Regime and depth of seasonal freezing and thawing of the ground

The classification of the types of seasonal freezing and thawing elaborated by Kudryavtsev takes account of the influence of the main geological-geographical factors, both individual and in their combination, on the processes of freezing and thawing as well as on the thickness of the seasonally frozen and seasonally thawed layers. One should bear in mind the mutual relationships of all elements of the environment. Thus, for example, variation of thickness or changes of species of vegetation cover will lead to changes of A0 and fmean parameters and also soil humidity and, perhaps, composition which, in the long run, will manifest itself in change of the depth of seasonal freezing (thawing). Thus, in order to evaluate the effect of one or another natural factor on the depth £ it is necessary to determine its influence on each of the classification parameters, i.e. A0, imean, A, C, <2ph. It should be especially emphasized that since the parameters imean and A0 used in the classification refer respectively to the bottom and top of the layer <!;, their determination will require quantitative data on the influence of surface cover (snow, vegetation, water, ice, etc.), i.e. one should properly treat mean annual temperatures and amplitudes of air temperature fluctuations to gain knowledge of fmean of the ground and A0 over the soil surface beneath the cover. Let us consider the effect of some of the most important factors of geological-geographical environment on fmean and <!;.

The effect of topography, aspect and steepness of slopes

The position of the site in the topography to a great extent defines the temperature regime of the ground and depth of seasonal freezing (thawing). Thus, as air temperature diminishes with altitude by about 0.4 —0.6 °C per 100 m, accordingly, the mean annual temperature of the ground is also reduced. This leads to reduced thickness of the seasonally thawing layer and increased thickness of a seasonally frozen layer. Composition of the soils (first of all grain size) varies with altitude as do soil humidity, thickness of snow cover, persistence and species of vegetation, etc.; i.e. there is variation of all the classification indices of the seasonal freezing (thawing) of the soils.

A substantial influence over fmean and £ is exercised by aspect (orientation of slopes with respect to cardinal points) and steepness of slopes (angle of incidence of solar rays on slopes of different steepness). Mainly, these are effective through the amount of inflowing solar energy absorbed by the surface. The values imean and A0 diminish (with other conditions being equal) from southern and south-western slopes towards north-eastern and northern ones. The difference between fmean and A0 on slopes of southern and northern aspects is mainly determined by the difference in the summer air temperature, since in winter with small inflow of solar energy southern and northern slopes are almost equally cold (provided there is the same snow and vegetation cover). In summer the southern slopes receive much more solar energy. Because of this the depth of thawing £"ha on the slopes of northern aspect is much less than that on slopes of southern aspect, i.e. always ¿;"ha < ^ha» with the exception of the regions of high (polar) latitudes, where f"ha « ^ha, since there is no solar radiation in winter, and in summer the Sun heats the slopes of all aspects more or less uniformly.

In the case of seasonal freezing, the thickness of the seasonally frozen layer is formed mainly on account of the heat cycle in the winter period, when slopes of different aspects are cooled practically in the same manner, as mentioned above. Therefore, the difference between depths of seasonal freezing on slopes of both southern and northern aspects will be insignificant (£fr ~ £fr)

The effect of slope steepness over the temperature regime and depth of seasonal thawing (freezing) follows from the different angle of incidence of solar rays and of shadowing, i.e. the different amount of the radiation absorbed by the surface of slopes. Higher temperatures of the ground in the summer period are typical of the slopes perpendicular to solar rays (about 30°) which promotes greater depth of seasonal thawing, while the depth of seasonal freezing remains similar to those on gently sloping hillsides, as there are small differences in their winter cooling. One should bear in mind that the effect of slope steepness (and aspect as well) on temperature regime may be complicated by other factors, for example, uneven distribution of snow cover, differences in vegetation cover, etc., which makes it difficult to distinguish the factors under consideration.

