Russian Journal of Earth Sciences
Vol 1, No. 1, July 1998
Translated December 1998

Postglacial tectonics of the Baikal rift

K. G. Levi1, V. D. Mats2, Yu. S. Kusner3, P. G. Kirillov1, A. M. Alakshin4, S. V. Tolstov3, E. Yu. Osipov2, I. M. Efimova2, S. Bak5

1 Institute of the Earth's Crust, Siberian Division, Russian Academy of Sciences,
2 Institute of Limnology, Siberian Division, Russian Academy of Sciences,
3 Institute of Geochemistry, Siberian Division, Russian Academy of Sciences,
4 GGP "Irkutskgeofizika'',
5 University of Potsdam, Germany



Many phenomena related to sedimentary basins and adjoining mountainous areas of the Baikal rift zone have not been adequately interpreted, as yet. These are widespread occurrence of continental sands in rifting basins of the Baikal region, (their formation was traditionally associated with the maximum glaciation in the middle Pleistocene, but now it is considered as a stratigraphic element encompassing the whole Pleistocene and including deposits of various genetic types and various climatic phases [Bazarov et al., 1982; Logachev et al., 1974; Mats, 1987; Olyunin, 1961]); moraines that occur at depths of 300-400 m below the water level the Lake Baikal on its Barguzin slope, which may imply a lower lake level existing at that time [Galkin, 1975]; terraces on the Ushkan'i Islands, much more numerous as compared with four terraces on the lake bank [Lamakin, 1968] (the origin of the "superfluous" ones need be explained, but the number of shore terraces may also exceed ten, in which case they should be correlated with the terrace levels of the Bol'shoi Ushkanii Island [Eskin et al., 1959]); and the unexplained presence of exotic boulders on the Olkhon and Ushkan'i Islands [Bukharov and Fialkov, 1996], as well as in certain eastern coastal areas of North Baikal. Also surprising are abnormally high velocities of recent vertical movements (RVM) of the Earth's surface in areas of a Pleistocene ice field [Logachev et al., 1974], tectonic fractures in glacier structures, and other phenomena. On the whole, these phenomena suggest that the North Baikal region may have experienced glacioisostatic movements similar to those presently observed in Fennoscandia and on the Canadian shield. We will discuss recent geological constraints on the Baikal glaciations and neotectonic deformations consistent with glacioisostatic movements and present a numerical model of the lithosphere at the postglacial stage of its evolution in the Baikal region.

Pleistocene Glaciation

The problem of glaciation in the Baikal region has been debated beginning from the works by V. D. Cherskii, P. N. Kropotkin, and others. Obruchev [1953] noted the presence of glacier structures in the Baikal region. Afterwards the presence of glaciers in the past was proven, and their dimensions and stages of their development have been established. More detailed data on these natural phenomena are reported in a number of works [Baikal Atlas, 1993; Bazarov, 1986; Bazarov et al., 1982; Kulchitskii, 1967; Lamakin, 1968; Logachev et al., 1974; Salop, 1964]. Geological and geomorphological studies revealed the existence of four glaciations, the most intense and oldest of which is the Samarovo Glaciation [Kulchitskii, 1967].

fig01 The maximum glaciation (300-250 ka) covered a vast territory in the Baikal region [Logachev et al., 1974], but a more or less monolithic ice cap that covered the northern Baikal basin and adjacent fig02 ranges (Figures 1 and 2) is most pertinent to the purposes of our paper. The ice field was not continuous.

The underlying surface was comparatively smooth. Fragments of the end moraine belt indicate the glacier to cover an area of more than 100 000 km 2. Ice tongues descended from mountains outward, toward the areas surrounding the Baikal rift zone, and inward the North Baikal depression.

fig03 Maximum glaciation moraines existed on the western and eastern flanks of the North Baikal basin (Figure 3), and their radiocarbon ages are presented in Table 1.

