Dniepr-Donets Rift: deep structure and evolution from DSS profiling

TECTONOPHYSICS ELSEVIER

Tectonophysics 268 (1996) 83-98

Dniepr-Donets Rift: deep structure and evolution from DSS profiling T. I l c h e n k o Institute of Geophysics National Academy of Sciences of Ukraine, 32 PaUadima av., 252680 Kiev, Ukraine

Received 15 May 1995; accepted 31 May 1996

Abstract

Forward modelling has been applied to refraction and wide-angle reflection traveltime data along two profiles crossing the intracratonic Dniepr-Donets Rift (DDR), a Devonian sedimentary basin located in the southwestern part of the Eastern European Platform between the Ukrainian Shield and the Voronezh Massif. The 2-D P-wave velocity models show that the crystalline crust beneath the rift basin is thinner and characterised by higher velocity values (of almost 1.0 km/s) than beneath the Shield. Differences along the strike of the DDR are also indicated. The two profiles cross the Dniepr and Donets segments of the DDR and reveal, respectively, sedimentary infill thickness of 7 and 20 km and crystalline crustal thicknesses of 31 and 17 km. In both cases, the crystalline crust comprises two distinct velocity layers. Beneath the Dniepr basin, the Moho lies at 38 km and the upper mantle velocity is 8.4 km/s. In the Donets region, the crustal layers are underlain by velocity boundaries at depths of 37 and 42 km with velocities of 7.6 and 8.0 km/s, respectively. The velocity models are interpreted to indicate that a key process in the formation and evolution of the DDR was the exchange of material between the crust and uppermost mantle. It is hypothesised that the extended and faulted pre-rift shield crust was intruded by mantle melts during rifting and that crust/mantle material exchange may also have occurred during the subsequent cooling and subsidence stage of development. Rifting was less intense in the Dniepr segment of the DDR compared to the Donets. To explain other differences in uppermost mantle structure between the two segments of the DDR, it is postulated that density reduction of some portion of the Donets subcrustal mantle occurred in response to compressional deformation and inversion during the Permian Uralian orogeny. Keywords: Dniepr-Donets Rift; deep seismic sounding; refracted wave phase; travel time; velocity modelling

1. Introduction

The intracratonic Dniepr-Donets Rift (DDR) is a Late Devonian aged geological structure located on the southwestern part of the Eastern European Platform, between the Ukrainian Shield and the Voronezh Massif (Fig. 1). Its basement is mainly of Early Proterozoic tectonic age. The rift strikes from northwest to southeast. Along strike to the southeast, the width and thickness of the basin, the intensity of deformation and degree of metamorphism of its sedimentary fill, and the degree of syn-rift volcanic

activity all increase. The DDR can be subdivided into two segments: the non-inverted Dniepr basin and the Donets Basin, partly inverted during the Permian Uralian orogeny (Popov, 1963; Chirvinskaya et al., 1972; Chirvinskaya and Sollogub, 1980). The crustal structure of the DDR has been studied extensively using refraction and wide-angle reflection techniques ('DSS'). In total, thirteen DSS lines traversing the rift every 50-150 km, one longitudinal profile, and several on the flanks of the rift have been acquired (Fig. 1). The DSS observation procedure comprised seis-

0040-1951/96/$15.00 Copyright © 1996 Elsevier Science B.V. All rights reserved. PII S0040- 195 1 (96)00221- I

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T. llchenko / Tectonophysics 268 (1996) 83-98

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mograph spacings of 100 or 200 m, recording frequencies of 7-30 Hz, shotpoint spacings of 15-50 km, and a maximum shot-receiver offset of 200-300 km. Such a high density of seismic trace acquisition allows an essentially unambiguous phase correlation for both first and second arrivals, even over distances of only a few kilometres. This significantly increases the confidence with which a resulting velocity model can be determined since the ambiguity of wave tracing, and misidentification of seismic phases, is the main contributor to non-uniqueness in velocity modelling. Similarly, the adopted frequency range of recording allows a highly resolved resolution of seismic phase arrival times; the traveltimes can be picked with an accuracy of -t-0.005 s with the possibility therefore of resolving seismic boundaries separated by only several kilometres. Shotpoint spacings <50 km provide sufficient resolution for the study of lateral crustal and uppermost mantle

heterogeneities and a maximum offset distance of 300 km allows, in principle, the Moho to be modelled as a refracting boundary. However, in practice along the DDR DSS lines, the maximum offset was often much less, approximately 200-220 km, such that Moho depth has been mainly determined from post-critical (wide-angle) reflections. In this paper, new velocity models of the Earth's crust and uppermost mantle along two DSS lines (A and B; Fig. 1) traversing the northwestern parts of the Dniepr and Donets segments of the rift, respectively, and the adjacent areas of the Ukrainian Shield and Voronezh Massif are presented. These models serve as the basis to frame some geological hypotheses on the origin and evolution of the DDR that account for the differences in crustal structure between the DDR and surrounding cratons and between the Dniepr and Donets segments of the rift.

