Lal R., Shukla M.K. Principles of Soil Physics
Подождите немного. Документ загружается.
Dryness C.T. (1969). Hydrological properties of soils on three small watersheds in the western
cascades. USD A Forest Service Research Note PNW-111.
FAO (1977). Guidelines for soil profile description. Food and Agriculture Organization of the
United Nations (F.A.O.), Rome, p 66.
Fisher R.F. and D.Binkley (2000). Ecology and management of forest soils. John Wiley & Sons,
New York.
Flint A.L. and S.Childs (1984). Physical properties of rock fragments and their effect on available
water in skeletal soils. In: J.D.Nichols, P.L.Brown and W.J.Grant (eds.), Erosion and
Productivity of soils containing rock fragments. Soil Science Society America, Madison, WI,
Special publication no. 13 p 91–101.
Gornat B. and D.Goldberg (1972). The relation between moisture measurements with a neutron
probe and soil texture. Soil Sci. 114:254–258.
Gras R. and G.Monnier (1963). Contribution de certains elements grossiers a l’alimentation en eau
des vegetaux. Soil Sol. 1:13–20.
Hanson C.T. and Blevins R.L. (1979). Soil water in coarse fragments. Soil Sci. Soc. Am. J. 43:813–
820.
Ingelmo F., S.Cuadrado, A.Ibanez, and J.Hernandez (1994). Hydric properties of some Spanish
soils in relation to their rock fragment content: implication for runoff and vegetation. In:
J.Poesen and H.Lavee (ed.) Rock fragments in soil: surface dynamics. Catena 23:87–97.
Jury W.A. and B.Bellantuoni (1976a). Heat and water movement under surface rocks in afield soil:
I. Thermal effects, Soil Sci. Soc. of Am. J. 40:505–509.
Jury W.A. and B.Bellantuoni (1976b). Heat and water movement under surface rocks in afield soil:
II. Moisture effects, Soil Sci. Soc. of Am. J. 40:509–513.
Lal R. (1974). The effect of soil texture and density of the neutron and density probe calibration for
some tropical soils. Soil Sci. 117(4): 183–190.
Lal R. (1979). Physical characteristics of soils of the tropics: determination and management. In:
R.Lal and D.J.Greenland (eds.) Soil Physical Properties and Crop Production in tropics. John
Wiley & Sons, New York, p 7–44.
Larmuth L. (1978). Temperature beneath stones used as daytime retreats by desert animals. J.Arid
Environ. 1:35–40.
Mansfield J.E. (1979). Land capability for annual rainfed arable crops in Northern Nigeria based on
soil physical limitation. In: R.Lal and D.J.Greenland (eds.) Soil Physical Properties and Crop
Production in tropics. John Wiley & Sons, New York, p 408–437.
McHenry F.R. (1963). Theory and application of neutron scattering in the measurement of soil
moisture. Soil Sci. 95:294–307.
Mehuys G.R., L.H.Stolzey, and J.Letey (1975). Temperature distributions under stones submitted
to a diurnal heat wave. Soil Sci. 120:437–441.
Miller F.T.and R.L.Guthrie (1984). Classification and distribution of soils containing rock
fragments in the United States. In: J.D.Nichols, P.L.Brown, and W.J.Grant (eds.), Erosion and
Productivity of soils containing rock fragments. Soil Science Society America, Madison, WI,
Special publication no. 13 p 91–101.
Montagne C., J.Ruddell, and H.Ferguson (1992). Water retention of soft siltstone fragments in a
Ustic torriorthent, Central Montana. Soil Sci. Soc. Am. J. 56: 555–557.
Moormann F.R., R.Lal, and A.S.R.Juo (1975). The soils of IITA.International Institute of Tropical
Agriculture, Tech. Bull. No. 3, Ibadan, Nigeria.
Nettleton W.E. Gamble, B.Allen, G.Borst, and F.Peterson (1989). Relict soils of subtropical
regions of the United States, Catena Suppl. 16:59–93.
Parsens A.J., A.D.Abrahams, and J.R.Simanton (1992). Microtopography and soil surface materials
on semi-arid piedmont hillscope, Southern Arizona, J. Arid Environ. 22:107–115.
