Water transfer and ice formation in freezing and thawing soils

There are basically two different ways that soils freeze: with and without water migration. Soil freezing without water migration occurs either in the case of soils of low moisture content or a sufficiently rapid advance of the freezing boundary. For example, when a soil sample is frozen through quickly at temperatures of — 60°C, — 70°C, water freezes in situ, since the temperature field is by a factor of 10 more active than the moisture content field. Usually freezing in nature proceeds slowly enough for water to migrate. Then the frozen part of the freezing soil displays a certain pattern, as described, of moisture migration and ice formation brought about by the temperature gradient.

Soil freezing (or thawing) results in a sharp disturbance of any fully formed, thermodynamic equilibrium system and is seen as a dynamic coexistence of frozen, freezing and unfrozen zones, and with the development of a movable boundary dividing the phases, i.e. a front of freezing (thawing). Note that it is the frozen (not unfrozen) part of freezing and thawing soils that causes and determines the moisture migration. This lies in the fact that the existence and development of the gradient of freezing temperature in the frozen zone results inevitably and naturally in the development of the considerable gradient of thermodynamic moisture potential and the gradient of partial pressure of water vapour (grad /.t0) and grad P), and consequently, the gradients of unfrozen water and vapour content (grad Wunf and grad dsU). The driving forces of moisture migration present in the frozen part of freezing (or thawing) soils cause the advance of liquid and vapour in the direction from the higher moisture potential (or moisture content) towards the lower, i.e. from regions of higher to those of lower freezing temperatures.

A moisture deficiency , arising thus in the high-temperature part of the frozen zone of a sample, will be replenished by migration out of the unfrozen part of a freezing or thawing soil. This proves more profitable energetically since the water here is less bound and more mobile than that in the frozen part. This in turn will result in the formation of gradients of thermodynamic potential of the moisture and of water content in the unfrozen part of the soil, which gradients in turn provide the frozen part with the necessary (for the frozen part) amount of liquid and vapour. The unfrozen part of the soil is thus a sort of'reservoir' or moisture source for the frozen part. The temperature gradients in the unfrozen part do not lead to moisture migration driving forces. The explanation is that thermodiffusiion moisture transfer in unfrozen soils only becomes perceptible and is noted in tests where the temperature gradients are greater than 2°C cm"1 to 4°C cm" In natural conditions grad t in the unfrozen zone of freezing (thawing) soils proves to be smaller by an order of magnitude.

At the freezing (thawing) boundary, i.e. in the transition from the unfrozen part of the soil to the frozen part (according to N.A. Puzakov), the continuity principle of water flow Im and, thus, of bound water films is to be obeyed. Experiments verify this and show the continuous nature of the distribution with depth of the major moisture transfer parameters, i.e. /.ijx), KJx), a(0(x), W(x) and IJx). The values of temperature l* and of water content JV( at the freezing-thawing interface are functions of the process and can be determined from combining equations for simultaneous heat and moisture transfer in the frozen and the unfrozen parts of soil. In the general case, with a decreasing rate of freezing of the soil u, value t* goes up due to thermal inertia and, on the contrary, when the rate increases, the value goes down (Fig. 2.11). Soil moisture content, W{, at the freezing front behaves in the opposite fashion.

Experiments prove that the origin and growth of ice layers occurs not on the very boundary of freezing (or thawing) but within the already frozen part of the soil and will be predetermined by both thermal-physical and physical-mechanical conditions of the soil system. Fig. 2.11 shows that the most intensive ice formation in the frozen part is noted on the sections of the sharpest inflexion of curves ¡.im = fix) and Wuni = fix), since this makes for a sharp change in the driving forces and thus in the intensity of moisture flow The curves of moisture distribution with depth demonstrate this all clearly enough (see Fig. 2.11b); visual observations of the freezing of soil samples of various composition and structure provide confirmation. Fig. 2.12 shows how a segregated ice layer starts growing and increasing in size during the freezing of kaolinite clay. Three distinct soil sections are conspicuous: I, the frozen section with an earlier formed schlieren cryogenic structure in which practically no formation of segregated ice occurs at

Digit Range Motion Measurments

Fig. 2.11. Change with depth and time in sample of freezing kaolinite clay: a - of freezing front £fr, of segregated ice formation £si and of temperature at these boundaries (i£fr and i£si); b - of the total moisture content Wlol and of unfrozen water content fVm(, at two times indicated by ^ and with the formation of ice schlieren (L1 and L2 respectively) on account of moisture migration for curves (VM, — AIVi and iVun 2 — AIV2 from the unfrozen zone.

