Introduction
Soil is eroded as surface water flows over bare or partially vegetated hillslopes. The quantities
and rates of erosion depend on the resistance of soil particles to detachment and the transport capacity of
the runoff. Soil particle detachment varies with inter-particle friction, bonding, and interlocking, and the
sediment transport capacity of runoff is a function of velocity and turbulence. The capability of a soil to
resist erosion depends on soil-particle size and distribution, soil structure and structural stability, soil
permeability, water content, organic content, and mineral and chemical constituents (Lal and Elliot,
1994). Low-density soils are more readily eroded, while compaction generally increases erosion
resistance. Because of significant seasonal changes in both soil density and moisture, Pall et al. (1982)
proposed erodibility as a time-varying rather than static soil characteristic. Thus, a combination of many
factors determines the volume of sediment eroded during a runoff event. Because of this complexity,
coefficients of the rill erosion equation used in the WEPP model have been found to vary by one to two
orders of magnitude (Elliot et al., 1990).
During periods of decreasing air temperature, heat is lost from the soil surface. When sufficient
heat is lost, water in the soil begins to freeze. The net effect of this ice formation on soil structure
depends on soil type, water content, and intensity of the freezing. Freezing and thawing of soils cause
movement of water and solutes in the profile (Radke and Berry, 1997; Gatto, 2000). Water can move
upward toward the freezing front, allowing ice lenses or layers to form. Three conditions must exist for
ground ice to grow and become a substantial component of a soil mass: an adequate supply of soil
moisture, sufficiently cold air temperatures to cause heat loss and freezing, and a frost-susceptible soil
with a significant silt component (Anderson et al., 1978). Silts absorb water rapidly because the particles
are small enough to provide high capillary rise and yet large enough so that pore spaces allow quick flow
of water through the soil (Jumikis, 1962). These characteristics promote a rapid increase in moisture
content within the pores near the location of freezing. Water movement through more fine-grained soils
is much slower, and coarse-grained soils do not develop sufficient capillary forces to move soil moisture
to the freezing front. Thus, silts with available soil water are most susceptible to the changes in strength
and erodibility caused by FT cycling. However, Janson (1963) reports that even sand may become frost
susceptible if it is well compacted, and Chamberlain (personal communication) has observed needle ice in
almost any soil type. In addition to soil texture, frost susceptibility depends upon vegetative cover; the
thickness and density of snow cover; initial soil temperature and temperature gradient; air temperature
regime; exposure to the sun; depth to the water table; overburden stress; and soil density (Jumikis, 1962;
Chamberlain, 1981). As ice crystals form within soil voids particles are forced apart, and ice pressure
may compress or rupture soil aggregates. These FT-induced, physical changes affect soil-particle
cohesion, density and strength, surface soil moisture, infiltration capacity, runoff production, and soil-
surface geometry, which, in turn, affect soil erodibility and the erosivity of subsequent runoff.
Benoit and Voorhees (1990) and Kok and McCool (1990) report that soil FT effects are some of
the least understood aspects of soil erosion processes. Several investigators have recognized that FT
generally increases soil erodibility (Bryan, 2000), and that the magnitude of this effect varies with soil
texture, moisture, and extent of freezing. Laboratory experiments of Formanek et al. (1984) found that
the shear strength of a silt loam was reduced to less than half its original value after one FT cycle, but
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