The geographical location of the site and, first of all, its location with respect to oceans is expressed through convective heat exchange of the atmosphere with the lithosphere, forming meridional sectors of climatic continentality. The maximum mean monthly temperature is higher and the minimum is lower with distance from the coast into the continent, i.e. the annual amplitude of monthly temperatures increases. Therefore, with the same mean annual temperature of ground surface in continental regions seasonal freezing and thawing are deeper, as is the penetration of annual temperature fluctuations. However, as a rule, more continental climate involves more severe permafrost conditions, with the lower mean annual temperatures of air and ground. Consequently, the outlines of the permafrost-temperature zones reflect both zonal and sectorial changes of heat-and water-exchange between soil and atmosphere. The southern boundary of the permafrost and, therefore, seasonal thawing, on the Kola Peninsula is confined to zones of tundra, in the north-east of the European part of Russia to the subzone of the northern taiga, in West Siberia to the southern taiga, and in the Zabaykal'ye and Mongolia to permafrost islands, with seasonal thawing encountered in steppes and even semi-deserts.

The influence of snow cover

Snow cover causes substantial changes of the heat exchange between ground and atmosphere. First, the albedo of snow is much higher than that of a bare surface or of vegetation. This leads to reduced absorption of solar energy and lowering of snow surface temperature relative to air temperature. According to meteorological data the mean winter temperature of a snow surface may be 0.5-2 °C lower than the mean winter temperature of the air.

At the same time, snow cover having a low thermal conductivity (varying from 0.12 to 0.46 Wm"1 K"1 and 5-10 times lower than that of mineral ground) prevents loss of heat by the ground in winter. The soil surface therefore, beneath the snow cover, can have a much higher temperature than that of the air. On the average, increase of snow thickness by 5-15 cm leads to a 1 °C increase of mean annual temperature of the ground. Therefore, with sufficient thickness of snow cover the mean annual temperature of the soil surface can be positive at low mean annual temperatures (to — 6 to — 8°C) of the air.

If the snow remains on the soil surface after the air temperatures become positive, it prevents heating of the ground as a considerable portion of the inflowing solar energy, first, is reflected and, secondly, is consumed for snow melting. Melting snow keeps a temperature of zero at the ground surface despite the positive air temperature at this time. This leads to a certain cooling of the ground and lowering of its mean annual temperature.

The influence of snow cover on the temperature regime of the ground is many-faceted. Both the value of the effect and its vector (heating or cooling) are dependent on thickness of the snow cover. Thus, its cooling effect prevails at small thickness of snow due to higher albedo. Then there is the warming effect of snow as a heat insulator. Increased thickness of snow cover (to a certain critical value) leads to an increased cooling effect due to slower melting of snow in summer. For the greater part of the seasonally frozen and permafrost areas the thickness of snow cover is such as to have a warming effect on the underlying ground. Snow cover not only causes an increase of mean annual temperature of the ground surface, but also leads to much reduced amplitude of the soil surface temperatures as compared with that of air temperatures and in some cases may cause changes in soil humidity.

Snow density is an important factor associated with the warming effect of snow. Thus, with snow cover density p equalling 75 kg m ~3 thermal diffusiv-ity K will be 0.36 x 10~3 m2 h_1; with p = 150kgm~3,K =0.72 10"3m2 h"1; with p = 225kgm-3, K = 1.08 x 10-3 m2 h"1; p = 300kgm"3, K = 1.44 x 10~3 m2 h_1; with p = 380kgm"3, K = 1.8 x 10~3 m2 h"1. Loose snow has a greater warming effect on the temperature regime of the ground compared to dense snow due to its low thermal diffusivity and thermal conductivity.

V.A. Kudryavtsev found the relationship between snow cover and values of heat cycles - the amounts of heat going through the soil surface for half-periods of heating and cooling. The bigger the heat cycles of the soil (with other conditions being equal) the more intense is the influence of the snow cover on mean annual temperature and amplitude of surface temperatures. Taking into consideration the patterns of heat cycle variations as determined by geological-geographical factors, one may conclude that the maximum effect of snow cover occurs at a mean annual ground temperature near to 0°C (i.e. in the vicinity of the southern boundary of perennially frozen ground) with a maximally continental climate and very wet soils in the seasonally thawing and seasonally freezing layers.