On the western North Baikal coast, a buried 90-m thick moraine was drilled through at depths of 17 to 106 m below the present level of Lake Baikal [Kulchitskii, 1967]. To the east, the Samarovo moraines, extending into the lake water area for more than 7 km, have been traced to depths of 350-400 m. They are supposed to have formed under subaerial conditions [Galkin, 1975], because morphologically expressed moraines of recent glaciation, developed under subaqueous conditions, are unknown. Then, the lake level considerably dropped some 300 ka.

Younger glacier structures include those of the late mid-Pleistocene (Taz stage) and early Late Pleistocene (early Ermakov stage, 80 ka and later) and have similar relationships with Baikal terraces. At greater distances from the shore zone, they form surfaces that preserve specific features of the glacier relief. In the near-shore zone, their surfaces experienced abrasion and show the plane relief of Baikal terraces. The latter are developed at levels of 150, 80, and 35-50 m. As is observed in the Tyya Promontory, in bank bluffs from the Kurla Head to Tyya River mouth, at both heads of the Frolikh Cove, in the left-hand divide area of the Biramya River near its mouth, and at several other places, intricate facial relationships characterize deposits that compose the terraces: lacustrine, glacial-lacustrine, and glacial boulder loams are overlain by lacustrine deposits and covering loam with relict lake pebble. Overall, this group of glacial deposits is characterized by close association with lake deposits. The latter include widespread frost involutions and clastic material with clasts as large as gigantic blocks (bearing obvious traces of glacial abrasion).

Thus, general synchronism has been established between glacial structures and Lake Baikal deposits (occasionally including endemic diatoms), and the moraine relief in the shore contact zone of Lake Baikal has been reworked as a result of abrasion. The penetration of lake Baikal facies into glacial ones appears to reflect specific relationships between the lake water and glacier tongues during the deglaciation period. The age of these formations is determined from geological and geomorphological evidence, findings of remnants of large mammal species specific to the middle Pleistocene in the section of a 80-m terrace [Bazarov, 1986], and radiocarbon datings ( ge57 ka) of 40-m terrace deposits, Tyya Promontory (Table 1).

Younger moraines are widely developed on the shores of North Baikal (Figure 3). Their characteristic feature is a well-preserved ensemble of bottom, side, and end moraines and fluvioglacial boulder-pebble and lake plains behind and ahead of the front of end moraines. They are evidently younger than the glacial structures associated with 30-50-m terraces; river and Baikal terraces as high as 20-25 m are leaned against or occasionally inset in the latter. The above glacial features may include two age groups: late Markov, as old as 50 ka (moraines of the Tampuda, Shegnanda, Kichera, and other rivers; a few dates are presented in Figure 3 and Table 1) and Sartan, younger than 25.88 pm 0.35 ka. The Sartan structures are mainly represented by cirque glaciers that only occasionally advanced into the Baikal shore zone.

As seen from the above data, the most important characteristic of glacial structures of the maximum glaciation is its discontinuous occurrence, weak (by far weaker than at the postglacial stage) differentiation of the underlying surface, and occurrence of the Baikal basin moraines below the present level of the lake. The latter is difficult to explain without the assumption on a lower position of the lake water level at the Samarovo time. However, the drilling results from deep sea holes on the submarine Akademicheskii Ridge [Kuz'min et al., 1997] show that this drop in the lake level may have been insufficient in order that the ridge area including the BDP-96 drillhole site at sea depths of 300-350 m emerged above sea level. The level drop was accompanied by tectonic subsidence of the lake bottom and uplift of its shores [Dem'yanovich et al., 1988; Logachev et al., 1974; Obruchev, 1953; Salop, 1964]. Post-maximum glacial features in upland cis-Baikal areas occur in deep (to 1000 m) valleys cutting the exaration surface of maximum glaciation, which implies a considerable tectonic uplift postdating the maximum. Study of the whole complex of glacial and lake deposits yields evidence of considerable post-glacial reworking of the relief.

