisotropy in the Tauern Window. We explain this finding by foliations and laminations generated by lateral extrusion. Based on the P-wave model we also localized almost 100 local earthquakes recorded during the 2-month acquisition campaign in 1999. Seismicity patterns in the North seem related to the Inn valley shear zone, and to thrusting of Austroalpine units over European basement. The alignment of deep seismicity in the Trento-Vicenza region with the top of the Adriatic lower crust corroborates the suggestion of a deep thrust fault in the Southern Alps. D 2005 Elsevier B.V. All rights reserved. Keywords: Crustal structure; Eastern Alps; Tauern Window; TRANSALP; Seismic tomography; P-wave velocities; Reflection; Refraction; Anisotropy

1. Introduction Deep Seismic Sounding (DSS) investigations in the Eastern Alps started more than 50 years ago with observations of quarry blasts and lake shots (Angenhe* Corresponding author. Present address: Department of Earth and Environmental Sciences, Geophysics Section, Theresienstr. 41, D80333 Munich, Germany. Tel.: +49 89 2180 4202; fax: +49 89 2180 4205. E-mail address: [email protected] (F. Bleibinhaus). 0040-1951/$ - see front matter D 2005 Elsevier B.V. All rights reserved. doi:10.1016/j.tecto.2005.10.028

ister et al., 1972), and were intensified in the 1970s by several explosion seismic campaigns (Alpine Explosion Seismology Group, 1976; Italian Explosion Seismology Group and Institute of Geophysics ETH Zu¨rich, 1981). These refraction seismic data were of high quality, but observations were sparse due to very large shot point spacing (Fig. 1). The complex geological setting and velocity structure in the crust left room for different phase correlations and models, particularly with respect to the crustal thickness of the Southern Alps and the transition from Adriatic to European Moho (Miller

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Fig. 1. Tectonic sketch map of the Eastern Alps indicating the TRANSALP profile. The solid part of the TRANSALP line marks the section displayed in Fig. 2. Older refraction seismic DSS profiles are indicated by dashed/dotted lines (reversed observations) and arrows (single-ended observations). Circles denote the corresponding shot points.

et al., 1977; Giese and Reutter, 1978; Italian Explosion Seismology Group, 1978; Italian Explosion Seismology Group and Institute of Geophysics ETH Zu¨rich, 1981; Giese et al., 1982; Miller et al., 1982; Yan and Mechie, 1989; Giese and Buness, 1992; Scarascia and Cassinis, 1997). These explosion seismic campaigns were followed by a break of 20 years, while seismic profiling concentrated on the Western and Central Alps. DSS (Deep Seismic Sounding) investigations in the Eastern Alps were resumed in 1998/99 by a combined reflection and refraction seismic profile between Munich and Venice (TRANSALP Working Group, 2001). This line crosses the orogen almost perpendicular to the main strike direction, in a place, where the northward extension of the South Alpine suggests the strongest compression (Fig. 1). 2. Data Measurements were performed using a combination of explosion and Vibroseis sources. Compared to the

older DSS surveys, charges were significantly smaller, but dense source spacing (5 km for shots, 100 m for sweeps) provides the basis for highly resolved models, especially for the upper crust. For active source tomography, shot and Vibroseis data were recorded by an array of up to 100 short-period three-component stations (Fig. 2). When selecting the station locations, our major concern was to avoid the high ambient noise in Europe’s most touristy regions, which led to offline distances of up to 10 km. Including short-period data from a passive array (TRANSALP Working Group, 2001) resulted in even larger lateral offsets. Station coverage varies remarkably, because acquisition was carried out in two separate campaigns. Shots north of the Inn Valley (55 km) were recorded in 1998 by up to 100 stations from different subprojects, including a mobile array configured for dense recordings of critical Moho reflections (Fig. 2). Most of the data was acquired in 1999 by 40 stations distributed along the central 200 km of the profile, corresponding to 5 km station spacing on average. A more detailed description

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of the variable acquisition geometry is given by Bleibinhaus (2003). The Vibroseis data was diversity stacked, correlated and subsequently diversity binned to enhance the signal, and to reduce the number of traces. Bin locations follow the crooked Vibroseis line with an average spacing of 250 m, and are centered at the corresponding Vibrator points to minimize the necessary time corrections. Picking of the zero-phase Vibroseis signal was adjusted to the physical minimumphase onsets of explosions comparing travel times from nearby shots. A summary of the picked travel times of explosions and Vibroseis signals is displayed in Fig. 3. 3. Tomographic inversion We applied a weighted damped least squares inversion scheme from local earthquake tomography after Thurber (1983) to obtain a P-wave velocity model. We adapted the accuracy of this method to the observation distances of our data, and extended it to simultaneously invert reflection and refraction travel times for reflector depths and velocities. The inversion of Moho reflections is crucial to constrain the Moho geometry as well as the lower crustal velocity structure, particularly due to the absence of Pn observations, and because the penetration depth of the direct wave does not exceed 10 km due to the very small vertical velocity gradient. Three-dimensional computations are necessary to avoid the geometrical distortions from a 2D projection, although the limited resolving power of the data does not allow an inversion for 3D structure at deeper crustal levels. 3.1. Initial model First we constructed an initial 1D model by averaging velocities and Moho depths derived from the refraction seismic study by Miller et al. (1977). We divided the Moho in this model (Fig. 4) to independently constrain its Adriatic and European part. We also left a gap in the Moho, because there are indications from the near vertical reflection data (Lu¨schen et al., 2004) and from the wide-angle data itself (Fig. 3) that the top of the mantle is not reflective below the Alpine Fig. 2. Map of seismic sources and three-component stations. FB— Foreland Basin, NCA—Northern Calcareous Alps, QPZ—Quartzphyllite Zone, TW—Tauern Window, UAG—Upper Austroalpine gneisses, D—dolomites, TC—Tertiary clastics, PL—Periadriatic Lineament, VSB—Valsugana Thrust Belt. Station (AAS, IGS, IRO) and shot labels (2, Q3W, Q4W) are referred to in the text.

