Engineering Geology 53 (1999) 151–159
Chemical osmosis in compacted clayey material and the
prediction of water transport
Th.J.S. Keijzer *, P.J. Kleingeld, J.P.G. Loch
Utrecht University, Institute of Earth Sciences, Department of Geochemistry, P.O. Box 80021, 3508 TA Utrecht, The Netherlands
Abstract
Compacted clay membranes are semi-permeable if the double layers of the clay particles overlap, thereby restricting
the passage of ions. Semi-permeability is quantified by the reflection coefficient s. In the design of clay liners for
waste contaminant water, transport as a result of coupled transport is rarely taken into account. Where large salt
concentration differences exist across natural or man-made clay barriers, water may be transported as a result of
chemical osmosis.
In a flexible wall permeameter two samples of a commercially available Wyoming Na-bentonite were subjected to
a chemical gradient in order to monitor water transport and to obtain values for the reflection coefficient. In both
experiments water transport by chemical osmosis was observed, and reflection coefficients of 0.003 and 0.001 were
obtained, which are significantly lower than those predicted by the Fritz–Marine model and values obtained from
Bolt (1982). However, the values are in good agreement with those obtained by Bresler (1973). Both experiments
showed a period of 50 h of linear pressure increase as a result of chemical osmosis, after which the pressure difference
dropped, i.e. after reaching a maximum pressure difference the water flow was reversed. The reversal of the water
flow is consistent with diffusion osmosis, which is the transport of water as a result of the diffusion of ions in the
absence of an externally applied electrical field. However, diffusion osmosis is limited to clays of low cation exchange
capacity with high pore water concentrations and porosities. © 1999 Elsevier Science B.V. All rights reserved.
Keywords: Bentonite; Coupling; Disposal barriers; Osmosis; Solute transport
1. Introduction
It is well established that clays can act as semipermeable membranes and are therefore capable
of inducing coupled transport (Mitchell, 1993).
Clays exhibit membrane properties when the
double layers of the adjacent clay particles overlap
and both cations and anions are excluded from
the pores. On the other hand, water and noncharged solutes are freely admitted to the mem* Corresponding author. Fax: +31-30-2535030.
E-mail address: [email protected] ( Th.J.S. Keijzer)
brane. The osmotic pressure difference induced by
differences in chemical composition across a semipermeable membrane can be calculated using the
Van’t Hoff equation:
RT
a
ln fresh
(1)
V
a
9
w
salt
in which Dp is the osmotic pressure difference
across a membrane, R the gas constant, T the
absolute temperature, V
9 the mean partial molar
w
volume of water and a the activity of the water at
either side of the membrane. The ability of a
membrane to restrict solute transport is expressed
Dp=
0013-7952/99/$ – see front matter © 1999 Elsevier Science B.V. All rights reserved.
PII: S0 0 1 3 -7 9 5 2 ( 9 9 ) 0 0 02 8 - 9
152
Th.J.S. Keijzer et al. / Engineering Geology 53 (1999) 151–159
as the reflection coefficient, s. For ideal membranes
s is 1, for porous media without membrane properties, e.g. sand, s equals 0 ( Fritz, 1986).
Laboratory experiments on compacted, monomineral clay membranes, mostly bentonite or
kaolinite, have resulted in the description of
coupled transport with different driving forces,
e.g. electro-osmosis (Mitchell, 1993; Grundl and
Michalski, 1996), thermo-osmosis (Dirksen,
1969), chemical osmosis ( Kemper, 1961; Olsen,
1969; Fritz and Marine, 1983) and reverse osmosis
( Kharaka and Smalley, 1976). Experimental data
on the semi-permeable behaviour of naturally
occurring clayey materials are rare, and limited to
shales and siltstones ( Young and Low, 1965;
Yearsley, 1989). Field evidence on these transport
mechanisms is also limited. Marine and Fritz
(1981) reported anomalously high pore water pressures as a result of different pore water compositions separated by a shale. Only recently Neuzil –
cited in Horseman et al. (1996) – reported an in
situ chemical osmotic experiment on the Pierre
Shale of South Dakota ( USA), showing convincing evidence of chemical osmosis on a field scale.