Qualitatively, the effect of snow cover on the depth of seasonal thawing and freezing can be analyzed with the help of graphs as plotted in Fig. 11.7. As snow is a thermal insulating layer, after its removal the amplitude A0 of surface temperature fluctuation increases and mean annual temperature usually diminishes (with exception of the cases when snow has a cooling effect). Higher A0 always causes increase of depth of seasonal thawing £tha and seasonal freezing £fr. Quite different is the effect of tmean: in the area of perennially frozen ground its lowering leads to lower £tha, while in the area without permafrost (unfrozen or thawed ground) its lowering leads to greater £fr. Thus, in the first case, variation of A0 and imean after removal of snow (reduced thickness) is compensated by their different direction and has an insignificant effect at the depth of £tha (see Fig. 11.7a), whereas in the second case, the effects are additive and the depth £fr changes substantially (see Fig. 11.7b).

The influence of vegetation cover

There are several aspects to this effect on the temperature regime of ground and depth of seasonal freezing and thawing. Vegetation cover causes a change of reflective capacity compared with that of the underlying surface,

Fig. 11.7. Diagram of the effect of snow cover on depths of seasonal thawing (a) and freezing (b) of ground according to S.Yu. Parmuzin):fmeani, A0i, ^, -respectively, mean annual temperature of the ground, amplitude of annual temperature fluctuations at soil surface, depth of seasonal thawing or freezing, depth of penetration of annual temperature fluctuations, with snow cover;fmean2, A02, m2,h2- the same characteristics after removal of snow cover; A£tha, A£fr -variation of depths of seasonal thawing and freezing.

Fig. 11.7. Diagram of the effect of snow cover on depths of seasonal thawing (a) and freezing (b) of ground according to S.Yu. Parmuzin):fmeani, A0i, ^, -respectively, mean annual temperature of the ground, amplitude of annual temperature fluctuations at soil surface, depth of seasonal thawing or freezing, depth of penetration of annual temperature fluctuations, with snow cover;fmean2, A02, m2,h2- the same characteristics after removal of snow cover; A£tha, A£fr -variation of depths of seasonal thawing and freezing.

absorbs solar energy, causes evaporation of moisture from its whole volume, makes air flow turbulent above the level of biomass development or, on the contrary, causes stagnation of air in it. Conditions of snow accumulation and properties of the snow layer are to a great extent determined by vegetation cover which also has an influence on moisture content and thermal-physical properties of the soil.

Compared with the effect of snow, it is much more difficult to evaluate qualitatively the effect of vegetation cover (as a thermal insulator) on temperature regime and depth of seasonal freezing or seasonal thawing of the ground. This is explained by the fact that vegetation cover insulates the soil both from cooling in winter (as snow) and from heating for the whole summer period. The double effect of these two influences is dependent on duration of summer and winter seasons, continentality of climate, depth of snow cover, moisture content of the underlying soil, etc., i.e. on a number of factors and conditions that determine the role of vegetation cover in heat exchange between the soil surface and atmosphere, on the one hand, and between soil surface and underlying soil and rocks, on the other.

As a first approximation it can be concluded that in the area of perennially frozen ground the influence of vegetation cover on the seasonal thawing depth is greater than that on the depth of freezing in the area of unfrozen or thawed ground (Fig. 11.8). In both cases removal of vegetation cover leads to higher amplitude of annual temperature fluctuations with both lower imin and higher imax (unlike snow that influences only imin). Variations of imean are dependent on the new values of imin and imax. If Aimax > | Aimin | it is apparent that imean increases. In the regions of seasonal thawing (see Fig. 11.8a) both higher A0 and imean promote greater values of ctha (i.e. the effects are additive). In the area of seasonal freezing (see Fig. 11.8b) greater A0 leads to greater gfr while an increase in imean would compensate this effect. As a result, gfr increases insignificantly which is well seen in Fig. 11.8. It is known, however, that in the northern regions of the permafrost zone where the thickness of the snow cover is not large (0.20.3 m) vegetation cover has a warming effect, i.e. it leads to higher mean annual temperature of the ground compared with the regions having no vegetation cover. But this does not lead to deeper thawing as the amplitude of the temperature fluctuation under the vegetation cover is always less than when it is absent.