Peck A.J. and J.D.Watson (1979). Hydraulic conductivity and flow in non uniform soil, In:
Workshop on Soil Physics an Field Heterogeneity. CSIRO Div. Of Environmental Mechanics,
Canberra, p 31–39.
Principles of soil physics 574
Poesen J.W. and F.Ingelmo-Sanchez (1992). Runoff and sediment yield from topsoils with different
porosity as effected by rock fragment cover and position. Catena 19:451–474.
Poesen J.W. and H.Lavee (1994). Rock fragments in top soils: significance and processes. In:
J.Poesen and H.Lavee (ed.) Rock fragments in soil: surface dynamics. Catena 23:1–28.
Poesen J.W., D.Torri, and K.Bunte (1994). Effects of rock fragments on soil erosion by water at
different spatial scale: a review. In: J.Poesen and H.Lavee (ed.) Rock fragments in soil: surface
dynamics. Catena 23:141–166.
Poesen J.W., F.Ingelmo-Sanchez, and H.Muler (1990). The hydrological response of soil surfaces
to rainfall as effected by cover and position of rock fragments in the top layer. Earth Surf.
Process, Landforms 15:653–671.
Ravina I. and J.Magier (1984). Hydraulic conductivity and water retention of clay soils containing
rock fragments. Soil Sci. Soc. Am. J., 48:736–740.
Soil Survey Staff (1951). Soil Survey Manual. Agric. Handb. No. 18, USDA, U.S. Government
Printing Office, Washington, D.C.
Soil Survey Staff (1993). Soil Survey Manual. USDA-SCS Agric. Handb. No. 18, U.S.
Government Printing Office, Washington, D.C.
Stewart V.I., W.A.Adams, and H.H.Abdullah (1970). Quantitative pedological studies on soils
derived from Silurian mudstones II. The relationship between stone content and apparent
density of the fine earth. J. Soil Sci., 21:248–255.
Takijima Y. and H.Sauma (1967). The classification of paddy soils based on their suitability for
reclamation. 2. Effect of sand and gravel content of the soil on the development of the root
system of rice plants. J. Sci. Soil Manure, Tokyo 38: 313–318.
Torri D., J.Poesen, F.Monaci, and E.Busoni (1994). Rock fragment content and fine soil bulk
density: In: J.Poesen and H.Lavee (ed.) Rock fragments in soil: surface dynamics. Catena
23:65–71.
Unger P. (1971). Soil profile gravel layers: I. Effect on water storage, distribution, and evaporation.
Soil Sci. Soc. Am. Proc. 35:631–634.
Unger P.W. (1971). Soil profile gravel layers. Effect on water storage, distribution and evaporation.
Soil Sci. Soc. Am. Proc. 35:631–654.
Valentin C. (1994). Surface sealing as effected by various rock fragment covers in West Africa: In:
J.Poesen and H.Lavee (ed.) Rock fragments in soil: surface dynamics. Catena 23:87–97.
Webster R. (1985). Quantitative spatial analysis of soil in the field. Adv. Soil Sci. 3:1–70.
Physical properties of gravelly soils 575
20
Freezing and Thawing Effects, Swelling
Soils, and Other Special Problems
The theory and experimental verification of isothermal moisture storage and movement
in soils are well developed. Some of the processes such as freezing and thawing, presence
of macropore flow channels, swelling and shrinking of soils, and salinity complicate the
flow and transport processes through soils by modifying the physical and water
transmission properties of porous media. It is generally recognized that during freezing
and thawing processes, soil moisture and thermal states of a soil system are coupled.
Inadequate understanding of this complex problem of phase changing processes during
freezing and thawing and sparsity of development in regions where these processes are
significant are some of the reasons why enough progress has not been made in this field.
This chapter discusses some of the mathematics of coupled heat and moisture transport
models as well the effects of freezing and thawing on soil’s physical properties. The
modeling of macropore flow and its effect on infiltration and solute transport through soil
profile, characteristics, and water flow in swelling soils, and pressure potentials in saline
soils, are also discussed. Another important topic covered in this chapter is water-
repellent soils. This is an important research topic to which an increasing attention has
been paid in view of its. This chapter describes hydrological processes in water-repellent
soils.
