SPECIAL REPORT 95-12

Freeze–Thaw Processes and

Soil Chemistry

Giles M. Marion March 1995
Abstract

This review broadly examines the interactions between freeze–thaw processes

and soil chemistry, focusing on 1) the effect of solutes on physical properties

such as freezing-point depression, unfrozen water and frost heaving, 2) the

effect of freeze–thaw cycles and low temperatures on soil chemistry, and 3)

modeling of freeze–thaw processes and chemistry. The presence of solutes

causes a freezing-point depression, which increases the amount of unfrozen

water in soils. Liquid films on soil particles provide the dominant route for the

flow of water and associated solutes in frozen soils. In general, salts reduce

the hydraulic conductivity and water flow to the freezing front, which reduces

frost heaving. Solute exclusion during freezing leads to supersaturated solutions, which promotes the precipitation of secondary minerals in soils. At the

water-shed level, ionic concentrations in early meltwaters are often 2–9 times

higher than snowpack concentrations. Temperature is the dominant factor controlling decomposition rates, with minimal detectable rates occurring at temperatures as low as –10°C; both bacteria and fungi are physiologically active

at subzero temperatures. Extracellular enzymes are active in soils at temperatures as low as –20°C; this activity is thought to occur in unfrozen water on

surfaces of soil particles. Nitrogen mineralization is reported at temperatures

as low as 1°C and is promoted by freeze–thaw cycles. There are strong and

complex interactions among soil properties that control solute and water flows

along concentration, temperature and hydrostatic gradients in freezing and

frozen soils. These com-plex interactions necessitate the development of computer simulation models that can integrate physical–chemical properties and

processes.

For conversion of SI units to non-SI units of measurement consult Standard

Practice for Use of the International System of Units (SI), ASTM Standard E380-

93, published by the American Society for Testing and Materials, 1916 Race St.,

Philadelphia, Pa. 19103.

This report is printed on paper that contains a minimum of 50% recycled

material.
Special Report 95-12

Freeze–Thaw Processes and

Soil Chemistry

Giles M. Marion March 1995

Prepared for

OFFICE OF THE CHIEF OF ENGINEERS

Approved for public release; distribution is unlimited.

US Army Corps

of Engineers

Cold Regions Research &

Engineering Laboratory
PREFACE

This report was prepared by Giles M. Marion, Research Physical Scientist, Geochemical Sciences Division, Research and Engineering Directorate, U.S. Army Cold Regions

Research and Engineering Laboratory.

Funding was provided by DA Project 4A161102AT24, Research in Snow, Ice, and Frozen

Ground, Tasks EC and SC, Work Units B03, Soil Solute Interactions at Low Temperature, and

F02, Chemical Processes in Frozen Soil; Strategic Environmental Research and Development Program (SERDP) Work Unit, Fate and Transport in Seasonally Frozen Soil and Discontinuous Permafrost; and AT25-EC-B03 Work Unit, Air-Snow-Ice-Soil Contaminant Interactions in Cold Regions.

The author thanks Dr. Steven A. Grant and Dr. Patrick B. Black for technically reviewing earlier drafts of this report.

ii
iii

CONTENTS

Preface .................................................................................................................................... ii

Introduction ........................................................................................................................... 1

Effects of chemistry on soil freezing and thawing ........................................................... 1

Soil moisture ..................................................................................................................... 1

Soil strength ...................................................................................................................... 5

Frost heaving .................................................................................................................... 6

Effects of freezing and thawing on soil chemical properties.......................................... 7

Chemical reactions ........................................................................................................... 7

Chemical transport .......................................................................................................... 8

Decomposition and nutrient availability .......................................................................... 10

Site remediation .................................................................................................................... 13

Modeling ................................................................................................................................ 15

Recommendations ................................................................................................................ 17

Literature cited ...................................................................................................................... 18

Abstract .................................................................................................................................. 25

ILLUSTRATIONS

Figure

1. Freezing-point depression for three salt solutions ................................................... 2

2. Stability diagram for the NaCl–H2O system at subzero temperatures ................. 2

3. Schematic of a freezing soil system............................................................................. 5

4. Ion exchange equilibria in a soil system .................................................................... 7

5. Nitrogen cycle in a terrestrial ecosystem ................................................................... 11

6. Cumulative nitrogen mineralization with time and temperature for an organic

tundra soil ............................................................................................................... 12

7. Concentrations of major seawater constituents during freezing............................ 15

8. Unfrozen water during seawater freezing ................................................................. 16

9. Distribution of sodium and chloride during seawater freezing ............................. 16

TABLE

Table

1. Matrix of direct and coupled transport phenomena in frozen porous media ...... 4
INTRODUCTION

This review will broadly examine the interactions between freeze–thaw processes and soil

chemistry, relying primarily on the empirical

record; the theoretical foundations for freezing and

chemical interactions were discussed in another

report (Grant, in prep.). In the first part of this report, I examine the effect of soil chemical properties on physical factors associated with freezing

and thawing, such as freezing-point depression,

unfrozen water content and frost heaving. The

second part of this report will examine the effect

of physical factors such as freeze–thaw processes

and subzero temperatures on chemical reactions

and chemical transport in soils. Then the effect of

freeze–thaw processes and subzero temperatures

on soil biology and soil remediation will be examined. And finally I will examine theoretical models for freezing and thawing processes that

explicitly include chemistry.

EFFECTS OF CHEMISTRY ON

SOIL FREEZING AND THAWING

Two important ways in which soil chemical

properties influence soil physical properties in

freezing and thawing environments are through

freezing-point depression and unfrozen water content. These properties, in turn, affect water and

solute movement in frozen soils, which influences soil strength, frost heaving and contaminant

transport.

Soil moisture

Solutes lower the freezing point of water. The

freezing-point depression in dilute solutions is

given by the Van’t Hoff equation (Lewis and Randall 1961):

dT/dmB ≈ –RT2/(55.5 ∆H) (1)

where mB = molality of solute B (mol kg–1)

T = temperature (K)

R = gas constant (J K–1 mol–1)

∆H = enthalpy of fusion of ice (J mol–1).

This equation simplifies to:

∆T ≈ –1.860 (νmB) (2)

where ∆T is in °C and ν is the number of aqueous

species resulting from the dissolution of the solute

(e.g., 1 for sucrose, 2 for NaCl, 3 for CaCl2) (Lewis

and Randall 1961). Therefore, the freezing point of

a 0.1-mol kg–1 NaCl solution would be 0.37°C lower

than the corresponding freezing point of pure

water.

Equation 2 is general for “dilute” solutions and

holds for all chemical species regardless of their

degree of dissociation. However, departures from

the limiting slope of –1.86°C/νm are apparent at

concentrations > 0.2 νm (Fig. 1). These departures

are most significant for electrolytes that dissociate

into ions of higher valences (i.e., MgSO4 > Na2SO4

> NaCl). Part of the explanation for the deviations

from the limiting slope are attributable to ion pairing between oppositely charged species, which decreases the true concentration of chemical species

in solution (e.g., one 

MgSO4

0 ion pair instead of two

separate Mg2+ and 

SO4

2– ions). In general, the

strength of ion pairing is proportional to the product of the ionic charges (Davies 1962). Therefore,

di-divalent ion pairs such as 

MgSO4

0 are much

stronger than uni-univalent ion pairs such as

NaCl0.

For a few salts, primarily highly soluble chlorides and sulfates, it is possible to construct complete phase diagrams at subzero temperatures. A

Freeze–Thaw Processes and Soil Chemistry

GILES M. MARION
Solute Molality (νm)

01234

0

Temperature (°C)

Limiting Slope = –1.86 °C/νm

– 8

– 6

– 4

– 2

MgSO4

Na2SO4

NaCl

Figure 1. Freezing-point depression for three salt solutions. The data were taken

from Linke (1965).

phase diagram for the H2O–NaCl system at subzero temperatures (Fig. 2) is very useful in describing the change in the solution phase as ice and

NaCl precipitate during the freezing process.

Assume an initial NaCl concentration of 1.00 mol

kg–1 and an initial temperature of 0°C. As this solution cools below 0°C, the NaCl concentration will

remain constant until the temperature reaches

–3.3°C, when ice, largely a pure phase, will begin

to precipitate, concentrating NaCl in the remaining unfrozen solution. If equilibrium is maintained

between the solution and ice phases during freezFigure 2. Stability diagram for the NaCl–H2O system at subzero temperatures.