Forest and shrub vegetation reduces the inflow of solar energy to the surface due to the effects of shadowing, which lead to less warming of the surface in summer compared to bare sites and slower melting of snow. According to the observations on thermal balance obtained by A.V. Pavlov near the town of Yakutsk, Igarka and the village of Syrdakh, the albedo of the forest on the permafrost is less than that of bare sites; the effective radiation balances of forested and unforested regions do not differ substantially year by year and the annual sum of the radiation balance for the forest exceeds that for unforested sites.

The influence of the forest vegetation on temperature regime of the ground is closely connected with the geobotanical zones. The larger the phytomass surface of the forest depending on height, density and closeness of its layers, the less the solar rays penetrate the soil surface. Consequently, with more close crowns in the direction from north southwards the role of forests in the formation of mean annual temperature of the ground changes substantially. In the open woodland of the forest-tundra area and the light forest and shrubs of the northern taiga zone, reduced inflow of radiation at

b

Fig. 11.8. Diagram showing the influence of vegetation cover on depth of seasonal thawing (a) and freezing (b) of ground (according to S.Yu. Parmuzin):imean , A0i, hl - respectively, mean annual temperature, amplitude of annual temperature fluctuations at the surface, depth of seasonal thawing or freezing, depth of penetration of annual temperature fluctuations, with vegetation cover; t , An , iru, lu - the same characteristics after removal of vegetation

cover; A£tha, A£fr - variations of depth of seasonal thawing and freezing.

Fig. 11.8. Diagram showing the influence of vegetation cover on depth of seasonal thawing (a) and freezing (b) of ground (according to S.Yu. Parmuzin):imean , A0i, hl - respectively, mean annual temperature, amplitude of annual temperature fluctuations at the surface, depth of seasonal thawing or freezing, depth of penetration of annual temperature fluctuations, with vegetation cover; t , An , iru, lu - the same characteristics after removal of vegetation

cover; A£tha, A£fr - variations of depth of seasonal thawing and freezing.

the soil surface is compensated by reduced turbulent heat exchange, and under conditions of strong winds more loose and thick snow cover is accumulated compared to woodless regions. As a result, the mean annual temperature of the ground in the northern forest exceeds that of woodless sites.

With closer crowns the incoming radiation is so much reduced that diminished turbulent exchange cannot compensate for it. With weak winds typical of the forested zone (especially in dense, dark coniferous forest) the depth of snow cover is much less than in treeless sites. Therefore, in the central and southern taiga zone and in the area of unfrozen ground in the south of the country, forest serves as a cooling factor. This is confirmed by regular observations in the area of the town of Zagorsk where the mean annual temperature of the soil surface in the coniferous forest is lower by 2°C than on bare sites. In Central Yakutia the mean annual temperatures of the ground under forest differs from those on bare sites by 1-2 °C. In West Siberia, in the vicinity of the southern boundary of the permafrost zone, islands of frozen ground are confined to mixed and dark coniferous forest with crown closeness of 0.7-0.8.

Grass cover to a lesser degree causes changes in the heat exchange and temperature regime of soil surface and atmosphere. The total thermal effect of grassy vegetation on imean of the ground can be either warming or cooling, but it does not exceed fractions of a degree. The amplitude of mean monthly temperatures is also reduced insignificantly. Of importance is the influence of soil covers (mossy, moss-lichen, lichen, mossy-peaty) which are thermal insulators, preventing heating of the soil in summer and reducing heat yield from the surface in winter.

