Physical properties of tilled soils

ELSEVIER Soil & Tillage Research 43 (1997) 41-63

soil& Tillage Research

Physical properties of tilled soils

A.R. Dexter * Silsoe Research Institute, Wrest Park, Silsoe, Bedford MK45 4HS, UK

Revised 8 December 1996

Abstract

The key areas of soil physics: structure, water, aeration, and strength and stability are discussed. Recent developments in experimental techniques and theory are reviewed. In addition, some older research is discussed where it is thought to be worthy of re-examination. It is concluded that for studies of tilled and other structured soils, measurements of bulk density and gravimetric water content are not ideal, and a system based on specific volumes and voids ratios is proposed. Soil structure or spatial heterogeneity dominates all the physical properties of soil and hence, its functioning. For research of agriculture and environmental problems, it is essential to characterize the soil structure if we are to develop the improved understanding of soil which is necess‘ary for its proper management. 0 1997 Elsevier Science B.V.

Keywords: Soil aeration; Soil water; Soil strength, Soil stability; Soil structure; Soil physics

1. Introduction

It is not easy to define exactly what are physical properties in contrast with chemical or biological properties. There is no clear boundary between these different disciplines. The boundaries which are often imposed are artificial and are, at best, an administrative convenience. These different disciplines cannot usually be identified uniquely with discrete sets of phenomena or problems in the real world. All we can say is that soil physics tends to be what physicists do when studying soil, soil chemistry tends to be what chemists do when studying soil, and soil biology tends to be what biologists do when they study soil. Phenomena and problems in real soils usually require more than one of these components for their adequate study.

* Corresponding author.

0167-1987/97/$17.00 0 1997 Elsevier Science B.V. All rights reserved. PII SO167-1987(97)00034-2

42 A.R Dexter/Soil & Tillage Research 43 (1997) 41-63

Nevertheless, there are some areas of soils research which are more physical than the others and which provide the central core of the subject which we call soil physics. For the present purposes, we can use the definition of soil physics given by Dexter and Young (1992):

Soil physics deals with the energy status of the different phases (soil, liquid and gas) of the soil system and aims to quantify the fluxes of these phases that are produced by energy gradients. Central to soil physics is the study of the energy state and flux of water in relation to the spatial heterogeneity of the different phases. This heterogeneity is usually referred to as soil structure.

However, soil physics is concerned with much more than water. Fluxes of heat and gas through the soil surface and within the soil are also considered. Mechanics is the oldest branch of physics and in the case of soil mechanics deals with how soil systems change their size, shape and heterogeneity in response to different imposed mechanical potentials.

This rather short review attempts to describe some recent developments in measurement techniques which are relevant to the characterisation of the physical properties of highly-structured, tilled soils.

2. Structure

Soil structure is the key to the functioning of soil. Many key aspects of soil structure have been reviewed by Dexter (1988) and Kay (1990). Except in the coarsest of soils, it is because of the structure that water and gases can move readily through the soil and that aerobic life can occur within the soil. Yet many soil scientists measure only average or bulk soil properties which give little or no information about its structural state.

It is conventional to measure the dry bulk density, y, and the gravimetric water content, w, of soils. However, these are not good quantities to use when thinking about the physical status or structure of soil or when developing equations for its state or behaviour. These two quantities do not have a common basis and are not additive. It is much more rational to use a system where the quantities of the different phases are additive and where the denominator is constant. Such a system is provided by the use of specific volumes and voids ratios. We can write the total volume of a soil sample as

I& = v, + VW + v, (1) where V,, V, and V, are the volumes of solids, water and air respectively. It is then possible to render this dimensionless by dividing by the volume of solids to obtain the specific volume as

R,,,,=l+R,+R, (2) where R, and R, are known as the water ratio and the air ratio, respectively.

A.R. Dexter/Soil & Tillage Research 43 (1997) 41-63 43

It will be noted that

Rsoil = 1 + e (3)

where e is the usual voids ratio which is used in soil mechanics. What we have done is simply to split the total voids into water-filled and air-filled voids. In terms of the conventionally-measured parameters,

R,=wh,and PW

(4)

where ps, pw and y are the densities of the soil solids, water and of the total soil, respectively.