Loch and Keijzer (1996) concluded that in coastal
regions, where large differences in salt concentrations are often separated by clay layers or lenses,
water and solute transport should be governed not
only by hydraulic but also by osmotic pressure
gradients. In the western parts of the Netherlands
different aquifers, containing either fresh or saline
water, are separated by clay or clayey layers. The
hydrology of this region may well be influenced
by chemical osmosis (De Haven, 1982). In the
predictions of water and contaminant fluxes from
waste storage sites, e.g. harbour sludge depots or
landfill sites, the ability of a clay layer to act as a
semi-permeable membrane has never been taken
into account, although models are developed that
recognise the semi-permeability capabilities of clay
barriers ( Yong and Samani, 1987; Yeung and
Mitchell, 1993).
To determine whether clay layers can act as
semi-permeable membranes under near surface
conditions and to assess their ideality as semipermeable membranes, a laboratory experiment
was designed to subject a commercially available
Wyoming Na-bentonite to
gradients.
osmotic pressure
2. Materials and methods
2.1. Bentonite
A
commercially
available
bentonite,
Ankerpoort Colclay A90 batch 61203, was used in
these experiments. This bentonite is used in the
Netherlands in both sand–bentonite and cement–
bentonite liners as a low permeable barrier to
prevent the transport of water and contaminants
from landfill sites. Some relevant physical and
chemical properties of the bentonite are listed
in Table 1.
2.2. Experimental design
The experimental design consists of a clay permeameter connected by two solution reservoirs. In
the flexible-wall permeameter a sample of 50 mm
diameter and 100 mm maximum thickness can be
fitted. The permeameter is located in a cylindrical
cell, with a polyoxymethylene (POM ) wall,
115 mm i.d. and 141 mm o.d., fitted between two
stainless-steel plates. The top and bottom plates
can be bolted together with three stainless-steel,
all-thread rods, see Fig. 1. A sample between two
porous stones and 15 mm nylon filters was placed
on a pedestal on the bottom plate. On top of the
sample a cap was fitted. A latex membrane was
then placed over the sample, the porous stones,
Table 1
Physical and chemical properties of Ankerpoort bentonite
Particle density (g cm−3)
Primary specific surface area (m2 g−1)
Organic matter (%)
Carbonate content (%)
Total fraction <2 mm (%)
Mineralogy <2 mm (%)
CEC (meq/100 g)
Cations (meq/100 g)
Na+
K+
Ca2+
Mg2+
2.65
120
0.2
0.3
98
>99 smectite
traces quartz
64
40
0
26
3
Th.J.S. Keijzer et al. / Engineering Geology 53 (1999) 151–159
153
mined prior to the experiments and calibrated at
regular intervals. The conductivity of the solution
in the reservoirs was recorded using a Tacussel
CD180 conductivity meter. All signals were logged
on a computer and the experimental set-up placed
in a temperature controlled bench-top chamber.
2.3. Experimental procedures
Fig. 1. Schematic representation of the experimental design.
Only the fresh water reservoir is shown. Drawing not to scale.
the top cap and the pedestal. Two o-rings were
placed over the latex membrane on the cap and
pedestal to provide a seal. The top cap is connected
to the bottom plate by two stainless-steel capillary
tubes. The cell can be filled, normally with de-aired
tap water, or drained through a separate influent
port in the bottom plate. A maximum pressure of
700 kPa can be applied to the cell liquid. For a
detailed description of the design and usage of
flexible wall permeameters the reader is referred
to Daniel et al. (1984).