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Fig. 3. Travel time picks vs. profile coordinate of station or vibrator point. First break observations are restricted to the crust (crosses). Some picks from the main line were added to ensure a continuous coverage of near offsets. Reflection observations from the European (circles) and Adriatic Moho (squares) show a significant gap between 170 and 200 km by which they can be distinguished. Moho observations in the North from shots in the South do not show this gap, as there were no dynamite shots south of 200 km, which would have been necessary to obtain reversed observations from the Adriatic Moho. Triangles denote reflections from an intracrustal reflector. Refractions from the upper mantle (Pn) are not observed at all due to a limitation of the maximum offset to 200 km.

root, at least with respect to the bandwidth of the corresponding signals. We further included a delimited bfloatingQ (not tied to the velocity field) reflector south of the Inn Valley at 12 km depth, from which we observed strong reflections at several local stations (e.g. Fig. 8b) and on a cross-line (Fig. 16). 3.2. Grid adaptation In the next step, we locally tailored the model parameterization to the resolving power of the data. The principles of this method are described by Thurber and Eberhart-Phillips (1999) and Bo¨hm et al. (2000). Adapting the inversion grid to the resolving power helps to avoid artifacts in regions with poor coverage, thereby improving the stability of the inversion and the general reconstruction quality. In particular, it allows the combination of high resolution in the upper crust with low resolution in the middle and lower crust in one

model. Thus, the projection of travel time residuals, which are caused by the highly heterogeneous upper crust, to deeper levels can be avoided. Another advantage of the tailored grid is a reduction of the influence of damping on the solution, because damping numbers can be smaller, when there are less underdetermined nodes. In order to establish the adapted grid, we followed an approach proposed by Kissling et al. (2001). First we determined an initial grid, the minimum axis spacing of which corresponds to the maximum possible resolution. We chose 2.5 km in the N–S-direction, which is half of the average station spacing, but still 10 times larger than the Vibroseis bins. In the vertical direction, the node axis spacing increases from 0.5 km at the surface to more than 10 km in the lower crust. In the E–Wdirection, the velocity model is extended by constant extrapolation. We also calculated inversions for a model, which is 3D in the upper crust, with 10 km

Fig. 4. Initial velocity distribution derived from an older refraction seismic model (Miller et al., 1977). Reflectors are indicated by solid lines. Regions without potential ray coverage are masked. Adriatic and European Moho are separated by a gap and can be inverted independently. Due to this gap and a delimited reflector between 50 and 100 km we refer to this initial model as bpseudo 1DQ.

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Fig. 5. Seismic inversion grid nodes (crosses – velocities, circles – reflector depths) and RDE values (grey coded) of the initial (a) and the final grid (b). The resolution values displayed for the initial grid were computed in the starting model (Fig. 4), while those for the final grid represent the solution (Fig. 7). As expected, the variable depth spacing of the horizontal grid axis alone is not sufficient to compensate for the decrease of resolution with depth (a). RDE values indicate a well-constrained solution only in the central part of the uppermost crust. In order to increase RDE values in the other parts of the model, nodes were iteratively omitted from the grid. In the final grid (b) all RDE values are above 0.2. RDE values for Moho depth nodes (circles) range between 0.3 for the southernmost node and 0.6. (Nodes of the intracrustal reflector at 12 km depth are distributed around this section and have not been projected into it for the clarity of this figure. Their final RDE values range from 0.2 to 0.6) This means that velocities and depths are well constrained throughout the model. Therefore, primarily the distance between the nodes defines the local resolution of the final model.

node spacing in the E–W-direction, and 2D in the lower crust. This resulted in minor lateral variations only, and for the sake of simplicity we will show results of the 2D model only. This 2D grid was further refined by evaluating the Resolution Diagonal Elements (RDE) after a complete inversion (Fig. 5). The RDE values represent the independence of a model parameter (Crosson, 1976; Thurber, 1993). Poorly resolved nodes, which are indicated by small RDE values, were omitted from the grid. Their corresponding velocity value is then interpolated from neighbouring nodes, or explicitly linked to a specific neighbour. This process of grid adaptation was iterated until all RDE values were greater than 0.2. This number is commonly accepted as an indicator of reliable results (e.g. Haslinger et al., 1999). The reflector depth grids were set up and tailored in the same way. Adriatic and European Moho finally consist of only two nodes each, i.e. both interfaces are flat and have no offline dip. The above-mentioned intracrustal reflector south of the Inn Valley (60–90 km) was modeled with x
y = 2
2 nodes to take a possible offline dip into account. 3.3. Inversion In order to stabilize the inversion, several other parameters have to be adjusted. Weighting and damping numbers are crucial to compensate for data heterogeneity and to steer the inversion. In local earthquake tomography (LET), weighting numbers are important to take into account the accuracy of individual picks and to suppress the influence of mispicks. As we use travel time data from correlated seismic sections, our picking is generally more accurate, and we refrained