There are numerous advantages in using Eq. (2). Firstly, when a non-swelling soil is wetted, water displaces an equal volume of air and

~~~~~ = constant, and AR, = -AR, (6)

Secondly, with a swelling soil, it can be assumed that as a first approximation, R, does not change much with water content in the normal-swelling range, and therefore that

AR,,, = AR, (7) The swelling and shrinkage characteristics of soils are given by plots of e against R, (Crescimanno et al., 1995). Smiles (1974) has used water ratios in the analysis of water infiltration into swelling soil. Thirdly, if a soil is compressed (compacted) with a reduction in its air-filled pore space and without expulsion of water, then

ARsoi, = AR, (8) Eq. (2) can be further extended by splitting R, into the air ratios for cracks, R,;

worm holes, R,; inter-aggregate pores, R,; intra-aggregate pores, R,; etc. R, can be similarly sub-divided. All of these sub-categories are additive.

In this way, the usual bulk properties of soil can be transformed and extended into a single system which contains useful structural information. Because of the large-scale macrostructure of tilled soils it is essential to use large samples to measure the bulk density accurately. The measuring frame of 0.707 X 0.707 m described by H%ansson (1990) is ideal for this purpose.

Structure will not be reviewed further here because it requires a full review to cover its different aspects. However, several important aspects are covered in the following sections on water, aeration and on strength and stability.

3. Soil water

3. I. Statics

When there is no gradient in total water potential, there is no flow of water and the situation can be described as static. Water contents in the field are often changing only

44 A.R. Dexter/Soil & Tillage Research 43 (1997) 41-63

slowly and the water is at or close to its static, equilibrium or minimum free energy state. Soil water needs to be characterised in terms of its content and its potential. The relationship between content and potential is known as the water characteristic.

Water content can be measured in various ways, and it can be expressed on a dry weight basis as the gravimetric water content, w; on a volumetric basis (volume of water per unit volume of moist soil), 8; or as a water ratio (volume of water per unit volume of soil solids), R,, as defined in Eq. (2).

The traditional method of drying the soil in an oven at 105°C until no further weight change occurs can be considered to be the standard method for measuring gravimetric water content, w. This can then be converted to

8= wy,or PW

R,zwP, PW

(10)

as required. This and other methods become increasingly difficult when measurements are required on small samples in hot, dry climates because significant water losses can occur very rapidly during handling and weighing.

In the field, the neutron probe provides a good method for monitoring water content in the subsoil. It is of little use in tilled top-soils because it is sensitive to protons (hydrogen nuclei) within a radius of about 0.15 m in wet soil and about 0.3 m in dry soil. It is also influenced to a limited extent by the presence of some other nuclei. Therefore, within the normal tilled layer, the readings are affected by the proximity of the soil surface. Since organic matter in addition to water contains a high density of protons, the readings are also influenced by soil organic matter content, and for accurate work the probe needs to be calibrated for each horizon of each soil. Measurements are non-destructive (after the initial installation of aluminium access tubes), but the method is time consuming, labour intensive, and has safety problems associated with its use.

In recent years, electromagnetic or dielectric methods have been used increasingly. These include time-domain reflectometry or TDR (Whalley, 1993) capacitance probes (Whalley et al., 1992) and a standing wave method (Fig. 1). These are all based on the concept that water has a higher dielectric constant (or relative permittivity), E, , than soil solids, es. At the usually measurement frequencies (or effective frequencies) E, = 80 and typically E, = 4. Therefore, by measurement of the permittivity of a mixture, such as moist soil, it is possible to estimate the water content. Soil density influences the readings. Calibration is desirable also because different soil minerals (and organic matter) have different values of E,. However, for most mineral soils, the effects of composition are not large. The size scale of the measurements is of the order of 50 mm. Problems can arise in shrinking and swelling soils where air-gaps can form around the electrodes. Air-gaps of only a fraction of 1 mm can have a large effect on TDR readings obtained (Annan, 1977; Whalley, 1993). These electronic devices have the enormous advantages that they can be readily multiplexed and data-logged for continual monitoring of soil water in complex field trials. There are no safety problems.

A.R. Dexter/Soil & Tillage Research J3 (1997) 41-63 45

Fig. 1. A commercially-produced device for measurement of soil water content. The device measures soil water through electrical impedance measurement and produces an output suitable for data-logging. The metal rods are 60 mm long. Photograph by permission of Delta-T Devices. Cambridge, England.

The measurement of manic water potential in the field is easy in principle but is often problematic in practice. Conventional porous cup tensiometers are satisfactory in the range from 0 to about -80 kPa. They can be fitted with pressure transducers and data-logged. De-airing of the filling water (e.g., by boiling it for 20 min) and de-airing the porous ceramic cups by placing them in de-aired water under vacuum can greatly increase the reliability of operation. All air bubbles must be removed or the expansion and contraction of these bubbles with temperature changes will give rise to erroneous pressure readings. The size scale of tensiometer tips ranges from about 50 mm to microtensiometers of about 1 mm (Gunzelmann et al., 1987).

Ridley and Burland (1993) of Imperial College, London have developed a new concept in tensiometry. They recognised that the tensile strength of water is around - 500 MPa (Tabor, 1979) and built tensiometers which take advantage of this fact. The Imperial College tensiometers contain an extremely small volume of water and they are stored full of water and surrounded by water at a hydrostatic pressure numerically greater than the modulus of the matric potential which they subsequently wish to measure. Their tensiometers at present cover the range from 0 to -700 kPa although they see no reason in principle why this range can not be extended. When these tensiometers are placed in contact with soil, they reach an equilibrium reading in about 2 min. Unfortunately, their tensiometers are not suitable for permanent field installation


