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Deep structure of the East European platform according to seismic data
Physics of the Earth and PlanetaryInteriors, 25 (1981) 21-31 Elsevier Scientific Publishing Company, Amsterdam - Printed in The Netherlands
21
DEEP STRUCTURE OF THE EAST EUROPEAN PLATFORM ACCORDING TO SEISMIC DATA *
LP. VINNIK ’ and V.Z. RYABOY * 1 Institute of Physics of the Earth, Moscow (U.S.S.R.) * Institute of Geophysical Exploration, Moscow (U.S.S.R.)
(Received January 12,1979; accepted for publication June 2,1979)
Vinnik, L.P. and Ryaboy, V.Z., 1981. Deep structure of the East European platform according to seismic data. Phys. Earth Planet. Inter., 25: 27-37.’ The paper presents a review and analysis of new seismic data related to the structure of the mantle beneath the East European platform. Analysis of observations of long-range profiles revealed pronounced differences in the structure of the lower lithosphere beneath the Russian plate and the North Caspian coastal depression. The highest P-velocities found at depths around 100 km are in the range 8.4-8.5 km s Deep structure of the Baltic shield is different from the structures of both these regions. No evidence of azimuthal anisotropy In the upper mantle was found. A distribution of P-velocity in the upper mantle and in the transition zone consistent with accurate travel-time data was determined. The model Involves several zones of small and large positive velocity gradients In the upper mantle, rapid increases of velocity near 400 and 640 km depths and an almost constant positive velocity gradient between the 400 and 640 km discontinulties. The depth of the 640 km discontinuity was determined from observations of waves converted from P to SV in the mantle.
-r.
1. Introduction For better understanding of the structure and composition of the Earth’s interior, detailed velocity cross-sections of the mantle have to be determined in many regions. Such cross-sections could be obtained from abundant and accurate seismic data, which unfortunately are available in very few regions. The Precambrian East European platform occupies the area to the west of the Urals and includes the Russian plate and Baltic shield as the main structural units (Fig. 1). Until very recently this vast area was almost devoid of goodquality seismic data for the upper mantle. However, in recent years many new data have been obtained in this region.
* Presented at the symposium: ‘Velocity, Attenuation and other Geophysical Properties of the Upper Mantle”, held in Leningrad, U.S.S.R., October 2-l 1,1978. 0031-9201/81/0000-OOOO/$
Among new seismic studies, long-range refraction profiling has been undertaken in the U.S.S.R. (Burmakov et al.,1975;Egorkinet al., 1977).The analysis of later arrivals of P-waves propagating through the mantle of the East European platform and recorded by the NORSAR array (Ring and Calcagnile, 1976) is a valuable contribution. Absolute travel times assumed by Ring and Calcagnile were not accurate enough and later Vinnik et al. (1978) determined travel times more accurately. Another recent contribution to the studies of the deep interior is the detection of waves converted from P to SV in the mantle (Vinnik, 1977). Our paper presents a review and analysis of new seismic data related to the structure of the mantle beneath the East European platform. Section 2 contains a new analysis of observations of the long-range profiles. In Section 3 we describe P travel-time data in the distance range -1000-3500 km and an associated velocity model. Observations of converted waves are
02.50 0 1981 Elsevier Scientific Publishing Company.
28
The Kineshma—Vorkuta profile crosses the north.
~:
‘_~~
•-$~
____________________
Russian pLat~~k~j0~
_________
~
________________
____________ —
______
_________
______
400 km long. reversed profiles These of the profiles same were length. coupled Figures with 2 and three 3
=
_________ _________
______________
—~---
2
~llE~4 ____ ,,.‘ -.
_______________________________
Pechora depression (east of Timan). According to eastern margin of the East European and of DSS about data 40 thekm Mohoroviëi~discontinuity everywhere along theplatform profile. is at aThe depth thickness of the low-velocity sedimentary layer affecting average crustal velocity is about 3—A km in the region of the Russian plate while increases up to 10—12 km in the Pechora depression. The observations in the main profile 1400 km long were complemented by overlapping profiles 800 and
•ai
o
________________
Fig. 1. Map of the region and tentative scheme of P-velocity anomalies in the upper mantle: 1,seismograph stations; 2, long-range profiles (I, Kineshma—Vorkuta, II, Elista— Buzuluk); 3, normal velocity; 4, anomalous velocity at depths less than 200 km ; 5, anomalous velocity at depths less than 1). 100 km;6, residual velocity (km s
analysed in Section 4. In Section 5 lateral heterogeneities in the upper mantle are described. In the same section we touch on the question of seismic anisotropy. The principal fmdings are summarized in Section 6.
2. Fine structure of the uppermost mantle Detailed observations in profiles more than 700 km long were made in the U.S.S.R. in two regions of the East European platform (Kineshma—Vorkuta and Elista—Buzuluk profiles) (locations of profiles are shown in Fig. 1). Field observations were made with calibrated equipment having frequency bands from 1—2 to 10—15 Hz. Velocity sections of the crust were obtained along the same proffles (Egorkin et al., 1977).
show record sections of the main profile and of the overlapping 800 km profile. The end-points of both profiles coincide. In the record section of the main profile, two refraction lines are visible. The first line with an apparent velocity of 8.55 km s~’is observed in a distance range from 200 to 600 km. The second line emerges in later arrivals at a distance of about 700 km and becomes the first arrival at a distance of about 1100 km. Both lines are almost parallel but the second line is late relative to the first by about 3 S. Offsets of this kind are characteristic of low-velocity layers. First arrivalsin a distance range between 600 and 1100 km have an apparent velocity of about 8.0—8.1 km s~.This velocity is too low to be appropriate for any refraction line and the only reasonable explanation for these arrivals is scattering of primary waves by small-scale lateral heterogeneities. The same features are evident in the record section of the overlapping profile (Fig. 3). The only difference is a somewhat lower apparent velocity of the first refraction line (835 km s’’). To account for lateral velocity heterogeneities in the crust, allowances up to 1.0 S were applied to the observed travel times. A velocity model of the uppermost mantle consistent with the corrected traveltime data was derived by trial and error. A remarkable feature of this model (Fig. 4) is the well pronounced low-velocity layer at depths of 60—110 km. Of course, this model is only a representative of a family of models consistent with observational data. However, the existence of a low-velocity layer with properties qualitatively similar to those of our model is demanded by the data. The second area where data of long-range pro-
29 21
ti
t
20
-
0
______
________
iOO
___-~._
—
300
—
400
1
L
900
—__~
—
_J
- —
______
1(5)
600
_____
—
—
____________
700
Boo ~ (en)
4A~It
~
i~oo
IiOO
1000
ii ~ 1300
ILIOO
Fig. 2. Record section of Kjneshma—Vorkuta profile from southwest to northeast; amplitudes are normalized.
