Oerter-2014-Oxygen isotope fract
Oxygen isotope fractionation effects in soil water via interaction with cations (Mg,Ca,K,Na)adsorbed to phyllosilicate clay
minerals
Erik Oerter a ,?,Kari Finstad a ,Justin Schaefer b ,Gregory R.Goldsmith c ,Todd Dawson d ,Ronald Amundson a
a
Department of Environmental Science,Policy and Management,University of California,Berkeley,145Mulford Hall,CA 94720,United States b
Department of Earth and Planetary Science,University of California,Berkeley,United States c
Environmental Change Institute,Oxford University,Oxford,United Kingdom d
Department of Integrative Biology,University of California,Berkeley,United States
a r t i c l e i n f o Article history:
Received 5February 2014
Received in revised form 5April 2014Accepted 11April 2014
Available online 26April 2014
This manuscript was handled by Corrado Corradini,Editor-in-Chief,with the
assistance of Juan V.Giraldez,Associate Editor
Keywords:Smectite
Isotope hydrology Isotopic partitioning
Soil-atmosphere interaction Pedogenic carbonate Vertisols
s u m m a r y
In isotope-enabled hydrology,soil and vadose zone sediments have been generally considered to be iso-topically inert with respect to the water they host.This is inconsistent with knowledge that clay particles possessing an electronegative surface charge and resulting cation exchange capacity (CEC)interact with a wide range of solutes which,in the absence of clays,have been shown to exhibit d 18O isotope effects that vary in relation to the ionic strength of the solutions.To investigate the isotope effects caused by high CEC clays in mineral–water systems,we created a series of monominerallic-water mixtures at gravimetric water contents ranging from 5%to 32%,consisting of pure deionized water of known isotopic composi-tion with homoionic (Mg,Ca,Na,K)montmorillonite.Similar mixtures were also created with quartz to determine the isotope effect of non-,or very minimally-,charged mineral surfaces.The d 18O value of the water in these monominerallic soil analogs was then measured by isotope ratio mass spectrometry (IRMS)after direct headspace CO 2equilibration.Mg-and Ca-exchanged homoionic montmorillonite depleted measured d 18O values up to 1.55‰relative to pure water at 5%water content,declining to 0.49‰depletion at 30%water content.K-montmorillonite enriched measured d 18O values up to 0.86‰at 5%water content,declining to 0.11‰enrichment at 30%water.Na-montmorillonite produces no mea-sureable isotope effect.The isotope effects observed in these experiments may be present in natural,high-clay soils and sediments.These ?ndings have relevance to the interpretation of results of direct CO 2-water equilibration approaches to the measurement of the d 18O value of soil water.The adsorbed cation isotope effect may bear consideration in studies of pedogenic carbonate,plant-soil water use and soil-atmosphere interaction.Finally,the observed isotope effects may prove useful as molecular scale probes of the nature of mineral–water interactions.
ó2014Elsevier B.V.All rights reserved.
1.Introduction
In isotope-based approaches to hydrology,soil and sediment have been implicitly considered to be an inert matrix through which water passes.Yet,this assumption is inconsistent with the fact that soils contain a wide range of solutes and highly variable concentrations of chemically reactive clay particles (Sposito,2008),all of which may react with bulk water to create regions of water molecules with different coordination environments and varying isotope compositions.
Previous researchers have postulated the existence of various ‘‘pools’’of water with differing oxygen isotope compositions in soils (Ingraham and Shadel,1992;Araguas-Araguas et al.,1995;Hsieh et al.,1998),but it has only been recently that a growing body of evidence has emerged supporting this hypothesis (c.f.Brooks et al.,2010;Soderberg et al.,2012).The mechanisms for the formation and retention,as well as the locations,of these pools remains largely unexplored,particularly in light of both macro-and molecular-scale considerations of soil mineralogy and physical chemistry.
Soil solutions commonly have high solid to ?uid ratios in a matrix of minerals with diverse structures and electrical charges.Based on molecular scale modeling of the organization of water molecules at and near electronegatively charged particle surfaces (Bourg and Sposito,2011),we anticipate that isotope effects should be produced and be experimentally observable,especially at lower water contents where the ratio between the isotopically
https://www.wendangku.net/doc/644504904.html,/10.1016/j.jhydrol.2014.04.029
0022-1694/ó2014Elsevier B.V.All rights reserved.
?Corresponding author.Tel.:+13039901499.
E-mail address:erik.oerter@https://www.wendangku.net/doc/644504904.html, (E.Oerter).
‘‘perturbed’’water is high relative to free water.The magnitude of these effects should be most evident in soils with high cation exchange capacity(CEC)combined with speci?c cations adsorbed to the clay particles.Here we conduct laboratory experiments to explore the potential existence of such isotope effects as a?rst step in a more detailed consideration of the role of stable isotopes in mineral–water interactions.
2.Background
2.1.Ionic solutions
Water that passes through soils physically interacts with an array of solutes,as well as with solids(here,we consider colloidal suspensions as part of the water–solid interaction).The effects of solutes on oxygen isotopes have a long history of study in stable isotope geochemistry(Feder and Taube,1952,Taube,1954;Sofer and Gat,1972,1975).In general,oxygen isotope fractionation in aqueous saline solutions varies with dissolved cation concentra-tion,and the sign and magnitude of the effect is cation-speci?c,lar-gely correlating with the‘‘ionic potential’’(charge/ionic radius)of the ion in solution and its corresponding polar coordination with water molecules resulting in the formation of hydration spheres around each cation(O’Neil and Truesdell,1991).The isotopes of oxygen in the water molecules associated with hydration spheres around dissolved ions can be either enriched or depleted in the heavy oxygen isotopes of water,with a corresponding effect on the bulk water.This is the so-called salinity effect.