The bottom end of the sample is connected to
a reservoir containing a low concentration solution, this reservoir will be referred to as the fresh
water reservoir. The top end of the sample is
connected to a reservoir containing a saline solution, or salt water reservoir. All connections to
and from the sample exit the cell through the
bottom plate and are fitted with stainless-steel
Serto regulating valves. Both reservoirs are made
of stainless-steel tubing (6 mm o.d., 4 mm i.d.),
each containing: an Ismatech Reglo-ZS magnetically coupled gear pump, a manifold connecting
each reservoir to calibrated standpipes (6 mm o.d.,
4 mm i.d.) and a pressuring and de-airing system,
pressure transducers and an electrical conductivity
cell. Additionally, a differential pressure transducer
is fitted between the reservoirs. The conductivity
cells were fitted into the stainless-steel tubing;
therefore, the cell constant for each cell was deter-
A sample with thickness of approximately 2 mm
was prepared by weighing 5 g of air-dried bentonite
into a stainless steel mold (50 mm i.d.). The sample
was then subjected to a compacting pressure gradually increasing to 20.3 MPa, in order to reduce
the porosity and to obtain an easy to handle clay
sample. After compaction the sample was mounted
in the cell between the porous stones and filters,
and subsequently the latex membrane was placed
over the sample. It was then saturated with
de-aired tap water of atmospheric pressure, with a
100 kPa cell pressure applied. The stones and filters
were soaked in the same saturation solution prior
to assembly to remove air. All drainage lines to
and from the sample were also flushed to remove
trapped air.
The hydraulic conductivity of the sample was
measured prior to the chemical osmotic experiment. The hydraulic conductivity was determined
using a falling head permeability test. The sample
was subjected to a hydraulic gradient, typically
between 100 and 125 cm cm−1, and the water flow
was measured gravimetrically at regular time
intervals. The large gradient was needed to obtain
a measurable water flow. Detailed descriptions of
different types of test are beyond the scope of this
paper; the reader is referred to Benson and
Daniel (1990).
Under identical conditions the sample was subjected to a chemical gradient. The fresh water
reservoir was filled with a 0.1 M NaCl solution
and the salt water reservoir with a 0.6 M NaCl
solution. According to Eq. (1) this corresponds to
an osmotic pressure difference of 2.3 MPa or
235 m H O. The solutions were de-aired prior to
2
the filling of the reservoirs. In an open reservoir
experiment the flow of water was measured using
the calibrated standpipes. In a closed reservoir
experiment the pressure was monitored in the
154
Th.J.S. Keijzer et al. / Engineering Geology 53 (1999) 151–159
reservoirs using the pressure transducers. To check
for diffusion or convection of salt the electrical
conductivity was monitored in the fresh water
reservoir. All experiments were conducted under
isothermal conditions at 298.15±0.2 K.
3. Results and discussion
3.1. Sample preparation and hydraulic conductivity
The samples obtained after compaction had a
porosity of approximately 0.35, which increased
during the saturation period, resulting in a porosity
between 0.6 and 0.7. In Table 2, where the samples
are coded A and B, measured sample properties
are listed. This porosity is somewhat higher than
that for naturally occurring clays under near surface conditions, which lies typically around 0.5.
Man-made clay barriers are often better compacted
and therefore show a lower porosity. The hydraulic conductivity of the samples ( Table 2) is significantly lower than required in European,
10−7 m s−1 (Stief, 1995) and American,
10−9 m s−1 ( Teplitzky, 1995) environmental
legislation for liners in municipal landfill sites.
brane. In Fig. 2 the reflection coefficient as a
function of the porosity for montmorillonite, illite
and Ankerpoort bentonite, at a mean concentration of 0.35 mol l−1 is given based on the Fritz–
Marine model (Marine and Fritz, 1981; Fritz and
Marine, 1983; Fritz, 1986). From this figure theoretical reflection coefficients can be derived. Using
the values for the bentonite in Table 1 and properties of the samples listed in Table 2, reflection
coefficients can also be derived from Bolt (1982)
using fig. 11.1, p. 403, in which the reflection
coefficient is plotted as a function of the reduced
thickness of the mobile liquid layer, assuming the
lowest value for the mobile countercharge, and
from Bresler’s fig. 1 (Bresler, 1973), in which s is
plotted as a function of the half distance between
the particles times the square root of the solution
concentration. The obtained theoretical values for
s using these three methods are listed in Table 3.