T. llchenko/ Tectonophysics 268 (1996) 83-98

2. Seismic data

The DSS data along profiles A and B were acquired in the mid-1960s. Line A is 445 km long and Line B is 420 km and the DSS surveys consisted of seven near-offset (observations to 80 km) and fifteen far-offset (to 230-280 km) shotpoints and thirteen near-offset (to 80 km) and twelve far-offset (to 160-220 km) shotpoints, respectively (Fig. 1, faroffset shotpoints are shown). Charges increased in size gradually from several kilograms for recording the wavefield near the shotpoints to 1000-1500 kg to register the uppermost mantle refracted phase Pn and were detonated in holes of 30 m depth. Every explosion was recorded simultaneously by three seismostations, each of which was linked to fortyeight seismographs placed every 100 m. Thereby one explosion covered an interval of about 15 km length. The resonance frequency of registration was 10 Hz. Only the vertical component was recorded. The time scale of the analogue paper seismograms was 0.01 s/mm. Figs. 2 and 3 show copies of typical seismograms reduced approximately five times. The

85

sections in Figs. 4 and 5, assembled from selected traces at 5-km intervals (therefore approximately every fiftieth trace), demonstrate the basic types of phases used for seismic modelling. The principal components of the recorded wavefield are refracted and reflected waves. Refracted phases originating in the crystalline crust are characterised by smooth traveltime curves (TTCs) and their apparent velocities increase gradually with increasing distance from the shotpoint. In contrast, the TTCs of refracted phases originating in sedimentary strata are characterised by numerous comer points and breaks indicating the presence of positive and negative changes in the velocity structure of the transmitring medium. Pn is observed at crossover distances of 170-200 km from the shotpoint, after which crustal refractions are recorded as secondary arrivals that are not usefully correlatable. Consequently, only the upper 15-20 km of the shield crust and about 30 km of the basin crust are constrained by refracted traveltime data. Furthermore, Pn is of very small amplitude (cf. Fig. 4) and has only been recorded on intervals not exceeding 40 km on both profiles.

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T. llchenko /Tectonophysics 268 (1996) 83-98

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Fig. 3. Example 17 km fragmentof one shot record: profileB, shotpoint 10 at 322 km along profilefrom the southwest (cf. Figs. 1 and 8b), seismographspacing of 100 m. Pg is the crystallinecrust refractedphase; PmP is the Mohoreflectedphase. Locationof the reflector correspondingto PmP is shown by a dotted line in Fig. 8b. Reflected phases have been recorded as secondary and later arrivals. Most of these are post-critical although there some are pre-critical. Post-critical reflected phases recorded from the Moho discontinuity (PmP) are generally stable, with a high signal-tonoise ratio (Figs. 2 and 3). These appear on almost all of the shot records, starting at an offset distance of 70-100 km. Typically, there are several separate PmP phases for every shotpoint (e.g., Fig. 6) due to an inferred curvature of the Moho surface. A number of crustal and uppermost mantle reflections are also recorded and are assumed, in most cases, to be not associated with stable first-order seismic boundaries, but rather with irregular thin layers of high velocity contrast. The precision at which the traveltimes of the observed seismic phases can be determined depends upon errors in picking phase arrivals and shot moment signals on seismograms and inaccuracies in determining static corrections (for topographic variations and directional irregularities along the profiles). It is estimated to be 4-0.02-0.04 s on profile A and

-4-0.01-0.02 s on profile B for refracted phases, and -4-0.03-0.05 s on both profiles for reflected phases.

3. Modelling technique Velocity models have been developed by 2-D forward modelling of refracted and reflected wave traveltimes according to the technique of Ilchenko (1985). The principal feature of this technique is independent iterative processing for each pair of reciprocally correlated TTCs. This approach is applicable to observed data with any density of shots and receivers and the residual between the observed and calculated traveltimes can be reduced to any desired value. There is no necessity to assign a priori the correspondence between recorded phases and layers or interfaces in the model. Ambiguous identification of PmP, because of its own low amplitude or its being in the presence of higher amplitude non-PmP wide-angle reflected phases, therefore, will not necessarily lead to the derivation of a false model. In practice, the identification of the recorded phases is one of the results of