because they cannot sustain large negative pressures for very long periods because of problems of the formation of air bubbles.

A ‘stripper’ for removing air-bubbles from tensiometers has been described by Miller and Salehzadeh (1993), but this device works only in the range 0 to - 100 kPa.

For soils drier than - 80 kPa, it is possible to use several other methods, all of which have associated problems. Perhaps the most common is to extract soil samples, to measure their water content and to estimate the water potential from water characteristic curves obtained previously in the laboratory. Another technique is the ‘filter paper method’ (Deka et al., 1995) which involves burying pieces of filter paper in the soil. These absorb water until they reach equilibrium whereupon they are excavated and their water content measured. The water potential of the field soil is then obtained from the pre-determined water characteristic of the filter papers. These above two methods are unfortunately destructive.

An extension of the last technique is to bury blocks of porous material with known water characteristics and to measure their water contents by conductivity or by capacitance or TDR methods. This provides a method for continual data-logging of water potential non-destructively. Considerable care has to be taken with these methods to avoid errors which might arise because of hysteresis of the water characteristics of the materials involved.

A direct measurement of water activity (or equilibrium relative humidity) can be obtained by thermocouple psychrometry (Richards, 1969, 1971). In this, thermocouple junctions are cooled by Peltier cooling until water condenses on them. By measuring wet and dry bulb temperatures with the same thermocouple junction, the activity of the water and hence, the sum of its matric and osmotic potentials can be determined. The method is amenable to data-logging but is not very accurate for soils wetter than about - 100 kPa. It is very good in dry soils.

Hysteresis in the soil water characteristic results from, amongst other things, problems of pore connectivity. Because of the complicated way in which soil pores are connected, they do not fill in the order from the smallest to the largest, and they do not empty in the reverse of this order. In other words, because of these and other effects, the existence of water at a matric potential $ at the surface of a soil sample does not mean that all pores of diameter smaller than

(11)

are water-filled, as is often assumed. In Eq. (111, (T is the surface tension of a water meniscus. These complications lead to entrapped air on wetting and residual water on drying of soil.

This effect is illustrated in Fig. 2 which illustrates drainage of a random, 4-level fractal structure. If the pores drained according to Eq. (1 l), then the water retention curve would be of the power-law form as is shown by the lower curves in Fig. 2. However, because of connectivity effects, pores are emptying (becoming air-filled) at the sides of the sample first and this gives the upper curves in Fig. 2. These sigmoidal water-curves are similar in form to those observed in practice and produced by the empirical Van Genuchten equations (Van Genuchten, 1980) which are so useful. When

A.R. Dexter/Soil & Tillage Research 43 (1997) 41-63

bl

Fig. 2. A random, 4-level fractal structure where the pores are shown black and the solid is shown in white (a) Increasing levels of suction are applied at the base of the sample and air (shown grey) is allowed to enter from the top and the sides (as with a sample on a pressure plate extractor). With increasing suction (shown in b, c, d and e), the water retention would follow Eq. (11) (shown by the lower curves in b,, c2, d,, and e?) if there were no problems of pore connectivity. With the first suction step (shown in b), all the largest pores (shown in bl) would drain in the absence of connectivity problems. However, only the grey pores are able to drain. Similarly, for the increasing suction steps (shown in c, d and e). The resulting water retention curve is given by the upper curves in b,, c?, d, and ez (N.R.A. Bird, Silsoe Research Institute, personal communication).

breakthrough occurs and a continuous connected network of air-filled pores forms throughout the entire sample, the sample is said to ‘percolate’ with respect to air. This situation is illustrated in Fig. 2d.

Much can be learned about the factors that influence water retention characteristics and fluid displacement in general from the study of pore network models. This commenced with the classic work of Fatt (1956a). Among his important conclusions are that a pore size distribution determined from the water retention curve of a real porous medium by the use of Eq. (11) will have an average pore radius and spread of radii which are smaller than the true average pore radius and spread of radii. In particular, he stated that the spread of pore radii is about twice that which is obtained from a drainage curve and Eq. (11).

There have been enormous advances in the theory and simulation modelling of processes in pore networks in recent years (e.g., Celia and Ferrand, 1992; Dullien et al., 1992; Ewing and Gupta, 1993). This surge in activity has mostly been driven by the demands of the oil industry and by concern about environmental problems.


Consider a simple cubic lattice of pores with a number distribution of pores of size Y of f(r). At a water potential of $, with the assumption of Eq. (1 l), the bottleneck pore size, rb, can be calculated from standard percolation theory as

r* I() f Y du=pc rb

(12) where pC = 0.3117. This is important because it is the pore bottlenecks which control the water retention characteristics and the hydraulic conductivity.

More realistic network models have been considered by some authors (e.g., Daian et al., 1994; Neimark, 1989). The further development and use of these models will undoubtedly contribute to our understanding of what is happening in soils at the pore level.

These network models help us to understand the origins of hysteresis of the water characteristic curve. On a more pragmatic level, hysteresis loops can be fitted with modified forms of the Van Genuchten equation (Kool and Parker, 1987).

AS we have seen, problems of pore connectivity mean that we cannot determine pore size distributions using Eq. (11) either from water retention data or from mercury intrusion data for the same reason. However, it is possible to get some idea of the sizes of water-filled pores in a soil sample by the use of NMR proton-spin relaxation (Hills and Babonneau, 1994; Hills and Quantin, 1993). The rate of decay of proton excitation decreases with increasing distance of the proton from the pore wall. Therefore, by measurement of the decay spectrum it is possible to obtain an estimate of the distribution of sizes of the water-filled pores. Obviously, this is not a field technique.

