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III. Stability of the Pore System

III. Stability of the Pore System

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108



ANN P. HAMBLIN



containing entrapped air. Many disaggragation forces act at two scales, the

micro scale of the individual aggregate and the bulk-soil scale, which

spatially extends as far as the particular soil or land use.

Why is it that some soils can be mechanically’deformedmany times when

wet without loss of aggregate stability, yet others collapse very readily, either

with wetting alone or wetting combined with imposed stress? In the first

place, their colloidal behavior differs depending on the proportional clay and

organic matter contents, the clay mineral composition, the colloidal charge

density, the distribution of charge on clay surfaces, and the cation composition. Second, other bonding agents linking charged and uncharged materials

also vary in amount and composition.



B. CLAY AND AGGREGATEBONDINGIN AGRICULTURAL

SOILS

1. Exchangeable Cations and Electrolyte Efsects



Except in single-grain soils, where pedality cannot develop, because of the

small proportion of colloidal materials, matrix structures develop through

the ordering of clay minerals in an hierarchy as follows:

Clay crystals



0.01-0.5pm



Domains

2-5 pm



Microaggregates



Aggregates



5-lo3 pm



i03-i05pm



Complete dispersion and electrochemical stability of clay suspensions only

occurs where watersoil ratios are very much greater than 1, that is, in very

dilute suspension. In the field, therefore, dispersion is rare in subsoils, where

pore volumes are less than 0.5 of the total volume. In surface soils, however,

temporary dispersion may occur in ponded microsites and under conditions

of flood irrigation, perched water tables, or natural flooding. Such dispersed

clay will normally reflocculate within a period of hours and settle as a layer,

or skin, over the larger mineral grains in the soil. Thus what appears in the

field to be dispersion is often the separation of aggregates or microaggregates

into “domain”- and silt-sized material (1- 10 pm). Domains were described

by Aylmore and Quirk (1960) as clustered packages of flocculated clay

crystals, which act as single particles through long-range bonding. Very

stable microaggregates ( c 50 pm) can frequently be separated by mechanical

and chemical dispersion techniques (Edwards and Bremner, 1967). They

possess much greater stability and have different chemical composition,

especially with regard to their organic materials, than do larger aggregates.

Aggregation is enhanced by the presence of calcium cements (Rimmer and

Greenland, 1976) due to the slow solubilization of calcium from carbonate

concretions. In addition, many Australian, African, and South Asian Ultisols

have exceptionally stable microaggregation, which makes the clay behave as



THE INFLUENCE OF SOIL STRUCTURE



109



a “subplastic” material of silt-size. So stable are some of these soils that even

where their exchangeable sodium percentages (ESPs) are as high as 20 % (as

in examples from aeolian clay deposits in Australia), they do not disperse

when wetted and maintain rapid hydraulic conductivities in relation to other

soils which have as little as 5 % exchangeable sodium (McIntyre, 1979). This

pseudo-silt material is in part due to Fe- complexes but has not been fully

explained.

Temporary dispersion of agricultural soils is common when soils containing sodium (Na) and magnesium (Mg) are mechanically worked in a wet

state. This forms the basis of many tests of structural stability, but satisfactory

theoretical explanations have only been developed recently. When calcium

(Ca)-saturated clays are dialyzed free of electrolytes, they remain coagulated

on remolding, unless they have been washed in alcohol, in which case they

will disperse when rewetted (Rengasamy, 1982). Sodium-saturated clays

disperse upon rewetting whether they have been alcohol washed or not.

Rengasamy has explained the alcohol effect on Ca clays by the removal of the

two to three layers of solvation (water) molecules associated with the

hydrated Ca cations. Hydrogen bonding of the hydration shells normally

links the clay crystals in close-packed, Ca-saturated clays. Such water layers

are not removed by normal oven drying, but in Na clays the hydrated shell

size is so much larger that the hydrogen bonding is correspondingly weaker

and easily sheared. Similarly, the hydrated ion size of Mg is larger than Ca

and the resultant hydrogen bonding relatively weak. The ratio of Ca, Mg, Na,

and K (potassium) on the exchange complex is the initial determinant of the

flocculation stability of soil clays. However, work on gypsum-responsive soils

has led to an increased awareness of the importance of the electrolyte

concentration in suppressing dispersion in field soils which have cations

other than calcium in significant proportions on their exchange sites (Loveday, 1974). A “critical” ESP at which dispersion occurs should not be defined

irrespective of the total clay content, soil solution concentration ( C ) , or

sodium adsorption ratio (SAR). In the United States a critical ESP value is

often taken to be 6 %, whereas many Australian soils with higher proportions

of exchangeable Mg disperse at ESPs of only 3 % (Rengesamy et a/., 1984).