88

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Fig. 6. Top: Fragment of traveltime data observed at 0-158 km of profile B is displayed by thick solid and dashed lines for the crustal reciprocal refracted and the Moho reciprocal reflected waves, respectively, with the single Moho reflections shown by thin dashed lines. Pairs of reciprocal refracted "I~Cs are labelled 1-8 indicating the succession of forward modelling steps. Reciprocal points of reflected wave TTCs are labelled with A, B and C. Triangles note shotpoint locations (shotpoints 1, 2, 5, and 6 in Fig. 8b at 0.4, 22.6, 126.7 and 158.0 km along the profile). Bottom: Cells corresponding to the pairs of refracted TTCs are labelled with the same numbers and bordered by the reciprocal rays (solid lines) with the calculated values of velocity only at the turning points. Reciprocal reflection rays are shown by dashed lines. A, B and C are reciprocal reflecting points on the reflectors (thick solid lines with brackets on both sides). Vertical bars (6.6 and 6.7) indicate the limits of possible variation in the velocity function determined from the Moho reflection traveltime data. P-wave velocity values (in km/s) are shown in italics. the seismic modelling process. Further, as the modelling is under the user's constant control, the simplest velocity model, free of complexities not required by the observed data, can be constructed. Refraction TTCs are divided into pairs of reciprocal branches (as short as shotpoint spacing permits) corresponding to non-overlapping areas within the model cross-section (e.g., Fig. 6). Every pair of reciprocal branches is taken to be a separate object for the subsequent iterative processing procedure. First, an initial estimate of the velocity structure for each unit is assigned, adopting one of three elementary ve-

locity model classes determined from the TTC shape (Gurvich, 1960): no velocity contrast, a positive velocity contrast, or a low-velocity layer. The various traveltime pairs are then introduced into the modelling process successively, according to increasing shotpoint-recorder offset, adopting a set of rules designed to provide the simplest result, as follows: (1) a single branch is always transformed to a 1-D velocity function; (2) a 1-D model is tested first for a pair of reciprocal TTCs, and if it proves to be unsuitable, a 2-D velocity distribution is constructed; (3) a gradual velocity-depth change is determined, starting with a

87

T. llchenko / Tectonophysics 268 (1996) 83-98

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T. llchenko /Tectonophysics 268 (1996) 83-98

constant vertical velocity gradient, then with velocity gradient contrasts provided by the user only when necessary; and (4) low-velocity layers are assumed to have constant velocity. Reciprocally correlated and single reflection TTCs are processed in a different way (Puzyrev, 1959). A reciprocal pair is used to determine the position of the reflecting horizon as well as the average velocity of the overlying medium, whereas a single reflection TTC is generally used only for determining the reflector position given independently derived velocity information. An exception to this rule is any case where a reflected phase has an uninterrupted length 1.5-2 times the reflector depth (i.e., 60--80 km for PmP), although these are observed rarely. Unlike refraction TTCs, one pair of reciprocal reflection TTCs constrains neither a 2-D nor a 1-D velocity gradient but only a constant average value. The inclination of the reflector can also be modelled in such a case. In practice, additional data often allow the adoption of a 1-D velocity function with a constant gradient when processing a pair of reciprocal reflection TTCs. Reciprocal reflections are usually less regular their refraction counterparts. The resulting processing stream is not as uniform and usually corresponds to an increase in reflector depth or, for the same depth, an increase in distance along the profile. The practical application of the modelling method is illustrated by Fig. 6. The observed TTCs (top) are divided into pairs of reciprocal branches. For the refracted TTCs, 2-D iterative processing is carried out first for the shortest offset pairs (labelled 1). The derived 2-D velocity distribution is then fixed and the procedure carried out for pairs labelled 2 and so on. The refracted (diving) rays originating from shotpoints (indicated by arrows) outline model units corresponding to the reciprocal branch pairs and labelled accordingly (bottom). For simplicity, the derived velocities are shown only at the ray turning points. In principle, Pn TTCs can also be processed according to this scheme but in the present case they are observed only sporadically and, as such, were used mainly for estimating upper mantle velocity only after the Moho boundary had been already modelled as a reflecting horizon. The velocity model derived from crustal refraction traveltime data shown in Fig. 6 reaches a depth not exceeding 14 km but is smoothly extrapolated

89

downwards on the basis of reciprocal reflection TTC pairs, here, for example, labelled A, B and C. Three 1-D velocity functions (which happen to be the same in this case) are calculated by fitting the traveltime along every reciprocal ray to the corresponding reciprocal time. For convenience, the resulting velocity model is represented by velocity isolines, constructed by linking equivalent velocities in different cells of the cross-section along the profile. The vertical position of a velocity isoline can be considered as the simplest delineation of a lateral inhomogeneity. The modelling procedure is carried out in a stepwise fashion and the result at each step should be the simplest of all those possible. This requirement does not, of course, ensure model uniqueness. The calculation of apparent velocities of reciprocally reflected waves is required, for example, in order to find the velocity function and the position of the reflector (Ilchenko, 1990). The adoption of extreme minimum and maximum values of apparent velocity, instead of an average, leads to the calculation of two limiting velocity functions and positions of the reflector (Fig. 7). The difference between extreme values, and hence the model uncertainty at a given step, depends upon the precision of observed traveltimes and the distance along the observation and calculation baselines. For example, the vertical bars labelled 6.6 and 6.7 in Fig. 6 indicate the depth interval where the velocity may be equal to the same value given the extreme limits of the observed apparent velocity (hence the range of possible positions of the appropriate velocity isoline). The upper and lower positions of reflectors in Fig. 7 indicate the possible depth range accordingly (maximum depth corresponding to highest possible velocity and vice versa). The velocities and depths adopted at each step and used in the subsequent stages of modelling are actually averages of extreme values allowable given the uncertainties inherent to the observed data. Nevertheless, errors associated with estimating the velocity--depth function (or constant velocity, depending on the model 'class' of the overlying medium) and the reflector depth and incline are calculated for all reciprocal pairs of reflected phases. In the case of refracted phases, the very small difference between acceptable extremes under ordinary conditions make