NaCl Molality (mol kg–1)

01 567

0

Temperature (°C)

Eutectic Temperature 20

15

10

5

25

2 3 4

Ice

Solution

Ice + NaCl . 2H2O(cr)

NaCl . 2H2O(cr)

Experimental

Model

ing, then the NaCl concentration in solution will

follow the ice–solution equilibrium line, increasing in concentration as temperature decreases until

it reaches the eutectic composition at –21.2°C and

5.17 mol kg–1, at which point the residual solution

will solidify as a mixture of ice and NaCl.2H2O.

The phase diagram (Fig. 2) can be used to estimate the unfrozen water content at any point in

the freezing process. To facilitate these calculations,

the following polynomial equations were fitted to

the ice-solution and solution-NaCl.2H2O equilibrium data:

2
water interactions and as a consequence only demonstrates qualitatively what might occur in a soil.

For example, Frolov and Komarov (1993) demonstrated that NaCl solutions in soil remain unfrozen to temperatures lower (–24°C) than the eutectic

for pure NaCl (–21.2°C, Fig. 2), presumably due to

capillary effects.

Banin and Anderson (1974) examined solute

effects on freezing-point depression by comparing

the freezing points of soils, with and without NaCl

additions. They found excellent agreement between experimental measurements and a theoretical solute model at low NaCl concentrations and

a small but consistent underestimate (10–20%) of

the experimental freezing point at higher NaCl

concentrations. Suleimanov and Andronova (1990),

using a similar theoretical solute model, reported

a similar underestimate of the measured freezing

point. Both studies suggested that this discrepancy was due to the inhomogeneous distribution of

salt in water because of the negative adsorption of

anions (anion repulsion) near negatively charged

surfaces, which resulted in abnormally high salt

concentrations in the bulk solution.

Yong et al. (1979) developed an unfrozen soil

water model that explicitly accounts for negative

salt adsorption using diffuse double-layer theory.

This model showed excellent quantitative agreement with a montmorillonite soil sample but only

qualitative agreement for kaolinite and grundite

soils. This discrepancy was attributed to the model requirement for interlamellar migration of

water in soil freezing, which is a reasonable

assumption for montmorillonite but not for kaolinite or grundite. Yong et al. (1979) also pointed out

examples where pure solution theory contradicts

experimental measurements. For example, the

unfrozen water content at subfreezing temperatures first decreases with the addition of salt to a

minimum value around 10–3 mol L–1, then increases with increasing salt concentration, as would be

predicted from pure solution theory. There are

strong interactions among soil surface area,

charged surfaces, soil salts and moisture that control freezing-point depression and unfrozen water

content of soils. A model that can describe these

interactions quantitatively over a broad range of

salt concentrations does not exist at present.

A number of papers have reviewed the physics

of moisture movement in freezing and frozen soils

(Hoekstra 1969, Anderson and Morgenstern 1973,

Miller 1980, Perfect et al. 1991). This is a broad field,

much studied and for the most part beyond the

scope of this review. For the purposes of this re-

∆T = – 0.13 – 2.978 mNaCl

 – 0.211 

mNaCl

2 (ice–solution) (3)

 R2 = 1.000

∆T = – 42.37 – 12.082 mNaCl

 + 3.131 

mNaCl

2 (solution–NaCl.

2H2O) (4)

 R2 = 0.964

and

Wu = mBi/mBf (5)

where ∆T = freezing-point depression

mNaCl = NaCl molality

Wu = fraction of unfrozen water

mBi and mBf = initial and final molalities.

Equation 5 assumes the precipitation of a pure ice

phase. At –10° and –20°C, the corresponding NaCl

molalities are 2.77 and 4.94 mol kg–1, respectively

(eq 3). Therefore, the unfrozen water contents at

–10° and –20°C for a solution that was initially 1.0

mol kg–1 are 0.361 (1.0/2.77) and 0.202 (1.00/4.94),

respectively (eq 5). Conversely the frozen percentages are 63.9 and 79.8%. If the initial NaCl solution was 0.01 mol kg–1, then the unfrozen fraction

at –20°C would be 0.00202 (0.01/4.94), with a frozen percentage of 99.8%. Clearly the concentration

of salts in the solution phase during the freezing

process, because of salt exclusion from the ice

phase, can substantially alter the freezing point

and unfrozen water contents of solutions.

The NaCl–H2O phase diagram also demonstrates that an NaCl aqueous solution with an initial concentration < 5.17 m will not precipitate

NaCl.2H2O until the temperature drops to –21.2°C

and the NaCl molality increases to 5.17 mol kg–1

(the eutectic point). On the other hand, if the initial NaCl molality is > 5.17 mol kg–1, then

NaCl.2H2O will be the initial phase to precipitate

and ice will not form until the temperature drops

to –21.2°C and the NaCl concentration decreases

to 5.17 mol kg–1.

For soils the major factors influencing the freezing-point depression and unfrozen water contents

are temperature, adherence of water to soil particles and solute concentrations (Hoekstra 1969,

Anderson and Morgenstern 1973, Perfect et al.

1991, Frolov and Komarov 1993). The pure NaCl–

H2O phase diagram (Fig. 2) neglects soil solid–

3
4

Table 1. Matrix of direct and coupled transport phenomena in frozen

porous media. (After Perfect et al. 1991).

Driving force

Flux ∇T ∇pl ∇π ∇ε

jh FOURIER’S LAW* Thermofiltration [Dufour Effect]† Peltier Effect

jw Thermo-osmosis DARCY’S LAW Capillary osmosis Electro-osmosis

js Soret Effect Reverse osmosis FICK’S LAW [Electrophoresis]

jc [Seebeck Effect] Streaming potential [Diffusion potential] OHM’S LAW

* Direct processes are upper case; coupled processes are lower case.

† No reference to bracketed processes found in the soil freezing literature.

view, I want specifically to examine the roles of

soil chemical properties on moisture movement in

freezing and frozen soils. Before discussing chemical effects, however, we need a theoretical framework for the discussion.

Following Perfect et al. (1991), we will use the

nonequilibrium thermodynamic approach, where

fluxes are written as explicit functions of both

direct and coupled transport phenomena. For example,

 

j L T T L pl

L L

w wh ww

ws we

=− ∇ − ∇

− ∇− ∇

/

π ε

(6)

 

j L T T L pl

L L

s sh sw

ss sc

=− ∇ − ∇

− ∇− ∇

/

π ε

(7)

where jw and js =fluxes of liquid water and

solute

∇T, ∇pl, ∇π and ∇ε =gradients in temperature,

hydrostatic pressure, solute

concentration and charge

Lmn =transport coefficient relating

the mth flux (jm) to the nth

component.

Fluxes and their driving forces (gradients) in frozen porous media are summarized in Table 1.

Water moves in soil (jw) in response to changes

in the chemical potential of water, which is related

to gradients in the hydrostatic pressure (∇pl) (Darcy’s Law), temperature (thermo-osmosis), solute

concentration (capillary osmosis) and electrical

potential (electro-osmosis) (Table 1, eq 6).

At the freezing front in dilute solutions, anions

(generally) are preferentially absorbed into the ice

phase, leading to a measurable charge separation

(freezing potentials); this phenomenon is called the

Workman Reynolds Effect (Drost-Hansen 1967,

Murphy 1970, Hanley and Rao 1982). Freezing potentials are properties of dilute solutions and pure

water and disappear at high concentrations (DrostHansen 1967, Murphy 1970). Although Hanley

and Rao (1982) developed a model to quantify the

relation between freezing potentials and the migration of moisture and ions in freezing soil, the

overall significance of this phenomenon for soil

freezing is unclear. Another electrical phenomenon, electro-osmosis, can cause considerable water movement, but this is only important in cases

where induced electrical potentials are applied to

soils (Hoekstra and Chamberlain 1964). Outcalt et

al. (1989) monitored electric potentials in freezing

soils during diurnal and seasonal freeze–thaw

cycles. They concluded that the rapid and systematic pattern of electrical potential variation during freeze–thaw events demonstrates that the

effects of electrolyte concentration and dilution are

products of evaporation–distillation, melting of

frost-purified ice, soil water advection to the freezing region, electrolyte expulsion from the freezing region and infiltration of rain and snowmelt.

They further concluded that soil electrical potentials will yield valuable information concerning

the state and mobility of soil water in freezing and

thawing soils.

In general, water moves from warm to cold,

from regions of low solute concentration to highsolute regions and from high-moisture zones to

low-moisture zones (Perfect et al. 1991). Chemical

potentials of water due to gradients of hydrostatic pressure, solute concentration and temperature

interact additively to create a strong thermodynamic sink for liquid water at the freezing front

(Fig. 3). As soils freeze from the top downward,

the thermal gradient will induce an upward flow

of water to the freezing front. Solutes are largely

excluded in the freezing process, and maximum

solute concentrations are generally found imme-
5

Figure 3. Schematic of a freezing soil system.