3.2. Dynamics

The saturated hydraulic conductivity of the tilled layer of soil is of course enormous because of the preponderance of large inter-aggregate macro-pores. However, the infiltration rate in practice is limited by that of any plough pan or compaction pan immediately below the depth of tillage (Addiscott and Dexter, 1994).

Soil structure is dynamic and in almost all soils changes with time or rainfall after tillage. The structure may collapse progressively (Dexter, 1976; Dexter et al., 1983; Domial and Slowifiska-Jurkiewicz, 1987) and surface crusts or seals may form especially on soils with low stability resulting from low contents of clay or organic matter. A crust is often a barrier to water infiltration and can be a cause of ponding and run-off.

The saturated hydraulic conductivity of freshly-tilled soil is therefore, largely irrele- vant. What are more limiting and more important are the conductivities of crusts and pans, and the ability of the soil clods and aggregates to absorb water.

In the case of tilled soils, such as aggregate beds, water transport can occur in a range of ways which have been described in detail by Youngs and Leeds-Harrison (1990). The different situations arise from various combinations of whether the inter- and intra-aggregate pores are saturated or unsaturated.

When rain falls on the surface of a tilled soil, it may be that the exposed surfaces of the aggregates are able to absorb the rain. If rainfall is faster than can be absorbed by the aggregates at the surface, the excess runs down the inter-aggregate pores. However, even in a uniform bed of aggregates, this flowing water does not move in equal quantities

A.R. Dexter/ Soil & Tillage Research 43 (1997) 41-63 49

between all pairs of adjacent soil aggregates. The rivulets tend to combine into a smaller number of more rapid flows with increasing depth in the aggregate bed (Dexter, 1994).

A device for measuring the ability of soil aggregates to absorb water has been described by Leeds-Harrison et al. (1994). This is essentially a miniature disc permeameter which enables the sorptivity to be measured reliably.

Disc permeameters (or tension infiltrometers) are also very useful for measuring the absorptive power of the surface of structured soils (Perroux and White, 1988; Nachabe and Illangasekare, 1994). Since they apply the water under suction, the water does not enter the macro-pores and the hydraulic properties of the soil matrix can be characterised. Typically, disc diameters in the range 100-300 mm are used, and water is applied at a potential of around - 3 hPa. This means, from Eq. (1 l), that water will not enter pores larger than 1 mm diameter.

From the analysis of infiltration data from disc permeameters it is straightforward to obtain the sorptivity of the soil, S. However, to obtain hydraulic conductivities, the use of multiple disc diameters and/or a range of water suction is required (Hussen and Warrick, 1993). In the field, Jarvis and Messing (1995) found that for water supply suctions in the range 0 to - 10 hPa, the hydraulic conductivity decreased by 3 to 4 orders of magnitude. The analyses, of course, assume that the soil is homogeneous. However, Lin and McInnes (1995) found non-uniform distribution of water in a clay soil beneath a disc infiltrometer which demonstrates the necessity for care in the interpretation and use of calculated ‘effective’ values of sorptivity or conductivity.

In the laboratory, X-ray CAT scanning can be used to follow the absorption of water into soil samples. With most medical X-ray CAT scanners, the spatial resolution is typically about 1 mm, but this can still give better resolution of wetting fronts than can be obtained by the mechanical slicing-up of soil samples (Fig. 3). The use of X-ray CAT scanners is also non-destructive so that a series of scans showing the development with time of wetting fronts in the same sample can be obtained. It should be noted that adding

source of water under suction

0.5 wetting front

I I

volumetric water

0

0.4

0.3

0.2

0.1

0.0

Fig. 3. Infiltration of water into a horizontal soil column (top), and the corresponding profile of volumetric water content, 0, as a function of distance, x, after 1000 s as determined by X-ray CAT scanning (bottom).


water to a non-swelling soil increases its wet bulk density (which is approximately what

is measured by X-ray absorption) whereas, adding water to a sample of an unconfined,

normally-swelling soil reduces its wet bulk density. New developments in industrial X-ray equipment and image reconstruction algo-

rithms have recently enabled spatial resolutions of around 20 pm to be achieved which brings the X-ray CAT technique into an important range of soil pore sizes. The new technique is known as cone-beam microtomography (Cheng et al., 1993; Ham et al., 1993; Flynn et al., 1994; Machin and Webb, 1994; Wang et al., 1994; Likhachov and Pickulov, 1995).

Unsaturated hydraulic conductivity of soil as a function of water content can be determined from wetting profiles such as shown in Fig. 3 as described by Bruce and Klute (1956). Alternatively, hot air can be used to dry the surface of a soil sample at a rate limited by the ability of the soil to conduct water. The resulting profiles of water content can also be used to calculate unsaturated hydraulic conductivity (Arya et al., 1975) although problems can arise from the heating of the soil surface (Van Grinsven et al., 1985). Again, CAT scanning can give the shape of the drying front much more accurately than was possible previously. A very useful method which gives simultane- ously the water retention characteristic and the unsaturated hydraulic conductivity is the evaporation method of Wind (Halbertsma and Veerman, 1994). This method can be used on undisturbed soil cores from the field and the measurements are made under essentially isothermal conditions. Unsaturated hydraulic conductivity can be readily measured over the range of water potentials from approximately - 90 to - 500 hPa in sandy soils and from - 5 to - 800 hPa in loamy or clay soils.

Measurements of unsaturated hydraulic conductivity can also be made at constant matric water potential. Cook et al. (1993) have described an elegant method which can be used on soil cores taken from the field. Water at a given suction is applied at the top of the sample with a disc permeameter and at the base of the sample with a coarse sintered-glass funnel containing a hanging water column. This gives constant matric potential throughout the sample, but unit potential gradient arising from the gravitational component. Steady-state water flux is measured in a standard way. Hydraulic conductivity can be determined easily in this way over the approximate range of matric potentials from 0 to - 1 kPa.