T. llchenko/Tectonophysics 268 (1996) 83-98

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this mostly unnecessary. Velocity and depth limits are determined only for the worst cases, those combining the largest uncertainty of traveltimes with the shortest observation and calculation baselines, although possible ranges of velocity isoline vertical position and in velocities of the main refracting boundaries are estimated in most cases. 4. Velocity models of the DDR crust and upper mantle The seismic data distribution and an indication of the reliability of the lower crust and uppermost mantle velocity models derived along profiles A and B crossing the DDR can be seen in Fig. 8 and simplified versions of the actual velocity models are presented in Fig. 9. The depth limit of ray paths of intracrustal refracted phases used for velocity modelling is marked by a wavy line in Fig. 8a

and b. Below this depth, velocities have been determined from the traveltimes of seismic phases reflected from a laterally (mostly) continuous, highly reflective horizon observed on both profiles inferred to be the Moho. Velocities above this horizon were found to be typical of crustal rocks whereas an upper mantle velocity for the underlying material was estimated from considerations of ray incidence up to the critical angle for various shot records. (A more ambiguous situation beneath the Donets basin on profile B is discussed further below.) The uppermost mantle velocity was also constrained in places by an observable Pn phase and by sub-Moho reciprocal reflection traveltime data. In consideration of the precision of the observed traveltimes and the modelling methodology described above, the error associated with velocities in those parts of the models constrained by refraction data does not exceed 4-0.05 km/s (profile A) or

91

T. llchenko /Tectonophysics 268 (1996) 83-98

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4-0.02 km/s (profile B). The least accurate determination of velocity contrasts at refracting boundaries is 4-0.1 km/s on profile A and 4-0.05 km/s on profile B and the vertical uncertainty of any velocity isoline cannot be larger than 4-0.25 km (profile A) or +0.1 km (profile B). The horizontal uncertainty is estimated to be 4-2-4 km. The m a x i m u m error on crustal layer thicknesses is 4-0.5 and -t-0.2 km on profiles A and B, respectively. The m a x i m u m error on 'block' width (delimited by vertical segments of velocity isolines) is ± 4 - 8 km. In those parts of the velocity models constrained primarily from reflection data, the estimated errors vary depending on position along the profile. Outside of the area of the D D R (beneath the Ukrainian Shield and Voronezh Massif), possible errors in velocity

isoline depth and in velocity are 4-2.0 k m and 4-0.05 km/s to the base of the crust with the cumulative error to the depth of the Moho reflector estimated to be 4-0.25-2.0 km. Beneath the DDR, the respective estimates are 4-1.0 km, 4-0.02 km/s and -t-0.25-1.0 km. The former are larger because the velocity structure of only the upper 15-20 k m of the shield crust is constrained by refracted phases with the rest exclusively from reflections, whereas velocities in most of the crust underlying the basin is obtained from refractions (Fig. 8). For model segments characterised by constant velocities (no gradient), the error estimates are slightly different: ±0.1 km/s in layer 3 (Fig. 9b) and +0.15 km/s in the layers between the inferred Moho and sub-Moho reflectors at approximately 105 and 240 k m on profile B (Fig. 8b).

T. llchenko /Tectonophysics 268 (1996) 83-98

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Fig. 9. Simplified P-wave velocity models of the DDR crust and uppermost mantle along DSS profiles (a) A and (b) B. Thick solid lines are refracting horizons; thin solid lines are velocity isolines; units are km/s. Layers interpreted to represent (Devonian and younger) sedimentary cover are filled with dots. The velocity distribution within sedimentary cover along profile A is from Kozlenko (1975, 1994). The horizon underlying sedimentary cover corresponds to the top of Early Proterozoic basement except where marked by circles where it is hypothesised to be post Early Proterozoic (Riphean). In this case, the horizon marked with dots is interpreted to be the top of the Early Proterozoic. Deeper horizons are based primarily on wide-angle reflected phases (cf. Fig. 8). The line with a vertical ornament on profile B is interpreted to be the top of lightened mantle. The line with oblique hatching is the top of the normal mantle. Crustal layers discussed in the test are labelled 1, 2 and 3. Shotpoint locations are shown by triangles, with the larger ones corresponding to those plotted in Fig. 1.