200
300
400
500
600
Fig. 3. As in Fig. 2, for the overlapping profile.
700
800
30
/
II
0
______
/ ~
/
filing are available is the North Caspian coastal depression (Elista—Buzuluk profile). The crust in the
ve1oc~ty Km/sec) 8,,, 8~S
central part of the depression is almost devoid of a “granitic” layer and the sedimentary layer is up to
50
~N. N,,~
~-
/ /
20km thick. Record sections of the direct and reversed profiles 800 km long are shown in Figs. 5 and 6. In both sections two refraction lines are visible. Arrivals corresponding to the line with an apparent velocity of 8.2 km s’ are observed in the distance range 200—500 km. Arrivals corresponding to the second line are observed as first arrivals at distances over 600 km. The apparent velocity of the second line is 8.85 km s’~ in the direct profile and 8.55 km
50
500
1Q00
~ (Km~
Fig. 4. P-velocity model and corresponding reduced traveltime—distance curve for Kineshma—Vorkuta profile. The velocity in the crust is 6.2 kin 5i~
~
k
I~
L
t -~
~
-~
~~-‘:.~
I
25
800
700
800
500
400
500
200
~ (om) Fig. 5. Record section of Elista—Buzuluk profile from southwest to northeast; amplitudes are normalized.
: ~20
:10
0
J1I~~~ I I411 ~I ~WtJftJ!JijL ‘
_____________H 200
I
-
300
,,
400
500
~ (on) Fig. 6. As in Fig. 5, for the reversed profile.
600
700
31 Vetotitilom/sec)
______________
0
I
~ T ~
and the associated velocity model, which is presented in Table I.
___0
/
The relatively low velocity gradient an the uppermost layer of the mantle is terminated by the “dis-
~
/
~0 I
N. 500
-~
N
100 000
~ (~cm)
Fig. 7. P-velocity model and the corresponding reduced traveltime—distance curve for Elista—Buzuluk profile. The velocity in the crust is 5.8 km ~
s in the reversed profile. The value of 8.85 km s~ is strongly in error due to variable thickness of sediments in the northeastern end of the profile. If allowances for the crustal heterogeneities are applied to travel times, the apparent velocity of the second line in the direct profile becomes 8.55 instead of 8.85 km s~.Figure 7 shows a velocity model consistent with the corrected travel-time data. A thin low-velocity layer explains the very low amplitudes of the first arrivals in the distance range 500—600 km. Thus, for the two structural units of the East European platform markedly different velocity models of the uppermost mantle are obtained. These differences are associated with evident differences in the wave fields,
continuity” at depths around 100 km. This “discontinuity” corresponds to a distinct increase in apparent velocity from 83 km s’~at distances of less than 800 km to 8.5 km s’ at distances over 1000 km. In a distance range 200—1000 km many of the traveltime determmations shown an Fig. 8 were taken from the Kineshma—Vorkuta profile (Section 2) and thus are representative mainly of the Russian plate. This was decided in view of the dominating linear dimension of the Russian plate among other geological structures in the region. The upper part of the model is therefore a simplified and smoothed version of the more complicated model shown in Fig. 4. This more detailed model could be adopted as the upper part of the velocity cross-section given by Fig. 8 and Table I. However, adopting this could give rise to the false impression that similar fine details are definitely absent in the lower part of the mantle. Moreover, these fine details are obviously variable across the region. The very low velocity gradient shown by the model at depths of 130—180 km corresponds to observations in a distance range between approximately 1300 and 2200 km. Apparent velocity in this distance range is almost constant and first arrivals are usually weak. A small negative velocity gradient in this depth range is not ruled out by the travel-time .
.
3. P-wave velocity distribution in the mantle Several explosions used for long-range profiling in the U.S.S.R. were strong enough to be clearly recorded by ordinary seismograph stations at epicentral distances of 3000—4000 km (locations of seismic events were given by King and Calcagnile, 1976). The locations of the stations providing travel-time data are shown in Fig. 1. Some of the stations were used only to delineate lateral velocity heterogeneities. Arrival times were either read directly from the seismograms of the Soviet seismograph network or taken from the bulletins of Scandinavian stations. in several instances, times of later arrivalswere read from the record sections of NORSAR published by King and Calcagnile (1976). Travel-time data were corrected for effipticity. Figure 8 shows travel times
TABLE I P-velocity model ______________________________________________ Depth (km) 4 40 85
130 175 200 250
Velocity (km s~)
Depth (km)
823 8.29 8.43 8.44 8.45
285 380 390 410 540 640 650 780
8.59
Velocity (km s~)
8.65 867 8:68 9.35 9.75
10.08 10.70 11 00
______________________________________________________
32
70
Vetocit~j (icm/sec) b 6u
0 0
0
8
9
10
II
0
e 0
0
0
100
0
200
0000
•
00
C
000
o0
°
•
.•
•
500
d
~00
E
>,.
‘iO
300
~I
0
1000
2000 i~
~000
‘1000
(tcm)
Fig. 8. Reduced travel times ofP first arrivals (solid circles), of later arrivals (open circles) and the associated velocity model.
data but a small positive gradient is preferred for preserving simplicity of the model, The relatively rapid velocity increase below 200 km corresponds to the bend of the line a—b (Fig. 8) at distances over 2200 km. This line is very distinct in the NORSAR record sections published by King and Calcagnile (1976). Arrivals corresponding to the line are often reported in the bulletins of Scandinavian stations as clear and we easily identified them in a number of seismograms available to us. However, the details of transition from an apparent velocity of 8.6 km ~ near point a to 9.2 km s~’near point b are not very clear and transition to higher velocities below 200 km could be more rapid than that assumed in the model,
The extension of the line a—b is very sensitive ~o velocity gradient in the base of the upper mantle. In the model this gradient is very close to zero and the end of the line a—b is at a distance of about 3600 km. At the same time, according to NORSAR observations (King and Calcagnile, 1976) this branch effectively extends to at least 4000 km. A possible explanation for such long-range propagation is diffraction and scattering (King and Calcagnile, 1976). However, a similar effect could be caused by a negative velocity gradient in the base of the upper mantle. The rapid increase of velocity by approximately 0.7 km s~at a depth around 400 km corresponds to the line b—c. The arrivals near point c are often recorded as sharp onsets and can be reliably identi-
33
fled in a number of the seismograms. The results of extensive numerical testing imply that the mean depth of this “discontinuity” is 400 km with an error of about 10 km. A large positive gradient below 400 km seems necessary to explain travel times and relatively large amplitudes corresponding to the line c—d. A rapid increase of velocity near 640 km depth corresponds to the line d—e Unfortunately, this line comes too close to other lines associated with energetic arrivals. For that reason, good observations of line d—e are very few and reliable determination of the corresponding structure from P-wave data alone see~tisvery difficult. This part of the model is constrained by the PS-wave data (Section 4) while at the
same time it provides a good fit of calculated and observed P-wave travel times. A relatively large velocity gradient below 650 km is demanded by the curvature of the line e—f. A comparison of the travel times and of the model discussed above with previously published data for the same region (Fig. 9) reveals similar as well as different features. Speaking most generally, the travel times of first arrivals presented in this study are relatively close to those of Masse and Alexander (1974) while the configuration of refraction lines in later arrivals resembles that of King and Calcagnile (1976).