from using this kind of weighting. However, trial inversions have shown that it is useful to apply a gentle distance weighting, strengthening the influence of small offsets. The weight of reflections with regard to refractions was decreased by a factor of 0.5. Such weighting helps to develop the model from top to bottom. It yields a compromise between layer stripping and simultaneous inversion of all parameters. Layer stripping works well, if there is heterogeneity in the near-surface only. Simultaneous inversion is preferable, if there is complexity at all depths (Zelt, 1999). We expect overall complexity, but we do not have the resolving power at deeper levels. Damping numbers (Fig. 6) were computed after Eberhart-Phillips (1986). The damping for station delays, which are inverted too, was adjusted empirically, so that the final static shift for a station does not exceed 0.1 s. In total, we used 8600 observations to constrain 312 nodes (207 velocity nodes, 97 static shifts and 8 reflector depths). Initial velocity distribution, grid parameterization, weighting factors and damping numbers together determine the inversion course (Fig. 6c). In five inversion steps, the weighted RMS residual decreases from 218 ms to 53 ms. Further model changes are insignificant, and the inversion is stopped, and the final solution (Fig. 7) is obtained. 3.4. Resolution and error estimation Fig. 8 shows some examples of seismic sections with calculated travel times and ray paths to illustrate the quality of the data and the travel time fit. RDE values (Fig. 5b) indicate well-constrained nodes throughout the whole model. However, due to the

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Fig. 6. (a, b) Damping number evaluation for velocities and reflector depths. These diagrams show travel time residuals vs. heterogeneity of the pseudo 1D initial model after one inversion step for different trial damping values. As a measurement for heterogeneity we use the average of the standard deviation (sdv) of the velocities in each layer (a, c) and of the depth of each reflector (b). The residual for infinite damping is equal to the initial residual of all observations (a) and of reflections only (b). For the computation of the velocity damping, we fixed the reflector depths and vice versa. For the simultaneous inversion, we used values of 15 for velocities and 0.2 for depths. (c) Inversion steps for the pseudo 1D initial model and a 2D initial model based on the results of Miller et al. (1977). After 5 inversion steps (3 for the 2D initial model), velocity changes become insignificant. Both inversions converge to a similar travel time residual at similar velocity heterogeneity. This indicates the independence of the results from the starting model.

non-linearity of the inversion, the resolution values alone (be it the full matrix or its diagonal elements only) are an insufficient measure of the solution quality. In order to better assess the quality of the solution, and to investigate blurring and leakage, we computed and inverted synthetic data sets. One major problem in constructing adequate synthetic models is the adapted grid in combination with the strong non-linearity of refracted and reflected ray paths. Each synthetic data set has a different ray distribution and would therefore require a different grid adaptation. However, it is not meaningful to modify the inversion grid to match the resolving power of the synthetic data, because the resolution for a modified grid is not representative. Therefore, instead of using checkerboard tests, which may cause entirely different ray distributions, we con-

structed synthetic models focusing on specific problems. These models are similar to the solution (Fig. 7), and differ in one or two aspects only. We also separated the resolution tests for the upper and the lower crust. Here we present tests for the lower crust only. The first model (Fig. 9a) is a dcharacteristic modelT (Haslinger et al., 1999; Lippitsch et al., 2003), containing anomalies shaped similar to the solution, but with opposite sign: The velocities in the North are fast compared to the South. The Moho is symmetric. The node configurations of the synthetic model and the inversion grid are identical, i.e. the solution could theoretically reproduce the synthetic velocity distribution perfectly. The result shows good agreement for the large-scale structure as well as in detail, although se-

Fig. 7. Inversion result based on the pseudo 1D initial model. Apart from the foreland basins (b25 km and N220 km), typical Alpine velocities range from 5.8 to 6.8 km/s. Velocities in the middle and lower crust of the European (northern) part are generally slower than in the Adriatic part. Upper mantle velocities are unconstrained due to the lack of Pn observations. Vertical exaggeration is 1.75. Corresponding RDE values are displayed in Fig. 5b.

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veral RDE-values are below 0.2 due to the different ray distribution. The largest error pertains to the dip of the northern Moho and the velocities right above it. It is a result of the general velocity–depth ambiguity of reflec-

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tions in combination with a reflector depth node distribution, which was not tailored to the reflection points. The other synthetic models (b, c) were produced with a different node distribution and contain gradients

Fig. 8. Some examples of shot and Vibroseis registrations with calculated travel times and corresponding ray paths, plotted vs. profile coordinate. The offset from the shot point or station is marked at the top of the seismogram sections. (a) SP 2 at the northern margin of the Alps (top) and crossline-SP Q4W close to the central crest (middle). (b) Binned Vibroseis sections for station AAS in the NCA (top) and IRO in the Dolomites (middle). For locations see Fig. 2. Section AAS shows strong upper crustal reflections. In this trace-normalized display, their high amplitudes mask the first breaks beyond ca. 10 km offset. The shallower reflection seen between 10 and 30 km offset (corresponding to ca. 4 km depth) was not modeled, because it is observed only at this station. Note the small penetration depth of the direct wave in all sections, typical for the Alps. The middle and lower crustal velocities are therefore constrained by wide-angle reflections from the Moho only.

