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Structure of the crust in the Baikal rift zone and adjacent areas from Deep Seismic Sounding data
Tectonophysics 351 (2002) 61 – 74 www.elsevier.com/locate/tecto
Structure of the crust in the Baikal rift zone and adjacent areas from Deep Seismic Sounding data Vladimir D. Suvorov a,*, Zabina M. Mishenkina a, Gennadii V. Petrick a, Ivan F. Sheludko a, Victor S. Seleznev b, Victor M. Solovyov b a
Institute of Geophysics SB RAS, 3, Academician Koptyuga Pr., Novosibirsk 630090, Russia b Geophysical Survey SB RAS, 3, Academician Koptyuga Pr., Novosibirsk 630090, Russia
Abstract 2-D P-wave velocity models of the crust along DSS profiles have been used together with 2-D seismic tomography for constructing maps of the Moho topography and average P-wave velocity in the crust in Baikal rift zone (BRZ) and adjacent areas. There are significant changes of Moho depth between 35 and 50 km but there is no pronounced crustal thinning beneath the rift axis. Instead, changes in crustal thickness appear to be related to accretion entities, including the transition between the Archean craton and the tectonic collage of the Transbaikal fold belt. Taking into account the decrease in resolution of the Deep Seismic Sounding (DSS) method with depth, it was found that the crust at a depth of approximately 20 km could be divided into two parts. The interval velocities in the upper and lower parts of the crust are inversely correlated, both within the craton and in the Transbaikal belt. Locally, this relationship does not hold and as such may be used to discriminate tectonic blocks. Velocity anomalies coincident with the position of the Baikal rift zone are distinctly seen in the upper part of the crust only. D 2002 Elsevier Science B.V. All rights reserved. Keywords: Baikal rift zone; Crustal structure; P-wave velocity
1. Introduction The deep structure of the Baikal rift zone (BRZ) and adjacent areas of the Siberian craton and Transbaikal fold belt has been studied by Deep Seismic Sounding (DSS) since 1968 (Puzyrev et al., 1978; Krylov et al., 1981, 1995a). New data have been recently obtained by 2-D seismic tomography using travel-time delays of diving waves (Krylov et al., 1995b; Song et al., 1996). *
Corresponding author. Fax: +7-3832-333432. E-mail address: [email protected] (V.D. Suvorov).
In this article, we examine the 2-D P-wave velocity models of the crust along DSS profiles and estimate the depths to which the surface geological units can be traced. The variations in crustal thickness and the 3-D crustal velocity model are discussed. We use published DSS data obtained in the Baikal rift zone (Krylov et al., 1981, 1995b) and in the Yakutian kimberlite province (Suvorov et al., 1997, 1999) involving long-range DSS profiles (Egorkin, 1984). We use also 2-D seismic tomography obtained from old DSS profiles. Our results are compared with the recent tomographic models of Petit et al. (1998) and Kylakov (1999).
0040-1951/02/$ - see front matter D 2002 Elsevier Science B.V. All rights reserved. PII: S 0 0 4 0 - 1 9 5 1 ( 0 2 ) 0 0 1 2 5 - 7
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2. Geological setting The Baikal rift zone (BRZ) developed along the contact zone between the Siberian craton and the Transbaikal segment of the Sayan– Baikal orogenic belt (Fig. 1). The main geological structures of the region originated in the late Riphean – early Palaeozoic, as a result of subduction – accretion tectonics (Belichenko et al., 1994; Dobretsov, 1986; Dobretsov and Kirdyashkin, 1994; Parfenov et al., 1997). The basement of the Transbaikal belt probably consists of Proterozoic crustal, now largely invaded Palaeozoic to early Mesozoic granites. East of the North Baikal basin, up to the Vitim plateau, these granites form the Baikal – Vitim batholith. The Mongol – Okhotsk belt parallels the Transbaikal belt on its southeastern margin. It mainly has a late Permian to Mesozoic history, linked to the progressive closure of the Mongol – Okhotsk Ocean (Dobretsov, 1986). From the tectonic point of view, the basement of most parts of the BRZ belongs to the Akitkan oro-
genic belt and related Vilyui basin, emplaced in the Palaeoproterozoic (Melnikov et al., 1994). It was reactivated in the middle Palaeozoic by thrusting of the Barguzin terrane over the southern margin of the Siberian craton (Rosen et al., 1994; Masaitis, 1995). The existing models of the lithosphere beneath the Baikal rift zone and adjacent areas are primarily based on geological, gravity and magnetic data (Burov et al., 1994; Logachev et al., 1982; Pismennyi et al., 1984; Zorin, 1971; Zorin et al., 1994, 1998).
3. Data analysis Deep seismic sounding measurements have been carried out in the BRZ and surrounding areas along several profiles (Fig. 2). Refraction profiling was conducted for the regional survey of the crystalline basement, together with the recording of the PmP reflections. In the Yakutian kimberlite province, in addition to DSS profiling, a special seismic survey was carried out (Suvorov et al., 1997, 1999). This
Fig. 1. Geological setting of the southeastern part of the Siberian craton and adjacent areas. Long solid lines—limits of Precambrian terranes (after Rosen et al., 1994) and other major geological structures; dotted lines—Paleozoic sedimentary basins; double line—geological boundary of the Baikal rift zone (after Solonenko, 1968); shaded area—Baikal – Vitim granitic batholith (after Parfenov et al., 1997).
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Fig. 2. Location map of seismic profiles in the Baikal rift zone and adjacent areas of the Siberian craton and Transbaikal fold belt. Solid lines— DSS profiles (after Krylov et al., 1981); short dashed lines—profiles with only PmP reflections; long dashed lines—long-range DSS profiles (GEON Centre, after Egorkin, 1984); shaded area—region of DSS survey in the Yakutian kimberlite province (after Suvorov et al., 1999); solid triangles—seismological stations. The thick line locates the composite profile discussed in the text.
provided the most even distribution of measurements, which were used to estimate the crustal thickness from PmP reflections. A rough estimation of the crustal thickness beneath the northeastern outskirts of the BRZ (west of the Aldan shield) has been derived from local earthquake data (Suvorov and Kornilova, 1985). These data were used for constructing Moho maps and for averaging crustal seismic velocity maps. The crustal velocity distribution was studied, using a new 2-D tomographic technique and integrating also results of previous investigations. Krylov et al. (1981) proposed a three-layer model of the crust for the BRZ. The upper layer is 10 km thick and contains the majority of earthquake foci. The second layer, between 10- and 20-km depth, is characterised by a discontinuous seismic lower velocity zone. The third layer corresponds to the lower crust, beneath 20 km. In the upper 20 –25 km of the crust, velocities have been determined by 2-D seismic tomography using diving waves travel-times (Mishenkina et al., 1983; Krylov et al., 1995b). This method was used for areas with dense observations, which is not always the case
along the DSS profiles in the BRZ (Krylov et al., 1995a; Song et al., 1996). In addition, a new seismic cross-section along another composite transect (Boguchany – Ust-Kut –Nizhneangarsk– Chita) is presented (Figs. 2 and 5). The seismic velocity of the lower crust was derived from the average crustal velocity and from interval velocities in the overlying crustal layers. Calculations were performed taking into account the lateral differences in the upper crustal velocities for down-going and up-going rays of reflected PmP waves. The offsets used were on average about 150 km. For depth down to 20 km, this distance is shortened to about 50 km. So we assume that the calculated values of the interval velocities in the lower part of the crust are attributed to areas with horizontal dimensions of f 50 km. To enable comparison of the velocity variations determined at various depths, we converted the data to a uniform horizontal scale, using averaging cells of 50 50 km separately for the upper crust (0– 20 km deep) and the lower crust (20 km to the Moho depth). The same averaging was used to compare the velocity
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variation in the upper crustal intervals of 0– 10 and 10– 20 km. The averaged data were used to construct a 3-D velocity model of the crust. For some DSS profiles with sparse observations, we obtained the velocities in the upper crust from diving P-waves by the Wiechert– Herglotz method (Mishenkina, 1967; Krylov et al., 1981). These were used to interpolate the velocity anomalies between profile distances of about 200 km. Not enough data are available to examine the depth interval of 10– 20 km. Data characterising the crust are analysed at two different scales along the DSS profiles. As a first step, velocity distributions are compared between the upper and lower parts of the crust. In a more detailed scale, velocity distributions in the upper crust are compared between the 0 –10- and 10 – 20-km intervals. Indeed, calculations on model examples show that the accuracy of velocity estimation in the lower crust is affected by uncertainties in the average crustal velocity. For example, an uncertainty in average velocity of F 0.05 km/s produces an error of up to F0.1 km/s in lower crust with a 20-km thickness. If lower crustal thickness is decreased down to a 15-km and upper crust is increased to a 25-km thickness, this error grows up to about 0.15 km/s. It is assumed that the velocity error is in the order of F 0.1 –0.15 km/s for the lower crust and F 0.05 km/s for the upper crustal intervals of 0– 10 and 10 – 20 km (Mishenkina et al., 1983). This rather high accuracy is reached by smoothing the 2-D velocity cross-section in intervals of 10 km in depth and 50 km in horizontal length. Thus, with the data used, it is possible to give evidence for velocity anomalies with horizontal sizes larger than 50 km and with amplitudes that exceed 0.1 km/s in the upper crust and 0.2 km/s in the lower crust.