-
70
/ Vetocity (Km/sec
60
....-,
)
0
-
\k.. .-Ø#
100
...I,_,,_
1’1~’
.\~.
.~.
200
1’.. /
‘~~-‘::...
/
400
‘~:-.
2
‘S
~
100
1000
2000
3000
~
£1000
t~(icm) Fig.9. A comparison of P travel times and
associated velocity models: this study (solid lines); Masse and Alexander (1974)
(dotted lines); King and Calcagnlle (1976) (broken lines).
~
34
4. Analysis of converted waves is
r
r~ determinations from the observed t5t. Accuracy of r0 r~determination depends on the accuracy of the average values of P. and S-velocities in the corresponding depth range. The average value of P-velocity at depths of 400—650 km is so well error con1esponding strained byP travel times that the cori of r 0 r~can be neglected. For the P-velocity distribution of Table I, the time interval of 23.5 s between PS-arrivals corresponds to r0 r~values of 239, 252 and 266 km if the assumed values of Vp/Vs are 1.85, 1.80 and 1.75 respectively. In modern average Earth models (Dziewonski et a!., 1975; Hart et a!., 1977), the ratio of P- to S-velocities at depths of 400—650 km is close to 1.85. This value is appropriate for the East European platform since relatively low V~values associated with Vp/ Vs = 1.85 and transition from low Vs values to much higher values at the 650 km discontinuity would provide a reasonably high coefficient of conversion. The depth of the upper discontinuity is 400 ±10km (Section 3). Therefore the depth of the lower discontinuity is 639 ±10km. This value strongly deviates from depths of 710 and 690 km determined by Masse and Alexander (1974) and by King and Calcagnile (1976). The arrival time of 43.0 s after P of the 400 km PS-phase constrains possible V~models of the upper mantle. However, the problem of Vs distribution in the upper mantle is too complicated for a short discussion. According to Vinnik et al. (1979) this distribution involves two well pronounced velocity reversals. 0
—
—
Converted waves PS associated with velocity discontinuities in the mantle were identified in the long-period NORSAR records (Vinnik, 1977). The procedure axis rotation, transformation of of detection records toinvolved a standard form and stacking
—
of processed records from events of various epicentral distances. The procedure detected waves converted from P to SV in the depth ranges of about 400 and 650 km. For an apparent slowness of 6.4 s deg’, the arrival time of the 400 km phase was 45.5 s after P. Arrival times of the 650 km phase differed by 1 s for two data sets with an average value of 69.0 s after P. The amplitude of the 640 km phase was 3.7% of P for a 10 s period of oscillation, Recently this experiment was repeated with longperiod records from the Obninsk Observatory near Moscow (Vinnik et a!., 1979). PS-phases with arrival times 43.0 and 66.5 s after P were reliably identified. The discussion which follows is after Vinnik et al. (1979). To explain later PS-arrivals at NORSAR, one should take into account that the average travel-time residual of teleseismic P-waves in the NORSAR area is about +1.0 s relative to the rest of the Baltic shield (Noponen, 1977) while teleseismic P-residuals are close for the Baltic shield and for the Russian plate (Kogan, 1976). Thus, teleseismic P-arrivals are late by about 1 sand PS’arrivals by about 3.5 s at NORSAR relative to Obninsk. The ratio of PS-residual to P-residual at NORSAR is very close to the ratio of S-residual to P-residual arising from the lateral velocity variation in the upper mantle (e.g. Hales and Herrin, 1972). TheEefore later PS-arrivals at NORSAR are almost certainly due to a local velocity anomaly in the upper mantle while the discontinuities in the phase transition zone are at the same depths in both locations, The delay of the PS-arrival relative to P can be expressed as 2
r0
f
=r
—
p2r2)”2
—
(V~2 —
p2r2)”21 dr
[(V~
C
where p is the ray parameter, r~the radius of the converting interface and r 4, the radial distance of the observation point. This equation can be used for
—
5.
Lateral heterogeneity and azimuthal isotropy
The travel-time curve shown in Fig. 8 is a smooth mean of the observations. The scatter of travel-time determinations around the curve is caused mainly by lateral velocity heterogeneities in the upper mantle. Clear differences between velocity cross-sections of the uppermost mantle have been demonstrated in Section In this section we present further evidence of lateral2.P-velocity variation in the upper mantle of the East European platform and of adjacent regions The main points of this discussion follow from Yinnik et a!. (1978). Figure 10 shows examples of travel-time residuals with respect to the travel-time curve shown in Fig. 8.
35
4
~ o Baltic shield
-~
2
/
~
/0 °
~
W
~
t-Sbertan
P
Late
~‘\
o~
I-
E
S
I
~
-
\
j
°
~_~________i__,, 1000
0~
\ t
2000
Dtstance (Km)
8
-1
0
0
-2
Fig. 10. Examples of P travel-time residuals.
To make effects of mantle heterogeneities more evideil~,the crust was “stripped off’ and travel times were recomputed with reference to a depth of 40 km. The data for the Baltic shield were taken from several sources (references given by Vinnik et a!., 1978). Arrivals corresponding to the Baltic shield are systematically late at distances less than 1000 km and systematically early at distances over 1400 km. Late arri~ialsare explained by relatively low velocity in the uppermost layer of the mantle, perhaps 20 or 30 km thick. Early arrivals at distances over 1400 km are due to relatively high velocities at depths over 100 km. Another example in Fig. 10 is the travel-time residuals corresponding to the northern part of the West Siberian plate. In this area arrivals are late in the whole distance range implying a large vertical dimension of the low-velocity anomaly. Figure 1 shows a tentative scheme of P-velocity anomalies compiled from data similar to those shown in Fig. 10. Velocity residuals S V were found from the simplified expression
Anomalies shown in the scheme are correlated with certain tectonic features. For example, the highvelocity anomaly of the Baltic shield seems to correspond to a broad zone of very old rocks (up to 3500 My, after Khain, 1977). Low velocity in the upper mantle of the West Siberian plate is most likely related to its younger geological age. The same explanation is valid for the Caucasus and the adjacent Scythian plate. The scatter of travel times could be caused not only by mantle heterogeneity but by velocity anisotropy as well. Azimuthal amsotropy was recently suggested by several authors to be an important property of the lower lithosphere in Eurasia (e.g. Bamford and Crampin, 1977; Fuchs, 1977). Kogan (1976) found pronounced dependence of teleseismic P travel times on azimuth in Central and Northern Europe. According to her, the direction of the earliest arrivals in Europe is close to the direction of the highest velocity (N20°E)found by Bamford (1973) from refractionexperiments in West Germany. If the dependence described by Kogan were due to azimuthal anisotropy, a layer responsible for it would be 100—200 km thick with a coefficient of anisotropy
£10 0 •o~
•0 ..~
~
30
S
-z:~ <~ I
00
20
~,
(
SV/V
0=—5t/t0
where V0 is the normal velocity, t0 the normal travel time and St the travel-time residual. Unfortunately, the’data were not abundant enough to delineate anomalous regions accurately and some fairly large residuals could not be explained.