Factors controlling the upward flux of water into

the freezing and frozen soil zones are of critical

importance in understanding practical problems

associated with freezing and thawing phenomena

such as soil strength and frost heaving.

Soil strength

Freezing, per se, invariably increases the mechanical strength of soils because of the strong

bonding between ice and soil particles (Mahar et

al. 1983). Freezing and thawing can either increase

or decrease soil strength, depending on the degree

of thaw weakening. With respect to clays, Chamberlain (1989) concluded that strength increases can

be expected where there is an increase in consolidation and density during freezing and thawing;

decreases in mechanical strength can be expected

with highly cemented clays and clay soils that are

highly overconsolidated before freezing. Rowell

and Dillon (1972) found that clay aggregates were

produced by freezing. After thawing, these aggregates swelled and dispersed, depending on electrolyte concentration, in a manner similar to

aggregates formed by drying.

The presence of salts in soils generally decreases mechanical strength (Tyutyunov and Derbeneva 1970, Tsytovich et al. 1973, Mahar et al. 1982,

1983, Ogata et al. 1982, Chamberlain 1983). This

decline in mechanical strength with increasing salt

concentration is generally attributed to a lowering

of the freezing point and, as a consequence, an increase in unfrozen water content. In addition to

diately in front of the freezing front (Hallet 1978,

Miller 1980, Kay and Groenevelt 1983, Hofmann

et al. 1990). This solute (or osmotic) gradient will

generally induce an upward movement of water.

And last, freezing of soil water, which is analogous to soil drying (Lahav and Anderson 1973,

Miller 1980), creates a strong sink for water and

induces an upward movement of water to the

freezing front.

Under subzero temperatures some soil water

remains unfrozen as thin films around soil particles and serves as an avenue for the movement

of both water and solutes in frozen soil. From 0°to

–1°C there is an exponential decrease in the

hydraulic conductivity (Lww, eq 6) of a frozen soil

from 10–8 to 10–12–10–14 m/s; below –1°C the

hydraulic conductivity remains essentially constant (Perfect et al. 1991). Solutes can theoretically

increase the water flux, in part because of the direct effect on water diffusion (Lws in eq 6) but also

because of indirect effects on freezing-point depression, unfrozen water content and increases in

the thickness of soil moisture films (Anderson and

Morgenstern 1973, Henry 1988). Despite these theoretical considerations, it appears that the dominant effect of solutes is generally a reduction in

hydraulic conductivity (Lww) that overwhelms the

other terms in the water flux equation (eq 6). As a

consequence, water flux to the freezing front is

generally reduced by solutes (Kay and Scott 1973,

Sheeran and Yong 1975, Chamberlain 1983, Cary

1987, Ershov et al. 1992).
unfrozen water films around soil particles, larger

unfrozen brine pockets can form during especially rapid ground freezing of saline soils, which also

decreases the mechanical strength of soils (Mahar

et al. 1982, Chamberlain 1983).

Kostetskaya (1977) found that NaCl brines freezing at the eutectic point (–21.2°C, 5.17 mol kg–1;

Fig. 2) are sludge ice and have zero strength. However, the addition of Na2SO4 to the NaCl brines or

to gravel soils containing the same brines increased

strength. Ogata et al. (1982) found that the compressive strength for sands, silts and clays generally decreased with increasing salinity. An

exception to this generality occurred where soils

were frozen at –32°C, where the compressive

strength at 2 and 3% salinity was actually higher

than at 1% salinity. They attributed this anomaly

to the fact that –32°C was beyond the eutectic temperature of NaCl. Clearly complex interactions are

possible among specific solutes, freezing and thawing, and soil physical properties.

Frost heaving

From the classical papers of Taber (1929, 1930)

and Beskow (1935) to more recent reviews (Anderson and Morgenstern 1973, Sheeran and Yong 1975,

Gilpin 1980, Miller 1980, Chamberlain 1981, O’Neill

1983, Derjaguin and Churaev 1986, Nakano 1990,

Black and Hardenberg 1991, Perfect et al. 1991),

frost heaving has been the subject of intense interest because of its importance for engineering in

frost-susceptible environments. The stability of

soils for road, railroad, airfield, building, underground storage and pipeline construction in cold

regions are critically dependent on frost susceptibility.

Frost heaving is due primarily to the formation

of segregated ice lenses in soil and only secondarily to volumetric expansion of water during freezing (Anderson and Morgenstern 1973, Mageau and

Morgenstern 1980, Miller 1980, Chamberlain 1981).

Three conditions must occur simultaneously for

frost heaving to occur:

• A prolonged period of subfreezing temperatures;

• Frost-susceptible soils; and

• A source of water.

Several models have been proposed to explain frost

heaving (Miller 1980, Chamberlain 1981, O’Neill

1983, Black and Miller 1985, Horiguchi 1987,

Nakano 1990, Black and Hardenberg 1991), including:

• The capillary (or primary) heaving model;

• The secondary heaving model;

• The adsorption model; and

• The osmotic model.

Although these models differ in the details of the

mechanisms responsible for frost heaving, there

is general agreement on the factors influencing

frost heaving (Miller 1980, Chamberlain 1981,

O’Neill 1983, Perfect et al. 1991). These factors include:

• Soil texture (silts are more frost-susceptible

than either sands or clays);

• Pore size (pore size controls capillary rise, soil

suction and hydraulic conductivity, all of

which are important in the movement of

water to the freezing front);

• Rate of heat removal (there is a maximum rate

of frost heave at intermediate rates of heat

removal);

• Temperature gradient (this affects the thickness and hydraulic conductivity of the frozen fringe where ice lensing occurs) (Fig. 3);

• Moisture conditions (saturated soils with

high water tables are most susceptible to frost

heaving) and

• Overburden stress (the greater the overburden stress, the lower the frost heaving).

In general, salts reduce frost heaving (Beskow

1935, Kay and Scott 1973, Yong et al. 1973a,b, Sheeran and Yong 1975, Chamberlain 1983). Exceptions

to this generality exists. For example, Beskow

(1935) found that dilute NaCl solutions (≤ 0.05 M*)

accelerated frost heaving, but more concentrated

NaCl (0.1 M) as well as CaCl2 (0.025–0.5 M) decreased frost heaving. Because of the generally

ameliorative effect, solutes have been examined

for their ability to reduce frost heaving. Lambe and

Kaplar (1971) and Lambe et al. (1971) evaluated

more than 50 additives classified under four functional groups (void fillers and cements, aggregants, waterproofers, and dispersants) for their

ability to modify frost heaving. They found that a

dispersant, tetrasodium pyrophosphate (TSPP),

and an aggregant, ferric chloride, possessed good

frost heave modifying capabilities. Both additives

reduced frost susceptibility in soils from “medium to high” to “very low to low.”

Sheeran and Yong (1975) have proposed five

hypotheses to explain the effect of salts on frost

heaving. Salts may affect:

• The position and temperature of the freezing

front;

• The chemical condition of frozen soil;

• The permeability of frozen soil;

6

* M is molarity (mol L–1).
important phenomenon in all its complexity, coupled with judiciously selected empirical lab and

field studies, will ultimately lead to the ability to

accurately predict frost heaving.

EFFECTS OF FREEZING AND THAWING

ON SOIL CHEMICAL PROPERTIES

Freeze–thaw processes and low temperatures

influence chemical reactions, chemical transport

and nutrient availability in soils. These interactions, in turn, influence mineral weathering rates,

pedogenesis, contaminant remediation and revegetation of disturbed lands in cold regions.

Chemical reactions

Two of the most important chemical reactions

affected by freezing and thawing are precipitation–

dissolution and cation exchange. Solute exclusion

during ice formation leads to supersaturated solutions, which promotes the precipitation of secondary minerals in soil, alters solution-phase

compositions (which may promote the dissolution

of primary minerals) and shifts equilibrium toward increased mineral weathering (Fig. 4) (Hallet 1978, Zvereva 1982, Sletten 1988, Richardson

et al. 1990). Martynenko et al. (1992) subjected

ground primary minerals in oxalic acid solutions

to 70 freezing cycles. The interaction of acidic hydrolysis and cryogenic comminution enhanced

mineral grain fragmentation and chemical weathering.

The precipitation of silica in the lower horizons

of soil profiles has been attributed to freezing

(Slavnyy and Vorob’yeva 1962). The distribution

of Na and Mg sulfate minerals in soils is temperature dependent, with low temperatures favoring

the precipitation of Na sulfates (e.g., mirabilite:

Na2SO4

.