4. Thickness and average velocity in the crust All available seismic data were used to compile a crustal thickness map, showing large-scale structures that are identified over more than one profile (Egorkin, 1984; Krylov et al., 1981, 1995b; Song et al., 1996; Suvorov et al., 1997, 1999). For the investigated area, the crustal thickness ranges from 35 to 50 km (Fig. 3). The area with the shallowest Moho
(35 –37 km deep) is observed beneath the Baikal depression and the Vilyui basin and defines an elongated roughly northeast-striking uplift. It possibly reflects a common tectonic processes in both regions (Masaitis, 1995). The area of Moho uplift corresponds to the Proterozoic Akitkan orogenic belt (Borukaev et al., 1988; Rosen et al., 1994), which underwent tectonic reactivation in late Proterozoic and in the Palaeozoic. Also to be noted is that the existence of a pre-Cenozoic seismic basement with velocities of 4.8 –5.1 km/s at the depth of about 6 km beneath Lake Baikal has been shown by Krylov et al. (1993). In the Siberian craton, northwest of the Akitkan orogenic belt, the crustal thickness is 37 – 40 km in the south and 43 – 46 km in the north. A crustal thickness of 46 – 50 km is observed beneath the Tunka rift depression (located to the southwest of Lake Baikal), under the Aldan shield and locally south of Chita in the Mongol – Okhotsk belt. An intermediate crustal thickness of 37– 43 km is determined beneath the Cenozoic depressions in the northeastern part of the BRZ, without a clear correlation between the crustal thickness and the depressions (Krylov et al., 1981, 1995a). As a rule, the location of the BRZ is not clearly reflected by the Moho topography. The crustal thickness beneath the BRZ varies between 35 and 50 km. Abrupt local changes are observed beneath the Baikal basin, from 35 to 43– 44 km (Krylov et al., 1981; Fig. 3, insert). A Moho depression is located in the middle of the lake basin, transverse to it, and divides the crust under Lake Baikal in northern and southern parts, both with elevated Moho. For the southern part of the investigated region, the map of crustal thickness was previously constructed from DSS and gravity data using a regression technique (Zorin et al., 1986, 1989). These data show considerable differences with our DSS data. Estimations by Zorin et al. (1986) are about 4 – 5 km larger than estimations from seismic data alone. These authors obtained a crustal thickness of 45 km in the East Sayan massif and under the Tunka depression, where we show an increased crustal thickness up to 49 km. A similar situation occurs locally in the central part of Lake Baikal where our estimations of the crustal thickness are 5 –7 km greater than those from gravity data. The opposite situation is encountered in
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Fig. 3. Map of crustal thickness. Inset shows details for the Baikal region and PmP reflection profiles (short dashed lines, after Krylov et al., 1981). Dashed lines—limits of geological structures as defined in Fig. 1; sparse dotted line—boundary of low velocity anomalous mantle (7.7 – 7.8 km/s, after Krylov et al., 1981 and Suvorov et al., 1985).
the vicinity of Chita, where our estimations are 5 – 7 km less than the gravimetric ones. The map of average P-wave velocity in the crust is shown in Fig. 4 (data from Egorkin, 1984; Krylov et al., 1981, 1995b; Song et al., 1996; Suvorov et al., 1997, 1999). Within the BRZ and the Transbaikal belt, the crustal velocity varies insignificantly around 6.4 km/s (Krylov et al., 1981). In the Siberian craton, the velocity fluctuates between 6.4 and 6.6 km/s with a mean value of 6.5 km/s (Egorkin, 1984; Suvorov et al., 1999). The P-wave velocity map shows an elongated area with relatively low velocities in the northern extremity of the BRZ, striking in a northwest direction. Low velocity is also detected along the southern margin of the Siberian craton (Fig. 4). The comparison of Figs. 3 and 4 highlights the lack of correlation between crustal thickness and average seismic velocity in the crust. The crustal thickness in the BRZ and in the Transbaikal belt shows locally sharp variations, under a practically constant average
velocity. In the cratonic area, on the other hand, changes of crustal thickness are gradual and accompanied by significant velocity changes, although correlation between velocity and Moho depth variations is not observed.
5. Variations in crustal velocity along DSS profiles The seismic cross-section of the crust along the Boguchany – Ust-Kut – Nizhneangarsk – Chita profile (Fig. 2) starts from the Siberian craton, crosses the northeastern part of the BRZ and terminates in the Transbaikal belt (Fig. 5). In the Siberian craton, the crystalline basement is located at a maximum depth of 5– 7 km under the sedimentary cover and gradually rises towards the craton margin. The normal P-wave velocity at the top of crystalline basement is 6.1 –6.2 km/s. However, a drastic increase of velocity up to 6.8 km/s is revealed in the upper part of the crystalline
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Fig. 4. Map of average P-wave velocity in the crust. Thick solid lines locate transects with nuclear sources; thin solid lines DSS profiles; dotes lines only PmP reflections.
basement in a local zone beneath the Srednaya – Angara zone, rich in iron deposits. A relatively low velocity of the 6.1– 6.3 km/s borders the anomaly
whose base is at 12 –15 km deep. High velocity in the basement is observed in other regions of the Tunguska basin where Triassic basalts have been observed
Fig. 5. P-wave velocity model of the crust and uppermost mantle along a profile crossing the Siberian craton, the Baikal rift zone and the Transbaikal fold belt. Thick lines—seismic boundaries; thin lines—velocity isolines; dashed lines—less reliable velocity isolines and seismic boundaries.
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(Kyznetzov et al., 1987), and consequently could be associated with basaltic intrusions. A low velocity zone (4.8 – 6.0 km/s) appears under the 6 – 8-km deep Verkhne – Angara (Upper-Angara) depression (Fig. 5). At this place, a velocity of 6.2 km/ s is located at a depth of nearly 10 km. Further to the east, the 6.2-km/s velocity zone descends to 22 – 24 km. Another low velocity zone (6.2 km/s) is present at the same depth in the 1100 –1200-km profile segment. The maximum Moho depth (46 km) is located under the central part of the Siberian craton and is associated with a high velocity in the upper part of the crystalline basement, as it is also the case under the Baikal rift zone. The minimum Moho depth (35 km) is located at the margin of the Siberian craton. The subMoho velocity is 8.2– 8.4 km/s under the Siberian craton and 7.7– 7.8 km/s beneath the Baikal rift zone, where lower another seismic boundary is defined at a depth of 50– 58 km (called ‘‘Moho 1’’) (Krylov et al., 1981). This zone of anomalous mantle with a velocity of 7.7– 7.8 km/s (contoured in Fig. 3) extend over a large area than the geological expression of the rift zone (Krylov et al., 1981, 1993; Suvorov and Kornilova, 1985). Within the anomalous domain, Krylov et al. (1981) pointed a few segments of the uppermost mantle with a normal velocity of 8.2 km/s. Fig. 6 relates the velocity distribution between upper crust ( < 20 km) and lower crust. The crust can be divided into two distinct groups. One of them, with comparatively high velocities, characterises the cratonic areas. The other, with lower velocities, is related to the Transbaikal belt. In the latter, there is an inverse correlation between the velocity variations in the upper and lower parts of the crust: relatively high velocity in the lower crust ( f 6.5– 6.7 km/s) and low values for the upper crust (6.1 – 6.3 km/s). This inverse correlation, however, is not everywhere. This is exemplified by the Nizhneangarsk – Ust-Muya profile crossing the BRZ (Fig. 6), where a direct correlation is observed between velocities in the upper and lower crust. In addition, local areas in the Siberian craton at distances 100– 200 and 500 –700 km along the profile (Fig. 5) present unusual low velocity in both upper and lower parts of the crust. The ranges of velocity fluctuation in the upper and lower crustal layers are f 0.2 – 0.3 km/s for the Transbaikal belt, whereas the velocity variations increase up to 0.5 km/s in the upper crustal part of
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Fig. 6. Relation between velocity changes in the upper crust and in the lower crust. Crosses—Transbaikal fold belt; diamonds— Siberian craton; solid triangles—Nizhneangarsk – Ust-Muya profile (location in Fig. 2).