1000
2000
~ .
.
3000
(t~m)
Fig. 11. A comparisonof travel times ofP-waves propagatmg in azimuths 0—40°(open circles) and 90—130°(solid circles). The crust is “stripped oil”.
36
of the order of several per cent. Anisotropy of this magnitude would result in travel-time residuals of the order of several seconds at distances of 1000—2000 km. With available data this prediction can be easily tested (Fig. 11). Figure 11 shows travel times of P-waves propagating through the upper mantle of the East European platform in perpendicular directions. One of these directions corresponds to the direction of the earliest arrivals of teleseismic P-waves. A comparison shows no systematic difference between both sets of data. This observation implies that average azimuthal anisotropy in the upper mantle of the region does not exceed 0.5%.
6. Conclusions We have reviewed and analysed several kinds of seismic observations related to the structure of the mantle beneath the East European platform. Detailed observations of long-range profiles reveal complex and laterally variable P-velocity distribution in the lower lithosphere. Wave fields and associated velocity models in the regions of the Russian plate and of the North Caspian coastal depression are distinctly different, while according to available data deep structure of the Baltic shield is clearly different from structures in both these regions. These differences could haye bearings on the origin of continental uplift and subsidence. The highest P-velocities found at depths around 100 km were less than 8.5 km s~i. Lateral heterogeneity of the upper mantle of the platform is the principal source of scatter of traveltime determinations which amounts to several seconds at distances over 1000 km. P-wave travel-time data provide no evidence of azimuthal anisotropy in the upper mantle. However, the possibifity of other kinds of anisotropy or of azimuthal anisotropy confmed to small regions and very thin layers is not ruled out by these observations, The P-velocity model of the mantle derived in this study is characterized by the following features: (1)a relatively rapid increase in velocity at a depth of about 100 km; (2) a very small velocity gradient at 130—180 km; (3) a very small velocity gradient at 300—400 1cm; (4) rapid increase in velocity by about 0.7 km s~near 400 km; (5) a rapid increase in
velocity by about 0.6 km s~near 640 km; and (6) an almost constant velocity gradient between 400 and 640 km depths. The model is based on fairly accurate travel-time data and is different in this respect from previously known models for the same region. A comparison of this model with numerous models for other old platforms reveals similar as well as different features. However, the question to what extent these differences reflect real variations in the structure of the mantle remains highly obscure. The depth of the 640 km discontinuity was derived from the observations of PS-phases on the assumption that Vp/ V~= 1.85 between 400 and 640 km depths. The depth thus obtained is 30 km less than that adopted in the recent average Earth models (Dziewonski et al., 1975; Hart et a!., 1977). A depth consistent with these models can be obtained only on the assumption of very high S-velocity (Vp/Vs = 1 .75) in the transition zone. The impossibility of reconciling the PS-data and properties of the average Earth models could be due tp large-scale velocity heterogeneities in the transition zone. To clarify this problem similar studies are necessary in other regions.
Acknowledgements -
We are grateful to Dr. E. Bessonova and G. Sitnikova for their very valuable help with the computations.
References Bamford, D., 1973. Refraction data in Western Germany — a time-term interpretation. Z. Geophys., 39: 907—927. Bamford, D. and Crampin, S., 1977. Seismic anisotropy — the state of the art. Geophys. J. R. Astron. Soc., 49: 1—8. Burmakov, Yu.A., Egorkin, A.V., Popov, E.A. and Ryaboy, V.Z., 1975. Upper mantle structure in the north-eastern regions of the European platform according to seismic data. DokI. Akad. Sd. USSR, 224: 84—87 (in Russian). Dziewonski, A.M., Hales, A.L. and Lapwood, E.R., 1975. Parametrically simple Earth models consistent with geo-
physical data. Phys. Earth Planet. Inter., 10: 12—48. Egorkin, A.V., Ryaboy, V.Z., Starobinets, L.N. and Druzhinin, V.S., 1977. Velocity profiles of the upper mantle according to DSS observations on land. Bull. Acad. Sci. USSR, Earth Phys., No. 7: 27—41 (in Russian). Fuchs, K., 1977. Seismic anisotxopy of the subcrustal litho-
37
sphere as evidence for dynamical
processes in the upper
mantle. Geophys. J. R. Astron. Soc., 49: 167—179. Hales, A.L. and Herrin, E., 1972. Travel times of seismic waves. In: The Nature ofthe Solid Earth, McGraw-Hffl International Series in Earth and Planetary Sciences, pp. 172—215. Hart, R.S.,Anderson, D.L. and Kanamori, H., 1977. The effect of attentuation on gross Earth models. I. Geophys. Res., 82: 1647—1654. Khain, V.E., 1977. Regional Geotectonics. Nedra, Moscow, 359 pp. King, D.W. and Calcagnile, G., 1976. P-wave velocities in the upper mantle beneath Fennoscandia and Western Russia. Geophys. J. R. Astron. Soc., 46: 407—432. Kogan, S.D., 1976. Experimental travel time curve and lateral heterogeneity of the mantle. Dokl. Akad. Sci. USSR, 230: 1318—1321 (in Russian).
Masse, R.P. and Alexander, S.S., 1974. Compressional velocity distribution beneath Scandinavia and Western Russia.
Geophys. J. R. Astron. Soc., 39: 587—602. Noponen, I., 1977. The Blue Road geotraverse: the relative residual of the teleseismic P-wave: application to the study of the deep structure in Fennoscandia. Geol. Fdren. Stockholm Fdrh., 99: 32—36. Vinnik, L.P., 1977. Detection of waves converted from P to SV in the mantle. Phys. Earth Planet. Inter., 15: 39—45, Vinnik, L.P., Ryaboy, V.Z., Starobinets, L.N., Egorkin, A.V. and Chernyshev, N.M., 1978. Velocities of P-waves in the upper mantle of the East European platform. Dokl. Akad. Sci. USSR, 244: 70—73 (in Russian). Vinnik, L.P., Mikhailova, N.G. and Avetisjan, R., 1979. Analysis of observations ofwaves converted from P to SV in the mantle, Dold. Akad. Sci. USSR, 248: 573—576 (in Russian).