In Fig. 3 the hydraulic pressure development
across sample A, and the change in electrical
conductivity of the fresh water reservoir, are presented. The development of the hydraulic pressure
difference between the two reservoirs, which is
presented as DP=P −P . The first 46 h,
fresh
salt
3.2. Chemical osmosis
Several authors have published models to
describe chemical osmosis and to predict the
behaviour of the clay as a semi-permeable membrane, concentrating on the reflection coefficient
s, because its value determines the maximum
expected osmotic water flux induced by a chemical
gradient. The reflection coefficient depends on the
type of clay, i.e. its surface charge and exchangable
cations, the porosity of the membrane and the
mean pore water concentration across the memTable 2
Measured properties for the two compacted Ankerpoort
Na-bentonite samples
Sample thickness, x (10−3 m)
Porosity
Hydraulic conductivity, K (10−12 m s−1)
d
A
B
2.3
0.638
7.6±0.9
3.4
0.671
2.9±0.9
Fig. 2. The reflection coefficient for a typical montmorillonite
(CEC=100 meq 100 g−1), Ankerpoort Wyoming bentonite
(CEC=64 meq 100 g−1) and illite (CEC=20 meq 100 g−1) as a
function of the porosity according to the Fritz–Marine model.
A mean salt concentration of 0.35 mol l−1 was used.
Th.J.S. Keijzer et al. / Engineering Geology 53 (1999) 151–159
Table 3
Theoretical and measured reflection coefficients, s, for the two
compacted Ankerpoort Na-bentonite samples
A
Theoretical reflection coefficient,
according to
Measured reflection coefficients
155
Describing the water flux by an analogue of
Darcy’s law modified by the reflection coefficient
s:
B
Fritza
0.270
0.212
Boltb
Breslerc
0.02
0.003
0.003
>0.02
0.002
0.001
a Fritz–Marine model, using eq. (10) of Fritz (1986).
b Derived from fig. 11.1, p. 403 of Bolt (1982).
c Derived from fig. 1 of Bresler (1973).
during which the experiment was run with open
reservoirs, is shown in the enlarged part of Fig. 3,
here a water flux of 0.019 cm3 h−1 from the fresh
to the salt water was measured. We assume this
flux to be the result of chemical osmosis.
Dp
J =sK A
w
d Dx
(2)
in which J
is the measured water flux
w
(m3 s−1), K the hydraulic conductivity (m s−1), A
d
the area of the sample (m2) and Dp the pressure
difference across the sample with thickness x, a
value for the reflection coefficient s of 0.003 is
obtained for the period with open reservoirs. After
this period the reservoirs were closed and the
hydraulic pressure difference was monitored using
the pressure transducers in the reservoirs. Note the
sudden decrease in DP after closing the reservoirs
( Fig. 3). A sudden reversal of the water flow took
Fig. 3. Development of the hydraulic pressure difference, DP=P −P , between the reservoirs (n) and the electrical conductivity
fresh
salt
($) in the fresh water reservoir for sample A. The first 50 h of the experiment is shown in the enlargement, where the measured
water flux (J ) is also given.
w
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Th.J.S. Keijzer et al. / Engineering Geology 53 (1999) 151–159
place after 67 h when a DP of −47 cm H O
2
was reached.
Sample B was subjected to an identical chemical
gradient and the hydraulic pressure difference and
electrical conductivity are shown in Fig. 4. The
complete experiment was run with open reservoir.
The first 50 h is also shown on a larger scale. From
the measured water flux, 0.016 cm3 h−1, in this
period a s of 0.001 is derived for this sample.
From Table 3 it is clear that the measured values
are in good agreement with those obtained from
the work by Bresler (1973), they are, however,
significantly lower than those predicted by the
Fritz–Marine model. After 50 h the DP decrease
dropped until the differential pressure reached a
value of −8 cm H O after 310 h. Both the rate of
2
change of hydraulic pressure and the absolute
value of the differential pressure are lower than
for sample A. This can be attributed solely to the
method used.
The overall trend for both experiments is the
same: initially the differential hydraulic pressure
increases linearly due to steady-state, water is
transported from the fresh to the salt water reservoir as a result of the applied chemical gradient.
After approximately 50 h the rate of change of
differential hydraulic pressure drops and DP
reaches a minimum value. Subsequently, DP
increases due to the reversal of the water flow and
the hydraulic pressure in the fresh water reservoir
increases. Because of the non-ideality of the clay
membrane it is not capable of completely
Fig. 4. Development of the hydraulic pressure difference, DP=P −P , between the reservoirs (n) and the electrical conductivity
fresh
salt
($) in the fresh water reservoir for sample B. The first 50 h of the experiment is shown in the enlargement, where the measured
water flux (J ) is also given. This experiment was run entirely with open reservoirs.
w
Th.J.S. Keijzer et al. / Engineering Geology 53 (1999) 151–159
restricting the diffusion of ions and the chemical
gradient across the membrane will slowly decrease.