10H2O) relative to Mg sulfates (Arndt and

Richardson 1989, Richardson et al. 1990); in these

papers it was hypothesized that freezing from the

top down may concentrate the more soluble Mg

sulfates at greater depth, leading to the development of surficial sodic (Na) horizons.

Freezing of Ca and Mg bicarbonate solutions

leads to the precipitation of the more insoluble

CaCO3, thereby increasing the solution-phase

Mg2+/Ca2+ ratio (Vlasov and Pavlova 1969, Ivanov

7

• The rate of water movement to the freezing

front; and

• The freezing process.

Salts are excluded in the freezing process, freezing points are depressed by salts, and unfrozen

water contents are increased in the presence of

salts. (See previous discussion on Soil moisture.)

These factors would reduce the water available for

forming segregated ice. Additionally, lower temperatures reduce the soil hydraulic conductivity

(Perfect et al. 1991). Salts in freezing soils reduce

the capillary rise and soil hydraulic conductivity

(Kay and Scott 1973, Sheeran and Yong 1975,

Chamberlain 1983, Cary 1987). These factors that

reduce water flow to the freezing zone are probably the most significant factors explaining lower

frost heaving in the presence of salts. The frozen

fringe generally thickens in the presence of salts

(Sheeran and Yong 1975, Yong and Sheeran 1978,

Chamberlain 1983), which could reduce the water flux to segregated ice lenses.

Solutes clearly affect the lensing process (Sheeran and Yong 1975, Chamberlain 1983). Chamberlain (1983) found plentiful but thin ice lenses

forming in the presence of salts. This type of ice

lens formation could reduce the flow of water in

the frozen fringe, where segregated ice forms. Salts

can interact with soil particles, causing them to

either aggregate or disperse, depending on solute

type and concentration (Lambe et al. 1971, Sposito 1989); this, in turn, will affect soil permeability

and the resultant water flux.

Salts may also affect the depth of frost penetration. The deeper the zone of freezing, the more

important overburden pressure becomes in limiting segregated ice formation. Sheeran and Yong

(1975) found that the depth to the freezing front

decreased with increasing salt concentration. On

the other hand, Mahar et al. (1983) found that the

rate of advance of the freezing front increased with

increasing salinity; they attributed this to the gradual release of the latent heat of fusion over a range

of temperatures and depths due to changing solute concentrations. Cary (1987), using a simulation model, predicted a greater depth of frost

penetration with increased salt concentration.

Contradictions abound in the literature on frost

heaving, probably because of the complex interactions among the driving forces controlling this

process. Frost heaving is clearly a complex process dependent on soil properties, moisture conditions, solute concentrations and energy balance.

Such complexity can not be explained by simple

models. Only a continuing effort to model this

Figure 4. Ion exchange equilibria in a soil system.

Exchange

Complex Ion Exchange SolubleIons

Precipitation

Dissolution

Soil

Minerals
8

and Vlasov 1973, Sletten 1988). Sletten (1988)

attributed the formation of aragonite, a relatively

rare polymorph of CaCO3 (Doner and Lynn 1977),

to a more favorable Mg2+/Ca2+ ratio in solution

brought about by the differential solubilities of Ca

and Mg carbonates. Hallet (1976) demonstrated

that freezing strongly concentrates solutes in

CaCO3 solutions, which leads to the precipitation

of CaCO3. Based on temperature-dependent solubility data, Hallet estimated the eutectic temperature for pure CaCO3 (–0.34°C). The relative

insolubility of minerals such as CaCO3 and CaSO4

make them difficult to study at subzero temperatures because of the limited temperature range in

which their single-salt solutions can co-exist in

equilibrium with solid phase ice. This is in marked

contrast to soluble chloride salts such as NaCl,

which has a eutectic temperature of –21.2°C (Fig.

2), or CaCl2, which has a eutectic temperature of

–50.4°C (Spencer et al. 1990, Marion and Grant

1994).

Ion exchange reactions play a major role in controlling many physical and chemical properties of

soils, such as aggregation, pH buffering and ion

transport. Hinman (1970) found that alternate

freezing and thawing increased exchangeable NH4-

N and decreased exchangeable K; there was no

change in the cation exchange capacity (CEC) or

exchangeable Ca and Mg. Pulubesova and Shirshova (1992) found no significant change in CEC

or exchangeable Ca, Mg, K and Na in two soils and

two clays (kaolinite and bentonite) following prolonged freezing (two months) and prolonged thawing (two months). Alternate cycles of freezing and

thawing have led to both fixation of fertilizer K

and release of crystal lattice K from K-depleted soils

(Graham and Lopez 1969). Freezing increased adsorbed bases (the pH increased), while thawing

increased soil acidity (the pH decreased) (Fedorov

and Basistyi 1974). Deep freezing of soils with liquid N2, which boils at –196°C, led to significant

increases in the concentration of Ca, Mg and K in

solution following thawing (Iskenderov 1976).

Surprisingly little work in soil science has examined the role of subzero temperatures on chemical thermodynamic equilibrium constants for

cation exchange. Tyutyunova and AntipovKaratatyev (1965) examined the Ca2+-K+ and the

Mg2+-K+ exchange on montmorillonite, kaolin and

soil over the temperature range from +20° to –17°C.

Equilibrium constants at negative temperatures

were determined in water–alcohol mixtures. There

was an increased adsorption of K with decreasing

temperature, which fell in the relative order:

kaolin > soil > montmorillonite. They reported standard Gibbs energies, enthalpies and entropies for

ion exchange.

Chemical transport

When aqueous solutions freeze, solutes are largely excluded from ice. As a consequence, solute concentrations are generally highest at the freezing

front (Fig. 3). Kay and Groenevelt (1983) derived a

simple equation:

Cf = Ci + 80 k Ci (8)

where Ci = original solute concentration

Cf = solute concentration at the frost front

k = solute inclusion coefficient.

If there is no solute inclusion in the frozen zone (k

= 1), then the solute concentration at the freezing

front could rise to 80 times the original concentration (assuming no salt precipitation). Such high

concentrations can cause the freezing front to leap

over solute pockets (Hallet 1978, Kay and Groenevelt 1983, Romanov and Levchenko 1989). These

solute pockets may ultimately freeze, resulting in

alternating bands of high and low concentrations

in frozen soils (Romanov and Levchenko 1989). Kay

and Groenevelt (1983) have argued that solute exclusion, which leads to narrow alternating highand low-solute bands in soils, may not be a significant mechanism for macroscale solute redistribution in soils.

The dominant gradients that control the movement of solutes through freezing and frozen soils

are concentration, temperature and hydrostatic

pressure (Table 1, eq 7) (Cary and Mayland 1972,

Baker and Osterkamp 1988, Qiu et al. 1988, Perfect

et al. 1991). Solutes will diffuse from zones of high

concentration (e.g., the freezing front) to zones of

low concentration (Ficks Law, Table 1). In soils, solutes will move from warm to cold regions (Soret

Effect, Table 1) (Cary and Mayland 1972, Qiu et al.

1988). However, the direct effect of temperature

gradients on solute movement in soils is generally

insignificant (Cary and Mayland 1972). Much more

significant is the movement of solutes with water

along hydrostatic gradients from warm zones to the

freezing front (reverse osmosis, Table 1). Freezing

of water creates a strong thermodynamic sink, and

water will move in both the vapor and liquid (carrying solutes) phases to the freezing front (Cary and

Mayland 1972, Gray and Granger 1986, Hofmann

et al. 1990). In this process, solutes may move

against a concentration gradient.
The redistribution of solutes depends strongly

on freezing rate, moisture content, soil texture and

time. There are cases, especially in fine-textured

soils, where maximum solute concentrations are

found in the frozen zone (Ershov et al. 1992). Qiu

et al. (1988) demonstrated that for single-salt solutions in moist sands, solutes migrate toward the

unfrozen zone (Fig. 3); on the other hand, in silt

and clay soils, solutes migrate toward the freezing zone. Over several freeze–thaw seasons, surficial salt applications may be leached to deeper

soil horizons, effectively moving in the opposite

direction of water during frost periods (Yong et

al. 1973a). Baker and Osterkamp (1988) found that

significant salt rejection and brine drainage occurred with downward freezing, but there was

none with upward freezing. The amount of salt

rejection increased with decreasing freezing rate.

