2011 Dexter Clay Dispersion from Soil as a Function of Antecedent Water Potential

Prévia do material em texto

444 SSSAJ: Volume 75: Number 2 • March–April 2011
Soil Sci. Soc. Am. J. 75:444–455
Posted online 16 Feb. 2011 
doi:10.2136/sssaj2010.0088
Received 23 Feb. 2010. 
*Corresponding author (guy.richard@orleans.inra.fr).
© Soil Science Society of America, 5585 Guilford Rd., Madison WI 53711 USA
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Clay Dispersion from Soil as a 
Function of Antecedent Water Potential
Soil Physics
The amount of clay that disperses from soil into water is of great agricultural 
and environmental signifi cance. Dispersion implies “zero strength”, and the 
soil structure can collapse with major reductions in water infi ltration rates and soil 
aeration status. Th erefore, clay dispersion is a cause of poor soil stability in water 
and this can contribute to soil erodibility and mud fl ows (Boardman, 2010). Th e 
presence of dispersed clay in soil can also cause the soil to become hard and non-
friable when it dries as a result of cementation (Kay and Dexter, 1992). In this pa-
per, we investigate the amount of RDC in soils. Readily dispersible clay is the part 
of the clay fraction of soils that is easily or potentially dispersible in water when 
small amounts of mechanical energy are applied to soil.
We consider only the phenomenon of clay dispersion and not the 
phenomenon of slaking. However, we mention slaking briefl y here because some 
of the discussion on eff ects of antecedent water potentials and pretreatments is also 
relevant to work on clay dispersion. Slaking is the breakdown of soil into micro-
aggregates (i.e., <250 μm) when it is immersed rapidly in water (Emerson, 1954; 
Panabokke and Quirk, 1957). Most research on soil (or aggregate) stability has 
involved measurements of the amount of slaking by the method of wet sieving of 
air-dried aggregates (e.g., Tisdall and Oades, 1980). Th is method involves the rapid 
wetting of the aggregates and then sieving the resulting fragments under water by 
Anthony R. Dexter
INRA, UR0272 Science du sol
Centre de Recherche d’Orléans 
CS40001 Ardon
45075 Orléans cedex 2, France
and 
Institute of Soil Science and Plant 
 Cultivation (IUNG-PIB)
 ul. Czartoryskich 8, 24–100 Pulawy 
Poland
Guy Richard*
INRA, UR0272 Science du sol 
Centre de Recherche d’Orléans 
CS40001 Ardon
45075 Orléans cedex 2, France
Ewa A. Czyz
INRAUR0272 Science du sol
Centre de Recherche d’Orléans
CS40001 Ardon
45075 Orléans cedex 2, France
and
Institute of Soil Science and Plant 
 Cultivation (IUNG-PIB)
ul. Czartoryskich 8
24–100 Pulawy, Poland
and
Rzeszów University
Faculty of Biology and Agriculture
Aleja Rejtana 16c
35–959 Rzeszów, Poland
Joëlle Davy
Michel Hardy
Odile Duval
INRA, UR0272 Science du sol
Centre de Recherche d’Orléans 
CS40001 Ardon
45075 Orléans cedex 2, France
Two standard tests for soil stability in water involve assessment of the readily dispersible clay (RDC) content of 
predried soil samples. It is shown that this predrying limits the applicability of the tests to very dry soils. Ways to 
avoid this limitation are developed. Undisturbed soil samples were collected from the 10- to 15-cm depth of arable 
topsoils at Villamblain and Faux Perche and from the 0- to 5-cm depth from plowed and direct-drilled plots at 
Boigneville. Th ese soils in north-central France were collected from the fi eld when moist (near fi eld capacity). A 
range of water contents and corresponding pore water potentials (or suctions h) was produced by slow drying in 
air. For moist soils at fi eld capacity (h = 100 hPa) the amount of RDC was large. Readily dispersible clay decreased 
progressively with increasing antecedent suction, h. For soil samples dried to beyond h > 10 MPa, the RDC was 
reduced to small values. With such intensely dried soils, rewetting for up to 2 wk did not reverse the eff ect on 
RDC. At h = 4500 hPa, re-equilibration to h = 100 hPa reversed the eff ect of this moderate drying on RDC. 
Two standard conditions (moist and dry) are proposed for measurement of clay dispersion from soil into water. 
Th e moist condition is relevant to soil at the end of a long, wet period (e.g., at the end of the winter in temperate 
climates), whereas the dry condition is relevant to dry surface crusts and dry topsoils. We conjecture that the soil 
“resets” during long, wet periods to values of stability and RDC at equilibrium with fi eld capacity. Experimental 
procedures used to determine the content of RDC in moist and dry soils are described.
Abbreviations: C, total clay; CC, complexed clay; CEC, cation-exchange capacity; COC, complexed 
organic carbon; CT, conventional tillage; DD, direct drilling; EC, electrical conductivity; GG, 
Groenevelt and Grant; NCC, non-complexed clay; NCOC, non-complexed organic carbon; OC, 
organic carbon; RDC, readily dispersible clay; T, turbidity.
Published March, 2011
SSSAJ: Volume 75: Number 2 • March–April 2011 445
 
the method originally described by Yoder (1936). In contrast, 
Chan and Mullins (1994) prepared aggregates at a range of 
antecedent water potentials by wetting them in a series of slow-
wetting stages to prevent disturbance of the structure before they 
subjected them to the slaking tests.
Th e RDC has been found to be sensitive to the hydraulic 
history (the history of dryings and rewettings) to which the 
soil has been subjected. Early in the 20th century, it was already 
observed that the amount of dispersed material was smaller when 
soil was initially drier (Keen and Hall, 1926). More recently 
Gaţe (2006) found that the content of RDC was reduced by a 
factor of about 20 by preliminary air-drying. Drying can occur 
when soil samples are stored in the laboratory and can also occur 
naturally in the fi eld by evaporation under the action of radiation 
and wind.
Clay particles in moist soil when aligned in face-to-face 
arrangement are not normally in direct contact, but have water 
(actually, electrolyte solution) between them. Such arrangements 
of particles are able to support mechanical stresses (Smith et al., 
2009). Whether solid-to-solid contact occurs when there are 
wedge-shaped pores and corner-to-face interactions was left an 
open question by Mitchell (1962) although his Fig. 5 shows an 
intervening water fi lm with the absence of contact.
Th e thickness of these interparticle layers of water is 
governed by the balance between the repulsive force due to 
the interactions between the electrical double layers and the 
attractive van der Waals force (Israelachvili and Adams, 1978). 
Th e repulsive force depends strongly on the nature of the particle 
surface, adsorbed cations, and the electrolyte solution, whereas 
the attractive force is essentially independent of these factors 
(Israelachvili and Adams, 1978). As the soil dries, the electrical 
double layer becomes progressively thinner. When the inter-
particle layer of water reduces to about a 5-nm thickness, then 
the surfaces fall into direct contact under the infl uence of the van 
der Waals forces. Th is eff ect is essentially irreversible due to the 
dominance of the attractive forces. Th e conditions under which 
soil particles come into direct mineral-mineral contact have also 
been discussed by Kemper et al. (1989).