4.1. Profile A The lower crust and uppermost mantle velocity and location of the reflecting horizon interpreted as the Moho discontinuity have been modelled with good reliability beneath the Dniepr segment of the DDR along profile A owing to three pairs of reciprocal reflectors 6-9, 7-10, and 8-11 (referring to shotpoints with reciprocal coverage illustrated in Fig. 8a) as well as Pn from shotpoints 3, 4, 5, 12 and 14. Refraction data have also constrained the Moho position in the distance ranges 148-165 and 240-245 km. In contrast, the Ukrainian Shield lower crustal velocity model is weakly controlled by partial superposition of single reflectors 1 and 2. Pn

from shotpoints 1 and 2 has provided data for upper mantle velocity and for Moho position in the range 100-120 km. The lower crustal velocity model of the Voronezh Massif is unconstrained. Single reflectors 13, 14 and 15 are assumed to have been generated at the Moho but do not overlap. Thus, the crustal velocity model at 20-30 km depth in this part of the model was merely extrapolated downwards with the same vertical gradient as derived above (0.02/s). The velocity for this part of the upper mantle was obtained from Pn from shotpoint 14. Profile A (Fig. 9a) transects the northwestern part of the Dniepr segment of the DDR, which is in this area 90 km wide, up to 7 km deep, and rather symmetric, with marginal faults of amplitude 4 km

T. llchenko /Tectonophysics 268 (1996) 83-98

93

(at 163 km) and 2.5 km (at 254 km). The sedimentary cover is characterised by a velocity increase from 2 to 5 km/s (Kozlenko, 1975, 1994) suggesting that high-grade metamorphic sediments are absent. The seismic boundary having velocity of 5.9 km/s is considered to be the top of the crystalline basement with the crust underlying the DDR divisible into two layers. The shallowest, layer 1 has a thickness of 3 km and velocity of 5.9-6.1 km/s. Layer 2, below this, comprises a zone of relative velocity increase compared to the flanking zones, marked by vertical steps in velocity isolines at 180 and 250 km. It is 28 km thick and 40 km wide. The top of layer 2 is the 6.4 km/s interface (indicated with solid dots) with velocities increasing with depth, rapidly at first and then with a smaller gradient, to an almost mantle value of 7.7 km/s. The existence of such a smooth velocity gradient suggests that layer 2 can be regarded as a single unit of common geological origin with the characteristic velocities (6.4-7.4 km/s in the upper half) suggesting an intracrustal origin. The crust-mantle boundary (line with oblique hatching), beneath the Dniepr segment of the DDR on profile A, lies at 38 km, shallower than beneath the Ukrainian Shield (where it is at 40 km) and the Voronezh Massif (44 km). The depth to Moho increases to 50 km beneath the flanks of the basin. The upper mantle velocity beneath the rift axis is 8.4 km/s, a value that is high compared to normal platform regions. In adjacent regions on profile A it is 8.0 km/s.

and two pairs of reciprocal TTCs from reflecting horizons within the mantle, 7-9 and 3-5 beneath the rift axis and the Ukrainian Shield, respectively. Profile B (Fig. 9b) transects the northwestern part of the Donets segment of the DDR. Its width of 170 km and thickness of 20 km are roughly two and three times as much, respectively, than for the Dniepr segment on profile A. High-grade metamorphic sediments with velocity of 5.4-5.7 km/s predominate and the basin geometry is characteristically asymmetric with the northeastern part of the basin displaying the higher velocities and a greater thickness. The sedimentary basin on profile B is bounded with the Ukrainian Shield by a marginal fault of 2 km amplitude (at 150 km) but with no basement displacement on its margin with the Voronezh Massif. The crystalline basement is represented in the model by the boundary with velocity 6.1 km/s (Ilchenko, 1991). The crystalline crust beneath the DDR on profile B consists of adjacent two units with different velocity structures (Ilchenko, 1994). The southwestern unit is characterised by a smooth vertical change in velocity from 6.1 to 6.8 km/s whereas the northeastern one is more complex, resembling the central part of the Dniepr segment crystalline crust on profile A. The velocity isolines on profile B define a 85 km wide zone of increased velocity bounded at its top by a horizon with velocity 6.6--6.8 km/s (marked by the contour with solid dots). Layer 1, overlying this horizon, is 5 km thick with a velocity increasing from 6.1 to 6.4 km/s. Below, layer 2 has a thickness of 12 km and displays velocities in the range 6.8-7.2

4.2. Profile B

km/s.