Deformations of Late Pleistocene and Holocene Deposits

Morphological variability of active faults and their spatial relationship with sedimentary sequences and relief forms, varying in age and including those of the glacial and postglacial origin, suggest their classification into three age groups [Dem'yanovich et al., 1988].

The first group includes deformations of relief elements related to faults; activation of these faults is dated at the beginning of the Late Pleistocene-Holocene stage. They are most clearly expressed as offsets in end moraine fronts marking a maximum advance of glaciers into the Baikal shore zone (Tyya Promontory) and in glacial deposits of the 80-m terrace [Dem'yanovich et al., 1988]. Because of limited occurrence of mid-Quaternary forms in the contemporaneous relief, recognition of such faults encounters considerable difficulties. Faults active since the late Pleistocene have been discovered in the delta area of the Verkhnyaya Angara River, near the Verkhnyaya Zaimka settlement, and within the Rel-Slyudyanskaya constructional plain. Moreover, echo-sounding survey showed that end-moraine amphitheaters occur at considerable depths all along the Barguzin shore and extend into the lake water area for more than 7 km, implying high activity of tectonic movements in the late Pleistocene [Sizikov and Levi, 1987]. Repeated reworking of existing benches by later movements explains the fact that relief forms associated with the faults of first group have been poorly preserved.

The second age group includes faults that deform terraces of the first Late Pleistocene glaciation; these terraces are widespread along the Baikal shore. Deformations of this group are observed in mouths of the Muzhinai and Molokon Rivers, in the mid-course of the Kichera River, in the Tompuda River mouth, along the southeastern boundaries of the Bolsherechensko- Davshinskaya and Sosnovsko-Tarkulikskaya depressions, in the Snezhnaya River mouth, and so on. The activation of fault motions in the Baikal shore zone was accompanied by reconfiguration of the shoreline and lake transgression that has left signatures in the second (6-8 m) Baikal terrace. Along shores of the Tyya-Goremykskoe plateau and Barguzinskii Range, the back suture of the terrace obliquely cuts the end-moraine amphitheaters and higher levels of fluvioglacial, alluvial, and lacustrine aggradation. At the foot of the Primorskii and Baikalskii Ranges, fault scarps of this age are overlain by piedmont alluvial fans. This age group may also include movements that offset deposits of the Karginsky-Sartan age (50-12 ka), including those dated by the radiocarbon method.

The third age group of discontinuous deformations includes "fresh'' ruptures on the northwestern lake shore (Kurla Head), associated with the epoch of the late Zyryanka (Sartan, 24-12 ka) glaciation dated from archaeological evidence [Endrikhinskii, 1982]. Existence of these movements is debatable in the topography of eastern shore zones, but they did take place in the Selenga River mouth, where a fault-line scarp bounding the contemporaneous river delta on the south began to form at that time. On the whole, the shoreline configuration changed insignificantly, and transgression features of the first (2-4 m) Baikal terrace formed in the Holocene are less pronounced than in terraces of higher levels. An exception is lowland aggradation banks in large river mouths, where active permafrost degradation was in progress at the Pleistocene-Holocene boundary time and large areas of Sartan (24-12 ka) terraces may have been completely reworked by thermal abrasion processes. Probably, some of the thermal abrasion and thermal erosion benches trace fault zones that are thermal water sources. Such a formation mechanism of benches that bound boggy basins may be responsible for the topography features observed at large piedmont promontories of the Baikal Range, in some delta areas of the Verkhnyaya Angara River between the Lake Baikal and Verkhnyaya Zaimka settlement, and on the isthmus of the Svyatoi Nos Peninsula.