Fig. 2.11. Change with depth and time in sample of freezing kaolinite clay: a - of freezing front £fr, of segregated ice formation £si and of temperature at these boundaries (i£fr and i£si); b - of the total moisture content Wlol and of unfrozen water content fVm(, at two times indicated by ^ and with the formation of ice schlieren (L1 and L2 respectively) on account of moisture migration for curves (VM, — AIVi and iVun 2 — AIV2 from the unfrozen zone.

present; II, the freezing section of intensive phase changes with massive or micro-schlieren structures where the initiation and development of ice schlieren occur; III, the unfrozen soil which is losing moisture. Sections / and II are light-coloured, while III is dark. This is due to the soil state, i.e. it is frozen in / and II, unfrozen in III. The presence of ice in section II is verified not only by temperature readings and movies and photographs, but also by microscopic investigations and special experiments with fluorescein.

Samples of montmorillonite clay, silty-clay and clayey, silt- rich sands produce similar results. They differ only in the intensity of ice formation in the frozen zone, the type of cryogenic structures formed and in some quantitative freezing indices. In all cases it was observed that with a reduction of the rate of freezing and with the freezing front subsequently stationary, the thickness of the actively freezing zone (zone of intensive phase changes) was reduced and the boundary of visible segregated ice formation approaches the freezing front and then merges with it (Fig. 2.11a).

With a linear pattern of temperature distribution in the frozen part of an

Unfrozen Water Fraction
Fig. 2.12. Origin and growth of streaks of segregated ice (schlieren) in freezing clays, kaolinite (a) and montmorillonite (b), at various times (t); I, II, III ~ frozen, freezing and unfrozen parts, respectively.

unsaturated freezing soil (G < 1), the intensity of moisture flow, as moisture migrates towards lower temperatures, decreases, which results in the freezing out of the excess amount of liquid and vapour to approach thermodynamic equilibrium. The intensity of free ice formation (with G < 1) at depth in the frozen part of unilaterally freezing soil j = AIJAx will be different and is determined by the form of the curve of j and the freezing velocity o. Ice formation i in any cross-section of the frozen zone in time interval Axwill be calculated from the expression AI Jo, i.e.

Water transfer and ice formation in freezing soils are determined both by soil composition and structure and by the conditions of freezing. The composition of freezing soils is a basic factor responsible for differences in moisture transfer and ice formation in deposits. Thus there is practically no liquid moisture migration in gravel-pebble and sandy deposits where water transfer occurs chiefly due to vapour. When the deposits are fully saturated (G = 1) freezing usually gives rise to a volume increase due to a 9% increase in water volume on its conversion to ice, and often produces the so-called 'piston effect', i.e. a pushing of the excess water downwards. Segregated ice formation is noted only with mineral particles under 1 mm in size, when the adsorption-film moisture transfer mechanism commences. Ice formation here is directly correlated with the values of moisture-transfer coefficients and thermodynamic potential gradients in the freezing zone. Moisture transfer coefficients (K£ and in frozen soils are reduced substantially over the range of 0 to — 1 °C, and from kaolinite clays to montmorillonite and frozen sands, while moisture potential gradients in the freezing zone increase as fineness increases, and from montmorillonite clays to kaolinite. Migration of moisture thus increases with an increase in fineness of soils and with a greater amount of kaolinite mineral. An increase in moisture migration results from higher silt content, and from an optimum combination of water conductivity properties and moisture transfer driving forces. Absorption capacity of cations increases also from kaolinite to montmorillonite. The cation-exchange influence on moisture migration depends on cation valency; thus moisture migration, ice segregation and heaving increase with the saturation of soil with multivalent cations and diminish with univalent.