the craton, with a maximum velocity of 6.4– 6.6 km/s (Fig. 5). No correlation between velocity and thickness is observed in the lower crust, probably due to the irregularities of the Moho topography. These features are illustrated in detail in Fig. 7a, for the Boguchany – Ust-Kut –Nizhneangarsk – Chita profile (location in Fig. 5). There is generally an inverse correlation between the velocity in the upper and lower parts of the crust. Under the BRZ, the velocity in the upper crust is 6.0 km/s and in the lower crust, 6.8 –6.9 km/s. In neighbouring areas, the velocity in the lower crust is 6.6– 6.7 km/s. Similar values were obtained along other profiles in this part of the BRZ (Krylov et al., 1995a). Beneath the BRZ, the lowest velocity (5.6 – 5.8 km/s) occurs in the upper crustal interval of 0– 10 km while 6.4 – 6.5 km/s characterises the 10 –20 km interval. This should be compared to the 6.2 km/s at the southeastern margin of the BRZ. An example of velocity variation in the Transbaikal belt is shown along the Ust-Uda– Irkutsk– Chita profile (Fig. 7b), (Krylov et al., 1981; Song et al., 1996). A clear inverse correlation between velocity in upper and lower crust is observed in the Transbaikal belt only. Along the 400 – 600-km segment of the profile, the velocity decreases from 6.3 to 6.1 km/s in the upper crust and increases from 6.5 to 6.7– 6.8 km/s in the lower crust.
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Fig. 7. Velocity changes in the crust along the Boguchany – Ust-Kut – Nizhneangarsk – Chita (a) and the Ust-Uda—Ulan-Ude—Chita (b) profiles.
For the 200– 400-km segment covering the BRZ, the lower crustal velocity increases slightly up to 6.7 km/s and is delimited by local velocity minima on both extremities. Thus, the high velocity in the lower crust is shifted northwest relative to the low velocity in the upper crust from 6.3 – 6.4 to 6.2 km/s. For the same segment, inverse velocity variations of 0.2– 0.3 km/s between the upper crustal intervals of 0 –10 and 10 –20 km beneath the BRZ are observed. The Nizhneangarsk – Ust-Muya profile (Fig. 6) shows that velocities in both the upper and the lower
crust may vary proportionally within both the craton and the fold belt. The highest velocity (6.9 km/s) in the lower crust corresponds to 6.1– 6.2 km/s in the upper crust. In the Transbaikal fold belt as a whole (Fig. 8), a relatively high velocity (6.2 –6.7 km/s) is observed for the 10 – 20-km upper crustal interval while in the upper 10 km, the velocity is 5.7 – 6.2 km/s. The opposite situation is found for the Siberian craton where mainly low velocities (6.0 – 6.4 km/s) are obtained in the 10 – 20-km crustal layer and the
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velocity increases from 5.9 to 6.0 –6.7 km/s in the upper part. The velocities in the intervals 0– 10 and 10 – 20 km are almost similar in the craton, not considering the peculiar high velocity spot (6.4 –6.7 km/s) in the upper crystalline basement (Fig. 5). The relation between velocities in the 0 –10- and 10 –20km upper crustal layers expresses mainly a positive correlation, except locally. This is on the contrary of the Transbaikal belt, where no correlation between the velocity values is present.
6. Lateral variations in crustal velocity
Fig. 8. Relation between velocity changes in the upper crustal invervals of 0 – 10 and 10 – 20 km.
Values of averaged interval velocities in the upper and lower crust along DSS profiles are mapped in Figs. 9 – 11. The sizes of reliably traced anomalies are
Fig. 9. Velocity map of the upper crustal interval 0 – 10 km. Solid lines—locations of DSS profiles; double line—geological boundary of the Baikal rift zone.
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larger than the average distance between the DSS profiles (150 – 300 km). In areas with greater distances, interpolation of the velocity values gives only a rough picture. 6.1. Upper crustal interval of 0– 10 km In the upper crust down to 10 km, the velocity varies laterally from 5.6 to 6.4 km/s (Fig. 9). Minimum values of 5.6 –5.8 km/s are shown in the central part of the BRZ on the eastern side of Lake Baikal. The area of low velocity extends southwest along the lake as a narrow band, and terminates before the western end of Lake Baikal (as no low velocity zone is observed under the Tunka depression). Velocity values of 5.6 – 5.8 km/s are typical for acidic rocks, and their extent coincides with the Baikal – Vitim massif of Paleozoic granites in the central part of the BRZ (Pismennyi et al., 1984; Parfenov et al., 1997).
The area with low velocities of 5.6– 5.8 km/s is surrounded by a zone of 5.8 –6.0 km/s which coincides distinctly with the contours of the BRZ derived from geological and seismological data (Solonenko, 1968), except for a band that extends southwards between Ulan-Ude and Chita. Velocities of 6.0– 6.2 km/s characterise the upper crustal layer of the Siberian craton. Such values are also detected in the Transbaikal belt, southeast and south of the rift zone, though we do not have sufficient data to discuss this in detail. To the northeast, at the vicinity of Olekma River, the BRZ is expressed by an area with velocities of 6.2 – 6.4 km/s revealed from aftershock data of the 1990 SouthYakutian earthquake (Suvorov et al., 1993). 6.2. Upper crustal interval of 10 –20 km The velocity in the upper crustal interval of 10– 20 km varies over the entire territory from 5.8 to 6.4 km/s
Fig. 10. Velocity map of the upper crustal interval 10 – 20 km. Solid lines—location of DSS profiles; double line—geological boundary of the Baikal rift zone. Solid lines are location of seismic profiles.
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Fig. 11. Velocity map of the lower crust. Solid lines—location of DSS profiles; double line—geological boundary of the Baikal rift zone.
(Fig. 10). A relative decrease in velocity down to 5.8– 6.0 km/s is observed under the central part of the rift zone but the size of this area is significantly smaller than that observed in the upper 10 km of the crust (Fig. 9). The area with velocities of 5.8– 6.0 km/s correlates with the location of the Baikal – Vitim granitoid massif. From gravity data, a thickness of this massif is estimated at 20 km (Pismennyi et al., 1984; Alakshin et al., 1990). A narrow band of velocity of 6.0 – 6.2 km/s bounds this area and extends south of the rift zone, between Ulan-Ude and Chita. In the north, the velocity isoline of 6.0 km/s is drawn assuming that velocities lower then 6.0 – 6.1 km/s are absent in the upper crust north of Lake Baikal and outside the BRZ, based on increased velocity obtained along three seismic profiles (located in Fig. 10) in the superficial part of the craton. Velocities of 6.2 –6.4 km/s bound the low velocity zone in the upper crust laterally, northeast of the BRZ (Fig. 10). In the vicinity of Olekma River, the upper crustal velocity is 6.2– 6.4 km/s (seismological data of Suvorov et al., 1993).
6.3. Lower crust In the crust beneath 20 km, the lateral velocity changes are in the range of 6.55 –6.85 km/s (Fig. 11). Velocities of 6.5– 6.7 km/s are observed in the central part of the territory, covering the southeastern margin of the craton and the southwestern part of the rift zone, as well as the Mongol – Okhotsk fold belt and the eastern edge of the rift zone. Rather high velocities of 6.7– 6.9 km/s are observed in an area about 200 km wide, stretching N –S for almost 500 km and crossing the rift zone between Nizhneangarsk and Chita. The northern limit of this area is not clearly determined, but the southern limit is confidently traced outside the rift zone. A large part of the craton is also characterised by such velocity values. The area under the Baikal – Vitim granitic batholith shows high velocities of up to 6.85 km/s. Comparison of Figs. 10 and 11 allows concluding that main velocity anomalies in the crust coincide with the rift zone and are distinctly visible only in the upper 20 km of the crust. As an exception to this rule, the
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roots of the Baikal – Vitim batholith is expressed by a relatively high velocity (6.85 km/s) in the lower crust (Fig. 11) and a relatively low velocity (5.8 – 6.0 km/s) in the 10 – 20-km interval and 5.6 –5.8 km/s in the upper 10-km interval (Fig. 10).