Th ese strong eff ects of drying have implications for the 
applicability of the results from standard tests of soil stability and 
clay dispersion that require intense drying as a pretreatment. Such 
tests include the Emerson test (Emerson, 1967; AS, 1997) which 
specifi es preliminary air-drying, and the Le Bissonnais test (Le 
Bissonnais, 1996; AFNOR, 2005) which specifi es preliminary 
drying in an oven at 40°C. Th ese predrying treatments have theadvantage of being reproducible, but questions remain about 
whether they are relevant to soils in the fi eld and about the 
consequences for subsequent parts of the test procedures.
Any predrying necessarily produces a requirement for 
rewetting because clay dispersion and aggregate stability tests are 
done in water. Th e literature on this subject is huge because of 
lack of standardization of the methodologies. We feel that it is 
not appropriate to try to review the whole subject here, but rather 
to give some key examples to illustrate the issues. Le Bissonnais 
(1996) used two rates of rewetting of his dried samples: (a) 
“slow” wetting at a water suction, h, of 3 hPa before analysis of 
clay dispersion and (b) rapid wetting by immersion in water for 
analysis of slaking. A study of the eff ects of the suction of the 
source of water used for rewetting showed that for remolded and 
dried samples of a silty clay loam and a silt loam, the water had to 
be at suctions, h, larger than 24 and 30 hPa, respectively, for the 
formation of micro-cracks to be avoided (Dexter et al., 1984). 
We therefore question whether wetting at 3 hPa as used by Le 
Bissonnais (1996) is slow enough to prevent structural changes 
within samples of all soils. In contrast, Panabokke and Quirk 
(1957) and Angers et al. (2008), used slow wetting at a 100-hPa 
suction. Chan and Mullins (1994) and Kay and Dexter (1990), 
who required fi nal suctions of 3 and 10 hPa, respectively, fi rst 
wetted their dried samples at 100 hPa for 1 wk before reducing 
the suction to the fi nal, required value over a further week. 
Th is procedure prevented structural disruption in the samples. 
However, Th orburn et al. (1989) found that rewetting of air-
dried aggregates of an intensely swelling-shrinking clay soil at 
100 hPa still induced small structural changes, and therefore that 
wetting at an even greater suction was required for this type of 
soil. Nevertheless, the results of Dexter et al. (1984) presented 
above together with the other observations, suggest that a suction 
of 100 hPa is appropriate for slow wetting of most soils because it 
avoids structural disruption in previously dried samples.
Rewetting of dried samples has also been done using mist 
(Kemper and Koch, 1966) to provide slow wetting and using 
wetting under vacuum to prevent aggregate “bursting” by 
entrapped air (although this does not prevent fragmentation by 
diff erential swelling). Th e whole subject of clay dispersion and 
soil stability is confounded by the absence of universally accepted 
standard methods for collection, storage and pretreatments of 
soil samples. Many laboratories have their own methods, and this 
usually makes comparison of results impossible.
Soils that contain RDC can form surface crusts or can 
“hard-set” on drying (Kay and Dexter, 1992). Th ese eff ects can 
prevent crop emergence and can increase the energy requirement 
for tillage. It has been shown that the tensile strength of soil 
when dry is approximately proportional to the amount of RDC 
that was present before drying (Chan, 1989; Shanmuganathan 
and Oades, 1982; Watts et al., 1996). Similarly, the soil friability 
is smaller when more RDC is present (Shanmuganathan and 
Oades, 1982; Dexter and Watts, 2000).
Th e content of RDC in soil has been found to be negatively 
correlated with a measure of soil physical quality, S (Gaţe, 2006). 
Soil physical quality is given by the slope of the water-retention 
curve at its infl ection point (when plotted as gravimetric water 
content against the logarithm of pore water suction). Soils with 
high S have been found to be more friable (Dexter, 2004), easier 
to till (Keller et al., 2007) and are not hard-setting (Dexter, 2004). 
Th ese eff ects of S on soil physical behavior are consistent with the 
eff ects of RDC on friability described above. Soils with high S 
also have lower penetration resistance (Dexter et al., 2007) and 
446 SSSAJ: Volume 75: Number 2 • March–April 2011
higher saturated hydraulic conductivity (Han et al., 2008) which 
are further attributes of soils with good physical quality.
Most previous work has been based on soil samples that are 
initially either dry or at (or close to) fi eld capacity. In this paper, 
the sensitivity of RDC to initial soil water potential over the 
whole range of potential from pF2 to pF6.47 is reported for the 
fi rst time. Here, pF is defi ned conventionally as pF = log10(h), 
where h is the pore water suction measured as a water column 
height (cm). Th is enables the behavior of soil in the laboratory 
and the fi eld to be better understood. It also enables problems 
associated with predrying (e.g., during storage or as a standard 
pretreatment) on the results of soil stability tests to be either 
avoided or taken into consideration. We propose procedures to 
avoid or to minimize these problems and thereby to improve the 
accuracy of comparisons of RDC content between soils.
MATERIALS AND METHODS
Soil Sampling
Samples of arable soils from northern France were collected with 
minimum disturbance in the spring of 2007 from soils that had not been 
tilled in that year. Th e soils were selected to have a range of clay and OC 
contents. Th e Villamblain and Faux Perche soils were on private farms 
and were managed conventionally including plowing to a 25-cm depth. 
Th e Boigneville A and B soils were located at the Institut Technique 
des Céréales and Fourrages at Boigneville. Th e Boigneville A soil was 
conventionally tilled by plowing (CT) whereas the Boigneville B soil was 
direct-drilled (DD). Th e plots with these treatments were established in 
1970. Th e Villamblain and Faux Perche soils were sampled at the 10- to 
15-cm depth. Th e Boigneville soils were sampled at the 0- to 5-cm depth 
to obtain the higher organic carbon content near the surface in the DD 
treatment. All the soils were cropped annually with cereals.
Th e soils were at or close to fi eld capacity at the time of sampling. 
Here, the term “fi eld capacity” is used in the sense of the water content 
to which a soil in the fi eld will drain over a period of 2 or 3 d starting 
from saturation. Measurements show that the pore water suction at fi eld 
capacity is around h = 100 hPa (corresponding to pF2). Aft er collection, 
the samples were kept moist in air-tight containers and were stored until 
required in a refrigerated room at 4°C to minimize any eff ects due to 
biological activity.
Characterization of the Soil Solids
Subsamples of the soils were analyzed by the Laboratoire 
d’Analyses des Sols d’Arras for particle-size distribution (by sieving and 
sedimentation) according to French standard NF X 31–107 and content 
of OC (by combustion at 1000°C) according to French standard NF 
ISO 10694. Correction for carbonate carbon was not necessary for these 
soils. Exchangeable cations were analyzed by the same laboratory using 
the cobaltihexamine chloride method according to French standard NF 
X 31–130. Full details of these experimental procedures are described 
at http://www.lille.inra.fr/las/methodes_d_analyse/sols/ (verifi ed 24 
Jan. 2011).
Th e clay fractions (<2 μm) were separated and prepared according 
to Robert and Tessier (1974). Th ree pretreatments were prepared: 
Mg–saturated with Mg; EG–saturated with Mg + ethylene glycol; and 
K– heated at 520°C. Th e mineralogy of the clay fractions was studied using 
an x-ray diff ractometer (Bruker AXS model D8 Advance, Bruker AXS 
Inc., Madison, WI) with a Cobalt anode giving a wavelength of 0.179 nm. 