Of great practical importance are the problems of moisture transfer and ice accumulation in soils freezing under different thermodynamic conditions. In a soil with freezing under 'open' system conditions the total migration of i = jAx = P^dr dx

moisture is due to a water exchange comprising an internal part on account of the redistribution of the moisture in the soil itself and an external part on account of water migration from an external water-bearing layer. At the start of freezing the external migration flow is absent in soil, but then it appears near the water-bearing layer. As the boundary ¿;fr nears the water bearing layer, the proportion of the external migration water flow greatly increases relative to the internal one. Soil freezing under 'closed' system conditions results only in the internal moisture redistribution between the frozen and unfrozen parts of the soil. Therefore the intensity of migration of moisture depends here on the water storage in the unfrozen part. Ice accumulation in freezing soils depends on their freezing regime and increases with higher grad t in the frozen zone. However the increase of grad t in the frozen zone results in a higher freezing velocity which, on one hand, produces an increase in grad ¡.im and thus in the intensity of migration water flow d towards the freezing front, and on the other, a reduction in ice formation on account of the shortening of the water migration period t. Therefore freezing soil has an optimum relationship of the parameters o and grad t, under which the maximum ice formation will be observed. With a freezing velocity of over 8-10cm day"1, ice formation is either barely perceptible or absent altogether, since the freezing front advance is so fast that even with high values of grad t and grad pm, ice formation in the frozen zone is not significant.

Also important in practice is the dependence of moisture transfer and ice accumulation in freezing soils on external load. The freezing of soil samples under 'open' system conditions, under pressure and previously consolidated in the unfrozen state, has shown that with a higher external pressure the density of moisture flow into the frozen part is decreased (Fig. 2.13). Depending on soil dispersion (amount of fine-grained material) there appear critical or limited values of external load Pcr under which moisture migration into the frozen zone and ice accumulation practically cease. For clayey silt-rich soils Pcr may be 0.5 MPa and for kaolinite clays 1.5 MPa. Hence moisture migration into the frozen zone of epigenetically freezing clay soils will not be significant at a depth of more than some 100 m where normally the load at the freezing front is over 1 MPa.

Moisture migration and segregated ice formation occur under temperature gradients in the frozen part of freezing and of thawing soils. Thawing of frozen soil is noted for simultaneous ice melting (with a part of the frozen soil passing into a thawed state) and ice formation in the frozen part of a sample near the thawing front. Moisture migration into the frozen part of thawing soil arises only if it has a temperature gradient. The thawed part of a

Fig. 2.13. Effect of ambient pressure P on migration of moisture /,„ into the frozen part of freezing soils (grad t = 0.2 - 0.4'Ccm ufr = 0.1-0.4 cm day"1). 1 - kaolinite clay; 2 - loam; 3 - montmorillonite clay; 4 - silty sand with some clay.

sample is then desaturated and consolidated and the total moisture content of the frozen part of the soil increases. When frozen soils are thawed rapidly the boundary of phase transitions is not always at 0°C. The temperature here is often above zero, but small ice layers (schlieren) are present for some time in that part of the soil where the temperature of mineral layers is already above zero. All this is due to the inertia of phase transition of ice inclusions. Ice thawing and water seepage sometimes turns ice layers into partly closed cavities.

The frozen part of a slowly thawing soil with initially massive structure (in the frozen state, prior to the experiment) can have an increase in its ice content and new ice layers under formation in a temperature range of 0 to

— 2°C. As the thawing front advances, the ice microlayers, formed under lower temperatures, find themselves in the region of higher negative temperatures. They grow much thicker, from fractions of a millimetre (under

— 2°C) to 2 cm, near the thawing front. The growth of ice layers, situated in the part of the sample with lower negative temperatures, occurs at the expense of other ice layers of higher temperature. Ice layers grow thicker most intensively near the thawing front where they are fed with water migrating from the thawed part of the soil. The thawing of frozen soils with original streaky (small ice layers) cryogenic structure results chiefly in an increase of thickness of the ice layers present in the frozen zone of the sample. Slow thawing results, as a rule, in soil 'collapsing' at the sites of the thawed-out ice layers without the forming of closed fissures in their place.

The influence of frozen soil composition and thawing conditions on water transfer and ice accumulation have not, so far, been studied sufficiently.