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Fig. 8 (continued).

and anomalies, which cannot be resolved by the inversion grid. Model (b) shows a similar large-scale velocity distribution as the drealT solution, but contains a strong horizontal gradient in the lower crust, and the Moho structure is completely different. In the solution, the gradient is smeared between the neighbouring nodes resulting in an absolute local error of more than 0.5 km/s, but with a wavelength, which is smaller than the local grid node distance. This smearing obviously produces some leakage towards the South. The general structure, however, particularly the Moho depth, is reproduced well. Model (c) is the most extreme model containing an unrealistic Low Velocity Zone (LVZ) (5.6 km/s) in the lower crust. The inversion grid contains no node in this area and the anomaly blurs the whole solution, producing amongst other errors an artificial LVZ in the middle crust. In the face of this, it is remarkable that the Moho is optimally recovered, and that the general trend of the velocities is reproduced, too. In all models, errors in the upper crust are very small for the most part. Only the foreland basins show large

absolute errors, and it was checked that they have little influence on wide-angle travel times. The intracrustal reflector at 60–90 km, which is flat in models b and c, shows a slight dip towards the North in the reconstructions. This error is caused by the strong velocity variations above it, and may also affect the solution (Fig. 7). RMS velocity errors amount to 0.1 km/s in the upper crust between profile 30 and 210 km for all three models, and to 0.1 and 0.2 km/s in the middle and deeper crust for models (a) and (b), respectively. Moho depth RMS errors of the three models are 1–2 km in average. We regard these values as representative for our solution. 4. Anisotropy These error estimations are based on isotropic calculations. However, from recordings of shear-wave splitting on a crossline at the northern margin of the Tauern Window (TW) (Fig. 10), it is clear that this restriction is at least locally incorrect. Anisotropy amounts to 10% in the upper 2 km, which can be explained supposing a Ndipping foliation of the TW rocks on its northern side

F. Bleibinhaus, H. Gebrande / Tectonophysics 414 (2006) 51–69 Fig. 9. Resolution tests by inversion of synthetic data. Travel times were computed for synthetic models a, b, c (top) and inverted with the same initial model (Fig. 4) and control parameters used to obtain the drealT solution. For each synthetic model we show the inversion results (middle row) and the absolute differences between the two (bottom row). RMS differences in Moho depth are printed below the reflectors, isolines contour velocity differences of F0.2 and F0.5 km/s. In all models, the initial RMS residual is similar to the real data. It decreases to 0.05 s after three (a) or five (b, c) iterations, respectively.

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Fig. 10. Cross-line recordings Q3 of a shot in the Tauern Window. The inset displays the cross line and the shot point Q3W on a tectonic map (Fig. 1), see also Fig. 2 for location. The P-wave and two distinct S-wave phases are observed. The latter are marked by dotted lines in the offset region 8–12 km, where the high ambient noise in the Ziller Valley suppresses the signal. The phases show almost constant velocities of approx. 6.0 km/s (P), 3.4 km/s (SO) and 3.1 km/s (S?) beyond 4 km offset yielding an anisotropy coefficient of ca. 10%.

(Fig. 11a). In order to better constrain the vertical and horizontal extent of anisotropy, the dependency of the average first break velocities on the azimuth of the observation was investigated for different tectonic units along the profile (Fig. 12). These data reveal that anisotropy in the upper crust is restricted to the TW. They confirm the value of 10% in the upper 2–3 km, which was determined by the analysis of shear-wave splitting. Anisotropy within the TW could be generated by folded foliations (Fig. 11b) or lineations, which have evolved during ascent and ductile lateral extrusion of the TW (Ratschbacher et al., 1991). Investigating anisotropy in deeper parts of the TW would require larger E–W-offsets as well as higher coverage to distinguish effects produced by heterogeneity. In our tomography, we simply neglected anisotropy, because the majority of observations is oriented N–S. Therefore, the velocity function for the TW is probably

distorted in the upper 3 km by the varying observational azimuths. At greater depths all velocities refer to N–Sdirected propagation. 5. Comparison with older models A comparison of our final model (Fig. 7), which we will refer to as model B in this context, with older refraction and wide-angle reflection seismic models by Miller et al. (1977) and Yan and Mechie (1989), in the TW (Fig. 13, models M and Y) provides further constraints on the quality of our solution. The models M and Y are relatively smooth and in good agreement down to a depth of 18 km. In contrast, the velocities we found are significantly smaller, differing by 10% at the surface, 0% at sea level and 5% at 17 km depth. The differences at shallow depths are clearly related to anisotropy, since the models M and Y are based on the E–W-oriented observa-

Fig. 11. Sketches illustrating mechanisms of anisotropy. The N-dipping foliation at the northern rim of the TW splits the shear wave into a fast, parallel and a slow phase (a). Both S-phases can be observed on the vertical component. The P-wave velocity is high within the foliation plane and slow perpendicular to it. This mechanism doesn’t hold for the bulk of the TW, which is internally folded and faulted, with an axis oriented E–W. The resulting anisotropy shows a fast P-wave in E–W-direction only (b). The same would be the case for an E–W-oriented lineation.

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Fig. 12. Average first-break velocities versus offset for different geological units (for abbreviations see Fig. 2). The azimuth of the observations is colour coded. The diagrams primarily reflect the near-surface velocity profile (see Figs. 7 and 15), emphasizing that the highest velocities are not found in the high-grade metamorphic core complexes of the Alps but in the NCA. In some areas, particularly to the South, no E–W-oriented observations are available at all, and no statement about anisotropy can be made. Most other diagrams reveal no significant dependency of the velocity on the azimuth. Only in the Tauern Window a significant difference of 10% between the fast (5.6 km/s) and the slow velocity (5.1 km/s) is observed. The apparent decrease of this difference with offset reflects a progressive turn of the azimuths from E–W to N–S with increasing offset and does not necessarily indicate a decrease of anisotropy with depth.