Although we talk about water moving through soil at a certain mean velocity, in fact it moves much more rapidly through percolating networks of large pores than through the small pores. This leads to the simplifying concept of ‘mobile’ and ‘immobile’ water. This concept is useful in considerations of, for example, the movement of solutes through soils. In these cases, the infiltrating fluid is the mobile phase and the original resident fluid may either be immobile or partly immobile and partly mobile. Exchange of solutes between the infiltrating and the resident fluid may be considered to occur by diffusion. Analyses and quantification of these mobile and immobile components of soil water and of interactions between them have been made by Clothier et al. (1992, 1995) and by Nachabe and Morel-Seytoux (1995).

Again, network models are an extremely valuable tool for gaining an understanding of the physical processes occurring at the pore level (e.g., Fatt, 1956b,c; Dullien, 1975; Daian, 1992; Ewing and Gupta, 1993; Neimark, 1989).


4. Aeration

Aeration is important for both the agricultural and environmental functions of soil. Plant roots and soil fauna require oxygen, and aerobic microbes are important decom- posers. Key aspects of soil aeration are reviewed in the excellent book by Glitiski and Stepniewski (1985).

A range of different measurements give information of varying usefulness about levels of aeration or soil oxygen. Some of these are now described briefly with their principal limitations.

Total porosity or air-filled porosity (or, indeed, the air-ratio from Eq. (5) are easily calculated from measurements of soil bulk density and water content. Although it is generally agreed that levels of air-filled porosity below 10 or 15% are often associated with reductions in crop growth and yield, this is not generally true. A simple measure of air-filled porosity, although better than nothing, gives no information about the quality of the soil air, about its rate of renewal or about its spatial distribution. It is the oxygen concentration where the roots or other organisms are located which is of greater importance than the mean level in the whole soil.

Air permeability is a measure of how easily air convection occurs through soil in response to pressure gradients. Pressure gradients can be generated naturally by air turbulence above the soil surface, and this can lead to air flows through the tilled layers of soils especially when they contain pores larger than about 5 mm (Farrell et al., 1966; Kimball and Lemon, 1971). Air flow will occur only if there is a continuous network of air-filled pores. Naturally-occurring air pressure gradients are not large enough to displace water from water-filled pores or to burst water films. Therefore, air permeability is a useful indicator of pore connectivity. This is well-illustrated by consideration of the effects of soil water potential on the air permeability of individual soil aggregates which is shown in Fig. 4. As a first approximation, it can be assumed that at any water potential, $, pores larger than D given by Eq. (11) are air-filled. So in progressively drier soil (larger values of I$ I>, there are large numbers of air-filled pores giving better connectivity and higher values of air permeability.

-10 -

"E

d -ll-

z * s

-12 -

-13 I d i

I -1 2 A B

L"g,~b+',kPa)

Fig. 4. Variation of air permeability, k,, of single soil aggregates with water potential, T, (B.M. McKenzie, personal communication).

52 A.R. Dexter/Soil & Tilluge Research 43 (1997) 41-63

Air permeability can be measured in the field in various ways. These usually involve pressing a cylindrical ‘permeameter’ with a closed top partially into the soil. The cylinder is then pressurized with a pressure only slightly above atmosphere, and the flow rate of the gas is measured. If accurate measurements are being attempted, then account has to be taken of the geometry of the gas flow paths (Liang et al., 1995).

In freshly-tilled soils, it can usually be assumed that the inter-aggregate macro-pores are filled with air with a level of oxygen which is at least sufficient for biological activity. The principal exception to this is where a cap, crust or seal has formed on the surface by the action of water as will be described in Section 5. When moist, a continuous non-cracked crust can cut off the oxygen supply to the underlying soil which may be left with only perhaps sufficient oxygen for 4 or 5 days of respiration at normal temperatures.

In the absence of convective fluxes, gas molecules move about by diffusion. The diffusivity of oxygen in air is about 10,000 times larger than in water. Therefore, diffusion occurs mainly through the connected (percolating) air-filled pores. The rate at which oxygen atoms arrive at a root can be measured by using a platinum micro-electrode. This is usually made of Pt wire of about 0.4 mm diameter to simulate a root. This electrode is set at an electrical potential of about - 0.7 V relative to a calomel reference electrode. The electrical current which flows is a measure of the rate at which oxygen molecules collide with the electrode and are reduced. This is used to calculate the oxygen diffusion rate (ODR), although it should be said that the method measures the rate at which oxygen atoms arrive at the electrode and does not distinguish (at least in its basic form) between diffusive and convective transport mechanisms. The technique is valid only in very moist or wet soils because the method requires that the electrode is covered with a film of water.

Rappoldt (1995) has further developed this technique by alternately applying air and nitrogen gas to the surface of a soil sample and using a Pt micro-electrode to measure the amplitude and phase shift of the ‘oxygen wave’ with depth in the soil sample. This gives accurate values of oxygen diffusivity. The high spatial resolution of the method enabled him to infer the presence of small anaerobic zones between individual air-filled pores in a moist clay soil.

Czyi and Tomaszewska (1993) have measured the effects of compaction of a loamy sand on ODR (Fig. 5a). Also shown in Fig. 5b are the values of the water air ratios calculated for this soil using Eq. (4) and Eq. (5). This illustrates how ODR decreases with decreasing values of the air ratio. ODR was negatively correlated with the yield of sugar beet. This emphasizes that reductions in oxygen supply are a major factor in the yield reductions which are often observed in compacted soils. Response surfaces showing how ODR and the oxygen diffusion coefficient vary with combinations of bulk density and water potential for four soils have been published by Stepniewski (1980, 1981).

Within individual aggregates or clods anaerobic conditions can occur. The demand for oxygen for respiration by microorganisms can exceed the ability of diffusion to supply oxygen from the aggregate surface. In these cases, the centres of the aggregates will be anaerobic. Theoretical aspects of this problem have been considered by Renault and Stengel(l994) whilst experiments by Zausig et al. (1993) showed that with artificial