21
DEEP STRUCTURE OF THE EAST EUROPEAN PLATFORM ACCORDING TO SEISMIC DATA *
LP. VINNIK ’ and V.Z. RYABOY * 1 Institute of Physics of the Earth, Moscow (U.S.S.R.) * Institute of Geophysical Exploration, Moscow (U.S.S.R.)
(Received January 12,1979; accepted for publication June 2,1979)
Vinnik, L.P. and Ryaboy, V.Z., 1981. Deep structure of the East European platform according to seismic data. Phys. Earth Planet. Inter., 25: 27-37.’ The paper presents a review and analysis of new seismic data related to the structure of the mantle beneath the East European platform. Analysis of observations of long-range profiles revealed pronounced differences in the structure of the lower lithosphere beneath the Russian plate and the North Caspian coastal depression. The highest P-velocities found at depths around 100 km are in the range 8.4-8.5 km s Deep structure of the Baltic shield is different from the structures of both these regions. No evidence of azimuthal anisotropy In the upper mantle was found. A distribution of P-velocity in the upper mantle and in the transition zone consistent with accurate travel-time data was determined. The model Involves several zones of small and large positive velocity gradients In the upper mantle, rapid increases of velocity near 400 and 640 km depths and an almost constant positive velocity gradient between the 400 and 640 km discontinulties. The depth of the 640 km discontinuity was determined from observations of waves converted from P to SV in the mantle.
-r.
1. Introduction For better understanding of the structure and composition of the Earth’s interior, detailed velocity cross-sections of the mantle have to be determined in many regions. Such cross-sections could be obtained from abundant and accurate seismic data, which unfortunately are available in very few regions. The Precambrian East European platform occupies the area to the west of the Urals and includes the Russian plate and Baltic shield as the main structural units (Fig. 1). Until very recently this vast area was almost devoid of goodquality seismic data for the upper mantle. However, in recent years many new data have been obtained in this region.
* Presented at the symposium: ‘Velocity, Attenuation and other Geophysical Properties of the Upper Mantle”, held in Leningrad, U.S.S.R., October 2-l 1,1978. 0031-9201/81/0000-OOOO/$
Among new seismic studies, long-range refraction profiling has been undertaken in the U.S.S.R. (Burmakov et al.,1975;Egorkinet al., 1977).The analysis of later arrivals of P-waves propagating through the mantle of the East European platform and recorded by the NORSAR array (Ring and Calcagnile, 1976) is a valuable contribution. Absolute travel times assumed by Ring and Calcagnile were not accurate enough and later Vinnik et al. (1978) determined travel times more accurately. Another recent contribution to the studies of the deep interior is the detection of waves converted from P to SV in the mantle (Vinnik, 1977). Our paper presents a review and analysis of new seismic data related to the structure of the mantle beneath the East European platform. Section 2 contains a new analysis of observations of the long-range profiles. In Section 3 we describe P travel-time data in the distance range -1000-3500 km and an associated velocity model. Observations of converted waves are
02.50 0 1981 Elsevier Scientific Publishing Company.
28
The Kineshma—Vorkuta profile crosses the north.
~:
‘_~~
•-$~
____________________
Russian pLat~~k~j0~
_________
~
________________
____________ —
______
_________
______
400 km long. reversed profiles These of the profiles same were length. coupled Figures with 2 and three 3
=
_________ _________
______________
—~---
2
~llE~4 ____ ,,.‘ -.
_______________________________
Pechora depression (east of Timan). According to eastern margin of the East European and of DSS about data 40 thekm Mohoroviëi~discontinuity everywhere along theplatform profile. is at aThe depth thickness of the low-velocity sedimentary layer affecting average crustal velocity is about 3—A km in the region of the Russian plate while increases up to 10—12 km in the Pechora depression. The observations in the main profile 1400 km long were complemented by overlapping profiles 800 and
•ai
o
________________
Fig. 1. Map of the region and tentative scheme of P-velocity anomalies in the upper mantle: 1,seismograph stations; 2, long-range profiles (I, Kineshma—Vorkuta, II, Elista— Buzuluk); 3, normal velocity; 4, anomalous velocity at depths less than 200 km ; 5, anomalous velocity at depths less than 1). 100 km;6, residual velocity (km s
analysed in Section 4. In Section 5 lateral heterogeneities in the upper mantle are described. In the same section we touch on the question of seismic anisotropy. The principal fmdings are summarized in Section 6.
2. Fine structure of the uppermost mantle Detailed observations in profiles more than 700 km long were made in the U.S.S.R. in two regions of the East European platform (Kineshma—Vorkuta and Elista—Buzuluk profiles) (locations of profiles are shown in Fig. 1). Field observations were made with calibrated equipment having frequency bands from 1—2 to 10—15 Hz. Velocity sections of the crust were obtained along the same proffles (Egorkin et al., 1977).
show record sections of the main profile and of the overlapping 800 km profile. The end-points of both profiles coincide. In the record section of the main profile, two refraction lines are visible. The first line with an apparent velocity of 8.55 km s~’is observed in a distance range from 200 to 600 km. The second line emerges in later arrivals at a distance of about 700 km and becomes the first arrival at a distance of about 1100 km. Both lines are almost parallel but the second line is late relative to the first by about 3 S. Offsets of this kind are characteristic of low-velocity layers. First arrivalsin a distance range between 600 and 1100 km have an apparent velocity of about 8.0—8.1 km s~.This velocity is too low to be appropriate for any refraction line and the only reasonable explanation for these arrivals is scattering of primary waves by small-scale lateral heterogeneities. The same features are evident in the record section of the overlapping profile (Fig. 3). The only difference is a somewhat lower apparent velocity of the first refraction line (835 km s’’). To account for lateral velocity heterogeneities in the crust, allowances up to 1.0 S were applied to the observed travel times. A velocity model of the uppermost mantle consistent with the corrected traveltime data was derived by trial and error. A remarkable feature of this model (Fig. 4) is the well pronounced low-velocity layer at depths of 60—110 km. Of course, this model is only a representative of a family of models consistent with observational data. However, the existence of a low-velocity layer with properties qualitatively similar to those of our model is demanded by the data. The second area where data of long-range pro-
29 21
ti
t
20
-
0
______
________
iOO
___-~._
—
300
—
400
1
L
900
—__~
—
_J
- —
______
1(5)
600
_____
—
—
____________
700
Boo ~ (en)
4A~It
~
i~oo
IiOO
1000
ii ~ 1300
ILIOO
Fig. 2. Record section of Kjneshma—Vorkuta profile from southwest to northeast; amplitudes are normalized.
200
300
400
500
600
Fig. 3. As in Fig. 2, for the overlapping profile.