The most favorable theory explaining the fractionation of water oxygen isotopes by cations in solution is based on the notion of cations with a high ionic potential(e.g.Mg2+,Ca2+herein)creating a more structured organization of water molecules around each cation,as compared to the bulk water.Conversely,cations with lower ionic potential(e.g.K+)disrupt the water structure around them.Frank and Wen(1957)?rst observed these macroscopic effects and discussed the structure-making and structure-breaking effects of electrolytes in aqueous solution.More recent studies have both con?rmed the effects of cations on the structure of water (Marcus,2010)and cast doubt on its molecular-scale underpin-nings(Turton et al.,2011).This uncertainty does not preclude the use of the structure making and breaking concept to describe the macroscale behavior of water oxygen isotopes in saline solu-tions.We adopt the operational de?nition of O’Neil and Truesdell (1991),who articulated that from an oxygen isotope perspective, structure makers have the effect of depleting the bulk water in 18O,while structure breakers yield bulk water with an enriched 18O composition.They justify this de?nition with the observation of positive isotopic fractionation between ice(highly structured) and liquid water(O’Neil,1968),and negative isotopic fractionation between vapor(no structure)and liquid water(Majoube,1971). These simple cases are analogous to water in hydration spheres around cations because these water isotopologues are sequestered away from the bulk water in a similar sense to how they would be in ice crystals.
While the speci?c molecular-scale mechanism(s)for the reorga-nization of oxygen isotopes around cations in solution remain uncertain,the prevailing theory(i.e.O’Neil and Truesdell,1991) is that structure making cations(Mg2+,Ca2+)create organized water by forming two hydration regions around the cation.In the inner sphere,water molecules are strongly bound by the cation’s ionic charge and are structured with their dipoles oriented radially to the cation by an intense electrostatic attraction.The more dif-fuse outer region is where water molecules are hydrogen bonded to each other and more weakly bound to the cation.The inner, strongly bonded region preferentially incorporates H218O with its shorter intramolecular bonds,while the outer,hydrogen bonded region concentrates H216O.It is the net balance of the effects of these hydration regions that determines the characteristic cation-speci?c fractionation effect,with the inner sphere dominating the net effect of the divalent Mg2+and Ca2+ions.In contrast,water structure breaking ions with lower ionic potential(K+)create a sin-gle hydration region of water molecules that are organized by hydrogen bonding.This region of hydrogen-bonded water appears to preferentially incorporate H216O molecules,thereby enriching the bulk water outside of the hydration region in H218O.
In the special case of Na+,which does not exhibit net oxygen-fractionating effects,there seem to be two potential mechanisms for the lack of fractionation:if Na+creates two hydration spheres around itself,they appear to have largely mutually cancelling effects(Phillips and Bentley,1987),or that Na+creates a region of very mild hydrogen bonded water that does not exert a strong enough in?uence on the bulk water surrounding it to be observable (Nag et al.,2008).In either mechanistic case,a macroscopic oxygen isotope effect is not seen for Na+in aqueous solution.
2.2.Clay mineral interactions with water
Soil is largely a heterogeneous mixture of silicate minerals.The sand(0.05–2mm)and silt(0.002–0.05mm)fractions tend to be dominated by primary silicates of low surface area to mass ratios and have a negligible electrostatic charge.In contrast,the clay frac-tion(<0.002mm)is commonly dominated by secondary phyllosil-icates with high surface area to mass ratios and a considerable permanent negative charge.This negative charge gives rise to a mineral’s cation exchange capacity(CEC),a measure of how many cations(typically Ca2+,Mg2+,K+,Na+)from the surrounding soil solution adsorb to the mineral surface to balance its negative struc-tural charge(Schoonheydt and Johnston,2013).The most common reactive mineral group within the phyllosilicates(clays)is the smectites,and montmorillonite is a common smectite.These cat-ions adsorbed to clay particles attract water in the form of hydra-tions spheres,and infrared spectroscopic studies of Mg-and Ca-montmorillonite reveal that the properties of the surrounding water are distinct from the bulk water and strongly in?uenced by the identify of the adsorbed cation(Burgess,1978;Johnston et al.,1992;Xu et al.,2000).
A comprehensive body of research exists on the effects of clays on water isotope fractionation at higher temperatures and pore water pressures than are found in soils.This work is intended to emulate underground settings such as con?ned aquifers and geo-thermal reservoirs and has concluded that diffusive gradients con-centrate the heavy isotopic species behind clay membranes, irrespective of cation identity and ionic potential(e.g.Coplen and Hanshaw,1973;Phillips and Bentley,1987;Hu and Clayton, 2003).These results differ markedly from observations of oxygen isotope fraction at ambient pressures which are the conditions found in soils and we therefore interpret these high temperature and pressure oxygen isotope effects to be due to mechanisms and processes not present in soils.
Given the large and predictable ionic-water isotope interactions described above,it follows that a more in depth investigation of oxygen isotope fractionation from water–solid interactions is war-ranted.These water–solid interaction processes likely manifest themselves in nature,particularly in soils with high salt contents, or more generally in any soil with a signi?cant proportion of charged clay minerals(e.g.Vertisols,soils with at least40% expandable clay and associated high CEC).
In this paper,we examine the effect of monominerallic homo-ionic smectite-water mixtures on the oxygen isotope composition of soil water measureable by CO2exchange.This research tests the assumption that soil is essentially an inert matrix for in?ltrating
2 E.Oerter et al./Journal of Hydrology515(2014)1–9
water.Our hypothesis is that cations adsorbed to the clay surface form isotopically organized hydration spheres of water around them and thereby sequester these water molecules away from the bulk water.In turn,this has possible implications for interpre-tations of the O isotope composition of plant waters and pedogenic carbonates,and their relationship to meteoric water.
3.Methods
To study the effects of smectite particles in soil on the oxygen isotope fractionation of soil water,we used Arizona montmorillon-ite sourced from near Cameron,Arizona(API clay#49-5107)with a CEC of60.75cmol c/kg and Texas montmorillonite sourced from near Gonzalez,Texas(Source Clay STX1-b)with a CEC of 69.5cmol c/kg(both measured by Ba-replacement at the University of California,Davis Analytical Laboratory).