20.1 FROZEN SOILS
Freezing and thawing can have a profound influence on the stability, hydrology,
chemistry, biology, and ecology of soils. Under frozen conditions the soils become less
permeable to rainfall/snowmelt, water within the frozen part of the soil profile becomes
immobile and therefore unavailable for leaching, chemicals within the soil profile are
redistributed due to the presence of temperature gradients and nonuniform freezing, and
microbiological activity is reduced due to lowered temperature. The mass and energy
balance conditions within the soil profile are also greatly affected by the existing surface
cover conditions. For instance, the presence of a snow pack can increase the intensity of
microbiological activity within the soil profile due to the insulation of the soil by the
snow.
Laboratory and field experiments can provide useful information about the winter
processes within the soil profile. However, experimenta-tion, especially under in situ
conditions, is very difficult. Freezing processes in soil are also complicated, and use of
simulation models can be advantageous. The processes of coupled water flow, thermal
energy transport, and solute transport within a variably saturated, variably frozen soil
profile can be described by a set of coupled conservation (mass and energy) equations,
appropriate constitutive laws (Darcy’s, Fick’s, Fourier’s), and equilibrium
thermodynamic relations. A number of numerical solution models have been developed
to solve these systems of equations, with some methods imposing numerous simplifying
assumptions, while others solving the full set of equations.
20.1.1 Frozen Soil Composition
Soil moisture conditions influence water and contaminant migration in frozen soils.
Frozen soils consist of four main phases including air, unfrozen water, ice, and soil
particles. The four volumetric components of frozen soil are illustrated in Fig. 20.1. The
conversion of water into ice results in increase in volume. The state of water in soil water
systems for ice-free soil can be expressed as the matrix pressure of water and air (Φ
mwa
)
as follows (Black, 1990)
Φ
mwa
=P
w
−P
a
(20.1)
where P
w
and P
a
are soil water and air pressures. Similarly the state of water in an air-free
frozen soil can be explained as
Φ
miw
=P
i
−P
w
(20.2)
FIGURE 20.1 Schematic of
volumetric distributions of frozen soil
components where V
a
, V
w,
Vi, and V
s
represent volumes of air, unfrozen
water, ice, and soil particles,
respectively.
Freezing and thawing effects 577
where P
i
is pressure of ice. The state of water in air-free frozen soils is often expressed in
terms of temperature (T). For a solute-free soil, the Clapeyron equation relates the ice and
water pressures to temperature, specific gravity and volumetric latent
heat of fusion (L
0
) as follows (Black, 1990)
(20.3)
20.1.2 Soil Freezing Process
Freezing process in soil can be explained as follows:
1. Water begins freezing when soil temperatures are below the porewater freezing point.
2. Soil temperatures drop below pore waters freezing temperatures until enough energy
exists to instigate pore-water nucleation.
3. Pore-water temperature then increases to its freezing point.
4. Soil temperatures remain constant at this temperature until all latent heat of fusion is
released.
5. Soil temperatures then decrease if ambient temperatures are below the pore-water
freezing point.
Soil-particle surface tension and capillary forces cause depression of pore-water freezing
temperatures adjacent to the soil grains in saturated soils. This phenomenon is most
evident in fine-grain soils such as clays and silts. The unfrozen water content decreases as
soil temperature decreases. The unfrozen water layer surrounding soil particles acts as a
conduit for contaminant transport. Unsaturated frozen soils contain little or no unfrozen
water content. Therefore, air voids in pore spaces act as conduits for contaminant
transport. This is in part due to a 9% volume increase, which occurs when water freezes
to form ice. Ice in the soil restricts water and contaminant pathways. In general, freezing
process induces both heat and mass transfer from warm regions to cold regions (Harlan,
1973). The water film on soil surface exists in equilibrium with ice at temperatures below
0°C. The film thickness is independent of total water content (water + ice) but dependent
on temperature (Low et al., 1968) and decreases with temperature falling below the
freezing point of water.