These differential responses are due to the role of

soil physical properties such as porosity, surface

area and hydraulic conductivity in controlling

both water and solute flows along concentration,

temperature and hydrostatic gradients. The complex interactions possible among soil physical and

chemical properties is a strong incentive for developing computer simulation models.

Liquid films exist on soil particles in frozen

soils, which provides the dominant route for the

flow of water and associated solutes in frozen soils

(Murrmann et al. 1968, Hoekstra 1969, Cary and

Mayland 1972, Murrmann 1973, Gray and Granger 1986, Hofmann et al. 1990). Murrmann (1973)

found surprisingly high diffusion rates for Na ions

in the temperature range of 0° to –15°C, which he

attributed to the existence of thin water films.

Murrmann (1973) concluded that the temperature

dependence of ionic diffusion at subzero temperatures is primarily a function of water film thickness. Even in Antarctic soils, which are

continuously frozen and relatively dry, the liquid

films at mineral surfaces are believed to be the

dominant avenue for the movement of solutes

(Ugolini and Anderson 1972, 1973).

Because many soil surfaces are charged and the

solubility of different soil minerals are highly variable, ionic mobility depends on the specific ions.

For example, Ugolini and Anderson (1972, 1973)

found that Cl– is more mobile in Antarctic soils

than Na+, presumably because of the attraction

between the negatively charged CEC and the positively charged Na ion. Czurda and Schababerle

(1988) found that the monovalent cations Na and

K were more mobile than the divalent cations Ca

and Mg in frozen clay columns. This is probably

because divalent cations have greater electrostatic

attraction to charged surfaces than monovalent

cations. More-soluble soil constituents migrate

more readily through freezing and frozen soils because the less-soluble minerals precipitate; this can

play a major role in horizonation and pedogenesis

of cold regions soils (Hallet 1978, Zvereva 1982,

Panin and Kazantsev 1986, Sletten 1988, Richardson et al. 1990).

The transport of solutes between water or snow

and soil during freeze–thaw processes can play an

important role in geochemical cycling in cold regions. Kadlec et al. (1988) found that the freezing

of shallow waters in peatlands drives a considerable portion of solutes into the topsoil. Ostroumov

et al. (1992) examined the flux of solutes from soil

into snow under laboratory conditions. The maximum concentrations in snow were found at the

snow–soil boundary. Solutes were transferred to

snow in the same direction as the heat flux. The

flux of ions in snow fell in the order: K > Cl >> Li >

Ca > Cu >> Cd = Pb. They attributed the greater

flux of K relative to Cl to greater anion adsorption,

relative to cation adsorption, on the surface of ice

particles. This charge separation phenomenon is

identical to the Workman–Reynolds Effect, which

is, however, usually thought to be due to absorption of anions into the ice phase.

Soil freeze–thaw processes can play a role in

controlling geochemical cycling at larger scales,

such as watersheds and geographic regions. Edwards et al. (1986) found that freeze–thaw cycles

have significant effects on the chemical composition of streams. Substantial amounts of many elements that dominate stream chemistry become

available upon thawing, especially Al, K and organic C. They also hypothesized that soil freezing

might also influence solute chemistry by altering

hydrologic pathways. Everett et al. (1989) concluded that snowmelt was the most important hydrologic and geochemical event in a small arctic

watershed in northern Alaska. Ion concentrations

were highest during the first 15% of the snowmelt

event. In all cases ion concentrations in the spring

runoff were four to nine times those of the snowpack. Potassium was present in surface waters only

during meltoff and for a short time thereafter. Edwards et al. (1986) also reported the mobilization

of K during the spring thaw. During winters with

little or no snowmelt before spring, Johannessen

and Henriksen (1978) found that 50–80% of the

winter pollutant load is released with the first 30%

of meltwater. Average solute concentrations were

2–2.5 times higher than snowpack concentrations.

9
The resulting increases in acid concentrations of

low-buffered waters occasionally lead to severe

physiological stress on aquatic organisms. They

hypothesized that these high concentrations may

be due to a freeze-concentration process during

snow recrystallization and melting in which solutes preferentially accumulate at surfaces of ice

particles. Clearly snowmelt is a major hydrologic

and geochemical event in regions subject to freeze–

thaw cycles.

Ivanov and Vlasov (1973) concluded that cryogenic processes have the effect of forming waters

that have low mineral concentrations and low

Ca2+/Mg2+ ratios in the Transbaikal region of

Siberia. Factors that influence low mineralization

are the freezing of soil solutions and groundwaters, the redistribution of salts between ice and liquid phases, the slow transition of salts into solution

during thawing, and the precipitation of carbonates of alkaline-earth metals.

DECOMPOSITION AND

NUTRIENT AVAILABILITY

How low temperatures and freeze–thaw cycles

alter the soil as a medium for microbial activity

and plant growth in cold regions soils is an important factor affecting revegetation, contaminant stability and carbon cycling. The availability of

nutrients is critical for plant growth and, as a consequence, revegetation of severely disturbed lands.

Organic contaminants can be directly decomposed

by microbial activity, while mineral contaminants

(e.g., Hg, Pb, Cd) can be absorbed by plants, reducing their mobility (Lagerwerff 1972, Page et al.

1987). Carbon fluxes from northern ecosystems

may play a role in global carbon balance, affecting

the global climate (Kvenvolden 1993, Marion and

Oechel 1993, Oechel et al. 1993, Zimov et al. 1993).

Several chapters in a recent book edited by

Chapin et al. (1992) discuss arctic ecosystems and

microbial processes (Nadelhoffer et al. 1992),

nitrogen fixation (Chapin and Bledsoe 1992) and

plant nutrient absorption (Kielland and Chapin

1992). In the present review, our primary focus will

be on the effect of low temperatures and freeze–

thaw cycles on decomposition and mineralization

processes. For fuller discussions of factors influencing microbial activity and plant growth in cold

ecosystems, see the Chapin et al. (1992) book.

Temperature is the dominant factor controlling

decomposition rates (CO2 production), with minimal detectable rates occurring at –10° to –6°C

(McCown et al. 1972, Flanagan 1978, Flanagan and

Bunnell 1980, Nadelhoffer et al. 1992, Gilichinsky

1992, 1993). Flanagan and Bunnell (1980) concluded that both bacteria and fungi are capable of

growth at subzero temperatures. Decomposition

rates increase with increasing temperature between

5° and 20–30°C, and overall decomposition rates

increase by 20% yr–1 for every 1000 degree-days

above 0°C (Nadelhoffer et al. 1992).

It is well known that freeze–thaw cycles cause a

respiratory burst of CO2 and CH4 following thawing (Mack 1963, Ivarson and Sowden 1970, McCown et al. 1972, Skogland et al. 1988, Christensen

1993). Skogland et al. (1988) attributed the respiratory burst to killing of bacteria, lysing of their cells

and utilization of liberated carbon compounds as

nutrients by the surviving bacteria. Ivarson and

colleagues found strong correlations between extractable free amino acids and sugars (nutrients)

and soil respiration, suggesting that microbial stimulation by nutrients was the cause of the CO2 burst

upon thawing (Ivarson and Gupta 1967, Ivarson

and Sowden 1966, 1970). Ross (1972) found increased dehydrogenase activity with thawed soil

samples in aerobic assays. He attributed this response to increased availability of substrates rather than multiplication of microorganisms because

in anaerobic assays with shorter incubation times,

the effects of increase substrate availability were

less evident. McCown et al. (1972) attributed a

spring CO2 burst to trapping of CO2 within the soil

matrix upon bidirectional freezing, with subsequent release during spring warming; they believed this to be primarily a physical effect and not

biologically mediated. In the latter study, soil surface temperatures beneath a snowpack were in the

range of –16° to –10°C; studies, previously cited,

suggest that biological production of CO2 does not

become important until soil temperatures rise

above –10°C. Christensen (1993) attributed the

spring burst of CH4 to the release of trapped CH4

in the frozen soil and possibly also to changes in

the production and consumption of CH4 as temperatures increased and the active layer deepened.

The seasonal patterns of CO2 and CH4 emissions

are important in assessing the role of these gases

in the global carbon balance. Most of the evidence

to date indicates that CH4 emissions in the winter

are insignificant in tundra soils (Whalen and Reeburgh 1992, Christensen 1993); however, evidence

to the contrary also exists (Panikov and Zelenev

1992). Kelley et al. (1968) suggested that increased

CO2 concentrations at the soil surface beneath a

snowpack at Barrow, Alaska, in the winter was due

to microbial respiration as well as other physical

10
and chemical processes. Significant CO2 emissions

from wet tundra soils extended through February–

March in some northern sites (Federov-Davydov

1993, Zimov et al. 1993). On the other hand, in other soils, CO2 fluxes went to zero when soils froze

(Federov-Davydov 1993, Zolotareva and Demkina

1993). Zimov et al. (1993) hypothesized that energy production by microbes was an important contributor to soil heat balance, preventing the freezing

of the entire soil profile, which allows significant

microbial activity to occur even during the winter.