7. Discussion and conclusions The 2-D seismic tomographic analysis of old DSS data (Krylov et al., 1981) shows new aspects of the relation between the deep seismic structure of the BRZ and the adjacent areas. The seismic interpretation is in a good agreement with other geophysical results (Pismennyi et al., 1984; Alakshin et al., 1990; Rosen et al., 1994). The average crustal velocity ranges between 6.4 and to 6.6 km/s in the Siberian craton. The velocity is almost constant at 6.4 km/s in the Transbaikal belt, in areas underline by geological units of Paleozoic, Mesozoic and Cenozoic ages. The accuracy in velocity determination is everywhere in the range of F 0.03– 0.05 km/s. The crustal thickness in both regions varies between 35 and 50 km. However, in the Transbaikal belt, the Moho depth changes sharply in narrow zones, whereas in the Siberian craton, the Moho topography is smoother. Thus, average velocity and crustal thickness are not correlated other. The transition zone between the craton and the fold belt is not clearly expressed in average velocity and crustal thickness, possibly owing to complex interactions of the pre-Cenozoic structures in these regions. In particular, thinning of the crust beneath the Baikal depression and Vilyui basin may indicate a common tectonic origin as proposed for the Palaeozoic period by Masaitis (1995). The crustal thickness derived from DSS data and those derived from gravity data by Zorin et al. (1986, 1989) may differ significantly. According to DSS seismic data, the crustal thickness of the Siberian craton, of the Mongol – Okhotsk belt in the vicinity of the Chita and under the Central Baikal basin is 4 –7 km less than that in the model of Zorin et al. (1986, 1989). Beneath the Tunka depression, the seismic crustal thickness is 4 – 5 km greater. A clear division of the crust into upper and lower parts can be postulated, taking into account
a reduction in resolution of the DSS method with depth. The relation between velocities in the upper and lower crust is different. In the craton and in the Transbaikal belt, a reverse correlation between velocity variations in these intervals was found. For the upper crustal depth intervals of 0 –10 and 10 –20 km, velocity changes present a direct correlation in the craton, while this relation is much more complex in the Transbaikal belt. For the latter, the velocities in the 10 – 20-km depth interval are f 0.2 km/s higher than in the same level in the craton, and in the 0– 10-km interval, they are 0.2– 0.3 km/s lower. Our DSS seismic model of the crust is not in accordance with the tomographic models of Petit et al. (1998) and Kylakov (1999) that used local and teleseismic data. The discrepancy may be due to the lower resolution of the tomographic method used by Petit et al. (1998) and Kylakov (1999). Those models are slightly different, although they were derived from the same data. The model of Kylakov (1999) does not show reduced velocity in the upper crust under Lake Baikal, as in our model and in the model of Petit et al. (1998). These tomographic models also show no low velocity anomalies under other rift depressions, although it is demonstrated from DSS data that low velocity anomalies beneath the large rift depressions can be traced down to depths of 6 – 14 km (see Fig. 5, Krylov et al., 1995b; Song et al., 1996). The existing tomographic models lack velocity anomalies under the large Baikal – Vitim batholith, even though its bottom is estimated at a depth of 15 – 20 km from seismic and gravity data (Pismennyi et al., 1984; Alakshin et al., 1990). Another discrepancy is a too large crustal thickness in the reference model of Kylakov (1999). The crustal thickness obtained from our model is similar to that in the model of Petit et al. (1998) under the Tunka depression (southwest Lake Baikal) and under the central part of Lake Baikal, but in other areas, it differs by up to 5– 13 km. Partitioning the crust into two velocity layers is justified by the fact that the velocity anomalies in the upper crust correlate with the geologically defined rift zone, but not those in the lower crust. Crustal layering was found for both the Archean craton and the Transbaikal fold belt, despite their very different composition and thermotectonic ages. This can be explained
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by their rheological characteristics (e.g., Lobkovsky, 1988; Avouac and Burov, 1996; Westaway, 1998).
Acknowledgements The authors express gratitude to the Novosibirsk Regional Center of Geoinformation Technologies SB RAS for the advices and help in drawing up the electronic maps. We thank an anonymous reviewer for his very useful comments and Prof. A. Khan for their hard work to improve English style of the manuscript. This work was supported by the Russian Foundation of Basic Research (Grant No. 98-05-65239). It is a contribution to the IGCP400 Project Geodynamics of Continental Rifting.
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Krylov, S.V., Mishenkin, B.P., Mishenkina, Z.R., Petrik, G.V., Sergeev, V.N., Sheludko, I.F., Ten, T.N., Kulchinskiy, Yu.V., Mandelbaum, M.M., Seleznev, V.S., Solovev, V.M., Suvorov, V.D., 1993. Detailed Investigations of Lithosphere on P- and SWaves. Nauka, Novosibirsk (in Russian). Krylov, S.V., Mishenkina, Z.R., Seleznev, V.S., Sheludko, I.F., 1995a. Methods and results of detailed seismic investigations of the earth’s crust in the Baikal rift zone. In: Mooney, W. (Ed.), 1993 CCSS (Commission on Controlled Source Seismology) Workshop Proceedings Volume, Menlo Park, CA, USA 94025, 9 – 33. Krylov, S.V., Seleznev, V.S., Soloviev, V.M., Petrick, G.V., Sheludko, I.F., 1995b. Study of Baikal rift depression by seismic tomography method on refracted waves. Reports of the Russian Academy of Sciences 345 (5), 674 – 677 (in Russian). Kylakov, I.Yu., 1999. Three-dimensional seismic heterogeneities beneath the Baikal region according to teleseismic and local tomography. Geologiya i Geofizika (Russian Geology and Geophysics) 40 (3), 317 – 330. Kyznetzov, V.L., Bgatova, G.F., Titarenko, V.V., 1987. The basement of the Siberian platform according to seismic data. Geologiya i Geofizika (Soviet Geology and Geophysics) 28 (2), 77 – 83. Lobkovsky, L.I., 1988. Geodynamics of spreading and subduction zones, and the two-level plate Tectonics. Nauka, Moscow (in Russian). Logachev, N.A., Zorin, Yu.A., Sherman, C.I., 1982. Geodynamics of continental borderland. Geologiya i Geofizika (Soviet Geology and Geophysics) 23 (12), 10 – 19. Masaitis, V.L., 1995. Connected divergent and convergent zones of the East Siberia in Devonian – Early Carboniferous. Reports of the Russian Academy of Sciences 341 (6), 781 – 784 (in Russian). Melnikov, A.I., Mazukabzov, A.M., Sklyarov, E.V., Vasiliev, E.P., 1994. Baikal rift basement: structure and tectonic evolution. Bulletin des Centres de Recherches Exploration-Production Elf-Aquitaine 18 (1), 99 – 122. Mishenkina, Z.R., 1967. Interpretation of the travel times of refracted waves at vertical and horizontal gradients of velocity. Fizika Zemli 4, 84 – 91 (in Russian). Mishenkina, Z.R., Sheludko, I.F., Krylov, S.V., 1983. Application of a linearized approach of inverse kinematic problem for 2-D time fields t(x,l) of refracted waves. In: Alekseev, A.S. (Ed.), Numerical Methods in Seismic Investigations. Nauka, SB, Novosibirsk, pp. 140 – 152 (in Russian). Parfenov, L.M., Bulgatov, A.N., Gordienko, I.V., 1997. Terranes and formation of the orogenic belts, Trans-Baikal region. Tikhookeanskaya Geologia 15 (4), 3 – 15 (in Russian). Petit, C., Kulakov, I.Yu., Deverchere, J., 1998. Velocity structure around the Baikal rift zone from teleseismic and local earthquake traveltimes and geodynamic implications. Tectonophysics 296, 125 – 144. Pismennyi, B.M., Alakshin, A.M., Pospeev, A.V., Mishenkin, B.P., 1984. Geology and seismicity of the main Baikal – Amur railway zone. Deep structure. In: Logachev, N.A. (Ed.), Geology and Seismicity of the Baikal – Amur Magistral Area. Nauka, Novosibirsk, pp. 58 – 102 (in Russian).