20.1.3 Water Flow in Frozen Soils
If soil water characteristic curve is expressed as per Eq. (20.4)
Φ
m
=Φ
0
(θ/θs)
−b
(20.4)
then the hydraulic conductivity (K(θ)) can be represented by
K(θ)=K
s
(θ/θ
s
)
(2b+2)
(20.5)
Principles of soil physics 578
where Φm is the soil water matric potential (cm of water), Φ
0
is the air entry value (cm of
water), K
s
is the saturated hydraulic conductivity (cmh
−1
), θ is unsaturated liquid water
content (cm
3
cm
−3
), θ
s
is the saturated water content (cm
3
cm
−3
), and b is a constant
(Campbell, 1974). When ice is present in the soil pores at atmospheric pressure, the water
potential can be approximated by the vapor pressure of pure ice as follows (Cary and
Mayland, 1972):
1.2*10
4
T=Φ
m
+Φ
π
(20.6)
where Φ
π
is the osmotic potential of the soil solution (cm of water), and 1.2*10
4
is a
factor that approximately converts the temperature of ice T (°C) to total water potential
(cm of water) (Cary et al., 1979). Combining and rearranging Eqs. (20.4), (20.5), and
(20.6) results in the following equation, which describes the unsaturated soil water flow
in a frozen soil system
(20.7)
where J
1
is flow in the liquid phase and z is soil depth. The above derivation assumes that
the water content, including ice, is less than saturation, and that water transport due to
anion exclusion, liquid ice interface phenom-enon, vapor diffusion, and plastic flow of
ice are all negligible. The value of Φπ can be approximated by osmotic pressure and
temperature relationship for the cases where solubility limits are not exceeded as follows
(20.8)
Combining Eq. (20.8) with Eqs. (20.4) and (20.6) provides the value of Φt though
iteration for T less than freezing point of soil solution (Cary et al., 1979)
(20.
9
)
For T < the freezing point of soil solution, Eq. (20.9) converges rapidly to π for an initial
π
0
. The osmotic pressure effects in Eq. (20.4) do not correspond to those observed in clay
soils and biological systems with selective semipermeable membranes. In these types of
systems, water flows from regions of low osmotic pressure to high. When ice is present,
the direction of flow is reversed. Solute increases the amount of liquid phase water,
which flows towards the thinner films away from the regions of high osmotic pressure
(Cary et al., 1979).
Freezing and thawing effects 579
20.1.4 Simultaneous Heat and Fluid Transport
The simultaneous transport of heat and fluid in a partially frozen porous medium requires
the knowledge of interrelationships among the laws of heat and fluid flow, equations of
continuity for mass and energy, and the characteristics of the fluids and medium
involved. The energy state of liquid water at temperatures less than 0°C in equilibrium
with ice is a function of temperature but not of water content (except for very dry
conditions). Hence, ice phase can be assumed to be behaving as a sink (water is added),
or source (water removed from storage). The mass transport equation for one-
dimensional steady or unsteady flow in a partially or fully saturated heterogeneous
porous medium with freezing or thawing can be written as (Harlan, 1973)
(20..0)
where x is the depth of soil under consideration (g), ρ
t
is density of liquid fraction (g
cm
−3
), θ is volumetric water content (cm
3
cm
−3
), K is hydraulic conductivity of soil, T is
temperature (°C), Φ
t
is total head (cm), Φ
m
is matrix or capillary head (cm), and ∆S
t
is
change in ice content (g cm
−3
min
−1
). The total head or Φ
t
is the sum of Gibbs free energy
(Φ
m
+Φ
π
), or soil water potential, pneumatic pressure head (Φ
p
), and elevation head (Φ
z
).