Two stocks of Arizona montmorillonite mineral powder were combined with monoionic cation-chloride aqueous solutions of MgCl2,CaCl2,and KCl at either0.05M or1M concentrations, mixed by hand and allowed to equilibrate and gravitationally settle overnight,after which the supernatant saline solutions were man-ually decanted.Three saline washes at either0.05M or1M were followed in a similar manner with three deionized water rinses at circumneutral pH.The mineral–water slurries were dried at 70°C,yielding dry homoionic montmorillonite powders with all available surface adsorption sites occupied by Mg2+,Ca2+or K+cat-ions.Iotaòquartz sand was similarly prepared as a control.While quartz does have a very small(but measurable)CEC due to unsat-is?ed hydroxyl bonds at mineral cleavage surfaces,it should have almost no ability to adsorb cations and therefore any observable isotope effect should be minimal,if present at all.These samples are referred to herein as‘‘stir and settle’’homoionic and were designed to determine if an adsorbed cation isotope effect was present and detectable.Additionally,these experiments test if the effect is still present when smectite clays are exposed to even weakly saline solutions,and therefore simulate saltwater inun-dated soils in coastal and estuarine settings.
To more precisely constrain and quantify the adsorbed cation isotope effect,a third stock of clay was prepared using the‘‘acidic washing’’method.Texas montmorillonite(Source Clay STx1-b) was employed when our supply of Arizona montmorillonite was exhausted and was prepared by washing the mineral powder with acidic(adjusted to pH3with HCl)monoionic cation-chloride aque-ous solutions(MgCl2,CaCl2,NaCl,KCl)at1M concentration multi-ple times(each separated by centrifugation to allow the supernatant to be completely decanted),followed by multiple rinses with acidi?ed0.01M saline solution,ethanol rinses to remove chloride anions and residual H+cations,and oven drying at70°C.By preparing this smectite material with acidic electrolyte washes in this way,any weathering products that may have been present on the clay surfaces composed mainly of amorphous Al oxyhydroxide materials(proto-boehmite and gibbsite)that are found ubiquitously in soils are dissolved away from the mineral surfaces and removed(Sposito,2008).
From each of these homoionic smectite materials,sets of min-eral–water mixtures were created in autosampler-speci?c sample vials with butyl rubber septum caps from homoionic dry mineral powder and deionized water of known isotopic composition at 5–32%gravimetric water content(Tables1and2),the same range of soil water proportions that are commonly observed in nature (n=1replicate per material for‘‘stir and settle’’preparation,3–4 replicates per material for‘‘acidic washing’’).Each mineral–water mixture was prepared with the mineral mass scaled to0.05g (50l L)water,thereby reducing size effects in IRMS analyses.The choice of this small amount of water was necessitated by the relatively large volume of clay powder required to make gravimet-ric mixtures at5%water–95%clay:0.95g clay with$2mL clay volume in a12mL vial is the largest amount of clay that allows for the autosampler needle to be reliably separated from contact with the clay-water mixture.Water volumes of50l L still have three orders of magnitude more moles of H2O for oxygen isotopic equilibration with CO2headspace(2.78?10à3mol H2O: 7.98?10à6mol CO2for driest samples),though do require longer equilibration times.Measured d18O values on50l L pure water samples maintain accuracy and precision similar to samples with larger water volume regularly analyzed on this instrument. Quartz-water mixtures also yield reliable d18O values with50l L water volumes(discussed below).
Sample vial headspace was purged with0.2%CO2in He and allowed to equilibrate in the vial at room temperature for ca. 5days(Epstein and Mayeda,1953)and analysis was subsequently conducted for d18O values on a Thermo Delta V mass spectrometer with Thermo Gas Bench II interface(Center for Stable Isotope Bio-geochemistry,UC Berkeley).Analyses were started with three injections of tank reference CO2to evaluate instrument stability and calculate d18O values.Sample gas was injected?ve times and d18O values were calculated on the average of the last four peak heights.Analytical runs included three lab standard reference waters with well constrained d18O(VSMOW)values every6th sample allowing calibration to the VSMOW scale.Long-term exter-nal precision on this instrument is±0.12‰for d18O of water.
Oxygen isotope d18O values are de?ned and presented in stan-dard notation:
d?
R sample
R standard
à1
?1000e1T
where R sample and R standard are the18O/16O ratios for the sample and standard,respectively.All d18O data are reported in per mil(‰)rel-ative to VSMOW.The difference between the measured d18O values of the homoionic mineral–water mixtures(d m)and that of the deionized water used to make the mixtures(d0)is calculated as in Sofer and Gat(1972):
D d18O?
ed0àd mT
e1000td mT
?1000e2T
D d18O values are somewhat counter intuitive because they are opposite the sign of the observed isotope effect.D d18O values are correction factors and are added to the measured d18O values to cor-rect for the isotope effect.Positive D d18O values mean that the mea-sured d18O values are depleted in18O relative to the input water, and negative D d18O values mean that the measured d18O values are enriched in18O relative to the input water.
4.Experimental results
With respect to the Arizona montmorillonite prepared with the ‘‘stir and settle’’homionic washes(n=1),Mg-montmorillonite shows the largest isotope effect with D d18O values of3.36‰at 8%water to0.90‰at32%water when prepared with1.0M MgCl2 washing solution,and3.14‰at8%water to0.54‰at32%water when prepared with0.05M MgCl2washing solution(Table1and Fig.1A).Ca-montmorillonite shows D d18O values of2.40‰at8% water to0.84‰at32%water when prepared with1.0M CaCl2 washing solution,and3.37‰at8%water to0.81‰at32%water when prepared with0.05M CaCl2washing solution.
Quartz shows a negligible isotope effect with very small D d18O values that are likely due to limits in analytical precision for the continuous?ow IRMS instrument,as well as subtle evaporative effects due to the high surface area to volume ratio of the water in the thin capillary?lm of water surrounding each quartz grain.