700
800
30
/
II
0
______
/ ~
/
filing are available is the North Caspian coastal depression (Elista—Buzuluk profile). The crust in the
ve1oc~ty Km/sec) 8,,, 8~S
central part of the depression is almost devoid of a “granitic” layer and the sedimentary layer is up to
50
~N. N,,~
~-
/ /
20km thick. Record sections of the direct and reversed profiles 800 km long are shown in Figs. 5 and 6. In both sections two refraction lines are visible. Arrivals corresponding to the line with an apparent velocity of 8.2 km s’ are observed in the distance range 200—500 km. Arrivals corresponding to the second line are observed as first arrivals at distances over 600 km. The apparent velocity of the second line is 8.85 km s’~ in the direct profile and 8.55 km
50
500
1Q00
~ (Km~
Fig. 4. P-velocity model and corresponding reduced traveltime—distance curve for Kineshma—Vorkuta profile. The velocity in the crust is 6.2 kin 5i~
~
k
I~
L
t -~
~
-~
~~-‘:.~
I
25
800
700
800
500
400
500
200
~ (om) Fig. 5. Record section of Elista—Buzuluk profile from southwest to northeast; amplitudes are normalized.
: ~20
:10
0
J1I~~~ I I411 ~I ~WtJftJ!JijL ‘
_____________H 200
I
-
300
,,
400
500
~ (on) Fig. 6. As in Fig. 5, for the reversed profile.
600
700
31 Vetotitilom/sec)
______________
0
I
~ T ~
and the associated velocity model, which is presented in Table I.
___0
/
The relatively low velocity gradient an the uppermost layer of the mantle is terminated by the “dis-
~
/
~0 I
N. 500
-~
N
100 000
~ (~cm)
Fig. 7. P-velocity model and the corresponding reduced traveltime—distance curve for Elista—Buzuluk profile. The velocity in the crust is 5.8 km ~
s in the reversed profile. The value of 8.85 km s~ is strongly in error due to variable thickness of sediments in the northeastern end of the profile. If allowances for the crustal heterogeneities are applied to travel times, the apparent velocity of the second line in the direct profile becomes 8.55 instead of 8.85 km s~.Figure 7 shows a velocity model consistent with the corrected travel-time data. A thin low-velocity layer explains the very low amplitudes of the first arrivals in the distance range 500—600 km. Thus, for the two structural units of the East European platform markedly different velocity models of the uppermost mantle are obtained. These differences are associated with evident differences in the wave fields,
continuity” at depths around 100 km. This “discontinuity” corresponds to a distinct increase in apparent velocity from 83 km s’~at distances of less than 800 km to 8.5 km s’ at distances over 1000 km. In a distance range 200—1000 km many of the traveltime determmations shown an Fig. 8 were taken from the Kineshma—Vorkuta profile (Section 2) and thus are representative mainly of the Russian plate. This was decided in view of the dominating linear dimension of the Russian plate among other geological structures in the region. The upper part of the model is therefore a simplified and smoothed version of the more complicated model shown in Fig. 4. This more detailed model could be adopted as the upper part of the velocity cross-section given by Fig. 8 and Table I. However, adopting this could give rise to the false impression that similar fine details are definitely absent in the lower part of the mantle. Moreover, these fine details are obviously variable across the region. The very low velocity gradient shown by the model at depths of 130—180 km corresponds to observations in a distance range between approximately 1300 and 2200 km. Apparent velocity in this distance range is almost constant and first arrivals are usually weak. A small negative velocity gradient in this depth range is not ruled out by the travel-time .
.
3. P-wave velocity distribution in the mantle Several explosions used for long-range profiling in the U.S.S.R. were strong enough to be clearly recorded by ordinary seismograph stations at epicentral distances of 3000—4000 km (locations of seismic events were given by King and Calcagnile, 1976). The locations of the stations providing travel-time data are shown in Fig. 1. Some of the stations were used only to delineate lateral velocity heterogeneities. Arrival times were either read directly from the seismograms of the Soviet seismograph network or taken from the bulletins of Scandinavian stations. in several instances, times of later arrivalswere read from the record sections of NORSAR published by King and Calcagnile (1976). Travel-time data were corrected for effipticity. Figure 8 shows travel times
TABLE I P-velocity model ______________________________________________ Depth (km) 4 40 85
130 175 200 250
Velocity (km s~)
Depth (km)
823 8.29 8.43 8.44 8.45
285 380 390 410 540 640 650 780
8.59
Velocity (km s~)
8.65 867 8:68 9.35 9.75
10.08 10.70 11 00
______________________________________________________
32
70
Vetocit~j (icm/sec) b 6u
0 0
0
8
9
10
II
0
e 0
0
0
100
0
200
0000
•
00
C
000
o0
°
•
.•
•
500
d
~00
E
>,.
‘iO
300
~I
0
1000
2000 i~
~000
‘1000
(tcm)
Fig. 8. Reduced travel times ofP first arrivals (solid circles), of later arrivals (open circles) and the associated velocity model.
data but a small positive gradient is preferred for preserving simplicity of the model, The relatively rapid velocity increase below 200 km corresponds to the bend of the line a—b (Fig. 8) at distances over 2200 km. This line is very distinct in the NORSAR record sections published by King and Calcagnile (1976). Arrivals corresponding to the line are often reported in the bulletins of Scandinavian stations as clear and we easily identified them in a number of seismograms available to us. However, the details of transition from an apparent velocity of 8.6 km ~ near point a to 9.2 km s~’near point b are not very clear and transition to higher velocities below 200 km could be more rapid than that assumed in the model,
The extension of the line a—b is very sensitive ~o velocity gradient in the base of the upper mantle. In the model this gradient is very close to zero and the end of the line a—b is at a distance of about 3600 km. At the same time, according to NORSAR observations (King and Calcagnile, 1976) this branch effectively extends to at least 4000 km. A possible explanation for such long-range propagation is diffraction and scattering (King and Calcagnile, 1976). However, a similar effect could be caused by a negative velocity gradient in the base of the upper mantle. The rapid increase of velocity by approximately 0.7 km s~at a depth around 400 km corresponds to the line b—c. The arrivals near point c are often recorded as sharp onsets and can be reliably identi-
33
fled in a number of the seismograms. The results of extensive numerical testing imply that the mean depth of this “discontinuity” is 400 km with an error of about 10 km. A large positive gradient below 400 km seems necessary to explain travel times and relatively large amplitudes corresponding to the line c—d. A rapid increase of velocity near 640 km depth corresponds to the line d—e Unfortunately, this line comes too close to other lines associated with energetic arrivals. For that reason, good observations of line d—e are very few and reliable determination of the corresponding structure from P-wave data alone see~tisvery difficult. This part of the model is constrained by the PS-wave data (Section 4) while at the
same time it provides a good fit of calculated and observed P-wave travel times. A relatively large velocity gradient below 650 km is demanded by the curvature of the line e—f. A comparison of the travel times and of the model discussed above with previously published data for the same region (Fig. 9) reveals similar as well as different features. Speaking most generally, the travel times of first arrivals presented in this study are relatively close to those of Masse and Alexander (1974) while the configuration of refraction lines in later arrivals resembles that of King and Calcagnile (1976).