The assumptions made to derive the heat transfer equation are: (i) soil is
homogeneous, (ii) free thermal convection is negligible, (iii) the temperature distribution
is smooth and continuous, (iv) the temperature of the fluid entering the medium is equal
to the temperature along appropriate boundary, and (v) thermal resistance between fluids
and soil matrix is small, and matrix temperatures are equal. Based upon these
assumptions, the one-dimensional steady or nonsteady convection–conduc-tion heat
transport equation may be written as (Harlan, 1973)
(20.11)
where K
T
is the thermal conductivity (cal cm
−1
°C
−1
min
−1
), T is tempera-ture (°C), C
w
is
bulk specific heat of water (cal g
−1
cm
−1
C
−1
min
−1
), v
x
is the fluid flow velocity in x
direction, C
v
is apparent volumetric specific heat (cal cm
−1
°C
−l
), and is defined as
follows:
(20.12)
where C′
v
is the volumetric specific heat capacity (cal cm
−l
°C
−1
) and is defined as the
sum of specific heat of the soil material, liquid water fraction, and ice fraction; L
0
is
latent heat of fusion (cal g
−1
); θ
s
is volumetric ice fraction (cm
3
cm
−3
); and ρ
ice
is the ice
density (g cm
−3
).
Principles of soil physics 580
20.1.5 Effect of Freezing and Thawing on Soil Physical Properties
The complexity of the hydrologic cycle is increased during winter conditions when the
temperature of the soil reaches the freezing point. Winter can be visualized as a static
period regarding the movement of water from precipitation through the soil profile and
into ground water recharge. Factors that affect infiltration are hydraulic conductivity,
saturation, soil makeup, and porosity. However, under frozen conditions, temperature
becomes the primary factor in determining a soil’s permeability. The contradictory
accounts on effect of freezing and thawing on soil’s physical properties are available.
Frost causes a breakdown of soil’s physical properties, i.e., decreased soil aggregation
and increased bulk density (Bisel and Nielsen, 1964, 1967; Leo, 1963). Frost action can
also improve soil’s physical properties, e.g., increase aggregation, and reduce bulk
density (Sillanpaa and Weber, 1961). In addition, changes in saturated water holding
capacity, hydraulic conductivity, and soil bulk density as a result of freeze-thaw activity
are significantly related to the type of soil, initial water content, degree of aggregation,
initial bulk density, degree and rate of freezing, and number of cycles. (Mostaghimi et al.,
1988; Benoit and Voorhees, 1990) Specific effects of freeze-thaw on soil properties are
described below.
Aggregation
Aggregate stability, a measure of a soil aggregate’s resistance to breakdown, influences
many soil physical and hydraulic characteristics, such as surface sealing, infiltration, and
hydraulic conductivity. Freeze–thaw cycles influence soil aggregation. The water
contained in the pores of the soil expands on changing to ice, and freezing of water inside
the soil pores affects soil aggregation. In coarse-textured soils, water freezes in situ,
whereas in fine-textured soils, water moves towards the freezing sites and forms ice
lenses. Because of the suction, the water is withdrawn from the areas where it is not
frozen and compression and heaving of the soil occurs as ice lenses grow (Young and
Warkentin, 1975). The compression of drier soil near the lenses leads to aggregation in
fine-textured soils. The freezing and thawing can also cause crumbling of cloddy soils. It
can produce cracks in the soils, which may open further during subsequent freezing and
thawing cycles (Richardson, 1976). Aggregate stability is inversely proportional to soil
water content at the time of freezing, constraint to expanding, and number of thaw-wet
cycles (Bullock et al., 1988; Mostaghimi et al., 1988; Lehrsch et al., 1990). Aggregates in
otherwise poorly aggregated soils become stable when frozen at intermediate water
contents.
Infiltration and Hydraulic Conductivity
The knowledge of water infiltration in frozen soils is important for water conservation,
runoff, or flooding and erosion in areas where freezing temperatures occur. The ice
content of soil and water infiltration rate is inversely related (Granger et al., 1984;
Thunholm and Ludin, 1990). The ice lenses on soil surface or at shallow depth impede
Freezing and thawing effects 581
water infiltration. Vertical freezing and thawing decreases soil hydraulic conductivity and
is related to water content, initial aggregate size, and freezing temperature. The greatest
decrease in hydraulic conductivity values in repacked soil columns is associated with
larger aggregate size and higher water content upon freezing (Benoit, 1973).