Th e pretreatments enable diff erent clay minerals to be identifi ed and 
separated. For example, the peaks from kaolin at 0.714 and 0.357 nm 
disappear aft er heating to 520°C; the peak for chlorite at 1.40 nm 
persists aft er heating; and smectite gives a wide peak at 1.72 nm aft er 
the EG treatment. Th e diff ractograms were analyzed to give qualitative 
estimates of the diff erent clay minerals present.
Water Retention Characteristics
The water retention characteristics of the experimental soils were 
required to enable us to calculate the water potentials corresponding 
to diff erent soil gravimetric water contents and also as part of the soil 
physical characterization. We used the pF scale for water potential as 
defi ned above. Because of the small errors involved, we assumed that 
potentials based on h measured in cm H2O and in hPa are equivalent 
(more accurately, 1 cm H2O = 0.906 hPa at latitude 50°N, which gives 
an error in pF values of 0.043 for all values of h, which may be neglected 
in the context of this paper).
Soil water retention was measured using two methods. In the 
range h = 10 to 15000 hPa (i.e., pF1.0–pF4.2) of pore water suction the 
water retention was measured on membrane and ceramic pressure plate 
extractors in conventional ways (e.g., Klute, 1986; Reynolds and Topp, 
2008). Th e samples were collected and stored moist, were then saturated 
for 1 d and then drained to 12 water potentials equally spaced over the 
above range of pF. Separate duplicate samples were measured for each 
soil and each suction. Th e mean values of the resulting gravimetric water 
contents were used in the curve-fi tting and calculations.
To extend the range of measurements to drier conditions, we 
equilibrated samples over saturated solutions of the salts KCl, NaCl, 
NaBr, MgCl2, and LiCl to produce values of relative humidity of 85.1, 
75.7, 59.1, 33.1, and 11.3%, respectively at 20°C (ASTM, 2002). Th e 
Kelvin equation shows that these values correspond to pF values of 5.34, 
5.58, 5.85, 6.18 and 6.47, respectively (Dexter and Richard, 2009).
Solution saturation was maintained by keeping each solution in 
contact with excess, un-dissolved salt. Th e samples together with their 
saturated salt solutions were kept in sealed plastic boxes which were 
kept in larger thermally insulated boxes which were kept in a constant 
temperature room. Every week, the boxes were opened and the solutions 
stirred to destroy any fi lms that might have formed at the liquid surface. 
At these times, samples were taken for measurement of the gravimetric 
water content. When no further weight loss was occurring, it was 
assumed that thermodynamic equilibrium had been reached in which 
case the samples were ready for use. Equilibration required about 1 mo 
over LiCl and 3 mo over KCl.
Th e choice of water retention equation was critical. We used 
the Groenevelt and Grant (2004) equation because it is based on 
thermodynamic principles (Groenevelt and Bolt, 1972) and is appropriate 
for systems in thermodynamic equilibrium. Th is is in contrast to other 
water retention equations, for example, the van Genuchten (1980) 
equation, which oft en predict fi nite residual water contents as pF →∞. 
Finite residual water contents are a characteristic of soils that have been 
dried by immiscible displacement (e.g., in which air displaces water) and 
SSSAJ: Volume 75: Number 2 • March–April 2011 447
 
which are not in thermodynamic equilibrium. With drying by diff usion 
(e.g., slow evaporation, as done here), there is no residual water content 
and the system is at (or close to) thermodynamic equilibrium.
Water retention data over the whole range of applied potentials (1.0 
< pF < 6.47) were fi tted to the Groenevelt and Grant (2004) equation:
w k k k
n n=
−
( )
⎛
⎝
⎜⎜
⎞
⎠
⎟⎟−
−
( )
⎛
⎝
⎜⎜
⎞
⎠
⎟⎟
⎡
⎣
⎢
⎢
⎤
⎦
⎥
⎥
1
0
0
0exp exp
pF pF
 [1]
where k1, k0, and n are adjustable parameters. Th is equation states that 
the water content will be zero at pF0.
Before we used Eq. [1], we determined the value of pF0. Th is was 
done using the water contents obtained by equilibrating samples of 
the soils over saturated salt solutions as described above. For each soil, 
the value of pF at which the water content is predicted to be zero was 
obtained by extrapolation. Th e mean of these values for the four soils 
was then used as pF0 in the calculations for all the soils.
Equation [1] has an infl ection point (a point where the curvature 
= zero) at:
pFi
n
nk
n
=
+( )
⎡
⎣
⎢
⎤
⎦
⎥0
1
1
 [2]
w k k
pF
n
ni n=
−
( )
⎛
⎝
⎜
⎜
⎞
⎠
⎟
⎟
−
− +( )⎛
⎝
⎜
⎞
⎠
⎟
⎡
⎣
⎢
⎢
⎤
⎦
⎥
⎥
1
0
0
1
exp exp [3]
Th e curve of w plotted against pF has a slope at the infl ection 
point given by:
d
d pF
w k
n
n
nk n
nk
n n
( )
=−
− +( )⎡
⎣
⎢
⎤
⎦
⎥
+⎡
⎣
⎢
⎤
⎦
⎥
+( )
1 0
0
1
1 1
exp *
[ / ]
 [4]
Th e slope at the infl ection point when w is plotted against 
ln(h) has been used as an index of soil physical quality, S. In previous 
work, S was calculated from fi tted parameters of the van Genuchten 
(1980) water retention equation. Here, the slope is calculated from the 
fi tted parameters of the Groenevelt and Grant (2004) equation, as in 
Eq. [4], above. Th is procedure gives values for the slope at the infl ection 
point which we designate S*:
S dw
d pF
*
ln
*=
−
( ) ( )
1
10
 [5]
In Eq. [5], the minus sign is added simply to make values of 
the index positive. Values of S and S* are very close and are considered 
to be equivalent for the purposes of this paper. Values of S > 0.035 are 
associated with good soil physical properties whereas values of S < 0.035 
are associated with poor soil physical properties. However, the changes 
in behavior are progressive and there is no sudden change at S = 0.035. 
A fuller description of S and of the soil physical behavior associated with 
diff erent values of S is given in Dexter and Czyz (2007).
Th e Groenevelt and Grant (GG) equation, Eq. [1], was inverted to 
enable the values of pore water suction, expressed as pF, corresponding 
to diff erent gravimetric water contents, w, to be calculated using
pF
n
k
k
pF
w
kn
= ⎛
⎝
⎜
⎞
⎠
⎟
−
−
( )
⎛
⎝
⎜⎜
⎞
⎠
⎟⎟−
⎡
⎣
⎢
⎢
⎤
⎦
⎥
⎥
⎧
⎨
⎪
⎪⎪exp ln
ln exp
1 0
0
0 1⎩⎩
⎪
⎪
⎪
⎫
⎬
⎪
⎪⎪
⎭
⎪
⎪
⎪
⎡
⎣
⎢
⎢
⎢
⎢
⎢
⎢
⎤
⎦
⎥
⎥
⎥
⎥
⎥
⎥
 [6]
Th e GG equation is appropriate for all soil-water systems that 
are in thermodynamic equilibrium irrespective of whether equilibrium 
has been attained through convective or diff usive movement of water 
(e.g., evaporative drying). We therefore used the GG equation for all 
calculations in this paper.