The faults that were active in the Late Pleistocene-Holocene have significantly affected the lake shoreline configuration. The latter appears to be controlled by the relative orientation of fault zones and along-shore drift movement, particularly at places where loose Pleistocene deposits are involved in the shore formation process. Thus, Baikal and delta terraces in the Tyya and Slyudyanka mouths and in the Selenga delta are truncated by linear, NE trending benches of tectonic and abrasion origin. All this is evidence of rather intense tectonic processes at the Baikal deglaciation stage of the middle and late Pleistocene, and we cannot exclude the possibility that vertical tectonic movements were affected by the glaciation.

Estimation of the Barguzin Glacier Parameters

In the light of the aforesaid, the thickness of the Barguzin half-covering glacier, which existed in the northern part of the Lake Baikal, is of particular interest. The knowledge of contemporary glaciers provides means for estimating this thickness. As shown below, it is convenient to consider ice as a viscous solid body which flows under the action of gravity [Landau and Lifshits, 1965] but rather strictly preserves the relation between its area dimensions and thickness; under this approximation, contemporary glaciers are described by equations of the type


where Sg is the glacier area (in km 2 ) and Hg is the ice shield thickness (in km), with the correlation coefficient being r2 = 0.625 at the sample size n = 70. Substituting an approximate value of the glacier area (in our case, determined from the end moraine belt of maximum glaciation) into equation (1), we obtain an approximate ice sheet thickness of about 700 m. This value is not much divergent from a geological-geomorphological estimate of about 400 m, obtained from the height of trough valleys [Logachev et al., 1974].

For comparison, we present the estimated sizes of present continental ice covers of Antarctic and Greenland [Dolgushin and Osipova, 1989]. The general area Sg of the Antarctic ice sheet is 13 589 000 km2, and its average thickness Hg is = 2450 m, with a maximum of 4700 m. For the Greenland ice sheet, the respective values are 1 726 400 km2, 1790 m, and 3416 m. Note that if these ice caps were removed from the Earth's surface, the internal sea at the place of Antarctic would have a depth of about 1500 m, and in the case of Greenland the depth would be about 800 m. Although the North Baikal glacier is not so large, it could ignificantly affect the underlying surface.

In our case there is enough evidence to suppose that, during the maximum glaciation period, the Lake Baikal level was lower by 300-400 m. This hypothesis is illustrated in Figure 1a. As seen from the figure, the Lake Baikal consisted of three connected or even isolated reservoirs at a time of 300-400 ka. Then, the water circulation in these reservoirs and therefore landscape environment essentially differed from those presently observed. A lower lake water level in the glaciation epoch is additionally supported by the fact that certain archaeological monuments (e.g. Ulan-Khada) have been destroyed as a result of their slow submersion below the lake level.

Large dimensions and total mass of the ice cover suggest the existence of glacioisostatic movements in the Baikal region. However, in order to verify the postglacial Baikal rift uplift, one should exclude the possible influence of gravity effects on the RVM rates of the Earth's surface. Below, we briefly consider the spatial pattern of both gravity and RVM velocity anomalies, based on geodetic data.