tions of the ALP 75 main line (Fig. 1). The differences at 5–17 km depth might be related to anisotropy in the TW as well. If so, one could assess its approximate depth extent. However, one has to take into account that the velocities in this depth region are averaged over several tens of kilometers in E–W- (models M and Y) and N–Sdirection (model B), respectively. This means that unresolved heterogeneity could account for these differences as well, and that there is no possibility to distinguish between these effects based on the present data. Velocities of models B and Y differ at greater depths by 0.1 km/s only. This agreement confirms the low velocities of the lower crust (6.4 km/s) in this area. Moho depth is 49 km in model B, 48 km according to

model Y and 47–48 km in model M. The depth distortion in the models along strike, caused by the southward dip of the European Moho of ca. 78, amounts to 0.4 km only and can be neglected. The vertical Moho two-way travel time (TWT) amounts to 16.8 s in model B and 15.8 s in models M and Y, respectively. These values can be compared to the reflection Moho as inferred from near-vertical data (Fig. 14). Mooney and Brocher (1987) showed in a study of coincident near-vertical and wide-angle DSS profiles that the refraction Moho generally coincides with the base of deep (i.e. lower crustal) near-vertical reflectivity patterns. Synthetic data for a laminated or random-media lower crust computed by e.g. Wenzel et

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model B is small, and may be attributed to the different lateral resolving power of the data. The smaller Moho TWT of 15.8 s, predicted by the models M and Y, does not match the near-vertical data. This discrepancy between the models is accumulated mainly in the upper crust, and is therefore partly caused by anisotropy. 6. Interpretation

Fig. 13. Comparison of v(z) functions in the Tauern Window at the intersection of TRANSALP and ALP 75 (profile coordinate 92 km; see Fig. 1 for location) between our model (B) and older refraction/ reflection seismic results by Miller et al. (1977) (M) and Yan and Mechie (1989) (Y), based on data from the ALP 75 main line. The model surface of the older models was related to 1.5 km asl. Models M and Y agree perfectly in the upper 20 km, where model B is 5–10% slower. Below 30 km depth, model B and Y are very similar, and Moho depth coincides within 1 km. Lower crustal variations of model M cannot be confirmed.

al. (1987) or Gibson and Levander (1988), respectively, can explain this finding, in that near-vertical data map small-scale perturbations in the lower crust rather than the classical refraction Moho. At the marked position, the bottom of the reflection Moho lies between 17.0 and 17.4 s. The difference to our

In order to facilitate interpretation of the velocity model, we have overlain a line drawing of a migrated reflection seismic section (Fig. 15). 6.1. The upper crust Generally, the correlation of upper crustal velocity structures with reflectivity is remarkably good. From north to south, the profile crosses the following units: the Bavarian Molasse (0–25 km), the Austroalpine (AA, 25–125 km) surrounding the Tauern Window (TW, 80– 110 km), and the Southern Alps (125–240 km). The AA comprises the Northern Calcareous Alps (NCA, 25–55 km), with velocities as high as 6.0 km/s already at the surface, a quartzphyllite zone (QPZ, 55–80 km), and various gneisses in Val Tures (UAG, 110–125 km). The folded and faulted S-dipping reflective sheets of the NCA are separated from the QPZ (b5.8 km/s) by a

Fig. 14. Detail from the time-migrated shot section (Lu¨schen et al., 2004). The vertical line marks the position of the velocity functions (Fig. 13). The reflective band at 16–18 s is interpreted as European lower crust north of the Alpine root. Two shot records of high quality basically produce this image, which is blurred by smiles. Two solid lines delimit the region where the lower crustal reflectivity ends. Their dip (78) corresponds to the average dip of the European Moho inferred from the whole shot section. (The section is scaled 1:1 for an average velocity of 6.4 km/s.)

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Fig. 15. Manual line drawing derived from a combination of the migrated Vibroseis and shot sections (TRANSALP Working Group, 2002) displayed together with the P-wave velocities. White lines denote wide-angle reflectors. Large arrows mark the position of the Sub-Tauern-ramp after TRANSALP Working Group (2002), small arrows indicate the approximate onset of lower crustal reflectivity in the Adriatic plate. A star denotes the projected position of a cluster of local earthquakes (see Fig. 18), circles mark projected positions of two magnitude 5 earthquakes after Kummerow et al. (2004). Mantle velocities are unconstrained. For geological units see Fig. 2.

significant velocity contrast south of the Inn Valley. The intracrustal wide-angle reflector at 12 km depth below these units deviates by up to 5 km from the underlying, slightly S-dipping reflecting elements. Detailed investigations of the near vertical reflectivity at the border QPZ/ TW by 3D prestack depth migration of crossline data resolve a series of reflectors at 8–14 km depth (Fig. 16). The largest continuous reflector at 12 km depth correlates well with the wide-angle reflector from tomography, except for the dip. The tomographic results, particularly the dip, might be distorted by wrong phase correlations, i.e. phases from different reflectors might have been merged into one. Despite such inconsistencies, the strong reflection amplitudes (see also Fig. 8b) emphasize the importance of these reflectors, which we regard as thrust faults separating the Austroalpine (AA) nappe stacks from the underlying autochthonous European basement. Due to the high velocities of the NCA, the typical basement velocities of 5.8 km/s form a LVZ from 12 to 20 km depth. In comparison to the shallow surrounding AA units (5.3–5.7 km/s), the TW shows slightly increased velocities (5.4–5.9 km/s) only. Velocities ranging from 5.6 to 5.8 km/s down to 15 km might indicate its depth extent. Further to the South, the Periadriatic Lineament (PL, 125 km) separates the AA from late Palaeozoic granitoids and metasediments (125–135 km) of the Southern Alps. The bulk of the Southern Alps comprises mainly Mesozoic carbonate deposits (dolomites and limestones) of varying thickness with intercalations of volcanic rocks. We interpret the slow P-wave velocities around paso Falzarego (155 km) as indicative of such extrusions. South of the Valsugana Thrust Belt (VSB), a high velocity layer (6.7 km/s in average between 3 and

Fig. 16. 3D-prestack depth migration of Q3 crossline data after Ha¨hle (2003). Q3 is located at profile coordinate 85 (see Fig. 2 and inset in Fig. 10). Five WE-oriented sections at 2 km distance through a 10
10
35 km3 cube are displayed from south (front) to north (back). They show a series of three slightly S-dipping reflectors at 8–14 km depth. Amplitudes are scaled by AGC. Migration velocities have been derived from our tomographic results with additional consideration of anisotropy.