Estimation of the Specifi c Surface Areas
Th e specifi c surface areas of the soils were estimated by two 
methods. Th e fi rst was the method of described in Kutilek and Nielsen 
(1994). In this method, the equilibrium gravimetric water content, wm, 
at a relative humidity of 20% is related to the specifi c surface area, Am, by
Am = 3160wm, m2 g−1 [7]
Th e second was based on the correlation between Am and 
the cation exchange capacity (CEC) that was obtained by Bruand and 
Tessier (2000) using data measured on 31 horizons of French soils. 
Th ey measured Am of the clay fraction by N adsorption. Th e regression 
equation given in their Fig. 1 can be inverted and rescaled to give:
Am = 2.13 CEC + 3.19, m2 g−1, r2 = 0.85 [8]
where the CEC values are in cmol kg−1 as in our Table 2. Although 
Bruand and Tessier (2000) measured only the clay fraction, we applied 
Eq. [10] to the whole soil on the assumption that most of the surface 
area and the CEC of the whole soil are associated with the clay fraction.
Slow Drying Experiments
We needed to dry soil samples from fi eld capacity (pF2) to the 
whole range of water potentials from pF2 to air-dry which is around 
pF6. Th e middle part of this range, from about pF4.2 to pF5.2 is 
diffi cult to produce using normal experimental techniques. Porous 
ceramic plates are reaching the limits of their eff ectiveness at around 
pF4.2 (Gee et al., 2002; Cresswell et al., 2008). Equilibration over 
osmotic solutions is diffi cult for values of pF < 5.3 because it requires 
very accurate temperature control to avoid problems with condensation 
and takes an excessively long time. For these reasons, we decided to 
produce the whole range of water potentials by slow-drying of initially 
moist soil samples in air.
Subsamples ofthe soils of about 20 g were crumbled by hand 
without shearing into 1- to 3-mm aggregates and were placed in plastic 
bags that had numerous small pin-holes in them. Th ese bags with soil 
were placed on horizontal drying racks in the laboratory at 22°C. Th e 
subsamples dried very slowly by loss of water vapor through the pin-holes. 
It was assumed that water would redistribute relatively rapidly in the vapor 
phase within the bags to give a relatively uniform water potential or pF. If 
the redistribution within the bags is more rapid than water loss from the 
bags, then the soil within the bags will be in quasi-equilibrium. Th e bags 
were weighed every day, and from knowledge of the initial mass of moist 
soil in each bag and its initial water content, it was possible to calculate 
448 SSSAJ: Volume 75: Number 2 • March–April 2011
the water content of the soil at that time. From the water content and 
the water retention curve, it was possible to calculate the pF of the water 
remaining in the soil. Th is slow drying of soil from fi eld capacity (about 
pF2) to air-dry (about pF6) took about 4 wk.
We required a range of pF values for each soil. Th erefore, we 
calculated the mass that each soil + bag should have to give the required 
values of pF. When a soil had dried so that the approximate required 
mass of soil + bag had been attained, then the soil samples were taken 
from the bags. About one-half was used for measurement of the 
gravimetric water content in two replicates and the other half was used 
for measurement of RDC and electrical conductivity (EC).
Measurement of Readily dispersible Clay
Th e content of RDC was measured at the diff erent water contents 
produced by slow drying of initially moist soil as described above. Readily 
dispersible clay was measured by the amount of white light scattered 
by the colloidal particles (mostly clay) in aqueous suspension using a 
turbidimeter which measures the turbidity, T (NTU/[g L−1]) of a 
suspension; T is proportional to the concentration of the particles in 
suspension. Th e principles and limitations of turbidity measurement are 
discussed by Sadar (1998). Th e method used in this work for measurement 
of the RDC content of soils is based on that described previously (Czyz et 
al., 2002; Czyz and Dexter, 2008). It was further developed and improved 
for this work and is described in full in the Appendix.
Th e EC of the solutions obtained aft er immersion of dried soil 
in 125 mL of deionized water was measured with a Hanna model 
HI99300 conductivity meter (Hanna Instruments, Woonsocket, RI) 
that was calibrated periodically with standard solutions.
Th e reversibility of the eff ects of drying on RDC was investigated 
in two ways as follows.
Method 1: low intensity drying with slow rewetting.
Samples of the moist soil were dried on ceramic pressure plate 
extractors to h = 4500 hPa (or pF3.65) for 4 d. Th ey were then taken 
from the pressure plate and wetted to h = 10 hPa for 2 d and then 
dried to “fi eld capacity” at h = 100 hPa (or pF2) for 4 d. Th e purpose 
of the wetting to 10 hPa before drying to 100 hPa was to ensure that 
the samples were on the “drying limb” of the water characteristic curve.
Method 2: high intensity drying followed by rapid rewetting
In the case of the samples that had been dried to h > 10 MPa (pF 
> 5), the test of reversibility of the eff ect of drying was to rapidly rewet 
the samples by immersion and to leave the 
samples in water for up to 2 wk.
Soil Drying in the Field
To relate our laboratory studies above 
to reality, we sought to answer the question 
“how dry do soils become in the fi eld”? 
Although we did not have data for the French 
soils described above, we had data for the top 
2 cm of an arable sandy loam at the Waite 
Institute in Adelaide, South Australia. Th e 
climate in South Australia is Mediterranean 
with cool, wet winters, and hot, dry summers. Th ese measurements 
were made in 1981 by the fi rst author and have not been published 
previously. Forty samples were collected each working day (20 at 09.00 
h and 20 at 15.00 h) and the gravimetric water contents were measured. 
Th is was continued for a period of about 9 mo fi nishing on December 
18. It should be noted that Australia is in the southern hemisphere, 
and therefore that mid-summer occurs on December 21. Th e mean 
gravimetric water content at each sampling time was calculated.
Th e water retention characteristic was measured using sintered-
glass funnels with hanging water columns and with ceramic pressure 
plate extractors. Th e pF values corresponding to the measured fi eld 
water contents were calculated with Eq. [6] using the value pF0 measured as 
described and were ranked within each of the months September-December.
RESULTS AND DISCUSSION
Soil Characteristics
Th e compositions of the experimental soils are given in 
Table 1. Information about the clay mineralogy and exchangeable 
cations is given in Table 2 which shows that calcium-saturated 
illite and chlorite were the principal clay fractions. Th e amounts 
of complexed and non-complexed clay and OCwere estimated 
as described in Dexter et al. (2008) and are presented in Table 3.
Th e water retention characteristics obtained by drying over 
saturated salt solutions are shown in Fig. 1. It can be seen that 
there is very little eff ect of organic matter on water retention 
Fig. 1. Water retention for values of pF > 5.34 as obtained by 
equilibration of samples over saturated salt solutions.
Table 1. Compositions of the experimental soils together with the USDA texture classes, 
contents of organic carbon (OC) and the measured values of pH.
Soil
 Villamblain Boigneville A (CT) Boigenville B (DD) Faux Perche (9P55)
clay, g kg−1 331 260 236 118
fi ne silt 326 285 299 342
coarse silt 326 375 387 487
fi ne sand 13 63 61 30
coarse sand 4 17 17 23
USDA texture class† si cl l si l si l si
OC, g kg−1 13.0 13.3 28.9 12.1
pH 7.90 6.33 4.86 7.91
† si = silt, cl = clay, l = loam.