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Puzyrev, N.N., Babayan, G.D., Bochanov, F.I., Egorov, G.V., Krylov, S.V., Kyznetchov, V.L., Mandelbaum, M.M., Mishenkin, B.P., Monastyrev, V.K., Rudnitchkyi, A.L., Suvorov, V.D., Chichinin, I.S., 1978. Technique and Equipment of the Regional Seismic Investigations in Areas of Difficult Access and their Application in Siberia. Nauka, Novosibirsk (in Russian). Rosen, O.M., Kent, C., Natapov, L.M., Nozhkin, F.D., 1994. Archean and early Proterozoic evolution of the Siberian craton: a preliminary assessment. In: Condie, K. (Ed.), Archean Crustal Evolution. Elsevier, Amsterdam, pp. 411 – 459. Chapter 10. Solonenko, V.P., 1968. Seismotectoinics and present structure evolution of the Baikal rift zone. In: Florensov, N.A. (Ed.), Baikal Rift. Nauka, Novosibirsk, p. 5771 (in Russian). Song, Y.-S., Krylov, S.V., Yang, B., Liu, C., Dong, S., Liang, T., Li, J., Xu, X., Mishenkina, Z.R., Petrick, G.V., Sheludko, I.F., Seleznev, V.S., Soloviev, V.M., 1996. Deep seismic sounding of the lithosphere on the Baikal – Northern China international transect. Geologiya i Geofizika (Russian Geology and Geophysics) 37 (2), 1 – 15. Suvorov, V.D., Kornilova, Z.A., 1985. The depth structure of Aldan shield on data from close earthquakes. Geologiya i Geofizika (Soviet Geology and Geophysics) 26 (2), 79 – 86. Suvorov, V.D., Kornilova, Z.A., Kchudaeva, I.J., 1993. Crustal velocities in the area of the South-Yakutian earthquake of the 1990 from epicentral observations. Vulkanologiya i Seismologiya, (4), 98 – 104 (in Russian). Suvorov, V.D., Yurin, Yu.A., Timurshin, K.V., Krylov, S.V., Parasotka, B.S., Cherny, S.D., Seleznev, V.S., Soloviev, V.M., Matveev, V.D., 1997. Structure and evolution of the crust and
uppermost mantle beneath the kimberlite province of Yakutia based on seismic evidence. Geologiya i Geofizika (Russian Geology and Geophysics) 38 (2), 518 – 527. Suvorov, V.D., Parasotka, B.S., Chernyi, S.D., 1999. Deep seismic sounding studies in Yakutia, 1999. Izvestiya, Physics of the Solid Earth 35 (7 – 8), 612 – 629. (Translated from Fizika Zemli, Nos. 7 – 8, pp. 94 – 113). Westaway, R., 1998. Dependence of active normal fault dips on lower-crustal flow regimes. Journal of the Geological Society (London) 155, 233 – 253. Zorin, Yu.A., 1971. A Recent Structure and Isostasy of the Baikal Rift Zone and Adjacent Areas. Nauka, Moscow (in Russian). Zorin, Yu.A., Mordvinova, V.V., Novoselova, M.R., Turutanov, E.X., 1986. Density inhomogeneity of mantle beneath the Baikal rift. Fizika Zemli 5, 43 – 52 (in Russian). Zorin, Y.A., Kozevnikov, V.M., Novoselova, M.R., Turutanov, E.X., 1989. Thickness of the lithosphere beneath the Baikal rift zone and adjacent regions. Tectonophysics 168, 327 – 337. Zorin, Yu.A., Belichenko, V.G., Turutanov, E.X., Mordvinova, V.V., Kozhevnikov, V.M., Khozbayar, P., Tomurtogoo, O., Arvisbaatar, N., Gao, Sh., Davis, P., 1994. Baikal – Mongolian transect. Geologiya i Geofizika (Russian Geology and Geophysics) 35 (1), 94 – 108. Zorin, Yu.A., Belichenko, V.G., Tyrytanov, E.X., Kozhevnikov, V.M., Sklyarov, E.V., Tomurtogoo, O., Khozbayar, P., Arvisbaatar, N., Byambaa, Ch., 1998. Terranes of the East Mongolia and Central Transbaikalia and development of the Mongol – Okhotsk folded belt. Geologiya i Geofizika (Russian Geology and Geophysics) 39 (7 – 8), 29 – 40.
Structure of the crust in the Baikal rift zone and adjacent areas from Deep Seismic Sounding data Vladimir D. Suvorov a,*, Zabina M. Mishenkina a, Gennadii V. Petrick a, Ivan F. Sheludko a, Victor S. Seleznev b, Victor M. Solovyov b a
Institute of Geophysics SB RAS, 3, Academician Koptyuga Pr., Novosibirsk 630090, Russia b Geophysical Survey SB RAS, 3, Academician Koptyuga Pr., Novosibirsk 630090, Russia
Abstract 2-D P-wave velocity models of the crust along DSS profiles have been used together with 2-D seismic tomography for constructing maps of the Moho topography and average P-wave velocity in the crust in Baikal rift zone (BRZ) and adjacent areas. There are significant changes of Moho depth between 35 and 50 km but there is no pronounced crustal thinning beneath the rift axis. Instead, changes in crustal thickness appear to be related to accretion entities, including the transition between the Archean craton and the tectonic collage of the Transbaikal fold belt. Taking into account the decrease in resolution of the Deep Seismic Sounding (DSS) method with depth, it was found that the crust at a depth of approximately 20 km could be divided into two parts. The interval velocities in the upper and lower parts of the crust are inversely correlated, both within the craton and in the Transbaikal belt. Locally, this relationship does not hold and as such may be used to discriminate tectonic blocks. Velocity anomalies coincident with the position of the Baikal rift zone are distinctly seen in the upper part of the crust only. D 2002 Elsevier Science B.V. All rights reserved. Keywords: Baikal rift zone; Crustal structure; P-wave velocity
1. Introduction The deep structure of the Baikal rift zone (BRZ) and adjacent areas of the Siberian craton and Transbaikal fold belt has been studied by Deep Seismic Sounding (DSS) since 1968 (Puzyrev et al., 1978; Krylov et al., 1981, 1995a). New data have been recently obtained by 2-D seismic tomography using travel-time delays of diving waves (Krylov et al., 1995b; Song et al., 1996). *
Corresponding author. Fax: +7-3832-333432. E-mail address: [email protected] (V.D. Suvorov).
In this article, we examine the 2-D P-wave velocity models of the crust along DSS profiles and estimate the depths to which the surface geological units can be traced. The variations in crustal thickness and the 3-D crustal velocity model are discussed. We use published DSS data obtained in the Baikal rift zone (Krylov et al., 1981, 1995b) and in the Yakutian kimberlite province (Suvorov et al., 1997, 1999) involving long-range DSS profiles (Egorkin, 1984). We use also 2-D seismic tomography obtained from old DSS profiles. Our results are compared with the recent tomographic models of Petit et al. (1998) and Kylakov (1999).
0040-1951/02/$ - see front matter D 2002 Elsevier Science B.V. All rights reserved. PII: S 0 0 4 0 - 1 9 5 1 ( 0 2 ) 0 0 1 2 5 - 7
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2. Geological setting The Baikal rift zone (BRZ) developed along the contact zone between the Siberian craton and the Transbaikal segment of the Sayan– Baikal orogenic belt (Fig. 1). The main geological structures of the region originated in the late Riphean – early Palaeozoic, as a result of subduction – accretion tectonics (Belichenko et al., 1994; Dobretsov, 1986; Dobretsov and Kirdyashkin, 1994; Parfenov et al., 1997). The basement of the Transbaikal belt probably consists of Proterozoic crustal, now largely invaded Palaeozoic to early Mesozoic granites. East of the North Baikal basin, up to the Vitim plateau, these granites form the Baikal – Vitim batholith. The Mongol – Okhotsk belt parallels the Transbaikal belt on its southeastern margin. It mainly has a late Permian to Mesozoic history, linked to the progressive closure of the Mongol – Okhotsk Ocean (Dobretsov, 1986). From the tectonic point of view, the basement of most parts of the BRZ belongs to the Akitkan oro-
genic belt and related Vilyui basin, emplaced in the Palaeoproterozoic (Melnikov et al., 1994). It was reactivated in the middle Palaeozoic by thrusting of the Barguzin terrane over the southern margin of the Siberian craton (Rosen et al., 1994; Masaitis, 1995). The existing models of the lithosphere beneath the Baikal rift zone and adjacent areas are primarily based on geological, gravity and magnetic data (Burov et al., 1994; Logachev et al., 1982; Pismennyi et al., 1984; Zorin, 1971; Zorin et al., 1994, 1998).