E.Oerter et al./Journal of Hydrology515(2014)1–93
This is as expected,given the near complete absence of a structural charge on which to adsorb cations.
Experiments with the‘‘stir and settle’’homoionic preparation method show a larger isotope effect than the homoionic samples prepared with acidic washing and centrifugation,and therefore may be complicated by the incomplete removal of residual salts from the electrolyte cation-loading solutions.Even with this addi-tional complication,these samples demonstrate the ef?cacy of cat-ion loading solution concentrations equivalent to those found in nature(seawater[Mg]=0.05M)and these soil analogs are similar to soil or sediment samples found in the?eld,i.e.clay-rich soils in semiarid climates or within coastline or estuarine settings that have been inundated by saline or brackish waters.
D d18O values for mineral–water mixtures composed of Texas montmorillonite prepared with the acidic homoionic washing are shown in Table2and Fig.1B.Mg-montmorillonite shows the larg-est isotope effect with D d18O values from0.96‰to 1.55‰(Ave=1.20‰,S.D.=0.31,n=3)at5%water content to D d18O val-ues of0.49‰to0.70‰(Ave=0.60‰,S.D.=0.11,n=3)at30% water.Ca-montmorillonite had an isotope effect nearly as large as Mg,with D d18O values from0.74‰to1.01‰(Ave=0.92‰, S.D.=0.15,n=3)at5%water content to D d18O values from 0.33‰to0.39‰(Ave=0.35‰,S.D.=0.03,n=3)at30%water. K-montmorillonite shows D d18O values fromà0.51‰toà0.86‰(Ave=à0.69‰,S.D.=0.17,n=3)at5%water content to D d18O val-ues ofà0.11‰toà0.27‰(Ave=à0.21‰,S.D.=0.08,n=3)at30% water.Na-montmorillonite shows D d18O values from0.01‰to à0.24‰(Ave=à0.09‰,S.D.=0.12,n=4)at5%water content to à0.03‰toà0.16‰(Ave=à0.10‰,S.D.=0.06,n=3)at30%water.
5.Discussion
These experiments af?rm the fact that smectite minerals are not an isotopically inert medium for water.Pure water that is exposed to homoionic smectite will undergo reactions with the solid phase that ultimately changes the stable isotope composition of the more energetically free water that remains.The reactions are clearly impacted by the concentration and valence of the adsorbed cation phases,and also likely re?ect additional,but poorly under-stood,effects of water–mineral interactions.
When pure water is added to the homoionic smectite,two processes occur simultaneously.First,the exchangeable cations on the clay surfaces hydrate and recon?gure into the inner-and outer-sphere complexes of the electrical double layer,while some portion of the cations become fully solvated and go into solution where they exert the isotope effects previously discussed as the ‘salinity effect’.Second,those cations that remain in the inner-and outer-sphere complexes that are adsorbed to the clay sur-faces and are not fully solvated,form hydration spheres around them and exert a different fractionation effect on the bulk water, see Fig.2for a conceptual depiction of this process.The D d18O values measured on mineral–water mixtures are a combination of these two effects and their respective contributions to the net observed effect.Separating out the magnitude of each compo-nent’s contribution to the net isotope effect can yield information on the nature of the adsorbed cation effect.
Following equilibration with pure water,the concentration of cations desorbed from the clay and now are solvated in solution is calculated by multiplying the initial quantity of cations adsorbed on the clay by the partition coef?cient(K D):K D=A/C,where A is the amount of cations adsorbed on the clay and C is the amount of cations in solution.K D values available in the literature for base cations in clay soils are shown in Table3(Baes and Sharp,1983).K D values for Na are unavailable and so an estimated value of9mL/g is used,based on the value for K and its relative similarity in ionic properties to Na.To express K D values as a dimensionless ratio (K?
D
),K D(mL/g)is multiplied by the density of montmorillonite (2.5g/cm3)(Appelo and Postma,1999).The molarity of the equili-brated water in each mineral–water mixture after cation reparti-tioning is shown in Tables1and2,and calculated as follows.
The concentration of adsorbed cations per unit mass of dry homoionic montmorillonite can be determined by reasoning that
Table1
Data for mineral–water mixtures prepared with neutral pH‘‘stir and settle’’homoionic preparation.The values for the last three columns were determined as discussed in the text.The pure water used to prepare these mixtures had d18O ofà12.33‰(VSMOW).
Cation electrolyte Mineral Mineral
mass(g)%H2O
(g/g)
d18O
(VSMOW)
D d18O Repartitioned cation
molarity in added H2O
D d18O due to repartitioned
cations in added H2O
%D d18O due to repartitioned
cations in added H2O
MgCl2[1.0M]Montmorillonite0.5807.9à15.63 3.360.1040.116 3.5
0.26515.9à14.00 1.700.0480.053 3.1
0.10831.7à13.210.900.0190.022 2.4
MgCl2[0.05M]Montmorillonite0.5758.0à15.42 3.140.1030.115 3.7
0.26515.9à13.71 1.400.0480.053 3.8
0.10632.0à12.860.540.0190.021 3.9
CaCl2[1.0M]Montmorillonite0.5768.0à14.70 2.400.1430.067 2.8
0.26515.9à13.72 1.410.0660.031 2.2
0.10632.2à13.160.840.0260.012 1.5
CaCl2[0.05M]Montmorillonite0.5788.0à15.64 3.370.1430.067 2.0
0.26415.9à14.34 2.040.0650.031 1.5
0.11131.1à13.130.810.0280.013 1.6
MgCl2[1.0M]Quartz0.5778.0à12.380.05ààà
0.26415.9à12.24à0.09ààà
0.10731.8à12.31à0.02ààà
MgCl2[0.05M]Quartz0.5797.9à12.09à0.24ààà
0.26415.9à12.23à0.10ààà
0.10931.5à12.13à0.20ààà
CaCl2[1.0M]Quartz0.5748.0à12.06à0.27ààà
0.26216.0à12.18à0.16ààà
0.10532.2à12.08à0.25ààà
CaCl2[0.05M]Quartz0.5768.0à12.23à0.10ààà
0.26216.0à12.07à0.26ààà
0.10632.1à11.94à0.39ààà4 E.Oerter et al./Journal of Hydrology515(2014)1–9
every available charge site to adsorb a cation is occupied after the homoionic clay preparation.For monovalent cations this amount is simply the measured CEC of the mineral(Arizona Montmorillonite:60.8cmol c/kg;Texas montmorillonite:69.5cmol c/kg),and for divalent cations it is half this number,as each divalent cation bal-ances two structural charge sites.The total amount of cation
Table2
Data for Texas montmorillonite mineral–water mixtures prepared with acidic homoionic preparation.The values for the last three columns were determined as discussed in the text.The pure water used to prepare these mixtures had d18O ofà11.74‰(VSMOW).