Because of the high temporal and spatial variability in winter gas fluxes (Federov-Davydov 1993,

Zimov et al. 1993), their overall significance for the

global carbon balance is unclear at present.

Extracellular enzyme activities (urease, phosphatase, sulfatase) have been detected in soils at

temperatures as low as –20°C (Bremner and Zanuta 1975); these low-temperature activities are believed to be occurring in the unfrozen water at

surfaces of soil particles. Freeze–thaw events in

tundra water tracks in the fall can cause rapid decreases in soil redox potentials; associated with

these changes are concomitant increases in extracellular enzyme activity (cellulases, phosphomonoesterase, proteases) (Linkins 1987). Enzyme

activity continues to increase over three or four

freeze–thaw cycles. It appears that the number of

freeze–thaw events, rather than the duration of any

one event, is the principal factor determining total

enzyme activity. Linkins (1987) hypothesized that

11

Figure 5. Nitrogen cycle in a terrestrial ecosystem. (After Marion 1987.)

higher plants may obtain a significant proportion

of their annual nutrients during the fall as the result of increased soluble sugars, phosphorus and

amino acids in solution.

The effect of freeze–thaw processes on nutrient

availability is a critical factor in revegetation of

severely disturbed lands in cold regions. Because

nitrogen (N) is the element that most frequently

limits terrestrial plant growth (Raven et al. 1986),

it has been the most intensively studied nutrient

in cold regions.

The dominant processes controlling the cycling

of terrestrial N (Fig. 5) include:

• N2 fixation, which directly converts atmospheric N2 gas into organic N via N-fixing

bacteria;

• Ammonification, which converts organic N

into ammonium ( 

NH4

+);

• Nitrification, which converts 

NH4

+ into nitrate

(

 

NO3

– ) through the nitrite ( 

NO2

– ) intermediary;

• Cation-exchange reactions between 

NH4

+ and

exchange complexes;

• 

NH4

+ fixation, which can render 

NH4

+ difficultly exchangeable;

• Plant uptake of 

NH4

+ and 

NO3

– ;

• Immobilization of 

NH4

+ and 

NO3

– by microbial processes;

• Denitrification, which converts 

NO3

– into gaseous nitrogen (N2) and nitrous oxide (N2O);

• Ammonia (NH3) volatilization, which converts 

NH4

+ into NH3 gas;

• Adsorption of NH3 gas by soils; and

• Leaching, which is a particular problem for

the negatively charged 

NO3

– ion (Clark and

Rosswall 1981, Marion 1987).
Many of the reactions controlling nitrogen

cycling (Fig. 5 ) are controlled by biological mechanisms. Examples include N2 fixation (bacteria),

ammonification (bacteria and fungi), nitrification

(bacteria), denitrification (bacteria), immobilization

(bacteria and fungi) and uptake (plants) (Clark and

Rosswall 1981, Marion 1987). Some of the nitrogen

cycling processes can also be accomplished by

strictly abiotic mechanisms [e.g., ammonification

via fire (Marion et al. 1991) and chemodenitrification (Christianson and Cho 1983)]. Nevertheless, it

is clear that biological processes play a dominant

role in controlling the flow of nitrogen as well as

other nutrients through terrestrial ecosystems.

The responsiveness of nitrogen cycling processes at low temperature is critical to assessing the

nitrogen supplying power of cold regions soils.

Field measurements indicate net nitrogen mineralization at temperatures as low as 1°C (Gersper et

al. 1980). Laboratory incubations have demonstrated that nitrogen mineralization (ammonification

and nitrification) is less temperature sensitive at low

temperatures (3–9°C) than at high temperatures

(15–35°C) (Marion and Black 1987, Nadelhoffer et

al. 1991, 1992). For example, at 35°C, nitrogen mineralization follows an exponential curve that shows

a high initial rate of nitrogen mineralization followed by a gradually declining rate with time (Fig.

6); on the other hand, at 5°, 15° and 25°C, nitrogen

mineralization follows a parabolic curve that shows

a low initial rate of nitrogen mineralization followed by a gradual increase with time. The exponential mineralization curve is typical of temperate soils (Stanford and Smith 1972, Marion et al.

1981) but atypical of tundra soils in the temperature ranges normally encountered in cold regions.

The slow release of nitrogen at low temperatures

(Fig. 6) may be due to the generally unfavorable

C/N ratios of the organic soils, which in one study

ranged from 18 to 82 (mean = 40) (Marion and Black

1987). Flanagan and Bunnell (1980) reported a similar response for decomposers (CO2 production)

that responded more linearly at low temperatures

and more exponentially at high temperatures.

Freeze–thaw cycles generally promote mineralization (ammonification and nitrification) of organic nitrogen, which leads to increased concentrations

of 

NH4

+ and 

NO3

– (Soulides and Allison 1961, Mack

1963, Hinman 1970, Honnolainen and Reppo 1975,

Malhi and Nyborg 1979, 1986, Gersper et al. 1980).

An exception to this generality was reported by

Read and Cameron (1979), who monitored mineral nitrogen forms in several soil types over 10 years.

Between fall and spring, they found some increase

in NO3-N but a larger decrease in NH4-N, resulting in a net decrease in mineral nitrogen during

the winter. Malhi and Nyborg (1986) found large

increases in mineral N when soils froze in the winter and large decreases in early spring when soils

thawed; they attributed the spring loss to denitrification and not to leaching of 

NO3

– . Christianson

and Cho (1983) found that the maximum production of N2 gas via chemodenitrification occurred

at –3.5°C over the temperature range from –20° to

12

Incubation Time (weeks)

Nitrogen Mineralization (mg kg –1)

0 2 4 6 8 10

800

600

400

200

 5°C

15°C

25°C

35°C

Figure 6. Cumulative nitrogen mineralization with time and temperature for an

organic tundra soil. (After Marion and Black 1987.)
+20°C; they attributed this freezing effect to the

concentration of 

NO2

– into the unfrozen soil water.

Goodroad and Keeney (1984) found that N2O concentrations and fluxes during the spring thaw were

some of the highest values observed during the

entire season. They attributed these high N2O fluxes

to the physical release of N2O upon thawing and

to production by denitrification. Edwards and Killham (1986) found that rates of both denitrification

and NH3 volatilization were increased by freeze–

thaw cycles, especially in the presence of soils previously fertilized with urea (NH2CONH2).

Campbell et al. (1970) found large and sudden unexplained decreases in exchangeable ammonium

during the winter months following a steady fall

buildup. Although they were unable to explain this

phenomenon, possible explanations include enhanced denitrification, NH3 volatilization and NH4

fixation. These few papers indicate that freeze–thaw

cycles affect mineral nitrogen forms, which are the

forms utilizable by plants and also the forms that

are most easily subject to gaseous losses (Fig. 5).

Hinman (1970) found that freeze–thaw cycles

significantly increased the availability of extractable P in soils. On the other hand, Read and Cameron (1979) reported little change in the extractable

P of soils from fall to spring over a ten-year period.

Exchangeable K, Ca and Mg typically either decrease in concentration or are unaffected by freeze–

thaw cycles (Hinman 1970, Cheng et al. 1971). A

decrease in exchangeable cation concentrations

could be caused by the precipitation of soil minerals, which is promoted by soil freezing, causing a

shift in the equilibria (Fig. 4).

Potassium and NH4 are special cases among

exchangeable ions because both are subject to fixation reactions (Fig. 5). Fixation is caused by the collapse of layer-lattice clay minerals, which renders

ions in the interlayers slowly exchangeable. Fine et

al. (1940) reported net releases of K subjected to

freeze–thaw cycles in 75% of examined soils and a

fixation in the remaining 25% of soils. Fine et al.

(1940) also found that most clay minerals (bentonite, nontronite, Putnam clay, montmorillonite)

released nonexchangeable K when subjected to

freeze–thaw cycles. An exception was illite, a nonexpanding layer-lattice clay mineral, which fixed

K. Graham and Lopez (1969) found that severely

K-depleted soils will release K from fixed positions

when subjected to freeze–thaw cycles. On the other hand, when K was added to soils, repeated freezing and thawing led to K fixation.