3. Data analysis Deep seismic sounding measurements have been carried out in the BRZ and surrounding areas along several profiles (Fig. 2). Refraction profiling was conducted for the regional survey of the crystalline basement, together with the recording of the PmP reflections. In the Yakutian kimberlite province, in addition to DSS profiling, a special seismic survey was carried out (Suvorov et al., 1997, 1999). This
Fig. 1. Geological setting of the southeastern part of the Siberian craton and adjacent areas. Long solid lines—limits of Precambrian terranes (after Rosen et al., 1994) and other major geological structures; dotted lines—Paleozoic sedimentary basins; double line—geological boundary of the Baikal rift zone (after Solonenko, 1968); shaded area—Baikal – Vitim granitic batholith (after Parfenov et al., 1997).
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Fig. 2. Location map of seismic profiles in the Baikal rift zone and adjacent areas of the Siberian craton and Transbaikal fold belt. Solid lines— DSS profiles (after Krylov et al., 1981); short dashed lines—profiles with only PmP reflections; long dashed lines—long-range DSS profiles (GEON Centre, after Egorkin, 1984); shaded area—region of DSS survey in the Yakutian kimberlite province (after Suvorov et al., 1999); solid triangles—seismological stations. The thick line locates the composite profile discussed in the text.
provided the most even distribution of measurements, which were used to estimate the crustal thickness from PmP reflections. A rough estimation of the crustal thickness beneath the northeastern outskirts of the BRZ (west of the Aldan shield) has been derived from local earthquake data (Suvorov and Kornilova, 1985). These data were used for constructing Moho maps and for averaging crustal seismic velocity maps. The crustal velocity distribution was studied, using a new 2-D tomographic technique and integrating also results of previous investigations. Krylov et al. (1981) proposed a three-layer model of the crust for the BRZ. The upper layer is 10 km thick and contains the majority of earthquake foci. The second layer, between 10- and 20-km depth, is characterised by a discontinuous seismic lower velocity zone. The third layer corresponds to the lower crust, beneath 20 km. In the upper 20 –25 km of the crust, velocities have been determined by 2-D seismic tomography using diving waves travel-times (Mishenkina et al., 1983; Krylov et al., 1995b). This method was used for areas with dense observations, which is not always the case
along the DSS profiles in the BRZ (Krylov et al., 1995a; Song et al., 1996). In addition, a new seismic cross-section along another composite transect (Boguchany – Ust-Kut –Nizhneangarsk– Chita) is presented (Figs. 2 and 5). The seismic velocity of the lower crust was derived from the average crustal velocity and from interval velocities in the overlying crustal layers. Calculations were performed taking into account the lateral differences in the upper crustal velocities for down-going and up-going rays of reflected PmP waves. The offsets used were on average about 150 km. For depth down to 20 km, this distance is shortened to about 50 km. So we assume that the calculated values of the interval velocities in the lower part of the crust are attributed to areas with horizontal dimensions of f 50 km. To enable comparison of the velocity variations determined at various depths, we converted the data to a uniform horizontal scale, using averaging cells of 50 50 km separately for the upper crust (0– 20 km deep) and the lower crust (20 km to the Moho depth). The same averaging was used to compare the velocity
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variation in the upper crustal intervals of 0– 10 and 10– 20 km. The averaged data were used to construct a 3-D velocity model of the crust. For some DSS profiles with sparse observations, we obtained the velocities in the upper crust from diving P-waves by the Wiechert– Herglotz method (Mishenkina, 1967; Krylov et al., 1981). These were used to interpolate the velocity anomalies between profile distances of about 200 km. Not enough data are available to examine the depth interval of 10– 20 km. Data characterising the crust are analysed at two different scales along the DSS profiles. As a first step, velocity distributions are compared between the upper and lower parts of the crust. In a more detailed scale, velocity distributions in the upper crust are compared between the 0 –10- and 10 – 20-km intervals. Indeed, calculations on model examples show that the accuracy of velocity estimation in the lower crust is affected by uncertainties in the average crustal velocity. For example, an uncertainty in average velocity of F 0.05 km/s produces an error of up to F0.1 km/s in lower crust with a 20-km thickness. If lower crustal thickness is decreased down to a 15-km and upper crust is increased to a 25-km thickness, this error grows up to about 0.15 km/s. It is assumed that the velocity error is in the order of F 0.1 –0.15 km/s for the lower crust and F 0.05 km/s for the upper crustal intervals of 0– 10 and 10 – 20 km (Mishenkina et al., 1983). This rather high accuracy is reached by smoothing the 2-D velocity cross-section in intervals of 10 km in depth and 50 km in horizontal length. Thus, with the data used, it is possible to give evidence for velocity anomalies with horizontal sizes larger than 50 km and with amplitudes that exceed 0.1 km/s in the upper crust and 0.2 km/s in the lower crust.
4. Thickness and average velocity in the crust All available seismic data were used to compile a crustal thickness map, showing large-scale structures that are identified over more than one profile (Egorkin, 1984; Krylov et al., 1981, 1995b; Song et al., 1996; Suvorov et al., 1997, 1999). For the investigated area, the crustal thickness ranges from 35 to 50 km (Fig. 3). The area with the shallowest Moho
(35 –37 km deep) is observed beneath the Baikal depression and the Vilyui basin and defines an elongated roughly northeast-striking uplift. It possibly reflects a common tectonic processes in both regions (Masaitis, 1995). The area of Moho uplift corresponds to the Proterozoic Akitkan orogenic belt (Borukaev et al., 1988; Rosen et al., 1994), which underwent tectonic reactivation in late Proterozoic and in the Palaeozoic. Also to be noted is that the existence of a pre-Cenozoic seismic basement with velocities of 4.8 –5.1 km/s at the depth of about 6 km beneath Lake Baikal has been shown by Krylov et al. (1993). In the Siberian craton, northwest of the Akitkan orogenic belt, the crustal thickness is 37 – 40 km in the south and 43 – 46 km in the north. A crustal thickness of 46 – 50 km is observed beneath the Tunka rift depression (located to the southwest of Lake Baikal), under the Aldan shield and locally south of Chita in the Mongol – Okhotsk belt. An intermediate crustal thickness of 37– 43 km is determined beneath the Cenozoic depressions in the northeastern part of the BRZ, without a clear correlation between the crustal thickness and the depressions (Krylov et al., 1981, 1995a). As a rule, the location of the BRZ is not clearly reflected by the Moho topography. The crustal thickness beneath the BRZ varies between 35 and 50 km. Abrupt local changes are observed beneath the Baikal basin, from 35 to 43– 44 km (Krylov et al., 1981; Fig. 3, insert). A Moho depression is located in the middle of the lake basin, transverse to it, and divides the crust under Lake Baikal in northern and southern parts, both with elevated Moho. For the southern part of the investigated region, the map of crustal thickness was previously constructed from DSS and gravity data using a regression technique (Zorin et al., 1986, 1989). These data show considerable differences with our DSS data. Estimations by Zorin et al. (1986) are about 4 – 5 km larger than estimations from seismic data alone. These authors obtained a crustal thickness of 45 km in the East Sayan massif and under the Tunka depression, where we show an increased crustal thickness up to 49 km. A similar situation occurs locally in the central part of Lake Baikal where our estimations of the crustal thickness are 5 –7 km greater than those from gravity data. The opposite situation is encountered in
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Fig. 3. Map of crustal thickness. Inset shows details for the Baikal region and PmP reflection profiles (short dashed lines, after Krylov et al., 1981). Dashed lines—limits of geological structures as defined in Fig. 1; sparse dotted line—boundary of low velocity anomalous mantle (7.7 – 7.8 km/s, after Krylov et al., 1981 and Suvorov et al., 1985).
the vicinity of Chita, where our estimations are 5 – 7 km less than the gravimetric ones. The map of average P-wave velocity in the crust is shown in Fig. 4 (data from Egorkin, 1984; Krylov et al., 1981, 1995b; Song et al., 1996; Suvorov et al., 1997, 1999). Within the BRZ and the Transbaikal belt, the crustal velocity varies insignificantly around 6.4 km/s (Krylov et al., 1981). In the Siberian craton, the velocity fluctuates between 6.4 and 6.6 km/s with a mean value of 6.5 km/s (Egorkin, 1984; Suvorov et al., 1999). The P-wave velocity map shows an elongated area with relatively low velocities in the northern extremity of the BRZ, striking in a northwest direction. Low velocity is also detected along the southern margin of the Siberian craton (Fig. 4). The comparison of Figs. 3 and 4 highlights the lack of correlation between crustal thickness and average seismic velocity in the crust. The crustal thickness in the BRZ and in the Transbaikal belt shows locally sharp variations, under a practically constant average
velocity. In the cratonic area, on the other hand, changes of crustal thickness are gradual and accompanied by significant velocity changes, although correlation between velocity and Moho depth variations is not observed.