Consequently, the hydraulic pressure difference
will slowly diminish, however, the rate at which
the pressure difference decreases in these experiments is faster than that observed by other authors
( Kemper, 1961; Elrick et al., 1976).
The sudden drop in differential hydraulic pressure in both experiments is not a result of collapse
of the double layers due to the increased equilibrium concentrations within the clay membrane, or
of failure of the samples as a semi-permeable
membrane. If this were so DP would change until
it equalled 0. In that situation the salt gradient
would diminish rapidly, resulting in a rapid
increase in electrical conductivity (Ec) in the fresh
water reservoir. In both experiments the
differential pressure changed into high positive DP
values, and the observed electrical conductivity did
not increase rapidly. In both experiments the rate
of increase in Ec slowed down in the fresh water
reservoir during the last stage of the experiment.
In Fig. 5 the schematic representation of the
observed changes in DP and the Ec for the experiments and their interpretation is given, and the
expected signal if the sample would fail as a semipermeable membrane.
The changes in DP and Ec in the fresh water
reservoir are consistent with what is described by
Olsen et al. (1989, 1990) as diffusion osmosis.
Diffusion osmosis describes the transport of water
in response to the diffusion of dissolved solutes.
Olsen et al. (1990) described it as follows: ‘‘…
solute diffusion in response to a concentration
gradient imposes drag on, or momentum transfer
to, the pore fluid and thus tends to move the pore
fluid in the direction of decreasing solute concentration.’’ This phenomenon was, according to
Olsen et al. (1990), also observed but not recognised by other authors, e.g. Kemper and Quirk
(1972), Elrick et al. (1976), Veder (1979) and
Yearsley (1989). Diffusion osmosis should be more
dominant in loosely compacted clayey material of
low cation exchange capacity (Olsen et al., 1990).
Our samples exhibited a high porosity, which
according to Olsen et al. (1990) is one of the major
properties for a clay to show diffusion osmosis,
however, other observed factors like a low
157
Fig. 5. Schematic representation of the observed (——) changes
in DP and the electrical conductivity in the fresh water reservoir
for the experiments and their interpretation, and the expected
signal (- - - -) if the sample would fail as a semi-permeable
membrane.
exchange capacity and high pore water concentrations were not met in these samples.
The consequences of chemical osmosis for water
transport through clay liners, clayey waste or in
the shallow subsurface where large aquifers of
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Th.J.S. Keijzer et al. / Engineering Geology 53 (1999) 151–159
different chemical composition are separated by
clay layers, can be very large even if these layers
have a low value of s. When not taken into
account, the water flux from a site might be
underestimated, resulting in inadequate drainage.
Convective transport of contaminants may be
enhanced by osmotically induced water transport.
In present models for engineering applications
water flow as a result of a chemical gradient is not
taken into account. The influence of diffusion
osmosis in a field situation is harder to predict.
Man-made clay barriers are probably compact
enough to suppress this process and make it insignificant in the overall water and solute transport.
4. Conclusions
In laboratory samples of compacted Wyoming
Na-bentonite water transport was induced by a
chemical gradient and reflection coefficients (s)
were found to be 0.003 and 0.001, respectively.
Both values are significantly lower than predicted
by the Fritz–Marine model and the values calculated according to Bolt (1982), however, they are
in good agreement with values obtained from the
work by Bresler (1973). During the course of the
experiment the samples lost their semi-permeability, however, they did not fail as a low permeable
barrier. Another transport mechanism, presumably
diffusion osmosis, became more dominant.
Although the samples had relatively high porosities, 65% by volume, diffusion osmosis was not
expected to occur in these high surface charge
clays.
Acknowledgements
The authors wish to thank the Directorate
General for Public Works and Water Management
(Rijkswaterstaat) for partly supporting this work.
This is contribution 980804 of the Netherlands
Research School of Sedimentary Geology.
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