SSSAJ: Volume 75: Number 2 • March–April 2011 449
 
for pF > 5.3 as shown by the comparison of the results for the 
Boigneville A and B soils. Th e results shown in Fig. 1 were fi tted 
to Eq. [9] and the resulting parameters are given in Table 4.
w B C pF= + ( ) , kg kg−1 [9]
Equation [9] is not applicable at saturation (h = 0), where the pF 
= −∞, but gives a good fi t over the experimental range 5.34 < 
pF < 6.47. As can be seen in Fig. 1 and can be 
calculated from Eq. [9] with the appropriate 
coeffi cients, extrapolation of these results 
gives the prediction that the water content, w, 
would be zero at pF = 6.65 ± 0.02 for these 
soils. Th is is the value that we used for pF0 in 
Eq. [1], [3], and [6].
Th e fi tted parameters of the Groenevelt 
and Grant (2004) water retention equation 
for the four soils are given in Table 5. As 
an example, the 17-point water-retention 
data for the Boigneville A soil are shown 
in Fig. 2 together with the smooth curves 
corresponding to the fi tted Groenevelt 
and Grant (GG) water retention equation 
described above. Th e results and fi ts for the 
other three soils were equally good.
Th e specifi c surface areas calculated by 
the two methods are given in Table 6. It can 
be seen that the results are similar except for 
the Boigneville B soil that has a large content of 
organic matter. Th ese values put our soils into 
Class II of Dogan et al. (2007).
Amounts of Readily Dispersible Clay
Th e amounts of RDC obtained by the 
slow drying treatment and the application of 
Eq. [6] are shown in Fig. 3. It can be seen that 
RDC decreases progressively with increasing antecedent pF 
reaching approximately constant or asymptotic values for pF > 
5, approximately. Th e RDC data for the slowly air-dried samples 
were fi tted to
RDC = k + Eexp[-F(pF-2)], NTU/(g L−1) [10]
where k, E, and F are adjustable coeffi cients. Th e term X = (pF– 
2) was chosen so that X would have the value 0 at 
fi eld capacity (heretaken to be h = 100 hPa). At fi eld 
capacity, Eq. [10] shows that RDCfc = k + E whereas 
the asymptotic value of RDC for high values of pF 
is given by k. It should be noted that for saturated 
soil, for which h = 0, the pF value is −∞. Th e fi tted 
parameters for Eq.[10] are given in Table 7 which 
shows that the value of the exponent F is similar for 
all the soils with an approximate average value of F 
= 1. In other words, an increase in the initial pF of 
the soil water by 1 reduces the amount of RDC by a 
factor of 1/e = 0.37, approximately. Th is is in spite 
of the diff erences in the compositions of the four 
experimental soils. For our experimental soils, Eq. 
[10] can be approximated by
RDC pF RDC
RDC RDC
edry
fc dry
pF( ) ( )−
−
⎡
⎣
⎢
⎢
⎤
⎦
⎥
⎥
= − −2 [11]
Table 2. Mineralogy of the clay fraction and exchangeable cations in the 
experimental soils.
Soil 
Villamblain Boigneville A (CT) Boigneville B (DD) Faux Perche (9P55)
Mineralogy
illite low dominant dominant some
kaolinite medium high high high
vermiculite medium some some some
smectite medium low low none
chlorite high very low very low medium
quarts very low very low very low very low
Exchangeable cations
Ca, cmol kg−1 18.6 13.2 6.84 9.50
Mg, cmol kg−1 0.86 0.73 0.60 0.44
K, cmol kg−1 0.59 0.71 1.09 0.43
Na, cmol kg−1 0.060 0.038 0.042 0.024
CEC, cmol kg−1 20.1 14.7 9.9 9.3
Table 3. Measured amounts of clay (C) and organic carbon (OC) with estimated 
amounts of complexed clay (CC), non-complexed clay (NCC), complexed organic 
carbon (COC) and non-complexed organic carbon (NCOC) in the experimental soils.
Soil 
Villamblain Boigneville A (CT) Boigneville B (DD) Faux Perche (9P55)
C, g kg−1 331 260 236 118
OC, g kg−1 13.0 13.3 28.9 12.1
CC, g kg−1 130 133 236 118
NCC, g kg−1 201 127 0 0
COC, g kg−1 13.0 13.3 23.6 11.8
NCOC, g kg−1 0.0 0.0 5.3 0.3
Table 4. Parameters of the water retention equation (Eq. [9]) for soils dried 
over saturated salt solutions to values of pF in the range 5.34 < pF < 6.47. Also 
given are the values of the intercept, pF0, at which the water content, w, is 
predicted to be zero.
Soil 
Villamblain Boigneville A (CT) Boigneville B (DD) Faux Perche (9P55)
B, kg kg−1 0.3567 0.2650 0.2475 0.1190
C, kg kg−1 -0.05395 −0.03987 −0.03713 −0.01780
r2 0.997 0.986 0.998 0.990
pF0 6.61 6.65 6.66 6.68
Table 5. Fitted values of the parameters of the Groenevelt and Grant (GG) 
water retention equations and of the index of soil physical quality, S*.
Soil 
Villamblain Boigneville A (CT) Boigneville B (DD) Faux Perche (9P55)
k1, kg kg−1 2.4224 0.6508 0.4338 0.2765
k0 7.2667 6.8770 12.0482 80.1106
n 0.6459 1.0453 2.0658 3.9006
S* 0.0267 0.0245 0.0475 0.0576
450 SSSAJ: Volume 75: Number 2 • March–April 2011
for values of pF > 2. We suggest that Eq. [11], together with 
values of RDCfc and RDCdry measured as described in the 
Appendix, provides estimates of the content of RDC(pF) at any 
value of pF for any soil that has not been drier than pF since the 
last long wet period. However, Eq. [11] needs to be tested on a 
wider range of soils than were used in this study.
We suggest that diff erent soils should be compared at pF2. 
However, soils are oft en collected in the fi eld when drier than 
this. In this case, we suggest that the soil samples be wetted slowly 
at h = 100 hPa for 1 wk before measurement of RDCfc.
From the values given in Table 7, it is possible to calculate 
the ratio (k+E)/k which is the ratio of RDC from soil at fi eld 
capacity to that from intensively 
dried soil. Th is ratio has values of 28, 
27, 10, and 22 for the Villamblain, 
Boigneville A, Boigneville B, and 
Faux-Perche soils, respectively. 
Th ese values are similar to those 
found for Polish soils by Gaţe 
(2006). Th e Polish soils had smaller 
contents of clay, on average, than the 
French soils used here. Th e smaller 
value of this ratio for the Boigneville 
B soil is mainly because of the eff ect 
of organic matter in reducing the 
dispersion of clay from the soil when 
moist (i.e., at fi eld capacity).
Values of RDCfc calculated as 
above may be compared with the 
measured values of total clay (C) 
and estimates of non-complexed 
clay, NCC, given in Table 3. Th is 
comparison shows that when all of 
the clay is complexed with organic 
matter, then the amount of RDCfc 
is close to zero. Also given in Table 3 are the values of non-
complexed organic carbon, NCOC, calculated as described by 
Dexter et al. (2008).