However on present evidence it may be said that their influence is similar to that on freezing. So the intensity of moisture migration and ice formation in the frozen zone increases with increasing soil dispersivity and from mont-morillonite clays to those of kaolinite. The frozen zone of more compact and less wet soils usually has smaller ice accumulations. Decrease of grad t in the frozen part of the sample leads to reduced ice accumulation at the thawing front.

Let us consider in conclusion a water migration mechanism capable of acting in frozen and freezing soils and of forming injection and injection-segregation ice layers, lenses and sheets.

According to experiments the formation of a thick ice layer can occur only in cases where the hydrostatic pressure value of intruding water Pin is above the instantaneous rupture strength of frozen soil o™p and normal (ambient) pressure Pn. It is necessary for the formation of an injected ice layer that Pin > <7^sp + Pn. In an experiment with frozen ice-saturated silty sand (with some clay) with a recorded value of instantaneous strength ajup « 0.35 — 0.4 MPa at a temperature of — 1°C, no injection (using a metal needle) of water at a pressure of 0.1-0.3 MPa could be traced. It was only under a water pressure of 0.4 MPa that a hydrorupture of the sample occurred, 20min. after the experiment started. Water began entering the rupture zone and froze forming a compact injected ice layer. The upper part of the sample was displaced upwards by the amount of the intruded water volume and by its expansion in going from water to ice.

In the case of the frozen soil being unable to expand (deform) laterally or vertically, i.e. with heaving due to hydrorupture and injection being impossible, when the condition Pin < a'™ + Pn applies, the water coming under pressure seeps through frozen soil, fills ice-unsaturated cavities and freezes there. The necessary condition for water flow in frozen ice-saturated soil is here to overcome the initial gradient of unfrozen water seepage which will be determined by the ultimate shear strength <7sh of loosely bound water.

When the conditions for hydrorupture with the development of an injected ice layer are not fulfilled in a frozen ice-saturated sample and if the experiment lasts long enough (several days or weeks) there is a small water injection into the soil, some increase in ice content and the formation of ice microlayers in the sample which, as a whole, sustains little heaving (Fig. 2.14). This is all possible when the hydrostatic pressure of intruding water exceeds the long-term rupture strength <7'™ of frozen soil plus the value of normal load: Pin > <7'™ + Pn. Thus experiments made on frozen ice-saturated clay and clayey silty sand (at t ss 1 °C and Pn = 0) showed newly formed ice microlayers due to the freezing of water injected into the samples

Fig. 2.14. Cryogenic structure of clayey silt-sand (I) and kaolinite clay (II) prior to the experiment (a) and after water injection (b) into frozen samples: 1 - ice; 2 - mineral matrix.

under a pressure of some 0.2 MPa. The duration of tests was 3-7 days. The long-time strength of these soils was 0.1-0.18 MPa. The increase in the total moisture in the zone of injected streaky (fine lens) ice formation was 1020%. The injected water flow increased at first, but then as the ice content grew and the coefficients of water transmission went down correspondingly, it diminished to 0. Any subsequent rise in steps of the hydrostatic pressure of water resulted in cyclic formation of seepage-injected ice which must in the end bring about the hydrofracture of the sample.

All the above holds also for freezing soils in the case of the pressure-actuated water flow into the unfrozen part of the sample (below the freezing front). The formation and growth of injected ice layers are quite possible at the freezing front if Pin is in excess of the rupture strength of soil at the 'unfrozen-frozen soil' interface. Otherwise ice-formation in the freezing part of the soil will occur, due to the sum total of migration and injection flows. The higher Pin the greater will be the portion from the injection flow in the ice formation in the freezing part.

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  • razanur headstrong
    What force is responsible for the formation of ice during the freezing of water?
    2 years ago
  • Betty Schneider
    Is dative bond responsible for the formation of ice during freezing of water?
    2 years ago
  • benjamin
    What Force is responsible for the formation of ice during The freeing Of water?
    1 year ago
    What force responsible for formation of ice during frezing water?
    10 months ago
  • Euan
    What force is repsonsible for the formation of ice during the freezing of water?
    10 months ago
  • amina
    Why moisture content increase after freezing?
    3 months ago
  • Stig
    What happens when soil saturated with water and it freezes?
    3 months ago

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