Cation Mineral
mass(g)%H2O(g/g)d18O measured
(VSMOW)
D d18O Repartitioned cation
molarity into added H2O
D d18O due to repartitioned
cations into added H2O
%D d18O due to repartitioned
cations into added H2O
Mg2+0.960 5.0à13.27 1.550.1980.21914.1
0.940 5.1à12.690.960.1940.21522.3
0.950 5.0à12.80 1.080.1960.21720.2
0.4709.6à12.90 1.180.0970.10712.5
0.44010.2à12.78 1.060.0910.1019.1
0.4609.8à12.590.860.0950.1059.2
0.28315.0à12.92 1.190.0580.065 4.6
0.28714.8à12.690.970.0590.066 4.6
0.28415.0à12.680.960.0580.065 3.8
0.20419.7à12.680.950.0420.047 3.4
0.20119.9à12.580.850.0410.046 3.0
0.20119.9à12.620.890.0410.046 2.7
0.11330.7à12.220.490.0230.026 2.6
0.11330.7à12.430.700.0230.026 3.8
0.11430.5à12.340.600.0230.026 2.7
Ca2+0.953 5.0à12.470.740.1960.09217.2
0.951 5.0à12.74 1.010.1960.09212.5
0.952 5.0à12.73 1.000.1960.09212.7
0.45210.0à12.670.940.0930.044 6.4
0.4549.9à12.680.950.0930.044 6.4
0.4559.9à12.90 1.170.0940.044 5.2
0.28415.0à12.540.810.0580.027 4.7
0.28315.0à12.630.900.0580.027 4.2
0.28315.0à12.73 1.000.0580.027 3.8
0.20020.0à12.480.750.0410.019 3.5
0.20020.0à12.240.510.0410.019 5.2
0.20020.0à12.440.710.0410.019 3.8
0.12029.4à12.070.330.0250.012 4.8
0.11430.5à12.070.340.0230.011 4.5
0.11729.9à12.120.390.0240.011 4.0
K+0.949 5.0à10.89à0.860.391à0.0637.3
0.947 5.0à11.03à0.710.390à0.0628.8
0.947 5.0à11.23à0.510.390à0.06212.1
0.4539.9à11.17à0.580.187à0.030 5.2
0.44910.0à11.57à0.170.185à0.03017.1
0.44910.0à11.18à0.570.185à0.030 5.2
0.28115.1à11.42à0.320.116à0.019 5.8
0.28215.1à11.65à0.090.116à0.01920.7
0.28215.1à11.52à0.220.116à0.0198.4
0.19920.1à11.51à0.230.082à0.013 5.7
0.20020.0à11.70à0.040.082à0.01331.8
0.20020.0à11.62à0.130.082à0.01310.5
0.11829.8à11.51à0.240.049à0.008 3.3
0.11630.1à11.48à0.270.048à0.008 2.9
0.11530.3à11.63à0.110.047à0.008 6.7
Na+0.951 5.0à11.60à0.150.39200
0.950 5.0à11.50à0.240.39100
0.950 5.0à11.750.010.39100
0.950 5.0à11.740.010.39100
0.45210.0à11.960.230.18600
0.45110.0à11.70à0.040.18600
0.45010.0à11.890.150.18500
0.45110.0à11.800.060.18600
0.28315.0à11.810.070.11700
0.28215.1à11.840.100.11600
0.28415.0à11.830.090.11700
0.28415.0à11.73à0.010.11700
0.19820.2à11.900.160.08200
0.19920.1à11.820.080.08200
0.19920.1à11.68à0.060.08200
0.20020.0à11.70à0.040.08200
0.11829.8à11.67à0.070.04900
0.11829.8à11.58à0.160.04900
0.11729.9à11.71à0.030.04800
0.11630.1à11.63à0.120.04800
E.Oerter et al./Journal of Hydrology515(2014)1–95
adsorbed to clay particles in a gravimetric sample is then easily determined by multiplying the CEC by the mass of mineral.
The portion of the observed D d 18O values that are attributable to the isotope salinity effect,which is due to the fractionation of water molecules into hydration spheres around fully solvated cat-ions compared to the pure water used to make the solutions,can be determined by multiplying the equilibrated solution molarity by the salinity effect per molarity unit as characterized by Sofer and Gat (1972): 1.11‰for Mg 2+,0.47‰for Ca 2+,0‰for Na +,à0.16‰for K +.The component of D d 18O values for each min-eral–water mixture that is directly attributable to the salinity effect of repartitioned cations into the aqueous phase is shown in Tables 1and 2.
The magnitude of the isotope effect due to adsorbed cations can be determined by subtracting the salinity effect component from the observed D d 18O value.In general,the salinity effect exerted by cations in solution are on the order of 7–22%of the total observed for the driest mineral–water mixtures (5%water),and are less than 10%of the observed D d 18O values for wetter samples with 8%or more water.The balance of the isotope effect and observed D d 18O values is therefore attributable either solely to that exerted by adsorbed cations on the bulk water or by a combination of the adsorbed cations and some other factor(s),such as reactions with the mineral surface itself.