Th e van der Waals force decreases as the sixth power of 
distance from a surface whereas the pore water suction between 
parallel plates (such as clay plates) decreases as the fi rst power 
of the plate separation (Kutilek and Nielson, 1994). Th erefore, 
we can use estimates of pore water suction down to 5 nm 
spacing. Th e capillarity equation for parallel plates (Kutilek 
and Nielson, 1994) shows that the pore water suction at this 
spacing is 28 MPa which corresponds to pF5.4. However, the 
clay plates are not parallel (Bruand and Zimmer, 1992), so we 
shall make an informed guess that the clay particle surfaces 
begin to fall into solid-solid contact at pF5.0. For soil wetter 
than this, we hypothesize that the eff ect of drying on RDC 
is reversible whereas for soil drier than this, we hypothesize 
that the eff ect of drying is irreversible. Th is estimated value of 
pF5.0 is bracketted nicely by the values of pF3.65 at which we 
found reversibility and pF6.0 at which we found irreversibility. 
Th e exact value of pF at which the eff ect on RDC becomes 
irreversible remains unknown.
Figure 4 shows the values of RDC for soil samples that had 
been equilibrated over saturated salt solutions. Th ese results 
Fig. 2. Fit of the Groenevelt and Grant (GG) water retention equation 
for the Boigneville A soil. Experimental points are as follows: circles 
are from membrane and ceramic pressure plate extractors, and 
squares are from equilibration over saturated salt solutions.
Table 6. Estimates of the specifi c surface area (m2 g−1) for the 
whole soil as estimated by water adsorption at pF6.34 and cation 
exchange capacity (CEC).
Soil
 Villamblain
Boigneville A 
(CT)
Boigneville B 
(DD)
Faux Perche 
(9P55)
(a) water adsorption 51.9 44.0 43.5 22.0
(b) CEC method 46.0 34.5 24.3 23.0
Fig. 3. Amount of readily dispersible clay, RDC [NTU/(g L−1)], as a function of antecedent water potential 
(expressed as X = pF-2 where the origin, X = 0, has been shifted to correspond to fi eld capacity, pF2). 
Note that all values of pF were obtained using Eq. [6].
SSSAJ: Volume 75: Number 2 • March–April 2011 451
 
show a slight increase in RDC with increasing pF in the range 
5.34 < pF < 6.47. Th is eff ect is diff erent from that expected from 
Eq. [10]. We conjecture that this may be due to greater eff ects 
when drier soil is rewetted (c.f. heat of wetting). Comparison of 
the results in Fig. 3 and 4 shows that there must be some value of 
pF at which the amount of RDC is minimum. Th e results from 
the experiments reported here do not allow us to determine this 
value with any accuracy.
Th e values of RDC at high pF obtained by the air-drying 
method (RDCAD) were always larger than the values obtained 
by drying over saturated salt solutions (RDCSS). Th e average 
diff erence, ΔRDC was 0.04 NTU/(g L−1). We conjecture that 
this diff erence may be due to non-uniform drying of the soil in 
the perforated plastic bags. Th is will need to be tested in future 
experimental work. However, this small value of ΔRDC does 
not aff ect the conclusions from this work.
When soil was dried from pF2 to pF3.65, and then slowly 
rewetted to pF2 on a sand table (Method 1, described above), 
then the original value of RDC was obtained (100% regain) 
which is consistent with the discussion above.
Electrical Conductivity
Th e values of EC of the solutions produced during the RDC 
tests are shown in Fig. 5. It can be seen that for the Villamblain 
and Faux Perche soils the resulting EC is greaterwhen the soil is 
initially drier (with correlations of r = 0.56, p < 0.0001 and r = 
0.95, p < 0.0001 for these soils, respectively). Th is is consistent 
with the idea that when soil is initially drier, there will be greater 
disruption of the system on rewetting, for example by increased 
heat of wetting (Prunty and Bell, 2005). For the Boigneville A 
soils, there is no signifi cant correlation (r = 0.074, p = 0.77). 
In contrast, the Boigneville B soil shows a negative correlation 
(r = −0.60, p = 0.0089). Th is latter soil has a high content of 
organic matter and a content of NCOC of 5.3 g kg−1 as shown 
in Table 3. We can hypothesize that organic matter protects the 
system against the eff ects of rapid 
rewetting. Th is may be because the 
organic matter is more hydrophobic 
when the soil is drier (Bayer and 
Schaumann, 2007) and that greater 
hydrophobicity leads to slower 
rewetting (Doerr et al., 2000; Hallett 
et al., 2004). Th ese observations, 
although preliminary, enable us to 
propose an hypothesis for future 
testing: that is, that d(EC)/d(pF) 
is negative when NCOC is present.
To learn more about this 
EC phenomenon, we analyzed 
the supernatant in the 30-mL 
turbidimeter cells aft er measurement 
of turbidity for the cations Al, Ca, 
Fe, K, Mg, and Na. We found that 
the EC was strongly correlated with 
the concentrations of Ca and K in 
solution aft er the rapid wetting as 
shown in Eq. [12].
Table 7. Values of the parameters of Eq. [10] giving RDC [NTU/(g L−1)] 
as a function of antecedent water potential in the range 2 < pF < 6.
Soil 
Villamblain
Boigneville A 
(CT)
Boigneville B 
(DD)
Faux Perche 
(9P55)
K, NTU/(g L−1) 0.143 0.134 0.089 0.157
E, NTU/(g L−1) 3.906 3.499 0.838 3.339
F 1.128 0.719 1.017 1.031
r2 0.984 0.978 0.941 -
Fig. 4. Amount of readily dispersible clay, RDC [NTU/(g L−1)], as a 
function of antecedent water potential (expressed as pF) for values 
of pF > 5.34. Samples were equilibrated over saturated salt solutions.
Fig. 5. Electrical conductivity (EC) of the suspension after adding soil that had been dried to different 
values of antecedent water potential (expressed as X = pF-2 where the origin, X = 0, has been shifted to 
correspond to fi eld capacity, pF2). Note that all values of pF were obtained using Eq. [6].
452 SSSAJ: Volume 75: Number 2 • March–April 2011
EC = 0.3 + 4.40[Ca] + 5.02[K], r2 = 0.975 [12]
 (±3.2) (±0.43) (±0.48) 
where EC is in μS cm−1 and the concentrations [Ca] and [K], are 
in mg L−1. Th ere was no statistically signifi cant correlation with 
the other cations analyzed. Further work with a wider range of 
soils is needed to elucidate the mechanisms and processes that 
occur during rewetting of soil samples. However, the measurement 
of EC does seem to be one method of gaining useful information. 
In additional experiments, not described here, we showed that 
the changes in the electrolyte composition and concentration 
following immersion of dried versus moist soil had no signifi cant 
eff ect on the amounts of dispersed clay in suspension.
Th e water contents measured in the fi eld in South Australia 
were converted into pF values using Eq. [7] using the same value 
of pF0 = 6.65. Th e results are shown in Fig. 6 where it can be 
seen that soil can become dry enough under natural conditions 
to undergo the eff ects described above. It is likely that the soil 
will become even drier during January and February, although 
this was not measured.