The lower crustal velocity model and location of the reflecting horizon identified as the Moho on profile B are well constrained given the uniform and dense distribution of reciprocal reflection data. Beneath the Donets segment of the DDR (in the distance range 190-280 km) there is a complex situation involving two reflecting horizons both of which could be taken for the crust-mantle boundary. Reciprocal reflection pairs 6-8, 6-9 and 7-10 permit the calculation of the velocity between these horizons. Two sources of information constrain the velocity on the lower horizon and on its lateral continuations - - refracted phase traveltimes from shotpoint 8 (recorded in both directions) beneath the rift flanks

A significant distinguishing feature of the deep structure beneath the Donets segment compared to the Dniepr segment of the DDR, in addition to the asymmetry described above, is a lens-shaped layer 3. It has a velocity of 7.6 km/s, intermediate between that of crust and upper mantle. This, in combination with its structural disposition, isolated in terms of discontinuous velocity gradient from layer 2 (unlike the higher velocity parts of layer 2 on profile A), suggests crustal as well as mantle affinities. The top of layer 3 (line with vertical hatching) shallows to a minimum depth of 37 km beneath the DDR. The relief of the 8.0 km/s horizon (marked with oblique hatching), nominally the Moho (but discussed further below), is complex. It is convex downwards,

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T. llchenko /Tectonophysics 268 (1996) 83-98

from 34-36 to 42 km beneath the DDR and is slightly deeper beneath the basin flanks (Ukrainian Shield and Voronezh Massif, 40 and 44 km, respectively) decreasing in depth (to 37 km) to the southwest.

5. Geological implications of the velocity models 5.1. Crystalline crust of the DDR Velocity versus depth curves for the axial parts of the Dniepr (profile A) and Donets (profile B) segments of the DDR are displayed in Fig. 10. Fig. 10a shows the entire sedimentary basin and crustal section and Fig. 10b isolates the crystalline crustal sections, comparing these to velocity-depth curves for the adjacent shield areas. The crystalline crust beneath the DDR differs greatly from the shield crust, being thinner and, in general, characterised by a more complex velocity structure. It displays a greater velocity range than the shield, significant first-order seismic boundaries with velocity contrasts of up to 0.4 km/s, and, on average, has a velocity about 1 km/s higher than the shield at all depths. The observed difference in crustal velocity structure beneath the DDR compared to unrifted shield crust clearly must have implications regarding processes affecting the crust during rift formation. It has been proposed that the velocity differences may be due to the intrusion of fractures in the crust by mantle melts during the initial rifting phase (Milanovskiy, 1983; Chekunov et al., 1992; Chekunov, 1995). The observed velocity anomaly of +1 km/s beneath the rift, given the mainly acid and intermediate composition of adjacent crust (as corroborated by the observed velocities), would imply the presence of up to 60% mantle material (near the base of the crust) in this case. The actual composition of the mantle component, being either basic rocks partly or completely transformed to granulite and/or eclogite (Ringwood and Green, 1966; Sobolev and Sobolev, 1975; Sobolev, 1978) or a combination of basic and ultrabasic rocks (Buturlinov, 1974; Ramberg, 1976), cannot be constrained. Drilling and other geological investigations (Chirvinskaya et al., 1968; Gavrish et al., 1976; Chirvinskaya and Sollogub, 1980) indicate an Early Proterozoic age of the crystalline basement on the rift

flanks (cf. Shchipansky and Bogdanova, 1996). The shallowest velocity layer of the crystalline basement underlying the rift (labelled layer 1, with upper and lower bounds marked by open and filled dots, respectively, in Figs. 9 and 10) displays similar velocities and can be presumed to be of similar composition. Layer 1 is not characterised by significantly anomalous velocity compared to unrifted crust (Fig. 10b). It therefore overlies that part of the crust presumed to have been intruded with mantle melts during the rifting process. The thickness of layer 1, in view of data and modelling uncertainties discussed above, is in the range 2.5-3.5 km on profile A and 4.8-5.2 km on profile B, less, therefore, in the Dniepr than in the Donets segment of the DDR. Thus, there is an implication that mantle melts may have risen higher (base of layer 1) and, therefore, that the mantle intrusion process was more intense beneath the former compared to the latter. However, this is at odds with numerous geological and geophysical observations indicating that the thickness, deformation, and metamorphism of the Devonian and younger sedimentary fill of the DDR, as well as the intensity of the rifting process, all increase from northwest to southeast (Popov, 1963; Chirvinskaya et al., 1972; Chirvinskaya and Sollogub, 1980; Stovba et al., 1996; Wilson and Lyashkevich, 1996). Alternatively, the base of layer 1 has some other geological significance that predates the Late Devonian rifting. The velocities modelled in layer 1, 5.9-6.1 and 6.1-6.4 km/s for the two profiles, are consistent with those of metamorphosed and crystallised sediments. Based on this, and the occurrence in flanking regions of presumed Riphean aged metasediments, it has been proposed that layer 1 represents a Late Proterozoic (Riphean) rifting sequence (Levenshtain et al., 1971; Sollogub et al., 1975, 1977; Ermakov et al., 1988; Chekunov et al., 1992) that may have been a pre-cursor to late Palaeozoic rifting. The presence or absence of such a sequence, at the depths shown in the present models, cannot be determined from the available CDP data (e.g., Chirvinskaya et al., 1972; Kivshik et al., 1993; Stovba et al., 1996). If present, however, it would mean that the geometry of layer 2, and of the high-velocity zone as a whole in the DSS profiles, reflect the effects of Late Proterozoic as well as Late Devonian rifting processes.