-
70
/ Vetocity (Km/sec
60
....-,
)
0
-
\k.. .-Ø#
100
...I,_,,_
1’1~’
.\~.
.~.
200
1’.. /
‘~~-‘::...
/
400
‘~:-.
2
‘S
~
100
1000
2000
3000
~
£1000
t~(icm) Fig.9. A comparison of P travel times and
associated velocity models: this study (solid lines); Masse and Alexander (1974)
(dotted lines); King and Calcagnlle (1976) (broken lines).
~
34
4. Analysis of converted waves is
r
r~ determinations from the observed t5t. Accuracy of r0 r~determination depends on the accuracy of the average values of P. and S-velocities in the corresponding depth range. The average value of P-velocity at depths of 400—650 km is so well error con1esponding strained byP travel times that the cori of r 0 r~can be neglected. For the P-velocity distribution of Table I, the time interval of 23.5 s between PS-arrivals corresponds to r0 r~values of 239, 252 and 266 km if the assumed values of Vp/Vs are 1.85, 1.80 and 1.75 respectively. In modern average Earth models (Dziewonski et a!., 1975; Hart et a!., 1977), the ratio of P- to S-velocities at depths of 400—650 km is close to 1.85. This value is appropriate for the East European platform since relatively low V~values associated with Vp/ Vs = 1.85 and transition from low Vs values to much higher values at the 650 km discontinuity would provide a reasonably high coefficient of conversion. The depth of the upper discontinuity is 400 ±10km (Section 3). Therefore the depth of the lower discontinuity is 639 ±10km. This value strongly deviates from depths of 710 and 690 km determined by Masse and Alexander (1974) and by King and Calcagnile (1976). The arrival time of 43.0 s after P of the 400 km PS-phase constrains possible V~models of the upper mantle. However, the problem of Vs distribution in the upper mantle is too complicated for a short discussion. According to Vinnik et al. (1979) this distribution involves two well pronounced velocity reversals. 0
—
—
Converted waves PS associated with velocity discontinuities in the mantle were identified in the long-period NORSAR records (Vinnik, 1977). The procedure axis rotation, transformation of of detection records toinvolved a standard form and stacking
—
of processed records from events of various epicentral distances. The procedure detected waves converted from P to SV in the depth ranges of about 400 and 650 km. For an apparent slowness of 6.4 s deg’, the arrival time of the 400 km phase was 45.5 s after P. Arrival times of the 650 km phase differed by 1 s for two data sets with an average value of 69.0 s after P. The amplitude of the 640 km phase was 3.7% of P for a 10 s period of oscillation, Recently this experiment was repeated with longperiod records from the Obninsk Observatory near Moscow (Vinnik et a!., 1979). PS-phases with arrival times 43.0 and 66.5 s after P were reliably identified. The discussion which follows is after Vinnik et al. (1979). To explain later PS-arrivals at NORSAR, one should take into account that the average travel-time residual of teleseismic P-waves in the NORSAR area is about +1.0 s relative to the rest of the Baltic shield (Noponen, 1977) while teleseismic P-residuals are close for the Baltic shield and for the Russian plate (Kogan, 1976). Thus, teleseismic P-arrivals are late by about 1 sand PS’arrivals by about 3.5 s at NORSAR relative to Obninsk. The ratio of PS-residual to P-residual at NORSAR is very close to the ratio of S-residual to P-residual arising from the lateral velocity variation in the upper mantle (e.g. Hales and Herrin, 1972). TheEefore later PS-arrivals at NORSAR are almost certainly due to a local velocity anomaly in the upper mantle while the discontinuities in the phase transition zone are at the same depths in both locations, The delay of the PS-arrival relative to P can be expressed as 2
r0
f
=r
—
p2r2)”2
—
(V~2 —
p2r2)”21 dr
[(V~
C
where p is the ray parameter, r~the radius of the converting interface and r 4, the radial distance of the observation point. This equation can be used for
—
5.
Lateral heterogeneity and azimuthal isotropy
The travel-time curve shown in Fig. 8 is a smooth mean of the observations. The scatter of travel-time determinations around the curve is caused mainly by lateral velocity heterogeneities in the upper mantle. Clear differences between velocity cross-sections of the uppermost mantle have been demonstrated in Section In this section we present further evidence of lateral2.P-velocity variation in the upper mantle of the East European platform and of adjacent regions The main points of this discussion follow from Yinnik et a!. (1978). Figure 10 shows examples of travel-time residuals with respect to the travel-time curve shown in Fig. 8.
35
4
~ o Baltic shield
-~
2
/
~
/0 °
~
W
~
t-Sbertan
P
Late
~‘\
o~
I-
E
S
I
~
-
\
j
°
~_~________i__,, 1000
0~
\ t
2000
Dtstance (Km)
8
-1
0
0
-2
Fig. 10. Examples of P travel-time residuals.
To make effects of mantle heterogeneities more evideil~,the crust was “stripped off’ and travel times were recomputed with reference to a depth of 40 km. The data for the Baltic shield were taken from several sources (references given by Vinnik et a!., 1978). Arrivals corresponding to the Baltic shield are systematically late at distances less than 1000 km and systematically early at distances over 1400 km. Late arri~ialsare explained by relatively low velocity in the uppermost layer of the mantle, perhaps 20 or 30 km thick. Early arrivals at distances over 1400 km are due to relatively high velocities at depths over 100 km. Another example in Fig. 10 is the travel-time residuals corresponding to the northern part of the West Siberian plate. In this area arrivals are late in the whole distance range implying a large vertical dimension of the low-velocity anomaly. Figure 1 shows a tentative scheme of P-velocity anomalies compiled from data similar to those shown in Fig. 10. Velocity residuals S V were found from the simplified expression
Anomalies shown in the scheme are correlated with certain tectonic features. For example, the highvelocity anomaly of the Baltic shield seems to correspond to a broad zone of very old rocks (up to 3500 My, after Khain, 1977). Low velocity in the upper mantle of the West Siberian plate is most likely related to its younger geological age. The same explanation is valid for the Caucasus and the adjacent Scythian plate. The scatter of travel times could be caused not only by mantle heterogeneity but by velocity anisotropy as well. Azimuthal amsotropy was recently suggested by several authors to be an important property of the lower lithosphere in Eurasia (e.g. Bamford and Crampin, 1977; Fuchs, 1977). Kogan (1976) found pronounced dependence of teleseismic P travel times on azimuth in Central and Northern Europe. According to her, the direction of the earliest arrivals in Europe is close to the direction of the highest velocity (N20°E)found by Bamford (1973) from refractionexperiments in West Germany. If the dependence described by Kogan were due to azimuthal anisotropy, a layer responsible for it would be 100—200 km thick with a coefficient of anisotropy
£10 0 •o~
•0 ..~
~
30
S
-z:~ <~ I
00
20
~,
(
SV/V
0=—5t/t0
where V0 is the normal velocity, t0 the normal travel time and St the travel-time residual. Unfortunately, the’data were not abundant enough to delineate anomalous regions accurately and some fairly large residuals could not be explained.