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6 km depth) of limestones and dolomites is clearly outlined by reflections. It is overlain by tertiary clastic sediments, passing into the foreland (198–212 and 220– 240 km), where this sequence has been drilled (Fantoni et al., 2003).

6.2. The Periadriatic Lineament The PL is part of the Periadriatic/Insubric fault system, separating the Southern Alps from the Western, Central and Eastern Alps over a length of 700 km (Mu¨ller

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et al., 2001). It has accumulated significant amounts of dextral displacement (possibly up to 300 km). Its western continuation, the Insubric line, forms the suture between the European and Adriatic plate (Schmid et al., 1996). On that basis, it is often speculated that the PL represents the suture in the Eastern Alps too. The PL could have been detected in our model, if it would separate units with different velocities. There are indeed differences in the upper crust between the AA in the North (5.4–5.8 km/s) and the Dolomites to the South (5.7–6.5 km/s), but there is obviously no sharp transition. This can be explained by the similar degree of metamorphism on both sides of the PL. Although there is no significant velocity contrast across the fault, there might still be the chance to trace the fault by reflections of the fault gauge. However, the migrated stacked section after Lu¨schen et al. (2004) contains no reflection seismic image of the PL. From the position of nearby undisturbed reflecting elements, Lu¨schen et al. (2004) inferred two possible dips for the PL, either 608 north or 458 south. In order to check the wide-angle data for reflections from the PL, we inserted the fault in our final model and computed hypothetical reflection travel times. As we did not want to rule out the alternative of a subvertical fault, we assumed three different possible dips. We then compared the travel times to dip-filtered common receiver sections of Vibroseis data recorded on nearby stations. We found no indications for a north- or south-dipping, but rather for a near-vertical fault in the upper 5 km (Fig. 17). The corresponding amplitudes are weak, but this is typical for reflected refractions from near-vertical faults, when the vertical velocity gradient is small. Even at the San Andreas Fault, such observations are only slightly above the noise level (Hole et al., 1996). The small amplitudes limit the strength of our conclusion. However, the fact that there are no reflections from an inclined fault emphasizes the model of a near-vertical dip. 6.3. The deep structure Significant velocity differences in the middle and lower crust between the northern (5.8–6.6 km/s) and the southern part (6.1–6.8 km/s) are a major feature of

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the final model (Fig. 7). In the vicinity of the Alpine root (100–150 km), even lower velocities ranging from 5.8 to 6.4 km/s characterize the crust. This distribution indicates a weakened crust in between two stronger blocks, namely the European and the Adriatic plate. However, considering lateral resolution at mid to lower crustal levels (Fig. 5), the position of the suture between them is not well defined. A deep-seated vertical fault related to the PL agrees as well with our model as the inclined trans-crustal ramp proposed by the TRANSALP Working Group (2002) (Fig. 15). The European Moho is constrained by wide-angle reflections from ca. 43 km depth below the NCA to 50 km below the southern TW. Its position corresponds to the bottom of lower crustal reflectivity (Fig. 15). Minor differences may arise from the use of a different velocity model for the post-stack depth migration, which was assembled from an older version of our model B in the upper crust and a version of model M in the lower crust (Lu¨schen, personal communication). On the Adriatic side, a broad reflection pattern shows a variety of dips and depths cutting the refraction Moho (Fig. 15). The reasons for such deviations between refraction and reflection seismic results may arise from the different bandwidth of the data. Unlike wideangle data, near-vertical reflections are sensitive to small-scale heterogeneities, which often cannot be distinguished from the large structures (Long et al., 1994). In this context, it is important that the position of the refraction seismic Adriatic Moho coincides with the depth-migrated Moho derived from teleseismic data (TRANSALP Working Group, 2002; Kummerow et al., 2004). Both methods agree within F2 km on a depth of 40 km for the Adriatic Moho. Since wideangle reflections and teleseismic conversions directly map the crust–mantle boundary, we assume that the greater depth extent of the near-vertical reflectivity pattern in this area is caused either by side reflections or by heterogeneities within the uppermost mantle. At the Alpine root between 100 and 150 km, the refraction Moho shows a wide gap (Figs. 7 and 15) as a result of a lack of corresponding reflection observations (Fig. 3). The signal-to-noise ratio of our observations

Fig. 17. Receiver gathers of dip-filtered Vibroseis data recorded at station IRO (a, b) and IGS (c), see Fig. 2 for location. The dip filter was adjusted to reject apparent velocities, which do not match computed travel time curves of reflections from the PL. For station IRO two dips were tested, nearvertical (a) and 458S (b), and the corresponding raypaths are displayed below. Deviations between reflection points and the reflector in the ray diagrams arise from the 2D projection of the 3D geometry (fault strike is 778 with respect to our model). Crosses denote first-breaks and arrows indicate travel times of reflected refractions or reflections from the PL. While the travel time curves for a near-vertical reflector are almost a straight line (a, c), the 458-inclined reflector produces a bend, and the dip-filter was adjusted separately in two regions. Although the dip filter rejects the apparent velocity of the first-breaks, there are significant remains of direct P- and S-energy due to the lengthy wavelet, particularly at near offsets (c). We interpret the coherent amplitudes in the prolongation of the arrows in (a) and (c) as reflections of a near-vertical PL from up to 5 km depth. (The small deviation of the fault dip from vertical (a) has technical reasons.) There are no indications for an inclination of 458S (b).