T. llchenko /Tectonophysics 268 (1996) 83-98

a)

95

b) DNIEPR

Vp ~ 8

.7

basement

,6

5

5.~ 4

4

DONETS ,3

z o ~,, .3 4, I ....

s

. . . . .

~

DNIEPR

,6 5.2

?

e_v p

vp .8

7

intrudedtop 6 . 4 ~ 6.8

5.7

DONETS

basement 585 .6

7

.8 Vp

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D

--

8.4i11,Mi/m.i/,q~t/m 7.4

=

2 --

-

i

,

76

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2

_-

-

•

~.,

top

4///, "S","f/, ",#Sf, "/////. "//.~."l. "1//~.

....

,o '.2!221

................

depth (km)

.0

M-:-:_

8.0~/../7

30'

~-

40"

~

8.0

depth (km)

Fig. 10. Velocity--depth curves for (a) the whole crust in the DDR axial zone and (b) the crystalline crust only in the DDR axial zone [the zero-depth corresponding to the maximum crystalline basement depth in (a), 7 km in the Dniepr segment and 20 km in the Donets] compared to Ukrainian Shield and Voronezh Massif crust near the DDR unaffected by rifting processes (shaded envelopes). Other symbols and labels are the same as in Fig. 9.

5.2. Moho and uppermost mantle of the DDR

The transition from crust to mantle beneath the Dniepr basin is simple (Fig. 9a) but, in the Donets region, it is complicated by a lens-shaped layer 3 (Fig. 9b). Given its 7.6 km/s velocity, layer 3 can be interpreted either as a crust-mantle mixture (similar to the crustal high-velocity zone), developed during rifting, or as abnormally low density mantle. The concave shape of the base of the layer and considerations of surface geology suggest the latter. Layer 3 is not seen on DSS profiles, such as profile A here, that cross the Dniepr segment of the DDR (Demidenko et al., 1963; Chirvinskaya et al., 1968). In contrast, a similar geometry (with layer 3 or analogue) is shown by DSS profiles, such as profile B, in the Donets segment (Sollogub et al., 1977; cf. Stephenson et al., 1993). In particular, the thickest part of layer 3 on profile B (220-260 km, Fig. 9b) coincides with that part of the basin that has been folded during inversion subsequent to Late Devonian rifting. Thus, there is an implication that the present crustal and upper mantle structure is related to the post-rift evolution of the DDR, specifically the presumed Permian uplift and folding event, creating the

Donets foldbelt but hardly affecting the Dniepr basin (Popov, 1963; Chirvinskaya et al., 1972; Raznitsin, 1973). In such a case, a transformation would have taken place in the Permian that led to the reduction of uppermost mantle densities in a zone mirroring (in its lens-shape) the increasing densities of sedimentary rocks and uplift of topography at the Earth's surface, so as to maintain isostatic equilibrium. The pre-Permian Moho would have been approximately at the top of the present high-velocity layer 3 (line with vertical hatching in Fig. 9b) with layer 3 itself and the 8.0 km/s horizon marking its base (line with oblique hatching in Fig. 9b) developing thereafter. The actual process by which this would have occurred is not clear but it is speculated to have involved the downward displacement of heavier mantle components and it may be significant that, according to thermal modelling (Kutas, 1989; Kutas and Tsvyashchenko, 1993), the uppermost mantle was the shallowest level of plasticity within the Permian DDR lithosphere. Additionally, the upper mantle velocity beneath layer 3 on profile B is 8.0 krn/s compared to 8.4 km/s on profile A. The pre-rift upper mantle velocity

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T. llchenko /Tectonophysics 268 (1996) 83-98

was presumably about 8.0 km/s, as observed in the cratons near the DDR (Fig. 9). Thus, profile A is characterised by an anomalously high upper mantle velocity (e.g., Christiensen and Mooney, 1995), in common with numerous other rift basins worldwide (e.g., Mooney et al., 1983). However, rifting has been more intense on profile B than A. It is therefore suggested that the difference in upper mantle velocity is not a result of rifting processes directly but rather must be a consequence of events occurring during the post-rift evolution of the DDR. In particular, the Donets segment of the DDR, is characterised by post-rift inversion, as discussed above, and by a very high degree of post-rift (mainly Carboniferous) subsidence, significantly more than can be explained by conventional post-rift thermal subsidence models of basin evolution (e.g., Van Wees et al., 1996). It seems likely that the mechanisms responsible for the excessive post-rift subsidence may also have played a role in effecting the observed upper mantle velocity structure of this part of the DDR, for example, by introducing lower density crustal material into the upper mantle during the post-rift subsidence phase.