Manganese and Fe are also special cases among

exchangeable ions because these ions may be

present in both oxidized (Mn4+, Mn3+, Fe3+) and

reduced (Mn2+, Fe2+) forms, with reduced forms

being more soluble in soils (Lindsay 1979). In general, soil moisture saturation promotes reduced

forms and soil unsaturation promotes oxidized

forms. Cheng et al. (1971) found that increasing

soil moisture increased the availability of Mn and

Fe. Under water-saturated conditions, repeated

freeze–thaw cycles increased exchangeable Mn and

Fe, while repeated freeze–thaw cycles in unsaturated soils decreased exchangeable Mn and Fe. The

result for unsaturated conditions agrees with the

general tendency for exchangeable K, Ca and Mg

previously cited. The increased Mn and Fe concentrations under saturated conditions suggests that

freeze–thaw processes may promote reducing conditions in soil, perhaps by excluding O2 gas in the

freezing process or because of slow diffusion of O2

in frozen soils. This hypothesis might also explain

the previously cited increased denitrification under freezing conditions (Christianson and Cho

1983, Edwards and Killham 1986) and the increased

enzymatic activity under reduced conditions

(Linkins 1987). Little work has been done on the

effect of freezing and thawing on respiratory gas

(O2 and CO2) balance in soils. This is a critical factor in controlling both plant and microbial activity

in cold regions soils, with potential significance in

understanding the global carbon balance.

SITE REMEDIATION

There are a number of cases where chemical

composition and freeze–thaw processes interact,

affecting the remediation of problem soils. These

problem soils include saline soils, contaminated

soils and frost-susceptible soils. The effect of soil

salts in reducing frost susceptibility was discussed

previously. Here, our focus will be on natural and

artificial ground freezing as means for desalinating and decontaminating soils.

Kizilova (1959) examined the mobility and solubility of Na+ and 

SO4

2– ions under winter, but generally unfrozen, leaching conditions. The objective

of winter leaching was to remove Na2SO4 from the

soil profile; concern existed that Na2SO4

.

10H2O

(mirabilite) might precipitate at low temperatures.

However, mirabilite did not precipitate even at

subfreezing temperatures, and winter leaching was

effective in removing Na+ and 

SO4

2– ions from

these saline soils. Cary and Mayland (1972) examined a similar, but hypothetical, case. They pointed out that the lower solubility of CaSO4 and

MgSO4 relative to Na2SO4 could cause the former

13
minerals to precipitate more readily upon freezing,

causing an unfavorable Na/(Ca+Mg) ratio in the

solution, which could result in the dispersion of clay

minerals and a sharp decrease in infiltration, problems generally associated with Na-dominated soils.

Actually, at 0°C, epsomite (MgSO4

.

7H20) is four

times as soluble as mirabilite (Richardson et al.

1990). At freezing temperatures, CaSO4

.

2H2O (gypsum) should first precipitate from solution during

freeze-concentration of equal molal solutions, followed by mirabilite and finally epsomite. De Jong

(1981) examined the freeze-purification of saline

groundwater as a source of water for reclaiming

saline soils. The initial 30% of the meltwater had a

higher salt content than the original solution, and

the last 70% had a lower salt content. The results

suggest that freeze-purification of saline waters

could be used in reclaiming saline soils by:

• Collecting and disposing of the initial meltwater and leaching with the later meltwater; or

• Using all the meltwater, which is a procedure

similar to the high-salt-water dilution method

of reclaiming saline soils (De Jong 1981).

Freeze–thaw processes may have both adverse

and beneficial effects on contaminant transport

through soils (Iskandar and Jenkins 1985, Iskandar

1986, Henry 1988). Adverse effects include frost

heaving and reductions in the durability of clay liners, both of which can cause increased leakage from

hazardous waste landfills. Additionally, freezing

causes solute exclusion, which pushes contaminants

out of soils, potentially contaminating groundwaters. Ways in which freezing can be used for beneficial uses include dewatering hazardous materials,

building ice walls to contain contaminants, freezing of entire soil blocks to immobilize contaminants

and freezing to push contaminants out of soils.

Whether solute exclusion of contaminants is a beneficial or adverse effect of freezing depends on the

control of the unfrozen effluent.

Since solute exclusion has been suggested as a

possible mechanism for decontaminating soils

(Iskandar and Jenkins 1985, Iskandar 1986, Ayorinde et al. 1988), factors controlling the process are

critical for evaluating practical applications. A critical factor controlling the exclusion of solutes from

ice is the rate of freezing. Romanov and Levchenko

(1989) noted a marked increase in the effectiveness

of exclusion with a decrease in the cooling rate from

0.5 to 0.25°C/day. These workers also found that

the effectiveness of solute exclusion increased with

increasing solute concentration. In addition, the effectiveness of solute exclusion was much less in a

sand than in pure solutions. Baker and Osterkamp

(1988) also found that the amount of salt rejection

from the freezing region increased with decreasing freezing rate. Solute pockets that initially form

in soils during freezing may ultimately freeze, resulting in alternating bands of high and low salt

concentrations (Romanov and Levchenko 1989). On

the other hand, Baker and Osterkamp (1988) found

no evidence for salt banding in their freezing

experiments. Solute exclusion that leads to alternating bands of high and low salt concentrations

may not be a significant mechanism for macroscale

solute redistribution in soils (Kay and Groenevelt

1983). A more thorough discussion on solute exclusion is given in the section on Chemical transport.

Iskandar and Jenkins (1985) examined the potential use of artificial ground freezing for contaminant immobilization. Frozen metal-contaminated

soils eliminated metal leaching to groundwater.

Freezing the soil from the bottom apparently enhanced upward movement of volatile organics to

the soil surfaces, where losses occurred by volatilization. The amount lost depended on the mobility

of specific organics and ranged from 90% for chloroform, benzene and toluene to as low as 45% for

tetrachloroethylene.

Ayorinde and Perry (1990) examined the utility

of freezing to move explosive compounds through

soils. For one freeze cycle, they found a 40%

reduction in concentration for meta-nitrotoluene

(M-NT) and orthonitrotoluene (O-NT) and less than

a 20% reduction for 2,6-dinitrotoluene (2,6-DNT)

in frozen soil layers compared to unfrozen controls.

At an average freeze rate of 0.4 cm day–1, statistically significant movement was observed for MNT and O-NT but not for 2,6-DNT. They postulated

that for a given freeze rate, number of freeze–thaw

cycles, type of soil and level of soil moisture, the

ability to move contaminants through soils by freezing strongly depends on the type of contaminant,

the initial concentration level and soil–contaminant

interactions.

In another laboratory study, Ayorinde et al.

(1988) showed significant mobility of volatile

organics, such as benzene, chloroform and toluene,

through a Lebanon silty soil frozen from the bottom up. They found a 25–67% reduction in contaminant concentration in the frozen soil but no

corresponding increase above the freezing front.

They attributed this discrepancy to contaminant

losses through volatilization, biodegradation and

sorption.

14
MODELING

Solutes in saline environments play a critical role

in defining freezing and thawing of these systems.

A few mathematical models have been developed

that explicitly consider the effect of solutes on

freeze–thaw processes. These models can be divided into three classes:

• Chemical thermodynamic equilibrium;

• Solute segregation at the water–ice interface;

and

• Solute segregation in soil systems.

Classical chemical thermodynamic models only

provide information on the states of systems at

equilibrium; they do not describe the pathways to

the equilibrium states. Their utility depends on how

well natural systems that are constantly fluctuating can be approximated by the equilibrium states.

Spencer et al. (1990) developed a chemical thermodynamic model (Spencer–Møller–Weare model)

valid over the temperature range of –60° to +25°C,

for the system Na-K-Ca-Mg-Cl-SO4-H2O. This

model is parameterized with data from pure binary and ternary salt solutions. This model predicts

the formation of ice and the precipitation of 15 chloride and sulfate salts, which allows the accurate

prediction of the freezing point and chemical composition of simple and complex solutions, including seawater. Marion and Grant (1994) have

developed a Fortran version of the Spencer–Møller–

Weare model called FREZCHEM. The FREZCHEM

model has two reaction pathways:

• Freezing at variable temperature and fixed

water; and

• Evaporation at variable water and fixed temperature.

15

These two models can accurately estimate the

chemical composition (Fig. 7) and unfrozen water

(Fig. 8) of seawater at least down to –30°C. At lower temperatures some discrepancies exist [e.g., Mg

concentrations (Fig. 7) and unfrozen water (Fig. 8)].