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does not indicate that this is caused by insufficient data quality. However, the complex ray path geometry of overcritical reflections from the deep European Moho may account for the gap: Such rays would be scattered or deflected at the southern edge of the shallower Adriatic Moho. Therefore, the apparent Moho gap must not be a real feature. Indeed, the receiver functions, which are based on rays coming from below, provide a continuous image of the deep European Moho along the TRANSALP profile below the Alpine root down to 55 km depth at ca. 46.68N (135 km) (Kummerow et al., 2004). Since we observe wideangle reflections of the Adriatic Moho as far north as 150 km, a possible gap could not exceed 10–20 km width. In other words, the transition from the European to the Adriatic Moho must be rather steep or discontinuous, as it seems to be the case in the Swiss Alps (Pfiffner et al., 1997). 7. Local earthquakes Several local earthquakes and unknown blasts were recorded during the 2-month acquisition campaign from

10/99 to 11/99 by a network of 32 stations. 118 of these events are closer than 50 km to the profile, and were localized in the final model (Fig. 18). We did not use the events to simultaneously constrain velocities, because our network geometry is not really suited for local earthquake tomography. In the northern part of the profile, seismicity concentrates in the upper crust around the Inn Valley. At ca. 12 km depth, hypocenters align on the intracrustal reflector constrained by active seismics, emphasizing its importance as a basal thrust plane. The shallower events are located around the southern border of the NCA, and their distribution corresponds loosely to a series of steep faults in the Inn Valley shear zone after Ortner et al. (2005—this volume). Within the short acquisition period of two months, local seismicity is basically restricted to the AA nappes. Two events locate within European basement at 17 and 25 km depth, but their offline distances of 30–40 km might lead to large depth errors. Reiter et al. (2003) and Kummerow et al. (2004) report maximum focal depths of approximately 12 km for the Inn Valley area, mainly based on long-term observations of the Austrian Zentralanstalt fu¨r Meteorologie und Geodynamik in

Fig. 18. Local earthquakes and blasts recorded during two months (10/99–11/99) by a network of 32 stations (inverted triangles). 90% of the events on display have RMS residuals below 0.1 s, none of them is above 0.2 s. Each event is observed by 11 stations on average. Dashed lines on the map denote major faults in the Southern Alps after Bigi et al. (1990). The depth of events is grey-coded (hollow = surficial, light grey = deeper than 1 km bsl, dark grey = deeper than 15 km bsl). Three clusters of hollow circles correspond to known quarries at the northern and southern margin of the NCA (Q) and one cluster in the Southern Alps could be identified as a building site (B). On the side views, circles denote events closer than 30 km to the profile and stars mark more distant ones, the depths of which are not well constrained. Event magnitudes range between 1 and 2. The velocities displayed are identical to Fig. 7. Dotted lines indicate hypothetical faults.

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Vienna, who maintains four short-period stations in the area 25–60 km west of our profile. To the South, most events are located in or close to the VSB. Many of them are further than 30 km away from the profile, and their hypocenters are not well constrained. However, a group of nine events, observed by 15 stations in average, clusters in the SW around 46.188N–11.288E at 20–25 km depth. Such significant clustering of independently constrained hypocenters indicates a linear behavior of the inversion, despite large offline distances. Together with two nearby deep (18 and 23 km) magnitude 5 events in the TrentoVicenza region reported by Kummerow et al. (2004), the cluster was projected on the profile along the strike of the Eastern Alps (Fig. 15). The projected hypocenters align with the northward dipping top of the lower Adriatic crust, which is outlined by the onset of lower crustal reflectivity, which correlates roughly with the transition to velocities greater than 6.5 km/s, and which also appears as a prominent intracrustal converter in the TRANSALP receiver function image (Kummerow et al., 2004) south of 200 km. Based on the alignment of the seismic activity with the converter, Kummerow et al. (2004) interpret it as indicative of a thrust fault, which might accommodate crustal shortening. Our findings clearly corroborate the correlation of seismicity with the top of the lower crust, thereby supporting this interpretation. 8. Discussion Our model shows much detail in the upper crust and is in good agreement with surface geology. An anisotropy value of 10% for the upper 3 km of the western TW rocks indicates strong foliation or lamination, emphasizing the importance of ductile lateral extrusion. Comparisons with E–W-directed models suggest an anisotropic crust in the TW area down to 20 km depth, although the resolution is too low to distinguish between anisotropy and heterogeneity at mid-crustal levels. We interpret the seismic reflector south of the Inn Valley to indicate the contact between AA nappes and European basement. We suppose that the aligned seismicity is related to thrusting along the Sub-Tauern ramp, due to crustal shortening. At depths greater than 12 km, the Sub-Tauern ramp might act as a ductile shear zone without release of seismic energy. The lack of seismic activity at the PL, on the other hand, questions its connection to current orogenic processes. The large-scale P-wave velocity distribution correlates with temperature models for the TRANSALP line