6. Summary DSS data acquired in Ukraine, in particular over the DDR, are characterised by a high-frequency range of recording and a dense spacing of seismographs, providing the necessary conditions for construction of reliable velocity models: unambiguous phase correlation, highly accurate traveltime determination, and good resolution of seismograms. These kinds of data have made it possible to reveal layers of several kilometres thickness that are important for understanding the geological processes involved in the structural evolution of the crust and upper mantle. At the same time, the deduced seismic models are as simple as possible, free of elements or features not required by the observations. Velocity models of the crust and upper mantle, constructed using similar data and methodologies, of two distinct segments of the DDR - - the Dniepr and Donets - - reveal significant first-order differences that presumably reflect different tectonic processes active during the rifting and post-rifting phases of development. Model units interpreted to represent the Devonian and younger sedimentary succession are 7

and 20 km thick, respectively. Underlying crystalline crust, 31 and 17 km thick, respectively, comprises in both cases two distinct velocity layers, with the lowermost of these displaying significantly higher velocities than crust at equivalent depth in the adjacent shield areas. Beneath the Dniepr segment, the Moho lies at 38 km and the upper mantle velocity is 8.4 km/s. In the Donets region, the lower crustal layer is underlain by boundaries at depths of 37 and 42 km with velocities of 7.6 and 8.0 km/s, respectively. It is concluded that the present crystalline crust of the DDR axial zone consists of two complexes of different ages and geological history. The uppermost crustal layer is interpreted to correspond to greatly metamorphosed sedimentary rocks younger than Early Proterozoic but older than Late Devonian age. The second crustal layer is interpreted to represent the original Early Proterozoic crust intruded and densified by mantle material during rifting. Beneath the Dniepr basin, this layer is 28 km compared to 12 km in the Donets segment. The degree of mantle intrusion correlates with the degree of subsequent cooling and subsidence and therefore with the difference in interpreted sedimentary thicknesses. It is further concluded that the uppermost mantle structure beneath the Donets segment has also been influenced by the Permian (Uralian) orogeny by compacting and shortening the sedimentary fill and in so doing violating the previously established isostatic balance. It is proposed that this led to a process reducing the density of the uppermost mantle (perhaps, due to displacement of heavier components downwards), forming another mantle boundary at greater depth, and explaining the observed lower crust/upper mantle structure with two mantle major velocity discontinuities - - between the crust and the lightened mantle at 37 km and between the latter and the normal mantle at 42 km.

Acknowledgements The author is grateful to Randell Stephenson whose huge help made it possible to prepare this paper. The comments of the anonymous reviewers are much appreciated. This work was carried out in the framework of EUROPROBE, a programme of the European Science Foundation, and has been financially supported by INTAS Project 93-3346.

T. llchenko /Tectonophysics 268 (1996) 83-98

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Rift and Polish Trough basins. Sediment. Geol., 86: 159-175. Stovba, S.M., Stephenson, R.A. and Kivshik, M., 1996. Structural features and evolution of the Dniepr-Donets Basin, Ukraine, from regional seismic reflection profiles. In: R.A. Stephenson, M. Wilson, H. de Boorder and V.I. Starostenko (Editors), EUROPROBE: Intraplate Tectonics and Basin Dynamics of the Eastern European Platform. Tectonophysics, 268:127-147 (this volume). Van Wees, J.-D., Stephenson, R.A., Stovba, S.M. and Shimanovskyi, S., 1996. Tectonic variation in the Dniepr-Donets basin from automatedmodelling of back-stripped subsidence

curves. In: R.A. Stephenson, M. Wilson, H. de Boorder and V.I. Starostenko (Editors), EUROPROBE: Intraplate Tectonics and Basin Dynamics of the Eastern European Platform. Tectonophysics, 268:257-280 (this volume). Wilson, M. and Lyashkevich, Z.M., 1996. Magmatism and the geodynamics of rifting in the Pripyat-Dnieper-Donets rift, East European Platform. In: R.A. Stephenson, M. Wilson, H. de Boorder and V.I. Starostenko (Editors), EUROPROBE: Intraplate Tectonics and Basin Dynamics of the Eastern European Platform. Tectonophysics, 268:65-81 (this volume).