Under laboratory conditions clays swell increasingly as the proportion of

sodium on the exchange complex increases and as total electrolyte concentration decreases. Unstable aggregates then fail as differential pressures

within them cause them to rupture. Collapse of larger pores occurs, and K ,

is reduced (Lagerwerff et al., 1969). In Na smectites swelling without collapse

is sufficient to reduce K , considerably (Rowel1 et al., 1969), but in soils with

larger amounts of low-expansion clays, sesquioxides and amorphous materials, the swelling is less severe. Quirk and Schofield (1955) proposed a

“threshold concentration” value for C and the SAR at which laboratory K ,



110



A N N P. HAMBLIN



reduces abruptly ( V K ) through swelling, dispersion, and pore collapse. More

recently, Cass and Sumner (1982) have extended this concept to identify the

factors controlling pore structural stability in the presence of mixed Na-Ca

electrolytes on a wide range of South African, American, and European soils.

They used a “sodium stability model” to correlate a desired function of the

SAR, C, and V K with soil properties. Pedal, nonvertic soils had highly

significant (P < 0.01) correlation coefficients with smectite content, cation

exchange capacity, and specific surface area. Apedal soils and pedal soils of

high organic matter or sesquioxide content did not fit these regressions, nor

did Vertisols until sesquioxides were chemically removed. Vertisols displayed

greater sodium stability than predicted from their smectite content, but this

was due to the unconfined nature of the hydraulic conductivity test. Such

tests show the problems in achieving universality in classifying a wide range

of soils with different bonding agents.

2. Field Expression of the Esfects of Temporary Dispersion

In regions with a widely fluctuating water regime, the SAR varies in any

one soil between seasons. Threshold values for identifying dispersive soils

thus differ between different climatic regimes. Rengesamy et al. (1984) have

distinguished two separate dispersion SAR values which apply to Rhodoxeralfs in South eastern Australia: the first is at the level of spontaneous

dispersion, which occurs if dry aggregates are wetted without mechanical

agitation; the second is for a dispersion procedure which includes agitation

and represents the conditions encountered when soils are subjected to

raindrop impact and tillage. This distinction is based on a scheme of

Emerson’s (1967) for a test of aggregate stability. Many aggregate stability

tests still rely on the behavior of screened ( < 2 mm), dried aggregates to a

variety of wetting, wetting plus mechanical, or wetting plus chemical stresses,

despite their empirical and arbitrary nature. Emerson’s test differentiated

between long-range bonding mechanisms responsible for microaggregate

stability and short-range forces responsible for domain stability and flocculation. Eight categories were distinguished, starting with soils which are totally

stable, remaining unswollen, unslaked, and undispersed when immersed dry

into distilled water, and proceeding through those which merely slake in

water and those which disperse after remolding at $ = - 10 kPa, to those

which disperse completely when dry aggregates are placed in water. Many of

the arable soils used for intensive cropping in Eastern England (Greenland el

al., 1975) and for wheat production in southern Australia (the red brown

earths Haploxeralfs and Rhodoxeralfs) fall into category 3, the class in which

dispersion occurs after remolding. Such soils are subject to dispersion if

cultivated when wet or rewetted by heavy rain after tillage, mechanically



FIG.3. Coatings of clay stained with iron compounds having an extinction pattern characteristic of continuous orientation on the walls

of tubular voids. (a) Plain light; (b) crossed polarizers; x 24. (c) Plain light; (d) crossed polarizers; x 62 (Courtesy J. R. Sleeman, Division of

Soils, CSIRO, Canberra, Australia.)



112



ANN P. HAMBLIN



disrupting newly formed peds. Such crusts may have little mechanical

strength and may form no impedance to shoot emergence unless air-dried,

but their influence upon infiltration can be profound (see Section IV,A,l).

Micromorphological observations of soil crusts frequently reveal finelayered structures of graded silt-sized material blocking larger pores (Falayi

and Bouma, 1975; Chen and Banin, 1975). Clay illuviation (Lessivage) is also

a well-defined pedogenic process, being recognized as a criterion for identifying the argillic or B, horizon in the U.S. “Soil Taxonomy” (Soil Survey Staff,

(1975). Alfisols and Ultisols both contain argillic horizons, with the diagnostic criterion of an increase in clay content of at least 1.2 in the top 0.3 m; in

many instances the clay content doubles. These soils are generally composed

of low-activity colloids such as kaolinite, goethite, and hematite. Pores lined

with clay in this way may be stabilized by root exudates (Greenland, 1978).