5. Variations in crustal velocity along DSS profiles The seismic cross-section of the crust along the Boguchany – Ust-Kut – Nizhneangarsk – Chita profile (Fig. 2) starts from the Siberian craton, crosses the northeastern part of the BRZ and terminates in the Transbaikal belt (Fig. 5). In the Siberian craton, the crystalline basement is located at a maximum depth of 5– 7 km under the sedimentary cover and gradually rises towards the craton margin. The normal P-wave velocity at the top of crystalline basement is 6.1 –6.2 km/s. However, a drastic increase of velocity up to 6.8 km/s is revealed in the upper part of the crystalline
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Fig. 4. Map of average P-wave velocity in the crust. Thick solid lines locate transects with nuclear sources; thin solid lines DSS profiles; dotes lines only PmP reflections.
basement in a local zone beneath the Srednaya – Angara zone, rich in iron deposits. A relatively low velocity of the 6.1– 6.3 km/s borders the anomaly
whose base is at 12 –15 km deep. High velocity in the basement is observed in other regions of the Tunguska basin where Triassic basalts have been observed
Fig. 5. P-wave velocity model of the crust and uppermost mantle along a profile crossing the Siberian craton, the Baikal rift zone and the Transbaikal fold belt. Thick lines—seismic boundaries; thin lines—velocity isolines; dashed lines—less reliable velocity isolines and seismic boundaries.
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(Kyznetzov et al., 1987), and consequently could be associated with basaltic intrusions. A low velocity zone (4.8 – 6.0 km/s) appears under the 6 – 8-km deep Verkhne – Angara (Upper-Angara) depression (Fig. 5). At this place, a velocity of 6.2 km/ s is located at a depth of nearly 10 km. Further to the east, the 6.2-km/s velocity zone descends to 22 – 24 km. Another low velocity zone (6.2 km/s) is present at the same depth in the 1100 –1200-km profile segment. The maximum Moho depth (46 km) is located under the central part of the Siberian craton and is associated with a high velocity in the upper part of the crystalline basement, as it is also the case under the Baikal rift zone. The minimum Moho depth (35 km) is located at the margin of the Siberian craton. The subMoho velocity is 8.2– 8.4 km/s under the Siberian craton and 7.7– 7.8 km/s beneath the Baikal rift zone, where lower another seismic boundary is defined at a depth of 50– 58 km (called ‘‘Moho 1’’) (Krylov et al., 1981). This zone of anomalous mantle with a velocity of 7.7– 7.8 km/s (contoured in Fig. 3) extend over a large area than the geological expression of the rift zone (Krylov et al., 1981, 1993; Suvorov and Kornilova, 1985). Within the anomalous domain, Krylov et al. (1981) pointed a few segments of the uppermost mantle with a normal velocity of 8.2 km/s. Fig. 6 relates the velocity distribution between upper crust ( < 20 km) and lower crust. The crust can be divided into two distinct groups. One of them, with comparatively high velocities, characterises the cratonic areas. The other, with lower velocities, is related to the Transbaikal belt. In the latter, there is an inverse correlation between the velocity variations in the upper and lower parts of the crust: relatively high velocity in the lower crust ( f 6.5– 6.7 km/s) and low values for the upper crust (6.1 – 6.3 km/s). This inverse correlation, however, is not everywhere. This is exemplified by the Nizhneangarsk – Ust-Muya profile crossing the BRZ (Fig. 6), where a direct correlation is observed between velocities in the upper and lower crust. In addition, local areas in the Siberian craton at distances 100– 200 and 500 –700 km along the profile (Fig. 5) present unusual low velocity in both upper and lower parts of the crust. The ranges of velocity fluctuation in the upper and lower crustal layers are f 0.2 – 0.3 km/s for the Transbaikal belt, whereas the velocity variations increase up to 0.5 km/s in the upper crustal part of
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Fig. 6. Relation between velocity changes in the upper crust and in the lower crust. Crosses—Transbaikal fold belt; diamonds— Siberian craton; solid triangles—Nizhneangarsk – Ust-Muya profile (location in Fig. 2).
the craton, with a maximum velocity of 6.4– 6.6 km/s (Fig. 5). No correlation between velocity and thickness is observed in the lower crust, probably due to the irregularities of the Moho topography. These features are illustrated in detail in Fig. 7a, for the Boguchany – Ust-Kut –Nizhneangarsk – Chita profile (location in Fig. 5). There is generally an inverse correlation between the velocity in the upper and lower parts of the crust. Under the BRZ, the velocity in the upper crust is 6.0 km/s and in the lower crust, 6.8 –6.9 km/s. In neighbouring areas, the velocity in the lower crust is 6.6– 6.7 km/s. Similar values were obtained along other profiles in this part of the BRZ (Krylov et al., 1995a). Beneath the BRZ, the lowest velocity (5.6 – 5.8 km/s) occurs in the upper crustal interval of 0– 10 km while 6.4 – 6.5 km/s characterises the 10 –20 km interval. This should be compared to the 6.2 km/s at the southeastern margin of the BRZ. An example of velocity variation in the Transbaikal belt is shown along the Ust-Uda– Irkutsk– Chita profile (Fig. 7b), (Krylov et al., 1981; Song et al., 1996). A clear inverse correlation between velocity in upper and lower crust is observed in the Transbaikal belt only. Along the 400 – 600-km segment of the profile, the velocity decreases from 6.3 to 6.1 km/s in the upper crust and increases from 6.5 to 6.7– 6.8 km/s in the lower crust.
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Fig. 7. Velocity changes in the crust along the Boguchany – Ust-Kut – Nizhneangarsk – Chita (a) and the Ust-Uda—Ulan-Ude—Chita (b) profiles.
For the 200– 400-km segment covering the BRZ, the lower crustal velocity increases slightly up to 6.7 km/s and is delimited by local velocity minima on both extremities. Thus, the high velocity in the lower crust is shifted northwest relative to the low velocity in the upper crust from 6.3 – 6.4 to 6.2 km/s. For the same segment, inverse velocity variations of 0.2– 0.3 km/s between the upper crustal intervals of 0 –10 and 10 –20 km beneath the BRZ are observed. The Nizhneangarsk – Ust-Muya profile (Fig. 6) shows that velocities in both the upper and the lower
crust may vary proportionally within both the craton and the fold belt. The highest velocity (6.9 km/s) in the lower crust corresponds to 6.1– 6.2 km/s in the upper crust. In the Transbaikal fold belt as a whole (Fig. 8), a relatively high velocity (6.2 –6.7 km/s) is observed for the 10 – 20-km upper crustal interval while in the upper 10 km, the velocity is 5.7 – 6.2 km/s. The opposite situation is found for the Siberian craton where mainly low velocities (6.0 – 6.4 km/s) are obtained in the 10 – 20-km crustal layer and the
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velocity increases from 5.9 to 6.0 –6.7 km/s in the upper part. The velocities in the intervals 0– 10 and 10 – 20 km are almost similar in the craton, not considering the peculiar high velocity spot (6.4 –6.7 km/s) in the upper crystalline basement (Fig. 5). The relation between velocities in the 0 –10- and 10 –20km upper crustal layers expresses mainly a positive correlation, except locally. This is on the contrary of the Transbaikal belt, where no correlation between the velocity values is present.
6. Lateral variations in crustal velocity
Fig. 8. Relation between velocity changes in the upper crustal invervals of 0 – 10 and 10 – 20 km.
Values of averaged interval velocities in the upper and lower crust along DSS profiles are mapped in Figs. 9 – 11. The sizes of reliably traced anomalies are
Fig. 9. Velocity map of the upper crustal interval 0 – 10 km. Solid lines—locations of DSS profiles; double line—geological boundary of the Baikal rift zone.
