Discussion and Conclusion

[19]  One of the most noticeable features of the inferred structures is anomalously low velocities in the upper crust. Thus, Vpsim 4.1  km s -1 and Vssim 2.4  km s -1 in the uppermost layer about 1 km thick. Figure 4 shows the results of the average receiver function inversion from data of 12 earthquakes recorded at the stations C1, C3, and C5. If the receiver function is inverted from the data of the station C5 alone (Figure 2), the inverted velocities in the upper layer are even lower: Vp = 3.5 km s-1 and Vs = 2 km s-1. Apparently, these are the lowest velocities that have ever been observed in the upper crust of the Karelia region. In any case, such low velocities are absent in the velocity structures presented in [Sharov, 2004]. The velocity Vpsim 5.5 km s-1 is attained at a depth of about 3 km, and Vpsim 6.2 km s-1, only at a depth of about 10 km. Low velocities in the upper hundred meters are not surprising. All RUKSA stations were mounted on moraine deposits. However, anomalously low velocities are sometimes also observed in bedrocks of the upper crust [Vinnik et al., 2002]. Grad and Luosto [1992] reported relatively low velocities in the upper crust from the SVEKA profile in Finland. The velocities were determined from observations of short-period (0.5-2.5 Hz) Rayleigh waves. The lowest surface wave velocity is 2.7 km s-1, which yields an S wave velocity of 2.9 km s-1. Based on the analysis of O'Connel and Budiansky, 1977], Grad and Luosto attributed the relatively low velocities to the fracturing of upper crust rocks provided that the cracks are isolated and filled with water. Our observations yield much lower velocities that can also be accounted for by the presence of cracks filled with water except that all or some part of the cracks are interconnected. According to the calculations of O'Connel and Budiansky, an S wave velocity decrease in the case of isolated cracks reaches 25% compared to a monolithic rock. The concentration of cracks was assumed to be 0.5. If the cracks are not isolated, the velocity decrease is 60%. In our case, the observed S velocities amount to sim2.4 km s-1, whereas a value of 3.6 km s-1 is expected for rocks of the granite type. The velocity decrease amounts to 40% and, as distinct from the SVEKA profile observations, can accounted for by partially interconnected cracks. This inference could be supported (or disproved) by estimating the crack density from measurements of the number and sizes of fractures in the RUKSA area. Vinnik et al. [2002] noted a 30% decrease in S velocities at some points of the upper crystalline crust in central Tien Shan and interpreted this decrease in terms of the fracturing in the shallow layer. Our data are insufficient to assess to what extent the velocity decrease is affected by moraine deposits, the isolation or interconnectedness of cracks, and the possible presence of particularly low velocity rocks in the volcanogenic sedimentary series. All three factors appear to be significant.

[20]  In the lower and upper crust, at depths of 10 km to 40 km, the P wave velocity slowly rises to about 6.6 km s-1. Its average gradient at these depths is an order of magnitude smaller than in the upper crust at depths of 0 km to 10 km. The Moho depth is about 40 km. The P velocity immediately under the Moho is 8.4 km s-1. This value is much higher than in the IASP91 model but is consistent with modern models of ancient shields. Thus, the velocity structure obtained in this work consists of an essentially heterogeneous upper part where the velocity (converted into P wave velocity values) rises from 3.5 km s-1 at the surface to 6.2 km s-1 at depths of 10-12 km and nearly homogeneous middle and lower parts with minimal velocity contrasts in the corresponding layers. Velocity inversion zones are virtually absent in the final variant of the velocity structure.

[21]  Previously we presented and discussed the velocity structure beneath RUKSA obtained from a receiver function alone, without invoking data on surface waves and traveltimes of converted waves from the 410-km boundary [Sharov, 2004]. The main features of the structure, the Moho depth and average velocities in the crust and the mantle, are close to those inferred in our study. Distinctions are mainly observed in the upper part of the section and can be attributed to the instability of the inversion using data of the receiver function alone. The extent of misfit (the rms deviation) between the synthetic and observed waveforms of the published structure is 1.5 times greater than in our present work.

[22]  Reliable features of the presented models (Figure 4) are the presence of a sharp boundary at a depth of about 40 km (the Moho), anomalously low velocities in the depth interval 0-3 km that are untypical of ancient shields, a rapid rise in the velocity with depth in the upper crust, and weak differentiation in velocities of the middle and lower crust. Significant heterogeneities in the upper part of the crystalline crust to depths of 12-15 km were previously noted in studies of the fine structure of seismic wavefields from small aperture array data [Nevsky and Riznichenko, 1980].

[23]  Thus, using data of short-period instrumentation with improved characteristics, a detailed velocity structure is obtained beneath the RUKSA small aperture array. These data are necessary for remote location of seismic events. They can also be used for the construction of a 3-D lithosphere model of the European part of the Russian Federation. This approach is promising if mobile small-aperture arrays and natural sources (remote earthquakes) are used for local mapping of crustal boundaries from seismic data of relatively short period instrumentation.


Powered by TeXWeb (Win32, v.2.0).