RVM Velocity Anomalies as Constrained by Geodetic Data

The RVM in situ measurements performed in the early 1990s provided a basis for constructing an RVM scheme on the territory of the Baikal rift zone. A rather complicated pattern of RVM anomalies is observed in the area where the ice half-sheet lay. However, to a first approximation, two NE-trending bands of anomalies are recognizable; the anomalous RVM are positive in one of the bands and negative in the other. The first extends from the Svyatoi Nos Peninsula to the Verkhnyaya Angara basin, and the second extends from the Olkhon Island to northern slopes of the Verkhneangarskii Range and bounds the first band to the west. The first, positive-anomaly band is separated into several anomalies. Of those, three anomalies are most pronounced: the northwestern, most intense one is observed above the Verkhneangarskii basin and adjacent mountains of the Delyun-Uranskii and Severo-Muiskii Ranges (maximum rates are +27.4 mm/yr in the Verkhnyaya Angara head area and +16.3 mm/yr at the Yanchuya River source); central anomaly above the Kichera basin ( +8.9 mm/yr in the Kichera River head area); and southwestern anomaly above the Barguzin Range, Svyatoi Nos Peninsula, and southern Baikal Range (rates range from +0.2 to +8.8 mm/yr). Interestingly, the above maximums are confined to the most conspicuous traces of glaciation. The second band of negative anomalies is divided into at least four anomalies the largest of which trends northeast and is contiguous to the largest positive anomaly. This negative anomaly is observed above the eastern Verkhne-Angarskii Range, and its downward velocities reach 14.6 mm/yr in the Konkudera River head area. The rest of the anomalies, bending round the group of positive anomalies on the west, extends southwest through the Synnyr and Yngdar Ranges, crosses the axis of the Predbaikalskii basin, and ends at the Olkhon Island. Here, mean velocities of downward movements amount to -2 mm/yr. Obviously, these anomalies cannot be attributed to glacioisostatic movements, but they imply the presence of a complex mechanism in which both gravity-density inhomogeneities of the lithosphere and tectonic differences between various morphostructural elements may be main disturbing factors.

Many of the gravity anomalies spatially correlate with the anomalies of RVM velocities, and this correlation is either positive or negative. Undoubtedly some of the anomalies coincide with areas of isostatic adjustment, so that postglacial vertical movements are likely to occur in the Baikal region, but their intensity is by far smaller as compared with Fennoscandia and other regions with a similar Pleistocene history of development.

Deep Structure of the Crust and Upper Mantle Under the Baikal Rift

In order to estimate physical properties of the material on which the ice shield lay and which was involved (probably is still involved) in glacioisostatic movements, one should roughly assess the depth at which the ice load removal was accommodated. The accommodation depths of the crust and upper mantle are likely to have been associated with lower strength characteristics and higher ductility. Seismic soundings of the crust and upper mantle in the Baikal region yield evidence of at least three lower-strength layers that change seismic velocities [Artyushkov, 1993]. The shallowest inhomogeneity occurs at depths of 12 to 20 km and nearly coincides with the concentration area of seismic sources in the Baikal region; the next one is observed at depths of 35 to 50 km; and the deepest is the asthenosphere, whose roof occurs, according to various estimates, at depths ranging from 50 to 90 km. Another important fact is that earthquake sources in the Barguzinskii Range area are recorded down to depths of 55 km, whereas their usual depths in Baikal rift crust are as large as 35 km. Which of the inhomogeneities could have accommodated the ice load?

The answer to this question is not trivial, because the linear dimensions and mass of the Barguzin glacier are too small. The presence of abnormally deep (to 55 km) earthquake sources in the area where the ice thickness was highest and where abnormally high RVM rates are presently observed indicates that the subcrustal inhomogeneity was undoubtedly involved in the ice load adjustment process, because otherwise relatively ductile rocks at these depths would have inhibited the growth of seismogenic faults. However, from the standpoint of formal logic, the involvement of the intracrustal inhomogeneity in accommodation of the ice load also raises no doubts.

Thus, in our opinion, the available information is adequate for estimating the following important characteristics:

Model Estimates of Physical Parameters

In order to construct models of the postglacial uplift, one should estimate its characteristic linear dimensions L and height H. For this purpose, we represent the model postglacial uplift as a rectangular parallelepiped whose base area Sg = L2 and height H are such that gg H. According to empirical relation (1), this inequality is valid for all of the presently existing glaciers to a large degree of reliability.

To construct such a model, we consider the lithosphere as a very viscous fluid, and the process itself of the postglacial uplift, as a hydrodynamic attenuation of a strongly elongated disturbance of the length L on a plane boundary of an incompressible fluid of high viscosity h. Then, we obtain the equation


where r is the mean crustal density within the uplift area, g is gravity, and t is the postglacial uplift time in the area considered. In our case we have r = 3.03 g/cm 2, L = 320 km = 3.2 times 107 cm, and t = 30000 years = 9.5 times 1011 s. The substitution of these values into (2) gives h = 1.3 times 1020 poise, which is somewhat smaller than the estimates for Fennoscandia and Canadian shield [Artyushkov, 1993], but consistent with previous estimates for the Baikal region [Levi and Sherman, 1995].