(a) ‘*O) 100

\ yJ 80

E

-:\\\\

3

E 60

0

40

1.2 1.3 1.4 1.5 1.6 1.7 1.8

(b) 1.2 ,

am0.6 5

CE 0.4

0.2

0

1.; 1.3 1.b 1.6 1.7 1.8 Soil density (Mgm-3)

Fig. 5. Linear regression of oxygen diffusion rate with level of compaction of a heavy loamy sand in the field (a) from Czyi and Tomaszewska (1993). The bar graph (b) shows the corresponding water ratio, R,, and air ratio, R,, as calculated from Eq. (4) and Eq. (5).

homogenous soil aggregates, the oxygen supply dropped with increasing depth from the aggregate surface. Significant reductions were observed at depths beyond about 2 mm.

Inside natural, heterogeneous soil the oxygen concentration or supply can also be expected to be heterogeneous. This has been demonstrated by Bakker and Bronswijk (1993) who used computer assisted X-ray tomography (CAT scans) to determine non-destructively the internal geometry of the soil pore space, and then probed the samples with Pt ODR micro-electrodes. A good agreement was found between the soil porosity and oxygen concentration distributions.

Our understanding of the aeration situation within tilled soils has been extended by the work of Ball et al. (1988) who measured gas diffusivities, air permeability and hydraulic conductivities on the same soil samples as functions of water content. They were able to define useful indices of pore connectivity, and they also pointed out that bulk density is not a good indicator of soiI transport properties. Renault and Sierra (1994) modelled oxygen diffusion and consumption in beds of aggregates of varying shape and size. Gas exchanges between the inter- and intra-aggregate pores are modified by inter-aggregate contacts and by water films. Their simulations show that the

54 A.R. Dexter/ Soil & Tillage Research 43 (I 997) 41-63

relationship between the soil anaerobic fraction and water content is highly dependant on the soil structure.

In very wet soils, close to or at saturation, the redox potential becomes the appropriate measure of the oxidizing or reducing power of the soil (Glinski and Stepniewski, 1985). Zausig and Horn (1992) used measurements of redox potential to demonstrate anaerobic conditions in the central parts of natural soil aggregates. Biddle et al. (1995) showed that graphite rods (from electrical batteries) could be used instead of platinum for redox electrodes. These gave potentials which were typically 200 mV lower than those obtained with platinum, but graphite electrodes have the advantage of being enormously cheaper than platinum.

As a result of soil structure or heterogeneity, it is usually not sensible to talk about a soil as being either aerobic or anaerobic. Usually, there are regions of both present at the same time. The interest is in determining the proportions of the soil which fall into these two categories. For agricultural purposes, we need to know the aeration status where the crop roots are. For environmental purposes, we might need to know how to manage the soil to achieve activity by aerobic microbes at sites where pollutants, such as pesticides, are expected to be located.

5. Strength and stability

Strength and stability are necessary if soil is to retain its structure against imposed stresses. These imposed stresses may be natural such as raindrop impact, or may be anthropogenic such as those imposed by vehicular traffic. A complication with soil is that its strength must not be too great otherwise, plant roots and other organisms will not be able to penetrate.

The term ‘strength’ is used to describe the level of stress (force per unit area) that a soil can resist without undergoing irreversible deformation. The term ‘stability’ is used (in English) to d escribe the ability to retain a coherent structure in the presence of free water.

The strength of unsaturated soil is extremely sensitive to changes in water content. Therefore, the water content needs to be measured accurately every time that strength is measured. The pore water pressure is said to exert an additional stress on the soil. This additional stress is the level of external stress on a soil sample which would have the same effect as its internal water. Aitchison (1961) introduced this concept for unsaturated soils and estimated the value of this additional stress from the product of the manic potential of the soil water, $, and the degree of saturation, R,/e, which he denoted by x. For saturated soil, x = 1, but for unsaturated soil, 0 5 x < 1. Pore water pressure can be assumed to be isotropic and therefore, for a soil sample subjected to externally-applied principal stresses of (TV, CT* and cr3, the effective stresses were assumed to be

a;=cr,-,y*,~;=a, -x$,and a;=~,-x+ (13) The minus signs in Eq. (13) arise because $ is negative in unsaturated soil. It is these effective stresses which should be used when determining stress-strain relationships in soils, or when interpreting the strengths of unsaturated soils.


The Eq. (13) are useful, but are only an approximation because the contribution of the soil water to the effective stress is also due to factors other than the degree of saturation and the water potential. An understanding of the physics of the situation comes from detailed analysis of the forces between soil particles arising from water bridges and menisci. Fisher (1926) pointed out that there are three components of force acting on particles in unsaturated soil which arise from the water. First, there is the pore water pressure (or potential) as used in Eq. (13). Second, there are the forces due to surface tension in the water films and menisci; and third, there are buoyancy forces which can usually be ignored.

More recent works (e.g., De Bisschop and Rigole, 1982; Simons et al., 1994) have calculated the force due to water bridges between pairs of particles as a function of separation. Integration of this function yields the total energy for particle separation. This is important and useful because propagation of cracks in moist non-cemented soil can be thought of as occurring largely through the breaking of water bridges. The energy for fragmentation could also be estimated in this way.

The resistance of samples of moist soil to fracture is influenced strongly by the micro-cracks and other flaws within it. These structural features can arise in a number of ways as described by Dexter (1988). In order to investigate this, Hallett et al. (1995a) studied the effects of artificial cracks in otherwise homogeneous, artificial samples. They showed the large effects of flaws on soil tensile strength. In samples of natural soils containing many cracks the situation is very complicated and interactions between the micro-cracks are thought to influence the result. Although these interactions are not fully understood, it seems fairly certain that the propagation of the larger micro-cracks and the joining together of pre-existing micro-cracks are key processes in the fracture and fragmentation of soil.