Figure 3 shows thin sections under plain light (and with crossed polarizers) of

an upper horizon (at 0.17 m) of a red podzolic soil (chromic Paleudalf)

described by Brewer and Walker (1969), in which illuviated clay has

completely filled the smaller voids and has been deposited as cutans around

the walls of larger voids. The authors tentatively attributed this deposition to

human disturbance of the profile since it was a marked thin layer distinct

from the normal progressive illuviation in a chronosequence.

Clay illuviation which results from tillage and traffic working of wet soils

has been considered responsible for the deteriorating conditions of many

European arable soils (Greenland, 1977). The process is seen in its most

exaggerated form in the creation of paddy fields for wet rice cropping, in

which the puddling is deliberate. In temperate environments, where much

harvesting, tillage, and seeding must be done when soils are close to their

plastic limit (the minimum water content at which the soil deforms plastically), unintentional structural degradation is widespread and has been

identified as the cause of substantial crop losses in wet seasons (Strutt, 1970).

Surface and subsurface structural instability which result from tillage and

traffic cannot strictly be dissociated from the compaction, shearing deformation, and soil inversion which occur at the same time. These alterations are so

substantial that they have quite rightly been recognized by soil taxonomists

as deserving a special “anthro” horizon (Soil Survey Staff, 1975) and by most

modelers of crop water balance (Greacen and Hignett, 1976: Hanks and

Puckeridge, 1980), who allow for a surface soil layer with different physical

properties in their models. There are a number of comprehensive works

which describe different tillage operations and resulting soil deformations

(Barnes et al., 1971; M.A.F.F., 1975). Moldboard and rotary implements fully

invert and mix soil layers; disced and tyned blades lift, rupture, and drop

peds; plough pans result from tangential shearing compression with smearing

at the base of the disturbed layer in a ribbed or grooved fashion. Subplough

and wheeling deformations are mainly compressive, with maximum com-



T H E INFLUENCE OF SOIL STRUCTURE



113



paction when operations coincide with soil water contents just below the

plastic limit.



c. ORGANIC MATTERBONDING IN AGRICULTURALTOPSOILS

Organic materials are the most important stabilizing agents of many

topsoils. Other bonding agents, which are frequently associated with organoclay reactions, include hydrous metal oxides and exchangeable cations, but

their role is mentioned here mainly in association with organic matter. In

subsoils, on the other hand, the paucity of organic material increases the

importance of metal, hydroxy, and proton bonding. However, subsoils do not

normally experience the magnitude and frequency of stress deformations to

which surface soils are subjected.

Organic materials may be classified according to chemical composition

(FIaig et al., 1975; Hayes and Swift, 1975), age and humification (Jenkinson

and Rayner, 1977; Kononova, 1975), or strength of clay-humic bonds

(Edwards and Bremner, 1967; Mortland, 1970). Clay-organic complexes

result in very stable colloidal materials in which the humic substances are

greatly protected from microbial degradation and may be very durable, with

residence half-lives of 2000 years (Jenkinson and Rayner, 1977). These

complexes can account for more than 50% of the organic carbon in the

topsoil. Their bonding mechanisms and functional forms have been reviewed

by Greenland (1965, 1971) and Tate and Theng (1980). Humic substances,

which form the organic part of these complexes, are inportant ion exchangers,

with high exchange capacities (100 to 300 me4100 g at pH 7.0). However,

they are amphoteric (this is, their charge is pH dependent), and at soil pH

values the majority of their charge is negative. The charge density is similar to

that of soil clays, between 1.0 and 2.0 per square nanometer.

About 10 % of humified organic matter is present as polysaccharide, much

of which is derived from relatively recent plant material (Cheshire, 1977). As a

result, the amount and composition of soil polysaccharide fluctuates within a

single soil depending on season and management. Polysaccharides have

received substantial attention despite their relately low abundance because

they have been repeatedly implicated in aggregate stability, although the

evidence is often circumstantial. Greenland et al. (1962) used dilute periodate

oxidation and borate hydrolysis to extract soil polysaccharides. They found

much greater instability in long-cultivated soils than in pasture soils after

extraction. However, periodate-sensitive materials are of less significance to

the stability in younger soils containing more reactive (2: 1 and interstratified)

clays, in which polyvalent cation bonding of clay-organic complexes is often

more important (Hamblin and Greenland, 1977).