However, formula (2) is not advantageous for estimating the dimensions of the postglacial uplift area L, since it relates two poorly defined quantities which are L and h. Therefore, we try to derive an equation relating L to easily estimated physical characteristics of the lithosphere; for this purpose, the latter will be considered as a high-viscosity solid [Landau and Lifshits, 1965]. Then, the stress tensor of such a body is defined by a general formula


where G is the elastic modulus of the lithosphere, and Uik approx dL/L are strain components. For our estimates, it is sufficient to consider the simplest case of a homogeneous isotropic lithosphere and to replace the hydrodynamic derivative in (3) by the estimate


where we have tapprox L/n for the "hydrodynamic'' time and n is the characteristic rate of uplift. Then, equation (3) for an isotropic stress assumes an essentially simpler form:


In order that the viscosity and elasticity of the body in question might be values of the same order, the following condition must be satisfied:


Returning to the model adopted for the uplift area, the main condition L gg H gives


and hence the stress tensor can be written as


Then, the required dimensions of the uplift area can be estimated from the formula


Equation (9) is self-consistent, because formula (2) can be readily obtained through eliminating G from (5) and (6). Adopting G approx 2.8 times 1011 dyne/cm 2 and rapprox 3.03 g/cm 3 for rocks composing the lithosphere in the Baikal region, we obtain L approx 220 220 km, which agrees with observations. The same equations give an estimate of the characteristic minimum size of a glacier capable of producing glacioisostatic movements. Its characteristic size cannot be smaller than 25 km. On the other hand, the thickness of a very large glacier cannot exceed, on average, 3.5 km, because otherwise it would crush its own base.

To estimate the possible postglacial uplift amplitude, one should know the ice density rg, density of the underlying substrate r, and thickness of the ice shield Hg, which we assume, for simplicity, to be 1000 m. This is a reasonable estimate, because equation (1) gives only an average value of Hg. Then, the equation


gives an uplift height of 200-300 m, which agrees reasonably well with the amplitude estimates of Late Quaternary tectonic movements from North Baikal morphometric data based on the river system incision depth and sedimentation rates [Levi et al., 1981]. The uplift rate may be estimated from the equation


Deglaciation Rates of the Sartan Ice Sheets and Some Problems of Human Paleogeoecology

fig04 Presently, the deglaciation history of the Baikal region cannot be determined due to poor preservation of end moraine ramparts of maximum glaciation, which prevents the delineation of glaciation geographical boundaries. Based on the available geological and geophysical evidence, the time variation of deglaciation over the past 40 thousand of years can be, to an extent, reconstructed. Below, we shortly discuss this problem. Presently, the Kichera River valley (northernmost Lake Baikal area) is the only place where some ten ramparts of end moraines are well preserved. These ramparts are geomorphologically well expressed, and relative age determinations date them at the early to latest Pleistocene [Bazarov, 1986]. In 1989, S. S. Osadchii mapped the entire system of the end moraine ramparts in the Kichera River valley and established their number to be 11-13; he assumed that the lowermost moraine fixes the spatial position of the maximum glaciation boundary, the glaciation age being about 300 ka as mentioned above. In 1994 and 1995, P. G. Kirillov and S. Bak determined an age of 34350 pm 60 fig05 years from the Chalauta moraine (Table 1), which dramatically diverges from the relative geochronology data reported by previous researchers. The scheme of Figure 4 (constructed after Osadchii [1989]) clearly shows considerable variation in the spacing between the moraine ramparts, implying that the glacier retreat rate correlated with climatic variations in the air temperature. In order to estimate the relative age of each moraine rampart, one should admit that the temperature rose uniformly. Given the validity of this assumption, each moraine may be dated from spacings between moraine ramparts. These simple considerations showed that the deglaciation rate varied with time (Figure 5), indirectly reflecting the climatic variations in the air temperature.