Th e results presented above can be considered in relation to 
annual cycles of tillage and climate. For arable (i.e., plowed) soils 
the tilled layer becomes mixed every year. It can then be argued 
that over a period of many years, all of the soil in the arable layer 
will have been at the surface and will have been subjected to 
intense drying and rewetting many times. A logical conclusion 
from this is that all of the clay in the arable layer will have become 
stabilized and will have a low content of RDC as described above.
If the depth of plowing is P cm and the depth to which 
an intensity of drying suffi cient for the “irreversible” changes 
discussed above to occur is D cm, then we can model this very 
simply. We have done this with values of D/P = 0.05, 0.1 and 
0.2 and obtained the results that 90% of the change will have 
occurred aft er 45, 21, and 9 yr, respectively.
Our experimental results for moist arable soils show that 
there is a lot of RDC present at the end of a long, wet period 
such as the winter in temperate climates. Th erefore, the changes 
induced by drying and rewetting that we have reported above 
cannot be irreversible. Similar results were obtained with the 
tensile strength of arable soils in South Australia which has a 
Mediterranean climate (Kay and Dexter, 1992). At the end 
of the long, wet winter, the tensile strength of the air-dried 
soil samples was found to be high. Th is was attributed to the 
cementation produced by the particles of dispersed clay as the 
soil dried. However, this value reduced linearly with increasing 
number of natural wetting/drying events during the summer. 
Th ese wetting/drying events were easily countable during the 
predominantly hot, dry summers in that area. Each year, the 
tensile strength started from a high value again showing that 
during the long, wet winter of approximately 6 mo duration, the 
clay had dispersed and could again contribute to cementation 
when the soil dried.
We have not yet discovered either the mechanism or the 
dynamics of this “re-setting”. It seems that the soil “resets” its 
internal clock during long, wet periods (e.g., winter in temperate 
climates) as described by Kay and Dexter (1992). Th e evidence 
suggests that this resetting occurs in spite of the surfaces “falling 
into direct contact, which is essentially irreversibly”, as described 
in the introduction. Th is “resetting” clearly needs to be the 
subject of future research.
CONCLUSIONS
Th e experiments reported above involved the eff ects of 
drying of initially moist soil by vapor diff usion (evaporation). 
Th is is conceptually and physically diff erent from the immiscible 
displacement method of soil drying (or dewatering) that is 
normally done in the laboratory on pressure plate extractors for 
determination of the water retention characteristics.
We have shown that water contents resulting from equilibration 
through vapor diff usion are well-described over the whole range of 
pF by the Groenevelt and Grant (2004) equation. We found the 
value of the term pF0 in this equation to be pF6.65 ± 0.02.
Readily dispersible clay was found to decrease from RDCfc 
at fi eld capacity (pF2) to a minimum of RDCdry under air-dry 
conditions (typically pF6). Th is decrease was exponential (to 
within experimental error) with an increase in the initial pF of 
the soil water by 1 reducing the amount of RDC by a factor of 
1/e = 0.37, approximately. Equation [11] enables the eff ect of 
predrying to diff erent values of pF on the content of RDC of the 
four experimental soils to be estimated. Diff erences between soils 
aff ect mainly the term RDCfc and this illustrates the importance 
of measuring the properties of moist soil that has not been dried.
Th e tests of soil stability (including clay dispersion) proposed 
by Emerson (1967) and Le Bissonnais (1996) form the bases for 
Australian (Australian Standard, 1997) and French (Association 
Française de Normalisation, 2005) standard methods. Both of 
these tests involve a soil pretreatment that includes intensive 
drying of the soil. Th e Emerson test requires air-drying whereas, 
the Le Bissonnais test requires drying in an oven at 40°C. Under 
Fig. 6. Dryness of the top 2 cm of soil in the fi eld in South Australia. 
The curves show, for the months September–December 1981, 
the proportions of the time for which the soil was drier than the 
corresponding values of pF.
SSSAJ: Volume 75: Number 2 • March–April 2011 453
 
typical laboratory conditions,with an air temperature of 22°C 
and a relative humidity of 50%, these procedures produce 
intensities of drying of pF5.98 and pF6.40, respectively (Dexter 
and Richard, 2009). We have shown that this fi rst value can be 
reached in the fi eld whereas this latter value is drier than was 
found for the top 2 cm of soil in December (mid-summer) in 
South Australia (Fig. 6). Such high intensities of predrying are 
not appropriate for studies of clay dispersion from soils that are 
moist because the drying pretreatment changes the dispersibility 
of the clay fraction irreversibly (or at least for periods of >2 wk).
Th e Emerson (1967) and Le Bissonnais (1996) tests give 
only measures of the clay dispersion from soil that has been 
intensely dried before rewetting (equivalent to k in Eq. [10] and 
Table 7). We have shown that the dispersion of clay from initially 
moist soils is much greater and can be measured accurately by 
the methods described in this paper. We believe that an adequate 
understanding of clay dispersion from soil requires the analysis 
of initially moist as well as dried samples. Th is conclusion is 
consistent with the recommendation given by Pojasok and Kay 
(1990) and Kjaergaard et al. (2004) that soil samples should 
be collected when they are at fi eld capacity and should be kept 
moist during storage. We suggest that the RDC content of 
diff erent soils should be compared at pF2. However, when soils 
are collected drier than this, we suggest that the soil samples 
be wetted slowly at h = 100 hPa for 1 wk before measurement 
of RDCfc. Alternatively, Eq. [11] may be used to estimate the 
amount of RDC that may be expected.
A simple mixing theory suggests that all of the arable 
layer of the soil should have been intensively dried and rapidly 
rewetted many times in the past. Th erefore, we might expect that 
there would be a negligible content of RDC in the arable layer. 
Th e fact that this is not so, suggests that the eff ect of intensive 
drying and rewetting is not irreversible but perhaps only slowly 
reversible. However, in arable soils there is also an annual input 
of mechanical energy from tillage operations that may contribute 
toward reversal of the eff ect of drying (Watts et al., 1996).
Th e work described above on the eff ects of drying on 
soil physical properties leads to the proposal for having two 
pretreatments for determination of RDC and soil stability in 
general: one for soils that are dry in the fi eld and one for soils 
that are moist and that have not dried since the previous long, 
wet period. Th ese give indications of the instability of soils 
at diff erent times of year. Th e experimental procedures used 
for measurement of RDC are described in the Appendix. We 
support strongly the fi nal statement of Angers et al. (2008) 
which we simplify slightly as follows: “Because variations in 
sampling, storage, and pretreatment aff ect the results, all steps 
in the analysis should be described in great detail when research 
results are published”.
APPENDIX
Experimental Procedure used to Determine 
the Content of Readily Dispersible Clay in Moist 
and Dry Soil
A. Types of Test
We describe two types of test:Test 1. For moist soil to determine 
the content of RDC at fi eld capacity, RDCfc, where fi eld capacity is de-
fi ned as a pore water suction of h = 100 cm H2O = pF2 ≈ 98.6 hPa. 
Th is is relevant to erosion of soil or transport of colloids in the spring 
and during snow melt, etc. To determine RDCfc the soil must be ini-
tially wetter than a suction of 0.45 MPa, and must not have been drier 
than this on any occasion since the last long wet period (e.g., winter in 
temperate climates).