20.2 MACROPORE FLOW
The biological, chemical, and physical processes and their interactions in soils are
fundamental to the formation of preferential or macropore flow paths. In contrast to flow
through homogeneous soil profile, in which water penetrates the entire porous network of
soil, preferential flow occurs through distinct pathways, which constitute a small fraction
of total soil pore space but can conduct as much as 80% of water percolating through soil
profile. The macropore channels may be randomly or systematically distributed,
reflecting the influence of land use practices. Preferential flow paths have a significant
influence on water and solute transport in the soil. This influence is usually reflected in
reduced travel time, increased volumes of water and solute concentrations in drainage
water, or deeper penetration of chemicals into the soil profile, as compared to
corresponding predictions from conventional flow theory (Beven and Germann, 1982).
The mechanics of flow and transport through capillary pores in the soil matrix is well
documented, as this phenomenon has always been considered to be the dominant flow
process. However, since the mid 1970s, there has been an increase in the number of
research studies confirming the importance of macropore flow in water and solute
transport (McCoy et al., 1994), and several simple and complicated (able to account for
the spatial variability of macropores into the description of water and solute movement in
soils) models have been developed.
20.2.1 Preferential Flow Paths
Some of the terminologies used to describe macropore flow are confusing. Some
commonly used terms are subsurface storm flow short-circuiting, macropore flow, bypass
flow, channeling, fracture flow, preferential flow, or fingering (Bouma and Dekker,
1978). The preferential flow paths are often termed as macropores, which suggests that
size is a sufficient criterion to define them. In practice, however, it is difficult to put a
size limit on soil voids that function as macropore channels. The sizes ranging from 0.03
mm to 3 mm are being used as the lower limit for the equivalent diameter of a macropore
channel (Beven and Germann, 1982). The ability of a soil void to contribute to short
circuiting is dependent not only on size, but on pore structure and continuity, initial soil
water content, and water application flux (Beven, 1982). However, the flow dynamics
through macropore channels are more important than their size, and any size limit used in
a particular application needs to be explicitly stated. Preferential flow paths have been
traditionally separated in classes based on how closely they conform to a circle in cross
section (Bouma et al., 1978). A more functional system of classification is based on the
mode of formation and the persistence of a preferential flow path. According the
morphology, macropores are classified into four groups: (i) pores formed by burrowing
animals, (ii) pores formed by plant roots, (iii) cracks and fissures, and (iv) natural soil
Principles of soil physics 582
pipes (Beven and Germann, 1982). The first two groups are collectively called biopores,
and the last two, cracks. Brief description of these two categories follows.
Biopores
Biopores are formed as a result of the interactions between the soil and the soil biota.
This interaction is usually limited to the upper soil layers (the A and B horizons) in
mineral soils. Biopores consist of decayed root channels, living roots, wormholes, insect
burrows, and the burrows of small rodents, with the smaller holes being more prevalent.
They tend to be roughly cylindrical in shape and can extend up to 1/m depth (Edwards et
al., 1988). The void forming animals are classified as fossers who are equipped with
excavating limbs; miners who use their jaws to bite soil particles off of aggregates,
leaving the aggregates with jagged and pitted surfaces; and tunnelers that push or eat their
way through the soil (Hole, 1981). On an average, small soil animals form holes from 10
mm in diameter and up to 1 m to 2 m deep, and these channels tend to be very stable
(Hole, 1981).) Wormholes have no stress concentrations on their edges and as such do
not change shape in response to moisture content changes in the soil matrix (Dexter,
1978). In general, biopores are very persistent or stable and have relatively high transport
capacities. The size and frequency of biopores, which result from the action of soil flora,
are dependent mainly on plant species, on soil water status during the plant-growing
season, and on agrotechnical practices. Roots also increase macropore densities by
growing into microcracks and enlarging them.
Cracks
Cracks in a soil system are formed by the shrinkage in clayey soils or by chemical
weathering of bedrock material (Fig. 20.2). They can be formed by methods of
cultivation as well as freeze-thaw cycling. When clay soils, particularly those rich in
montmorillonite or other smectitic clays, dry out, they usually form desiccation cracks.
These soils are so widespread that soil cracking has been used as a differentia at the
highest level of soil
Freezing and thawing effects 583