A distinctive feature of humid, natural soil covers is a considerable change of their thermal conductivity while passing from the thawed state into the frozen. According to data from field observations thermal conductivity of moss-lichen covers in the thawed state is 0.1-0.7 W K-1, one third to one half of that in the frozen state. Therefore, the ability of moss covers to retard the entry of heat in the summer period is greater than its ability to restrict the yield of heat in winter by the same order. Thus, the layer of moss 2-3 cm thick reduces the sum of summer temperatures by two-thirds and more. In winter the influence of moss cover on temperature regime of the ground is much less, because of the drastic increase of thermal conductivity. On the average, in winter, temperature beneath moss differs only slightly from that of its surface. Thicker and less water-saturated mossy cover has a greater effect on fmean of the ground. Depending on difference of thermal conductivity of moss cover in thawed and frozen states, duration of summer and winter seasons, thickness of snow, etc., moss covers can have either a warming or cooling effect. One should bear in mind that an increase of mean winter (and minimum) and lowering of mean summer (and maximum) temperature of the soil surface resulting from the heat-insulating effect of mossy covers during a year leads to a sharp reduction of temperature amplitude. Moss covers 15-20 cm thick cause a reduction of temperature amplitude by 5-6 °C and lead to shallower depths, one-half to one-quarter those for seasonal thawing under a bare surface.

The influence of peat cover on the temperature regime of the ground should be dealt with separately. As shown by investigations, mean annual temperature of peat soil is lower than that of mineral soil. In the vicinity of the southern boundary of the permafrost zone, peat cover 0.1 m thick causes reduction of imean by 0.5-1 °C. Therefore, even with a positive mean annual temperature at the peat surface the underlying ground can be in the frozen state. On the West Siberian and north European plains the southernmost islands of permafrost are, as a rule, confined to peats. Thickness of the frozen strata in the south of the permafrost zone is greater in peat soil than mineral ground. The thermal conductivity coefficient of peat in the thawed state varies within the range of 0.23 to 0.93 W m"1 K"1 and in the frozen state between 0.93 and 1.28 W m-1 K"1. With higher moisture content of the peat, the difference of thermal conductivities of the frozen and thawed peat increases, the ability to prevent the underlying ground from heating in summer becomes greater than the ability to prevent the yield of heat in winter, and the cooling effect increases.

There are a number of calculation schemes and equations for the tentative quantitative estimation of the thermal influence of soil covers (mossy, peaty, sod, snow and others both natural and artificial) that serve as additional heat-insulating layers on the surface of the ground. For example, there is one enabling determination of mean annual air temperature variation Afmeana.r and of reduction of annual air temperature fluctuation AAail owing to different types of covers that exist during warm ts or cold tw periods or for the whole year T. It is assumed that phase transitions in covers do not occur. In the calculation, the equation of annual sinusoidal (or reduced to sinusoidal) variation of the cover surface temperature (with period T, mean annual temperature fmean and amplitude A0) is subdivided into two conditional harmonic fluctuations of temperature taking place near 0°C with periods 2ts and 2tw and amplitudes of summer and winter temperature fluctuation equalling, respectively, Aail and Aail (Fig. 11.9). Using the Fourier equation one can determine the reduction of the amplitudes A/ls and AAW of the temperature fluctuation for these two conventional sine waves, due to soil cover, respectively for the warm and cold periods of the year:

where z is the thickness of cover; Kth and K[r are the thermal conductivities of the soil cover under consideration, respectively in the thawed and frozen state; ts and tw are the duration of positive (warm or summer period) and negative (cold or winter period) temperatures of air. While calculating Aair and Aail the sign of the value fmean . is to be borne in mind.

Having knowledge of A^s and AAw it is simple to determine the true

Boulon Griffe Appareil Pression
Fig. 11.9. Diagram of temperature variation at the surface of ground cover broken down into two simple harmonic fluctuations with periods equalling the double duration of cold (2tw) and warm (2rs) periods of a year.

reduced value of annual amplitude AAair and the reduction and variation of mean annual air temperature Aimean . caused by soil cover by using the values of area AS, and AS'.