Considering the mass balance of water added to the mineral–water mixtures can help elucidate the observed trend of increasing adsorbed cation effect with decreasing water content.The total amount of water molecules coordinated in hydration spheres around cations adsorbed to clay surfaces can be found by multiply-ing the number of water molecules per hydration sphere (‘‘hydra-
6?69.5meq/100g =417mmol/100g.This quantity of coordi-nated water is then multiplied by the molar mass of water (18g/mol)and converted into mol from mmol,yielding 7.506g water per 100g clay,or 7.5%water content.Above this threshold,the increasing proportion of bulk to adsorbed water will diminish the in?uence of the adsorbed cation effect.
Though it is somewhat approximate,the relationship of the adsorbed isotope effect component to the concentration of cations remaining adsorbed (in a net equilibration sense)to the clay parti-cles in the mineral–water mixtures (which is,analogous to adsorbed cation molarity),is 0.96‰per mole of adsorbed cation for Mg 2+,1.19‰for Ca 2+and à0.9‰for K +(Table 4).Not surpris-ingly,and similar to its behavior in solution,Na +does not show an isotopic effect when adsorbed to smectite surfaces (Fig.3A and B).
The effects of Mg,both adsorbed and in solution,are relatively similar on a per mole basis.However,Ca shows a twice-greater effect when adsorbed,as compared to in solution,and is compara-ble to that of Mg.Xu et al.(2000),using infrared spectroscopy,found that as the water content of Mg-and Ca-montmorillonite decreases to <10water molecules per exchangeable cation ($12.5%water for Texas montmorillonite in this study),the water molecules are clustered around the adsorbed cation and are restricted in their ability to form hydrogen bonds with the sur-rounding bulk water,a condition that would seem to favor water structure making and the attendant isotopic effects.Interestingly,both Mg-and Ca-montmorillonite behaved in a similar manner in the study of Xu et al.(2000),suggesting that once the clay sur-face is saturated by adsorbed cations,divalent base cations regard-less of identity exert a similar effect on the structure of water.This relationship of D d 18O to gravimetric water content for (A)Mg-and Ca-montmorillonite-water,and quartz-water,mixtures prepared with the method with varying cation concentration,and (B)Mg-,Ca-,K-,Na-montmorillonite-water mixtures prepared with the acidic washing method.
cation repartitioning when homoionic clay is exposed to liquid water and the corresponding isotope effect (1972),adsorbed data this study.
(Fig.3B:à0.9‰moleà1for adsorbed K,Fig.1B:à0.16‰moleà1for K in solution).This disproportionate effect suggests that there is likely another mechanism present that enhances the inclusion of isotopically light water in the vicinity of K-montmorillonite sur-faces,thereby enriching the bulk water in H218O.That is,the net effect of the structure-breaking cation seems to be enhanced.An aspect of the K+ion that may be relevant is that its ionic size allows it to?t into the open crystal structure of phyllosilicate mineral sheets(Sposito,2008).This ability to create a relatively‘‘satis?ed’’surface both electrically and texturally may enhance the ability of the surface to create regions of hydrogen bonded water near to it, as discussed below.
The isotopic fractionation factor per unit mole between water surrounding desorbed cations in solution and water associated with adsorbed cations on clay surfaces can be calculated from the per mole d values by the relationship:
a aq-ad?e1000td aqT=e1000td adTe3Twhere d aq and d ad are the molar components of the measured d val-ues as described above.Fractionation factors for the cations in this study are listed in Table4,and expressed as1000ln/for utility.
There is some evidence that smectite surfaces themselves may exert an isotope effect on water by creating a very thin?lm of structurally broken water at the mineral surface.Bourg and Sposito(2011)conducted molecular dynamics simulations of the organization of water near smectite surfaces and found several regions with a higher density of water molecules in the electrical double layer at the mineral–water interface.Immediately adjacent to the clay(2.7?from surface),the density of water molecules in the?rst statistical water monolayer is more than three times larger than that in the bulk water,and that these molecules are hydrogen bonded to surface O atoms.In the region of the second statistical water monolayer(6.4?),water molecule density is between1 and3times that of the bulk water.Their calculations reveal that diffusion of both water molecules and ions within these high water density regions is rapid enough to indicate that this water is not rigidly structured in an‘‘ice-like’’manner(Bourg and Sposito, 2011).The presence of regions of hydrogen bonded water and the lack of rigid structure in the vicinity of smectite surfaces seem to indicate that there is the potential for water isotope fraction-ation between this high water density region and the bulk water in a manner akin to that observed in aqueous solutions of structure breaking cations like K+.The effects associated with smectite surfaces may be present in the interlayer space between phyllosi-licate crystal structures and may be contributing to the macro-scopic isotope effect that is observed in the bulk water. Moreover,the effects of the smectite clay surface may enhance the magnitude of the isotopic effects of the structure breaking cat-ions when adsorbed to smectite surfaces.
Beyond these fundamental physical interactions there are prac-tical implications of our observed effects on the use of stable iso-topes in terrestrial ecosystem hydrology.First,the analytical method chosen to assess soil water oxygen isotope composition must be considered in view of the intended‘‘pools’’of soil water to be measured.Direct equilibration,used here,clearly measures a portion of soil water that is free for equilibration with CO2,in contrast to water extracted by various methods which re?ect the total soil water isotope composition.While direct equilibration of CO2with in situ soil water will not re?ect the total soil water com-position,it may in turn be a more reliable guide to the composition of water that is biologically available or reactive in mineral precip-itation(such as calcite).Additionally,a spatial distribution of var-ious pools of soil water may be analytically accessible through the use of the screening effect that the limited CO2diffusive path length in various soil water environments presents.The use of direct equilibration techniques has greatly increased in recent years for laboratory IRMS(e.g.Scrimgeour,1995;Hsieh et al., 1998;Ignatev et al.,2013)and IRIS analyses(Wassenaar et al., 2008),as well as in?eld analyses(Breecker and Sharp,2008; Herbstritt et al.,2012;Rothfuss et al.,2013).