A fracture surface, then, is composed partly of the walls of micro-cracks which pre-existed in the soil before fracture and partly of ‘new’ surface area. Hallett et al. (1995b) investigated this effect by staining the walls of the micro-cracks with blue dye before fracturing soil samples. Measurement of the proportion of blue on the fracture surfaces gave a measure of the proportion of the fracture surface which already existed as walls of micro-cracks in the soil before fracture.

If a soil sample were completely homogeneous, without any internal micro-cracks or flaws, then the strength would be independent of sample size. However, in samples of real soils, with their internal micro-structure, strength is a function of sample size. In general, larger samples have a lower strength because they can (and do) contain larger flaws. This effect has been described by Braunack et al. (1979) and Perfect and Kay (1995).

This size-dependence of strength can be measured accurately and was used by Utomo and Dexter (1981) as a measure of soil friability. Friability is defined here as the tendency of large clods or aggregates to breakdown under an applied stress into a given size range of smaller aggregates such as might be required for a seedbed. Clearly, this is one of the objectives of tillage which will be much easier with a friable soil in which the large clods are relatively weak whereas, the small fragment aggregates are relatively strong and resistant to further breakdown. Soils in ‘good condition’ (e.g., with high contents of organic matter, non-compacted or calcic) are associated with high values of

56 A.R. Dexter/Soil & Tillage Research 43 (1997141-63

friability, whereas, degraded soils or soils in ‘poor condition’ (e.g., with low contents of organic matter, compacted or sodic) have low values of friability. Friability is negatively-correlated with the amount of readily-dispersible clay in soil (Shanmuganathan and Oades, 1982).

Soil instability can occur in several ways. These include dispersion of clay particles in free water and slaking upon rapid wetting. Clay particles (< 2 pm) form the fundamental building blocks of soil structure. If the clay particles disperse, then all the larger features of the soil structure are destroyed, The dispersion of clays depends on several complex mineralogical and physical-chemical properties, and on the quantity and disposition of organic matter (Quirk and Murray, 1991).

Slaking may be defined as the breakdown of soil aggregates into micro-aggregates which are stable compound particles in the approximate size range 20-250 pm. Slaking often occurs on rapid wetting, but only if the soil is initially drier than a matric potential of - 1 MPa (Grant and Dexter, 1989). Slaking is strongly negatively-correlated with the organic matter content of the soil.

The amounts of water-stable aggregates and micro-aggregates are readily determined by wet sieving. There are two very different wet-sieving techniques in common use which measure different things. In the first type (as used by Tisdall and Oades (1982) for example), air-dry soil aggregates are immersed directly in distilled water whereupon slaking may occur. A set of oscillating sieves is then used to determine the size distribution of the slaked fragments. Therefore, the proportion of the initial mass remaining on the 250 pm (top) sieve is the proportion of water-stable aggregates, and the proportion passing the 250 pm sieve but being retained on the 50 pm sieve is a measure of the water-stable micro-aggregation. In the second-type of test (as used by Kemper and Koch, 1966), soil aggregates are wetted slowly to saturation in a stream of water vapour and are then immersed in water with no slaking. This test measures the resistance of the immersed soil to breakdown under the action of the mechanical stresses imposed by the oscillating sieves. It is therefore a measure of the strength of saturated soil in the absence of slaking. This test is more appropriate for the extremely fragile soils of the northwestern USA which would always completely slake if wetted rapidly.

Another measurement of stability is the amount of readily-dispersible clay. This is usually determined after an unknown, but repeatable, amount of energy has been put into the soil by shaking it in a standard way with water. The amount of readily-dispersible clay can then be measured easily and rapidly in terms of turbidity (e.g., Kay and Dexter, 1990; Watts et al., 1996).

The effects of mechanical energy input on soil stability have been investigated by Watts et al. (1996). Some results are shown in Fig. 6. Soil aggregates of different initial water contents were subjected to different levels of mechanical energy input representa- tive of the energy inputs to soil during tillage operations. The soil aggregates were then shaken with water and the amount of clay dispersed was measured. It can be seen from Fig. 6 that for soil drier than the lower plastic limit, energy input did not increase the amount of dispersible clay. However, for soil at or wetter than the lower plastic limit, dispersible clay increased with increasing water content and mechanical energy input. These results illustrate the destabilisation of soil and the possible mobilisation of clay in the environment which can occur if soil is tilled intensively when it is too wet.

A.R. Dexter/Soil & Tillage Research 43 (1997141663

800

1.7 PL

0.2PL

0 200 400 600 800 1000

Specific energy J kg-l

Fig. 6. Effects of mechanical energy input on soil aggregate stability as measured by the content of readily-dispersable clay (measured in Nephelometric Turbidity Units, NTU). The different lines are for different initial values of soil water content which are expressed as proportions of the lower plastic limit, PL (after Watts et al., 1996).

As mentioned in Section 3.2, crusts (caps or seals) can form on the soil surface. Crusting has several adverse consequences including: Reduced rates of water infiltration resulting in increased run-off and erosion; reduced levels of soil aeration; and reduced crop emergence as a result of the mechanical strength of the crusts (e.g., Dcbicki and Lipiec, 1993; Debicki, 1994; Poesen and Nearing, 1993).

Crusts can form by one or more of the following mechanisms: Deformation of the surface by the impact of raindrops; clogging of macro-pores by dispersed clay or slaked fragments; and the deposition of transported material. The importance of the compactive action of raindrops has been shown by many researchers (e.g., Shainberg et al., 1992; Betzalel et al., 1995), and Loch and Foley (1994) have clearly separated this effect from other effects. The size of raindrops has a big influence on drop velocity and impact force and can be measured in a variety of ways (e.g., Bazzoffi, 1980). A protective cover of crops or crop residues can largely remove effects due to raindrop impact.