The relationships suggested for the calculation of A^air and Aimeanajr show that soil covers always cause reduction of the amplitude of temperature fluctuation, whereas variation of mean annual temperature Amean . can be both positive (cooling effect) and negative (warming effect) depending on which effect prevails: cooling in summer (A^sxs) or warming in winter (AAwtw). If the cover exists only part of the year (during the warm or cold period), for example, the snow cover period of the year, the equations (11.4) and (11.5) are simplified as the term AAsxs becomes equal to 0.

By using relationships (11.4) and (11.5) one may transfer from mean annual temperatures imean . and amplitudes of temperature fluctuation Aair of the air (measured at weather stations at an elevation of 2 m) to the values of tmean and A existing at the ground surface. With this aim in view one should determine the difference in mean monthly (or weekly) temperatures at the height of 2 m and at the ground surface during the warmest month of summer and coldest month of winter, i.e. experimentally find the values of and AAw and, substituting the values xs and xw into the equations (11.4) and (11.5), one may calculate the values A^air and Atmean . at the 2 m interval above the ground surface. In the same manner account should be taken of the influence of slope aspect and steepness on the temperature regime of the surface.

The influence of swamps and surface water

The influence of waterlogging on the temperature regime of ground is to a great extent determined by the general climatic setting and stage of development of swamps. So long as the surface of swamps is partially covered with water, which transmits shortwave radiation and retains longwave radiation, the mean annual temperature is higher than that of the adjacent sites.

In the course of bog evolution, when there is overgrowth by mosses and heaving of the freezing ground, individual sites are no longer covered with water and sedge-carex species are replaced by mosses and shrubs. The warming effect of snow on these sites is reduced due to the diminished thickness of the snow cover and lower moisture content of the peaty deposits. Mean annual temperatures of the ground of elevated sites are reduced compared to those of low-lying sites. Further drainage of bogs as a result of the surface uplift, leads to gradual dying off of sphagnum and replacement by lichen. Mean annual temperatures of the ground at these sites are, as a rule, much lower than in the surrounding terrain. In the vicinity of the southern boundary of the permafrost zone hillocky peats indicate the occurrence of perennially frozen ground.

Thus, depending on the stage of development the bogs can have either a warming or cooling effect on the permafrost and unfrozen ground.

The temperature regime of non-draining fresh-water bodies is dependent on their depth. If the depth of a lake Hw is greater than the maximum thickness of ice H{, which, in the most severe conditions, does not exceed 2-2.5 m, then bottom deposits are unfrozen year round. Depending on the size of the lake and the mean annual temperature of the ground on adjacent sites either a through, open talik (if the width of lake exceeds double the thickness of the frozen strata) or a closed talik is formed beneath the water body.

Fig. 11.10. Diagram of distribution of minimum, maximum and mean annual temperatures in a water body.

A water body having a smaller depth than the thickness of ice possible in the given region freezes to the bottom and bottom deposits can have either positive or negative mean annual temperatures. There is a depth of water body (Fig. 11.10) at which the mean annual temperature of bottom deposits will be equal to 0°C. This depth is called by Kudryavtsev a critical value Hcr. With a water body of depth less than the critical (iiw < HCI) and the mean annual temperature of bottom deposits below 0°C, there is permafrost and only in summer do the deposits thaw to a certain depth, i.e. seasonal thawing of the bottom deposits is observed. With depths of water bodies ranging from the critical HCI to that equalling maximum thickness of ice Hv i.e. HCI < < Hb mean annual temperature of the surface of bottom deposits will be positive, but in the winter period they will freeze to a certain depth. In this case seasonal freezing of bottom deposits will be observed. The critical depth of water body is mainly determined by climatic characteristics (air temperature and thickness of snow cover). Therefore, the characteristics of the climatic zones determine the critical depths of water bodies. Thus, according to data obtained in West Siberia Hcr is 0.2-0.3 m in the vicinity of the southern boundary of the permafrost. Northwards Hcr increases steadily reaching 1.6 m in the Yamal and Guydan peninsulas.