An especially vigorous area of research in ecohydrological sci-ence revolves around investigations into plant water use and its spatiotemporal variability.These approaches generally rely on var-iation in the isotopic pro?le of soil water as a function of depth in the soil horizon,which occurs when dry conditions drive soil sur-face evaporation(Barnes and Allison,1988;Braud et al.,2009). These dry conditions are where we observe the greatest adsorbed cation fractionation effect.Research on plant water use utilizing both oxygen and hydrogen isotopes has recently demonstrated the possibility of two distinct soil water pools:a more mobile pool of water that resembles the precipitation and stream water associ-ated with the local meteoric water line,as well as a second and less mobile pool of water that is consistent with plant water use (Brooks et al.,2010;Goldsmith et al.,2012).The adsorbed cation isotope effect presented here may be relevant to the isotopic com-position of plant available water in soils with some component of high-CEC clay and when soil water conditions are drying.Our experimental results combined with observations of aberrant soil water isotope signatures in?eld soils(Orlowski et al.,2013; Mei?ner et al.,2014)suggests that soil mineralogy may warrant greater consideration in isotopic studies of soil and plant water.
The results here are also relevant to interpreting the stable iso-tope composition of secondary,pedogenic carbonate in paleosols. Clay-rich soils(Vertisols,soils with a minimum of40%expandable clay)are found on outcrops of clay rich rock or sediment,coastal or estuarine environments,basin deposits,or outcrops of shale or mudstone(Soil Survey Staff,1999).Additionally,outcrops of basalt or ultrama?c rock commonly weather,especially in semiarid set-tings,into clay rich Vertisols.Fossil Vertisols(buried clay-rich soils)are especially prominent in the geological record since they form at the distal ends of large subsiding basins.Additionally,these fossil soils are commonly rich in pedogenic carbonate(CaCO3),and thus are similar to the Ca-rich smectite experiments in Table2and Fig.1B.It appears that the water recorded by the oxygen isotopes in carbonate forming in equilibrium with soil water may be about 1per mil more depleted than that which would be recorded by car-bonate forming in equilibrium with the meteoric waters that entered the soil,since pedogenic carbonate minerals precipitate when soils are at their driest(Breecker et al.,2009),soil moisture
Table3
Ionic properties for materials used in this study.Hydration numbers from Ohtaki and Radnai(1993).
Cation Hydration
#K D
(mL/g)
K?D Adsorbed cation molarity(mol/kg)
AZ
montmorillonite
TX
montmorillonite
Mg2+613.533.750.303750.3475
Ca2+69.824.50.303750.3475
K+4922.50.60750.695
Na+4922.50.60750.695
Table4
Measured isotope values and fraction factors per unit mole.All d units are‰per mole
of cation desorbed in solution(aq)or adsorbed to mineral surfaces(ad).
Cation d aq d ad1000ln/
Mg2+ 1.110.960.15
Ca2+0.47 1.19à0.72
K+à0.16à0.90.74
Na+000
E.Oerter et al./Journal of Hydrology515(2014)1–97
contents analogous to <15%in Fig.1B.While the net effect is likely modest,it may have signi?cance in certain paleo-interpretations.The biggest impact of these ?ndings may be on the evolving efforts to use the O isotope composition of atmospheric CO 2to serve as a probe of gross primary production and ecosystem exchange (e.g.Werner et al.,2012).One of the most complex parameters in these efforts is the net effect of soil respiration and abiotic exchange with the atmosphere (Stern et al.,2001).The additional issue of meteoric water reactions with soil minerals and salts adds yet an additional challenge for this approach,and even a few per mil shift in the soil water value on a global scale,has enormous impacts on global estimates of GPP (Stern et al.,2001;Yakir and Sternberg,2000).The in?uence of soil on the d 18O composition of the atmosphere has been investigated by ana-lytic and numeric modeling studies,and one conclusion of this work is that some component of the atmospheric oxygen isotope composition is determined by the global-scale direct equilibration of atmospheric CO 2with soil water (Tans,1998;Stern et al.,1999).As mentioned,soils that may be especially reactive in an oxygen isotope sense with atmospheric CO 2are Vertisols,which are rich in ?ne-grained,smectite clays and undergo intense shrink/swell behavior with desiccation cracks opening during the drying season to depths greater than 50cm (Soil Survey Staff,2010).These cracks increase the area of the soil-atmosphere interface,promoting iso-topic exchange between soil water and the atmosphere.This enhanced atmospheric perfusion to considerable depths,coupled with the occurrence of Vertisols across 2.4%of Earth’s surface (Dudal and Eswaran,1988),suggests that adsorbed cation isotope effects on the d 18O composition of the atmosphere may warrant further investigation.
6.Conclusions
We show that cations adsorbed to high-CEC clay mineral parti-cles can create isotopically distinct ‘‘pools’’of water that may not readily mix with each other or the bulk water in the soil solution.The measured d 18O values of soil water in the vicinity of high-CEC clay minerals may therefore only be re?ecting a portion of the total soil water.Investigations of oxygen isotope dynamics in soil water are of interest to a variety of disciplines ranging from soil science to stable isotope ecology to the study of global biogeochemical ?uxes.Our ability to observe and accurately predict the role of clay mineralogy on soil water oxygen isotopes is a critical step towards
the eventual evaluation of this isotope effect in natural soils and for incorporating the effect in biophysical applications.Acknowledgements
This research is an outgrowth of a project for the class Stable Isotope Ecology taught at the University of California,Berkeley (by T.Dawson,S.Mambelli and G.Goldsmith).We gratefully acknowl-edge isotope analyses provided by the Center for Stable Isotope Biogeochemistry at UC-Berkeley,as well as support from the Clay Minerals Society.We thank Paul Brooks for his assistance with IRMS analyses,and Garrison Sposito and Stefania Mambelli for enlightening conversations that improved this research.The reviews of Ian Bourg and two anonymous reviewers contributed to the improvement of this paper.References
Appelo,C.A.J.,Postma,D.,1999.Geochemistry,Groundwater and Pollution.Taylor &
Francis,536pp.