1000
2000
~ .
.
3000
(t~m)
Fig. 11. A comparisonof travel times ofP-waves propagatmg in azimuths 0—40°(open circles) and 90—130°(solid circles). The crust is “stripped oil”.
36
of the order of several per cent. Anisotropy of this magnitude would result in travel-time residuals of the order of several seconds at distances of 1000—2000 km. With available data this prediction can be easily tested (Fig. 11). Figure 11 shows travel times of P-waves propagating through the upper mantle of the East European platform in perpendicular directions. One of these directions corresponds to the direction of the earliest arrivals of teleseismic P-waves. A comparison shows no systematic difference between both sets of data. This observation implies that average azimuthal anisotropy in the upper mantle of the region does not exceed 0.5%.
6. Conclusions We have reviewed and analysed several kinds of seismic observations related to the structure of the mantle beneath the East European platform. Detailed observations of long-range profiles reveal complex and laterally variable P-velocity distribution in the lower lithosphere. Wave fields and associated velocity models in the regions of the Russian plate and of the North Caspian coastal depression are distinctly different, while according to available data deep structure of the Baltic shield is clearly different from structures in both these regions. These differences could haye bearings on the origin of continental uplift and subsidence. The highest P-velocities found at depths around 100 km were less than 8.5 km s~i. Lateral heterogeneity of the upper mantle of the platform is the principal source of scatter of traveltime determinations which amounts to several seconds at distances over 1000 km. P-wave travel-time data provide no evidence of azimuthal anisotropy in the upper mantle. However, the possibifity of other kinds of anisotropy or of azimuthal anisotropy confmed to small regions and very thin layers is not ruled out by these observations, The P-velocity model of the mantle derived in this study is characterized by the following features: (1)a relatively rapid increase in velocity at a depth of about 100 km; (2) a very small velocity gradient at 130—180 km; (3) a very small velocity gradient at 300—400 1cm; (4) rapid increase in velocity by about 0.7 km s~near 400 km; (5) a rapid increase in
velocity by about 0.6 km s~near 640 km; and (6) an almost constant velocity gradient between 400 and 640 km depths. The model is based on fairly accurate travel-time data and is different in this respect from previously known models for the same region. A comparison of this model with numerous models for other old platforms reveals similar as well as different features. However, the question to what extent these differences reflect real variations in the structure of the mantle remains highly obscure. The depth of the 640 km discontinuity was derived from the observations of PS-phases on the assumption that Vp/ V~= 1.85 between 400 and 640 km depths. The depth thus obtained is 30 km less than that adopted in the recent average Earth models (Dziewonski et al., 1975; Hart et a!., 1977). A depth consistent with these models can be obtained only on the assumption of very high S-velocity (Vp/Vs = 1 .75) in the transition zone. The impossibility of reconciling the PS-data and properties of the average Earth models could be due tp large-scale velocity heterogeneities in the transition zone. To clarify this problem similar studies are necessary in other regions.
Acknowledgements -
We are grateful to Dr. E. Bessonova and G. Sitnikova for their very valuable help with the computations.
References Bamford, D., 1973. Refraction data in Western Germany — a time-term interpretation. Z. Geophys., 39: 907—927. Bamford, D. and Crampin, S., 1977. Seismic anisotropy — the state of the art. Geophys. J. R. Astron. Soc., 49: 1—8. Burmakov, Yu.A., Egorkin, A.V., Popov, E.A. and Ryaboy, V.Z., 1975. Upper mantle structure in the north-eastern regions of the European platform according to seismic data. DokI. Akad. Sd. USSR, 224: 84—87 (in Russian). Dziewonski, A.M., Hales, A.L. and Lapwood, E.R., 1975. Parametrically simple Earth models consistent with geo-
physical data. Phys. Earth Planet. Inter., 10: 12—48. Egorkin, A.V., Ryaboy, V.Z., Starobinets, L.N. and Druzhinin, V.S., 1977. Velocity profiles of the upper mantle according to DSS observations on land. Bull. Acad. Sci. USSR, Earth Phys., No. 7: 27—41 (in Russian). Fuchs, K., 1977. Seismic anisotxopy of the subcrustal litho-
37
sphere as evidence for dynamical
processes in the upper
mantle. Geophys. J. R. Astron. Soc., 49: 167—179. Hales, A.L. and Herrin, E., 1972. Travel times of seismic waves. In: The Nature ofthe Solid Earth, McGraw-Hffl International Series in Earth and Planetary Sciences, pp. 172—215. Hart, R.S.,Anderson, D.L. and Kanamori, H., 1977. The effect of attentuation on gross Earth models. I. Geophys. Res., 82: 1647—1654. Khain, V.E., 1977. Regional Geotectonics. Nedra, Moscow, 359 pp. King, D.W. and Calcagnile, G., 1976. P-wave velocities in the upper mantle beneath Fennoscandia and Western Russia. Geophys. J. R. Astron. Soc., 46: 407—432. Kogan, S.D., 1976. Experimental travel time curve and lateral heterogeneity of the mantle. Dokl. Akad. Sci. USSR, 230: 1318—1321 (in Russian).
Masse, R.P. and Alexander, S.S., 1974. Compressional velocity distribution beneath Scandinavia and Western Russia.
Geophys. J. R. Astron. Soc., 39: 587—602. Noponen, I., 1977. The Blue Road geotraverse: the relative residual of the teleseismic P-wave: application to the study of the deep structure in Fennoscandia. Geol. Fdren. Stockholm Fdrh., 99: 32—36. Vinnik, L.P., 1977. Detection of waves converted from P to SV in the mantle. Phys. Earth Planet. Inter., 15: 39—45, Vinnik, L.P., Ryaboy, V.Z., Starobinets, L.N., Egorkin, A.V. and Chernyshev, N.M., 1978. Velocities of P-waves in the upper mantle of the East European platform. Dokl. Akad. Sci. USSR, 244: 70—73 (in Russian). Vinnik, L.P., Mikhailova, N.G. and Avetisjan, R., 1979. Analysis of observations ofwaves converted from P to SV in the mantle, Dold. Akad. Sci. USSR, 248: 573—576 (in Russian).