Test 2. For dry soil to determine RDCdry, which is relevant to the 
response of soil to rapid wetting in dry periods that may occur aft er seed-
ing or in summer thunderstorms, etc. Th is test involves conditions simi-
lar to those used in the Australian (AS, 1997) and French (Association 
Française de Normalisation, 2005) tests of soil stability. To determine 
RDCdry the soil can be collected at any degree of wetness.
B. Sample Collection and Storage
Soil samples must be taken from the fi eld with the minimum dis-
turbance possible. Th e sampling strategy (depths and spatial distribu-
tion of samples) will depend on the hypotheses to be tested. We nor-
mally take 5 or 10 replicates depending on the heterogeneity of the soil 
and on the desired accuracy of the mean values. Th e samples should be 
placed in air-tight containers for transport to the laboratory. Th e sam-
ples should be placed in the shade and kept as cool as possible. At the 
laboratory, put the samples in hermetically sealed containers and store 
in a constant temperature room at 4°C until required.
C. Standardization of Initial Conditions
For Test 1, samples for RDCfc should be placed on a sand table 
adjusted to a suction of h = 100 cm H2O. Th ey should be left for 4 d 
for the water and the soil to equilibrate. For Test 2, samples for RDCdry 
should be dried in a convection oven at 40°C for 2 d. With average labo-
ratory conditions of air temperature of 20°C and relative humidity of 
50%, this produces an equilibrium water potential of pF6.35 (Dexter 
and Richard, 2009).
Aft er equilibration, the samples should be moved rapidly to a vac-
uum desiccator (without desiccant) to prevent exchanges of water with 
the atmosphere (and to cool to room temperature in the case of Test 2). 
Samples for the RDCfc and RDCdry tests must not be placed together 
in the same desiccator.
D. Equipment
A turbidimeter is required to measure the amount of colloids 
(mostly clay) in suspension. It must measure the turbidity of suspen-
sions from the amount of white light scattered. It should not be sensitive 
to solution color (we used a Hach model 2100AN turbidimeter, Hach 
Company, Loveland, CO).
Also required are: a laboratory balance that weighs to 3-digit accuracy 
(that is to 1 mg), laboratory convection oven to work at 105°C, and a pi-
pette with the capacity of the glass turbidimeter cells (30 mL in our case).
454 SSSAJ: Volume 75: Number 2 • March–April 2011
Clean plastic bottles with large-opening screw tops, and Vb = 150 
mL capacity. We use about 100 of these. Each bottle must be numbered 
and calibrated by adding 125 mL of water (most easily determined by 
weighing on a balance) and drawing a horizontal line at the surface of 
added water. Th ese lines then enable water to be added to make 125 mL 
of suspension in the tests. Th e importance of having accurate volumes 
is because the energy input to the suspension depends strongly on the 
volume of the air bubble in the bottles. In this example, the volume of 
air is Va = 150 – 125 = 25 mL).
A large, clean plastic container to hold a supply (e.g., 20 L) of dis-
tilled or deionized water. Th is water must be allowed to stand in the 
laboratory for several hours before use for two reasons. First, it must be 
equilibrated at the laboratory temperature (20–22°C); and second, be-
cause dissolved gases may come out of solution producing microbubbles 
that can aff ect the turbidity readings. Th ese microbubbles may be invis-
ible to the naked eye.
E. Measurement Procedure
First, determine the gravimetric water content, w, kg kg−1 or g g−1, 
of a subsample of the soil by oven-drying at 105°C. While this is drying, 
weigh 4 to 5 g of moist soil (mass = Mm) and place it in one of the cali-
brated plastic bottles. Th en, add distilled or deionized water to make the 
volume of the soil + water in the bottle up to 125 mL. Screw the plastic 
top on the bottle. Record the time, T1.
Aft er 1 h (time = T2 = T1 + 1 h), agitate the soil + water in a stan-
dard way by holding the bottle in one hand and inverting it four times. 
Each inversion should take about 0.5 s. Th en loosen the plastic bottle top 
so that it can be removed later without disturbing the suspension or the 
sediment in the bottle. Leave the bottle to stand for 18 h (time = T3 = T2 
+ 18h), without any disturbance for sedimentation to proceed at room 
temperature (20–22°C). Th is time of 18 h is convenient because it enables 
samples to be prepared in the aft ernoon and measured the next morning.
Aft er the 18 h wait, take 30 mL of suspension from the center of 
the bottle with the pipette without disturbing the suspension or the sed-
iment in the bottle. Put the collected suspension in the glass turbidim-
eter cell by running it down the inside of the cell to avoid the formation 
of bubbles that can aff ect the turbidity reading. Measure the turbidity in 
units of NTU (Nephelometric Turbidity Units). Th is value of turbidity 
is a measure of the colloidal content (mostly clay) because the larger min-
eral particles will have sedimented from the initial suspension.
F. Calculations
Calculate the mass of dry soil particles, Md, in the sample from the 
moist mass, Mm, and the gravimetric water content, w, using
Md = Mm/(1+w) 
where Md and Mm are in g.
Normalize the measured turbidity to take account of the eff ects of 
diff erent initial water contents and diff erent initial masses of moist soil:
Normalized turbidity = NTU/(g L−1) 
where g L−1 is the mass concentration of soil particles in the suspension. 
For the example given above, we have:
Normalized turbidity = NTU/(1000*Md/125) 
G. Relating NTU to Absolute Clay Contents
If the suspension used in the turbidimeter cell is dried and weighed 
(this will require a 4 or 5 digit balance), then the mass concentration 
of clay giving a certain turbidimeter reading can be calculated directly.
Alternatively, if the agitation by four inversions in Section E, above, 
is replaced by complete dispersion as is done in standard methods of par-
ticle-size analysis (e.g., involving chemical dispersion with sodium meta-
phosphate, and intense mechanical dispersion using ultrasound and/or 
intense stirring), then the NTU value corresponding to the known total 
clay content of the soil may be determined. Knowledge of this enables the 
clay contents corresponding to other NTU values to be calculated. A fac-
tor K may be defi ned which, when multiplied by the turbidity in NTU/
(gL−1) gives the amount of dispersed clay in g (100 g soil)−1.
Calibration factors, obtained as described above, that relate tur-
bidimeter readings to absolute amounts of clay in suspension diff er 
slightly between soils. Th is is because the diff erent clay minerals present 
may have diff erent particle shapes, size distributions and diff erent light-
scattering properties.
H. Reporting of Results
Th e content of RDC can be reported either in terms of 
normalized turbidity units:
RDCfc = NTU/(g L−1) or 
RDCdry = NTU/(g L−1), as appropriate
or in terms of absolute clay contents
RDCfc = g (100 g soil)−1 or 
RDCdry = g (100 g soil)−1, as appropriate.
ACKNOWLEDGMENTS
A. R. Dexter thanks the leSTUDIUM Institute of Advanced Studies in 
Orléans, France for the fi nancial support that made his work in INRA, 
Orléans possible. Th e authors also thank the INRA Environment and 
Agriculture Division for the award of a “Projet innovant” and the French 
Agency for the Environment and Energy Management (ADEME) 
for the award of a grant. Dr. N.R.A. Bird of Rothamsted Research is 
thanked for mathematical assistance. Olivier Josière is thanked for his 
help with the mineralogical tests.
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