If the depth of water body is assumed to be comparable with the maximum thickness of ice H{, then the temperature regime of bottom deposits can be determined by using Kudryavtsev's scheme (see Fig. 11.10). In this figure imin and imean designate distribution of respectively, minimum (winter) temperature in the ice cover and the mean annual temperature, and imax is the maximum temperature of the water body in summer, which tentatively is assumed to be uniform for shallow northern lakes (owing to convection mixing). Stemming from this diagram the mean annual temperature at the surface of the bottom deposits (at the depth of the water body (i/w)) is imea„ = l/2[fm„ + (1 - HJH.^J (11.6)

The temperature regime of the bottom deposits of saline lakes is different as saline water, being heavier, moves downwards and cools without freezing at negative temperatures. Even in summer the temperature in the bottom of highly saline lakes can be negative. As a result, perennially frozen or cryotic bottom deposits are observed beneath such saline lakes as well as on shallow places off the northern seas coasts, while saline water in the liquid state, having a negative temperature, overlies such deposits.

The influence of convection flows of water and air

The temperature of ground can vary not only by conductive transfer of heat but also by convective transfer by seepage of water or air flows. The inflow of warm or cold air or water flows in the ground can lead to warming or cooling not only due to balancing of heat content between the convection flows and the soil or rock but also through heat released at phase transitions of moisture (freezing - thawing, evaporation - condensation, sublimation -ablimation).

Under natural conditions transfer of heat into ground occurs by infiltration of surface water and, primarily, of precipitation. Intensity of the process is dependent on the amount of infiltrating rainfall, and the temperature, thickness, seepage and thermal-physical properties of the layer of seasonal freezing and thawing. A tentative equation was suggested by Kudryavtsev for quantitative estimation of the increase of mean annual temperature of ground Atmean at the base of the seasonal freezing (thawing) layer as a result of infiltration of summer precipitation (17):

where Fis the amount of summer precipitation infiltrating the soil, kgm"2; tms is mean summer temperature,°C; £ is depth of seasonal freezing or thawing, m; T is time (year = 8760 h); Ar is reduced thermal conductivity equalling the average values (weighted for the year) in the frozen and thawed states, kJ (mh°C)_1.

According to the calculation data infiltration of the summer precipitation may cause an increase of mean annual temperature of the ground of 1.5-2°C. This effect will be greatest on sites composed of coarse-grained soil with high hydraulic conductivities. The presence of vegetation cover drastically reduces the infiltration. As a rule, the role of infiltration in the formation of mean annual temperature of the ground at these sites is insignificant and does not exceed 0.1 °C.

An important role in the formation of mean annual temperatures of porous loose and fractured bedrock is played by convection flows of air. In

At, mean mean

such materials there is a permanent exchange of gases with the atmosphere caused by variations of pressure and air temperatures near the soil surface. The process takes place in the following way: cold atmospheric air replaces warmer and lighter air in the cavities in the material and cools the latter. Such winter ventilation is clearly traced in boreholes and pits and occurs in porous rocks, cooling them intensely for a considerable depth.

In the regions with extensive rudaceous deposits, an additional warming factor is the condensation of water vapour. Thus, in the southern part of the permafrost zone the influence of vapour condensation can increase imean by as much as 2°C, in the northern part much less, while it is not observed on high interfluves and in the Arctic zone.

The temperature regime of the ground and the depths of seasonal freezing and thawing vary substantially under conditions of economic development. In large cities specific micro-climate is created with changes in air temperature regime, direction and velocities of wind, evaporation, etc. In the course of development there are substantial changes in vegetation cover, conditions of snow accumulation, drainage or waterlogging of the terrain surface, and artificial water bodies are created. The temperature regime of the ground is much affected by heat issuing from engineering structures. The accelerating tempo of development of the northern and Far East regions leads to deep changes of temperature regime and depths of seasonal freezing and thawing. There is accordingly a requirement for scientifically substantiated prediction of temperature regime and of changes in depths of seasonal freezing and thawing, with a possible need to prepare designs for the purpose-oriented alteration of these characteristics, i.e. management of the processes of seasonal freezing and thawing. Such tasks can be fulfilled only on the base of a thorough knowledge of the processes of seasonal freezing and thawing.

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