fig06 Now we briefly consider the intensity of soil formation in East Siberia during the past 50 k.y. The plot shown in Figure 6 was constructed using a collection of 14C datings (Appendix , Table 2). Furthermore, we assumed that, in the case of global development of a phenomenon, the number of datings that fall into a given "time window'' is maximum, whereas it is minimum if the process is depressed. The time window was taken to be wider than the standard deviation of the 14C determinations, and the windows overlapped, with a time step being 0.5. Simple calculations yielded a frequency curve, indirectly reflecting the climatic temperature variations in East Siberia; on the whole, this curve is consistent with the current geological notions (Figure 6).

fig07 Archaeological site datings in East Siberia (Appendix , Table 3) were processed in a similar way, and we obtained the site recurrence plot (which may show how often ancient peoples visited the Baikal region) illustrated in Figure 7.

Comparison between Figures 6 and 7 reveals an abnormally high recurrence of archaeological site datings during the Sartan glaciation in the Baikal region. The Holocene extremums may be explained by the general rise in temperature favorable for developing new lands by men. On the other hand, it is difficult to explain the frequent site occurrence during the Sartan time, when a rather cold climate dominated the Baikal region, and the ancient man was strongly dependent on environmental conditions. However, comparison of Figures 5 and 7 seems to resolve this contradiction. High deglaciation rates of the Sartan glaciers were characteristic of the time interval during which the Sartan archaeological sites were widespread. Thus, these facts, albeit indirect and somewhat insufficiently substantiated, yield evidence of anomalous climatic conditions that existed in the Baikal region over the past 50 k.y. We hope that joint geological and geophysical studies that have been conducted since 1997 will provide final verification of these climatic features.

Future Development of the Postglacial Uplift Problem

The next stage in the elaboration of the problem of postglacial uplift in the Baikal rift zone consists in the intensity estimation of postglacial tectonic events as a function of the distance to the local center of the North Baikal glaciation. Deglaciation features in the region should correlate with stages of the Lake Baikal water level variation; thereby our knowledge of the mechanism responsible for the lake water discharge through the Angara River will be improved and many questions concerning the Pleistocene evolution of the region and regional climatic changes will be answered. Supposedly, the interpretation of terrace-like benches and terraces of the Ushkun'i Islands [Lamakin, 1952link13] can be beneficial to the solution of these and many other problems of the postglacial Baikal history.


Numerous traces of neotectonic movements and postglacial active tectonics have been reliably established in the Baikal rift zone. Substantial relief transformations postdating the maximum glaciation epoch are evidence of considerable tectonic uplift events with a highly probable postglacial component, which is additionally supported by relationships of the Bikal terraces of intermediate (35-50 and 80 m) and high (150 m) levels with glacial features. Subsynchronism of lake and glacial deposits and abrasion planation of the glacial relief at various levels can be most reasonably explained in terms of the glacier slide into the lake shore zone and, at the initial stage of deglaciation, abrasion truncation of the glacial relief which did not experience significant glacioisostatic uplift at that time. The glacioisostatic hypothesis is consistent with theoretical estimates obtained in this work.


We are grateful to Academician N. A. Logachev for a preliminary discussion and valuable comments. This work was supported by the International Science Foundation, grant No. RLG300; the International Association for the Promotion of Cooperation with Scientists from the Independent States of the Former Soviet Union, grant No. 93-014; the Russian Foundation for Basic Research, grant Nos. 95-05-14211, 96-05-64187, and 97-05-96529; and the Siberian Division of the Russian Academy of Sciences, grant No. IGSORAN-97-22.


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