Given the difficulties in making experimental measurements at low temperatures, it is impossible with

the presently available data to determine whether

the model or the experimental measurements are

most accurate at low temperatures (Marion and

Grant 1994). Note that even at temperatures as low

as –50°C, a small fraction of seawater (0.3%) remains unfrozen. The importance of salt precipitation during the freezing process is demonstrated

for Na and Cl salts (Fig. 9), where only 0.3% of Na

and 4.5% of Cl remain in the solution phase at

–50°C.

These papers make it clear that to accurately

interpret or model solute effects during freezing

and thawing of saline solutions such as seawater,

we need accurate chemical thermodynamic data on

mineral solubilities at subzero temperatures.

Apparently the databases for handling chloride

salts down to the eutectic temperature of seawater

(–54°C) are adequate (Spencer et al. 1990, Marion

and Grant 1994). However, there are problems in

modeling some sulfate salts at subzero temperatures, which may be due to poor-quality solubility

data, to lack of data for some salts such as CaSO4

and perhaps to an inadequate solution-phase model

for concentrated sulfate solutions (Bukshtein et al.

1953, Linke 1965, Spencer et al. 1990, Marion and

Grant 1994). The quantity and quality of data for

the bicarbonate and carbonate salts of Na, K, Ca

and Mg at subzero temperatures are even more

sparse (Bukshtein et al. 1953). More research is clearFigure 7. Concentrations of major seawater constituents during freezing.

0

–10

–20

–30

–40

–50

02468

Molality (mol kg– 1)

Temperature (°C)

Cl

Na

Mg

Ca

Experimental Model
1 10 100 1000

Water (g)

Temperature (°C)

0

10

20

30

40

50

Experimental

Model

16

Figure 8. Unfrozen water during

seawater freezing.

0 –5 –10 –15 –20 –25 –30 –35 –40 –45 –50

Temperature (°C)

100

80

60

40

20

0

% of Total Sodium

Solution Na2SO4 NaCl

0 –5 –10 –15 –20 –25 –30 –35 –40 –45 –50

Temperature (°C)

100

80

60

40

20

0

% of Total Chloride

Solution KCl MgCl2 NaCl

a. Sodium.

b. Chloride.

Figure 9. Distribution of sodium and

chloride during seawater freezing.
As this review has shown, many factors with

complex interactions contribute to the differential

movement of water and solutes in freezing soils.

Because of the complexity of solute-freezing phenomena, modeling is an ideal approach for quantifying these relationships. However, quantifying

aqueous-solute properties at subzero temperatures

poses many experimental difficulties. As a consequence, our ability to model these systems is frequently limited by lack of basic data. The models

reviewed in this section, however, do indicate that

significant progress has been made in the past decade, which bodes well for the future.

RECOMMENDATIONS

Much research in cold regions is ultimately

prompted by a quest for solving practical problems.

Five major problems of cold regions associated with

freezing and thawing are soil strength, frost heaving, revegetation of severely disturbed lands, global carbon balance and contaminant transport. This

review has identified several important aspects of

chemistry and freezing–thawing processes that

impact on these problems.

• Chemical thermodynamic equilibria of aqueous electrolyte solutions at subzero temperatures are poorly understood. This is especially

true of the solubilities of Ca, Mg, K and Na

sulfates, bicarbonates and carbonates, important salts in many saline soils and seawater.

Chemical activities and mineral solubilities

play a fundamental role in controlling the

chemical potential of water and ice in soils, the

freezing-point depression and the unfrozen

water content, which, in turn, play important

roles in controlling the fluxes of water and solutes in soils.

• There are strong and complex interactions

among soil properties that control solute and

water flows along concentration, temperature

and hydrostatic gradients in freezing and frozen soils. These complex interactions necessitate the development of computer simulation

models that can integrate physical and chemical properties and processes. The ultimate goal

is to develop models that can accurately address practical problems such as soil strength,

frost heaving and contaminant transport in

freezing and frozen soils.

• Does the freeze–thaw process significantly alter soil gas (O2, CO2, N2O, CH4) concentrations? This has important ramifications for carbon and nitrogen balance; nutrient availability;

microbial, enzymatic and plant activity; soil

ly needed to accurately assess the solubilities of

the sulfate, bicarbonate and carbonate salts of Ca,

Mg, K and Na at subzero temperatures before

these constituents can be accurately integrated into

geochemical models and ultimately into soil physical–chemical models.

Leung and Carmichael (1984) developed a

model to describe solute partitioning between

water and ice during freezing. This model can be

used to estimate partition coefficients from experimental data. Working with slurries, Hanley and

Rao (1982) developed a model based on freezing

potentials that takes into account the diffusion of

cations in the unfrozen portion of the sample, the

separation of ions at the freezing front, and the

migration of moisture and ions towards the freezing front.

Several models have been developed that

describe the dynamics of freeze–thaw processes

and solute redistribution in soils. Mahar et al.

(1983) incorporated solute effects into a variation

of the “Guymon” model to describe freezing of

saline soils. They demonstrated that analytical

modeling of freezing front penetration underpredicts the rate of freezing when no account is made

for the effects of salinity. They also found that the

rate of advance to a given depth of freezing increased with increasing salinity; they attributed

this to the gradual release of the latent heat of fusion over a range of temperatures and depths due

to changing solute concentrations. Osterkamp

(1987) developed an analytical solution for freezing and thawing of soils containing water or

brines. This model used empirical data on seawater brine concentration as a function of temperature to parameterize the model. Osterkamp (1987)

found that maximum ice penetration is greater in

the present of brines, which is similar to findings

of Mahar et al. (1983).

Kadlec et al. (1988) developed a mathematical

model to describe the solute segregation process

at the freezing front and solute transport in the

unfrozen water of peatlands. This model demonstrated the downward movement of solutes,

which is important for establishing the geochronology of deposits and determining the nature of

pollutant burial. Cary (1987) developed a numerical model for calculating frost heave that couples

flows of heat, water and solutes as unsaturated

soils freeze. This model demonstrated, as noted

previously under Frost heaving, that increasing

solutes can decrease frost heaving by reducing

water flows to ice lenses. To my knowledge, this

is the only frost heaving model that explicitly includes solute effects.

17
oxidation–reduction status; and mineral solubility.

• Microbial activity exists to temperatures as

low as –10°C. The ability of microbes to decompose contaminants or plants to absorb

contaminants (immobilization) is poorly understood for low temperatures. Recent work

suggests that the winter “dormant” season

may be a misnomer in describing microbial

activity and carbon balance. Microbes and

plants adapted to cold regions have been subjected to different selective pressures during

evolution and may not respond the same as

their better-studied temperate counterparts.

• What are optimal conditions for solute exclusion in freezing soils? This has important implications for decontaminating soils.

To further our understanding of chemistry and

freeze–thaw processes in soils, future work should

be directed along these lines, with major foci on

both simulation modeling and empirical studies.

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Freeze–Thaw Processes and Soil Chemistry PR: 4A161102AT24

TA: EC and SC

WU: AT25-EC-B03

Giles M. Marion

U.S. Army Cold Regions Research and Engineering Laboratory

72 Lyme Road Special Report 95-12

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This review broadly examines the interactions between freeze–thaw processes and soil chemistry, focusing on

1) the effect of solutes on physical properties such as freezing-point depression, unfrozen water and frost heaving,

2) the effect of freeze–thaw cycles and low temperatures on soil chemistry, and 3) modeling of freeze–thaw processes and chemistry. The presence of solutes causes a freezing-point depression, which increases the amount of

unfrozen water in soils. Liquid films on soil particles provide the dominant route for the flow of water and associated solutes in frozen soils. In general, salts reduce the hydraulic conductivity and water flow to the freezing

front, which reduces frost heaving. Solute exclusion during freezing leads to supersaturated solutions, which promotes the precipitation of secondary minerals in soils. At the watershed level, ionic concentrations in early meltwaters are often 2–9 times higher than snowpack concentrations. Temperature is the dominant factor controlling

decomposition rates, with minimal detectable rates occurring at temperatures as low as –10°C; both bacteria and

fungi are physiologically active at subzero temperatures. Extracellular enzymes are active in soils at temperatures

as low as –20°C; this activity is thought to occur in unfrozen water on surfaces of soil particles. Nitrogen mineralization is reported at temperatures as low as 1°C and is promoted by freeze–thaw cycles. There are strong and

complex interactions among soil properties that control solute and water flows along concentration, temperature

and hydrostatic gradients in freezing and frozen soils. These complex interactions necessitate the development of

computer simulation models that can integrate physical–chemical properties and processes.

28

UNCLASSIFIED UNCLASSIFIED UNCLASSIFIED

Freeze–thaw cycles Soils

Frozen soil Soil chemistry