Jenkinson and Rayner (1977) distinguished five separate soil organic

categories on the basis of age and degree of decomposition. The respective



114



ANN P. HAMBLIN



half-lives vary in different environments according to the number of biologically active days per year, but the fractions have comparable equivalents in

other soils. Their “decomposable” and “resistant” plant materials (1.2 and

2.3 years, respectively, at Rothampsted, England) are equivalent to the “light

fraction” separated by Ford et al. (1969) using density fractionation. This

material is proportionally highest in acid or cold environments, where the

“microbial fraction” (1.7 years) is lowest. Physically stabilized but “noncomplexed” organic materials which form the fourth group (50-year half-life)

may be the group most vulnerable to oxidative loss when microaggregates

are ruptured by tillage, raindrops, or stock-treading of wet soil. Such noncomplexed organic materials have considerable significance to the physical

behavior of topsoils. Hamblin and Davies (1977) demonstrated that reduction in total organic matter in long-cultivated silty soils reduced the available

water-holding capacity and plastic limit but increased compactibility and

shear strength in comparison with paired soils of higher organic matter

content.

Oades and his co-workers have drawn particular attention to the interaction of roots and the rhizosphere environment, including pore stabilization

and aggregate formation. Turcheneck and Oades (1979) found that microaggregates were stabilized by aromatic humic substances and polysaccharides

of microbial origin from the rhizosphere. Tisdall and Oades (1979) presented

elegant confirmation of the function of the external hyphae of vesicular

arbuscular myccorhiza and root hairs in aggregation. Thus the initial “static”

view of bonding in soil crumbs suggested by Emerson (1959) has become a

more dynamic concept, in which aggregate development, stabilization, and

breakdown are intimately linked with the growth, decay, and subsequent

cultivation of roots in the microbially active topsoil. All cultivation tends to

reduce total organic matter by chemical and microbial oxidation, and

improvement of soil structural stability under short- and long-term pastures

has been widely documented. Alternatively, conservation tillage practices,

particularly where the use of stubble and other mulches encourages the

development of much greater macrofaunal activity, can counteract the

inevitable structural degradation of intensive cropping.



IV. WATER FLOW IN AGRICULTURAL SOILS

A. INFILTRATION



In the field we normally assume vertical infiltration from rainfall (or

rainfall-like sprinklers) and from ponded irrigation. In rigid porous media,

the infiltration rate I (the volume flux of water per unit area) will be less than

the material’s capacity to receive it until I = K , at the surface at time t, after



THE INFLUENCE OF SOIL STRUCTURE



115



which the process is controlled by sorptivity S, the rate at which water is

moved or redistributed away from the surface through the body of the

material. Early attempts at modeling the infiltration process used empirically

derived constants to link I with S (Green and Ampt, 1911). Later Philip

(1 957a,b) provided comprehensive mathematical solutions for idealized

infiltration. At large values of t, cumulative infiltration i was described in

terms of a power series of coefficients derived from K(8) and D(8), the soil

water diffusivity; which is defined as:



D(8) = K(B) d$/d8



(3)



where D is the product of K(B) and the reciprocal of the $(8) curve for any

value of 8 and has units of area per unit time (m2/s). For practical purposes

and short time intervals, Philip’s equation can be approximated by

1 = St’i2



(4)



The “sorptivity” S is the slope of cumulative infiltration against Jt. It

provides us with a convenient parameter of the soil’s absorption rate in units

of length per square root of time (m s-’”). S varies, however, both with the

initial water content Bi and with surface stability. Indeed, lack of variation inS

in any one soil is an excellent measure of stability. Philip’s work gave

solutions for vertical infiltration in situations where both suction and gravity

were operating. The boundary condition was for ponded water at the soil

surface.

More recent theories for nonponding infiltration at a constant rate also

assume a vertically uniform soil with uniform Oi, ignoring the effect of

raindrop impact on soil stability, and hysteretic effects, and assume air ahead

of the wetting front to be at atmospheric pressure. The rainfall rate V, is

assumed to be a constant. The basic assumption is that soil-water flow can be

described in terms of D(8) and K(8):

V(8,t ) = -D(O)(dO/dz



+ KO)



(5)



where V(8, t ) is the Darcy velocity of water at some position in the soil where

the water content is 8, at time t.

Because D(8) and K(8) are not easily measured and have high CVs, as we

have seen in Section II,A,3, the more sensible approach, which is now being

used in field conditions, is to simplify the theory and describe D(8) and K ( 8 )

in terms of S and K,. The conversion of K(8) to a scaled value of K, was

suggested by Brooks and Corey (1964):



~ ( 0 =) ~ s C ( e s- ei>/(es (6)

D(8) is expressed in terms of an exponential function of 8 (Brutsaert, 1979):



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III. Stability of the Pore System

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