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by Vosteen et al. (2003). Their inversions based on petrophysical heat-flow data reveal significantly higher temperatures at the Alpine root compared to the Adriatic. The low temperatures and the high velocities of the Adriatic crust also indicate a stronger, more rigid crust than on the European side. This can explain the seismicity on the Adriatic side, which is relatively deep for typical continental crust (Maggi et al., 2000). Analogue models of the East Alpine evolution by Willingshofer et al. (2003) brought evidence that the polarity of underthrusting depends on the relative strength of the plates involved. In all their experiments, the weaker plate was forced down. These results, together with the mutual position and dip of the refraction seismic Moho (Fig. 15), indicate southward subduction of Penninic oceanic crust, and are consistent with the first interpretations of the TRANSALP Working Group (2002). In the first instance, the tectonic situation in the Eastern Alps seems to be very similar to the Central Alps, where interpretations of DSS data by Pfiffner (1992) and Schmid et al. (1996) brought evidence for a S-directed subduction of Penninic oceanic crust. But new results from teleseismic tomography question this continuity (Lippitsch et al., 2003). Based on the position and dip of high velocity zones in the upper mantle, Lippitsch (2002) infers a reversal of subduction direction right below the TRANSALP line, from S-directed in the Central Alps to N-directed in the Eastern Alps. However, our model contains no hints for such a reversal. Position and dip of the crust–mantle boundary, which is in good agreement with the receiver functions (TRANSALP Working Group, 2002), indicate S-directed subduction. Moreover, ALP2002 shots (Bru¨ckl et al., 2003) recorded on the TRANSALP line showed continuity of the Moho structure east of the profile (Bleibinhaus et al., 2005—this volume). 3D gravity modeling by Ebbing (2004) also indicates lateral continuity. A reversal of subduction direction in time from northward to southward at a late stage of collision could explain the contradiction, but this question is beyond the reach of our wide-angle observations. Future interpretations of the ALP2002 DSS data will provide new constraints, revealing the crustal structure east of the TRANSALP line. Acknowledgements The TRANSALP programme is jointly financed by the Bundesministerium fu¨r Bildung und Forschung (BMBF, Bonn), the Bundesministerium fu¨r Wissenschaft und Verkehr (BMWV, Vienna), the Consiglio Nazionale

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delle Ricerche (CNR, Rome) and the company ENIAGIP (Milan). Seismological stations were provided by the Geophysical Instrument Pool Potsdam. Deployment and maintenance during the two campaigns was supported by T. Kraft. We thank M. Bopp, J. Kummerow and M. Tillmans for providing data from their subprojects. Helpful comments by K. Grubbe, J. Ansorge and an anonymous reviewer improved the text. References Alpine Explosion Seismology Group, 1976. A lithospheric seismic profile along the axis of the alps, 1975: I. First results. Pure Appl. Geophys. 114, 1109 – 1130. Angenheister, G., Bo¨gel, H., Gebrande, H., Giese, P., SchmidtThome´, P., Zeil, W., 1972. Recent investigations of surficial and deeper crustal structures of the Eastern and Southern Alps. Geol. Rundsch. 61, 349 – 395. Bigi, G., Cosentino, D., Parotto, M., Sartor, R., Scandone, P., 1990. Structural model of Italy and gravity map. Quad. Ric. Sci. 114 (3) (9 sheets). Bleibinhaus, F., 2003. Entwicklung einer simultanen refraktions- und reflexionsseismischen 3D-Laufzeittomographie mit Anwendung auf tiefenseismische TRANSALP-Weitwinkeldaten aus den Ostalpen. PhD thesis, Univ. Munich, 171 p. Bleibinhaus, F., Bru¨ckl, E., ALP 2002 Working Group, 2005. Wideangle observations of ALP 2002 shots on the TRANSALP profile: linking the two DSS projects. Tectonophysics. 414, 71–78 (this volume) doi:10.1016/j.tecto.2005.10.027. Bo¨hm, G., Galuppo, P., Vesnaver, A., 2000. 3D adaptive tomography using Delaunay triangles and Voronoi polygons. Geophys. Prospect. 48, 723 – 744. Bru¨ckl, E., Bodoky, T., Hegedu¨s, E., Hrubcova´, P., Gosar, A., Grad, M., Guterch, A., Hajnal, Z., Keller, G.R., Sˇpicˇa´k, A., Sumanovac, F., Thybo, H., ALP2002 Working Group, 2003. ALP 2002 seismic experiment. Stud. Geophys. Geod. 47, 671 – 679. Crosson, R.S., 1976. Crustal structure modeling of earthquake data: 1. Simultaneous least squares estimation of hypocenter and velocity parameters. J. Geophys. Res. 17 (81), 3036 – 3054. Ebbing, J., 2004. The crustal structure of the Eastern Alps from a combination of 3D gravity modelling and isostatic investigations. Tectonophysics 380, 89 – 104. Eberhart-Phillips, D., 1986. Three-dimensional velocity structure in the Northern California Coast Ranges from inversion of local earthquake arrival times. Bull. Seismol. Soc. Am. 76, 1025 – 1052. Fantoni, R., DellaVedova, B., Guistiniani, M., Nicolich, R., Barbieri, C., DelBen, A., Finetti, I., Castellarin, A., 2003. Deep seismic profiles through the Venetian and Adriatic foreland (Northern Italy). Mem. Sci. Geol. 54, 131 – 134. Gibson, B.S., Levander, A.R., 1988. Lower crustal reflectivity patterns in wide-angle seismic recordings. Geophys. Res. Lett. 15 (6), 617 – 620. Giese, P., Buness, H., 1992. Moho depth. In: Blundell, D., Freeman, R., Mueller, S. (Eds.), A Continent Revealed — The European Geotraverse. Cambridge Univ. Press. Giese, P., Reutter, K.-J., 1978. Crustal and structural features of the margins of the Adriatic microplate. In: Closs, H., Roeder, D., Schmidt, K. (Eds.), Alps, Apennins, Hellenides, pp. 565 – 588. Giese, P., Reutter, K.-J., Jacobshagen, V., Nicolich, R., 1982. Explosion seismic crustal studies in the Alpine Mediterranean region

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