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larger than the average distance between the DSS profiles (150 – 300 km). In areas with greater distances, interpolation of the velocity values gives only a rough picture. 6.1. Upper crustal interval of 0– 10 km In the upper crust down to 10 km, the velocity varies laterally from 5.6 to 6.4 km/s (Fig. 9). Minimum values of 5.6 –5.8 km/s are shown in the central part of the BRZ on the eastern side of Lake Baikal. The area of low velocity extends southwest along the lake as a narrow band, and terminates before the western end of Lake Baikal (as no low velocity zone is observed under the Tunka depression). Velocity values of 5.6 – 5.8 km/s are typical for acidic rocks, and their extent coincides with the Baikal – Vitim massif of Paleozoic granites in the central part of the BRZ (Pismennyi et al., 1984; Parfenov et al., 1997).
The area with low velocities of 5.6– 5.8 km/s is surrounded by a zone of 5.8 –6.0 km/s which coincides distinctly with the contours of the BRZ derived from geological and seismological data (Solonenko, 1968), except for a band that extends southwards between Ulan-Ude and Chita. Velocities of 6.0– 6.2 km/s characterise the upper crustal layer of the Siberian craton. Such values are also detected in the Transbaikal belt, southeast and south of the rift zone, though we do not have sufficient data to discuss this in detail. To the northeast, at the vicinity of Olekma River, the BRZ is expressed by an area with velocities of 6.2 – 6.4 km/s revealed from aftershock data of the 1990 SouthYakutian earthquake (Suvorov et al., 1993). 6.2. Upper crustal interval of 10 –20 km The velocity in the upper crustal interval of 10– 20 km varies over the entire territory from 5.8 to 6.4 km/s
Fig. 10. Velocity map of the upper crustal interval 10 – 20 km. Solid lines—location of DSS profiles; double line—geological boundary of the Baikal rift zone. Solid lines are location of seismic profiles.
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Fig. 11. Velocity map of the lower crust. Solid lines—location of DSS profiles; double line—geological boundary of the Baikal rift zone.
(Fig. 10). A relative decrease in velocity down to 5.8– 6.0 km/s is observed under the central part of the rift zone but the size of this area is significantly smaller than that observed in the upper 10 km of the crust (Fig. 9). The area with velocities of 5.8– 6.0 km/s correlates with the location of the Baikal – Vitim granitoid massif. From gravity data, a thickness of this massif is estimated at 20 km (Pismennyi et al., 1984; Alakshin et al., 1990). A narrow band of velocity of 6.0 – 6.2 km/s bounds this area and extends south of the rift zone, between Ulan-Ude and Chita. In the north, the velocity isoline of 6.0 km/s is drawn assuming that velocities lower then 6.0 – 6.1 km/s are absent in the upper crust north of Lake Baikal and outside the BRZ, based on increased velocity obtained along three seismic profiles (located in Fig. 10) in the superficial part of the craton. Velocities of 6.2 –6.4 km/s bound the low velocity zone in the upper crust laterally, northeast of the BRZ (Fig. 10). In the vicinity of Olekma River, the upper crustal velocity is 6.2– 6.4 km/s (seismological data of Suvorov et al., 1993).
6.3. Lower crust In the crust beneath 20 km, the lateral velocity changes are in the range of 6.55 –6.85 km/s (Fig. 11). Velocities of 6.5– 6.7 km/s are observed in the central part of the territory, covering the southeastern margin of the craton and the southwestern part of the rift zone, as well as the Mongol – Okhotsk fold belt and the eastern edge of the rift zone. Rather high velocities of 6.7– 6.9 km/s are observed in an area about 200 km wide, stretching N –S for almost 500 km and crossing the rift zone between Nizhneangarsk and Chita. The northern limit of this area is not clearly determined, but the southern limit is confidently traced outside the rift zone. A large part of the craton is also characterised by such velocity values. The area under the Baikal – Vitim granitic batholith shows high velocities of up to 6.85 km/s. Comparison of Figs. 10 and 11 allows concluding that main velocity anomalies in the crust coincide with the rift zone and are distinctly visible only in the upper 20 km of the crust. As an exception to this rule, the
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roots of the Baikal – Vitim batholith is expressed by a relatively high velocity (6.85 km/s) in the lower crust (Fig. 11) and a relatively low velocity (5.8 – 6.0 km/s) in the 10 – 20-km interval and 5.6 –5.8 km/s in the upper 10-km interval (Fig. 10).
7. Discussion and conclusions The 2-D seismic tomographic analysis of old DSS data (Krylov et al., 1981) shows new aspects of the relation between the deep seismic structure of the BRZ and the adjacent areas. The seismic interpretation is in a good agreement with other geophysical results (Pismennyi et al., 1984; Alakshin et al., 1990; Rosen et al., 1994). The average crustal velocity ranges between 6.4 and to 6.6 km/s in the Siberian craton. The velocity is almost constant at 6.4 km/s in the Transbaikal belt, in areas underline by geological units of Paleozoic, Mesozoic and Cenozoic ages. The accuracy in velocity determination is everywhere in the range of F 0.03– 0.05 km/s. The crustal thickness in both regions varies between 35 and 50 km. However, in the Transbaikal belt, the Moho depth changes sharply in narrow zones, whereas in the Siberian craton, the Moho topography is smoother. Thus, average velocity and crustal thickness are not correlated other. The transition zone between the craton and the fold belt is not clearly expressed in average velocity and crustal thickness, possibly owing to complex interactions of the pre-Cenozoic structures in these regions. In particular, thinning of the crust beneath the Baikal depression and Vilyui basin may indicate a common tectonic origin as proposed for the Palaeozoic period by Masaitis (1995). The crustal thickness derived from DSS data and those derived from gravity data by Zorin et al. (1986, 1989) may differ significantly. According to DSS seismic data, the crustal thickness of the Siberian craton, of the Mongol – Okhotsk belt in the vicinity of the Chita and under the Central Baikal basin is 4 –7 km less than that in the model of Zorin et al. (1986, 1989). Beneath the Tunka depression, the seismic crustal thickness is 4 – 5 km greater. A clear division of the crust into upper and lower parts can be postulated, taking into account
a reduction in resolution of the DSS method with depth. The relation between velocities in the upper and lower crust is different. In the craton and in the Transbaikal belt, a reverse correlation between velocity variations in these intervals was found. For the upper crustal depth intervals of 0 –10 and 10 –20 km, velocity changes present a direct correlation in the craton, while this relation is much more complex in the Transbaikal belt. For the latter, the velocities in the 10 – 20-km depth interval are f 0.2 km/s higher than in the same level in the craton, and in the 0– 10-km interval, they are 0.2– 0.3 km/s lower. Our DSS seismic model of the crust is not in accordance with the tomographic models of Petit et al. (1998) and Kylakov (1999) that used local and teleseismic data. The discrepancy may be due to the lower resolution of the tomographic method used by Petit et al. (1998) and Kylakov (1999). Those models are slightly different, although they were derived from the same data. The model of Kylakov (1999) does not show reduced velocity in the upper crust under Lake Baikal, as in our model and in the model of Petit et al. (1998). These tomographic models also show no low velocity anomalies under other rift depressions, although it is demonstrated from DSS data that low velocity anomalies beneath the large rift depressions can be traced down to depths of 6 – 14 km (see Fig. 5, Krylov et al., 1995b; Song et al., 1996). The existing tomographic models lack velocity anomalies under the large Baikal – Vitim batholith, even though its bottom is estimated at a depth of 15 – 20 km from seismic and gravity data (Pismennyi et al., 1984; Alakshin et al., 1990). Another discrepancy is a too large crustal thickness in the reference model of Kylakov (1999). The crustal thickness obtained from our model is similar to that in the model of Petit et al. (1998) under the Tunka depression (southwest Lake Baikal) and under the central part of Lake Baikal, but in other areas, it differs by up to 5– 13 km. Partitioning the crust into two velocity layers is justified by the fact that the velocity anomalies in the upper crust correlate with the geologically defined rift zone, but not those in the lower crust. Crustal layering was found for both the Archean craton and the Transbaikal fold belt, despite their very different composition and thermotectonic ages. This can be explained
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by their rheological characteristics (e.g., Lobkovsky, 1988; Avouac and Burov, 1996; Westaway, 1998).
Acknowledgements The authors express gratitude to the Novosibirsk Regional Center of Geoinformation Technologies SB RAS for the advices and help in drawing up the electronic maps. We thank an anonymous reviewer for his very useful comments and Prof. A. Khan for their hard work to improve English style of the manuscript. This work was supported by the Russian Foundation of Basic Research (Grant No. 98-05-65239). It is a contribution to the IGCP400 Project Geodynamics of Continental Rifting.
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