Structural crusts formed by the washing-in of dispersed clay or slaked fragments are probably the most common type of crust (Bresson and Boiffin, 1990; Bresson and Cadot, 1992; Poesen and Nearing, 1993). The importance of slaking has been demonstrated for the case of initially-dry soil by Loch and Foley (1994). They found a strong negative correlation between the proportion of aggregates < 0.125 mm after wetting by rain and the hydraulic conductivity and water infiltration rate of the resulting crust. An important observation was that, for this mechanism, it was the wetting rate of the soil which was important and not the size of the raindrops. Soil management methods which increase soil stability through, for example, increasing soil organic matter content, can reduce crust formation (e.g., Pikul and Zuzel, 1994).

The hydraulic conductivity of crusts in isolation is difficult to measure, However, the effects of crusts on hydraulic conductivity and water infiltration can be measured with

58 A.R. Dexter/ Soil & Tillage Research 43 (1997) 41-63

Force trysducer Capstan to retrieve beads

String with beads

\

- Ground wheel

Soil surface

Fig. 7. Schematic diagram of a device for continual measurement of crust strength or impedance to emergence in the field. The device moves very slowly along rails and extracts a succession of beads which have been buried previously at seed depth. Bead extraction forces are recorded with a data logger.

the crust still in contact with the underlying soil. This can be measured with disc permeameters or on soil cores (e.g., Ela et al., 1992; Bohl and Roth, 1993; Loch and Foley, 1994).

Seedbed aggregate size influences crusting. Braunack and Dexter (1988) and Free- bairn et al. (1991) found that larger aggregates at the surface of a seedbed reduced both the time to crusting and the time to the initiation of run-off (which were similar).

Crust strength varies significantly with crust composition, thickness and water content. Because a crust is at the soil surface, its water content can change rapidly in response to meteorological conditions. This means that it is desirable to measure crust strength continually over the emergence period of an experimental crop. Also, it is desirable to measure crust strength vertically upwards to simulate the emergence of plant shoots. A method of doing this automatically is shown in Fig. 7. Some typical results are shown in Fig. 8 where a soil which is drying slowly is giving steadily increasing values of crust strength as shown by the increasing forces for bead extraction. Similarly, a

loo0 -4 800 -

600 - 3 2 400- s

200 -

o-

I I I I I I I

0 200 400 600 800 1000 IiC lime (min)

)O

Fig. 8. An example of data obtained with the device for continual measurement of soil crust strength shown in Fig. 7. Here the soil was drying slowly and becoming stronger.


small amount of rain reduces the bead extraction force. The dynamics of crust strength are important in the interpretation of the emergence with time of crops in the field.

6. Concluding remarks

This short review has looked at methods which are relevant to the characterisation of the physical properties of tilled soils. There have been significant developments in measurement techniques over the last few years. New techniques have been developed which are faster, more accurate, have better spatial resolution than those previously available, and which are amenable to data logging. All of these improvements are especially important in the study of the tilled layer of soil where properties change significantly over small intervals of space and time.

The size scale of physical measurements must always be appropriate for the phe- nomenon being studied. For the tilled layer of soil, the appropriate size scale may be the depth of the layer (e.g., 200 mm); for processes within the tilled layer, measurements at the aggregate scale (e.g., 10 mm) may be appropriate; for processes involving root axes, measurements at the size scale of a root diameter (e.g., 1 mm) may be appropriate; whereas for processes at the size scale of microbes, measurements at the scale of bacteria (e.g., 1 pm> may be appropriate. At each of these size scales, specific measurement techniques are needed to determine the properties and heterogeneity of the physical environment at that size scale. Erroneous conclusions can be drawn if measurements at different size scales are combined, Care is necessary in defining each problem precisely so that measurements can be made which are relevant to that problem.

Certainly, our ability to collect vast quantities of data has increased dramatically. But has this solved all of our problems? I think not for several reasons. First, most soil physical measurements require the insertion of sensors. This always involves disturbing the soil and changing the properties which we are trying to measure. We must explore ways to minimise this disturbance. Second, often we do not know the details of the environment of our sensors. Certainly, we must have many replicate measurements of each physical property, not only to obtain an accurate mean value, but also to determine accurately the standard deviation as a measure of its heterogeneity. Third, there is the question of what we do with all of the data collected. For example, do we just look at it and discuss it, or do we use it either to develop or to evaluate simulation models.

Although simulation models have an important role to play in the study of complex systems, many basic mistakes continue to be made by modellers (Philip, 1991). These include the strong tendency of modellers to develop models which are too complicated and which are untestable.

Only the experimental study of the real world can increase our scientific knowledge through the testing of new hypotheses. Soil physics and indeed all research advances by the testing of hypotheses. This single process involves asking a question (posing a hypothesis) to which the answer is required, designing a simple experiment which will test the hypothesis unequivocally, and then performing the minimum number of measurements to test the hypothesis to the desired level of statistical significance.

I am convinced that experiments cannot be too simple, but that it is easy to make them too complicated. There are many questions which need to be answered regarding


the physical properties of tilled soils. Most of these revolve around the consequences of soil structure or heterogeneity, and many of them involve interactions with chemical and biological processes in the soil. Let us make sure that we use our measurements of the physical properties of tilled soils intelligently to test our hypotheses as effectively as possible.

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Physical properties of tilled soils