Araguas-Araguas,L.,Rozanski,K.,Gon?antini,R.,Louvat,D.,1995.Isotope effects
accompanying vacuum extraction of soil water for stable isotope analyses.J.Hydrol.168,159–171.
Baes,C.,Sharp,R.,1983.A proposal for estimation of soil leaching and leaching
constants for use in assessment models.J.Environ.Qual.12(1),17–28.
Barnes,C.J.,Allison,G.B.,1988.Tracing of water-movement in the unsaturated zone
using stable isotopes of hydrogen and oxygen.J.Hydrol.100(1–3),143–176.Bourg,I.C.,Sposito,G.,2011.Molecular dynamics simulations of the electrical
double layer on smectite surfaces contacting concentrated mixed electrolyte (NaCl–CaCl 2)solutions.J.Colloid Interface Sci.360(2),701–715.
Braud,I.,Biron,P.,Bariac,T.,Richard,P.,Canale,L.,Gaudet,J.,Vauclin,M.,2009.
Isotopic composition of bare soil evaporated water vapor.Part I:RUBIC IV experimental setup and results.J.Hydrol.369(1),1–16.
Breecker,D.,Sharp,Z.D.,2008.A ?eld and laboratory method for monitoring the
concentration and isotopic composition of soil CO 2.Rapid Commun.Mass Spectrom.22(4),449–454.
Breecker,D.O.,Sharp,Z.D.,McFadden,L.D.,2009.Seasonal bias in the formation and
stable isotopic composition of pedogenic carbonate in modem soils from central New Mexico,USA.Geol.Soc.Am.Bull.121(3–4),630–640.
Brooks,J.R.,Barnard,H.R.,Coulombe,R.,McDonnell,J.J.,2010.Ecohydrologic
separation of water between trees and streams in a Mediterranean climate.Nat.Geosci.3(2),100–104.
Burgess,J.,1978.Metal Ions in Solution.Ellis Horwood Chichester,481pp.
Coplen,T.B.,Hanshaw,B.B.,1973.Ultra?ltration by a compacted clay membrane—I.
Oxygen and hydrogen isotopic fractionation.Geochim.Cosmochim.Acta 37(10),2295–2310.
Dudal,R.,Eswaran,H.,1988.Distribution,properties and classi?cation of Vertisols.
Publication Soil Management Support Services,US Department of Agriculture,Natural Resources Conservation Service,Washington,DC,pp.1–22.
Epstein,S.,Mayeda,T.,1953.Variation of 18O content of waters from natural
sources.Geochim.Cosmochim.Acta 4(5),213–224
.
relationship of D d 18O to the concentration of adsorbed cation concentration for (A)Mg-and Ca-montmorillonite-water mixtures prepared by the acidic method and (B)K-and Na-montmorillonite-water mixtures prepared by the acidic homoionic washing method.
Feder,H.M.,Taube,H.,1952.Ionic hydration:an isotopic fractionation technique.J.
Chem.Phys.20(8),1335–1336.
Frank,H.S.,Wen,W.Y.,1957.Ion-solvent interaction.Structural aspects of ion-solvent interaction in aqueous solutions:a suggested picture of water structure.
Discuss.Faraday Soc.24,133–140.
Goldsmith,G.R.et al.,2012.Stable isotopes reveal linkages among ecohydrological processes in a seasonally dry tropical montane cloud forest.Ecohydrology5(6), 779–790.
Herbstritt,B.,Gralher,B.,Weiler,M.,2012.Continuous in situ measurements of stable isotopes in liquid water.Water Resour.Res.48(3).
Hsieh,J.C.C.,Savin,S.M.,Kelly,E.F.,Chadwick,O.A.,1998.Measurement of soil-water delta18O values by direct equilibration with CO2.Geoderma82(1–3), 255–268.
Hu,G.,Clayton,R.N.,2003.Oxygen isotope salt effects at high pressure and high temperature and the calibration of oxygen isotope geothermometers.Geochim.
Cosmochim.Acta67(17),3227–3246.
Ignatev,A.,Velivetckaia,T.,Sugimoto,A.,Ueta,A.,2013.A soil water distillation technique using He-purging for stable isotope analysis.J.Hydrol.498,465–473. Ingraham,N.L.,Shadel, C.,1992.A comparison of the toluene distillation and vacuum/heat methods for extracting soil water for stable isotopic analysis.J.
Hydrol.140,371–387.
Johnston, C.,Sposito,G.,Erickson, C.,1992.Vibrational probe studies of water interactions with montmorillonite.Clays Clay Miner.40(6),722–730. Majoube,M.,1971.Fractionnement en oxygene-18et en deuterium entre l’eau et sa vapeur.J.Chim.Phys.68(10),1423–1436.
Marcus,Y.,2010.Effect of ions on the structure of water.Pure Appl.Chem.82(10), 1889–1899.
Mei?ner,M.,K?hler,M.,Schwendenmann,L.,H?lscher,D.,Dyckmans,J.,2014.Soil water uptake by trees using water stable isotopes(d2H and d18O)àa method test regarding soil moisture,texture and carbonate.Plant Soil376(1–2),327–335.
Nag,A.,Chakraborty,D.,Chandra,A.,2008.Effects of ion concentration on the hydrogen bonded structure of water in the vicinity of ions in aqueous NaCl solutions.J.Chem.Sci.120(1),71–77.
Ohtaki,H.,Radnai,T.,1993.Structure and dynamics of hydrated ions.Chem.Rev.93
(3),1157–1204.
O’Neil,J.R.,1968.Hydrogen and oxygen isotope fractionation between ice and water.J.Phys.Chem.72(10),3683–3684.
O’Neil,J.R.,Truesdell,A.H.,1991.Oxygen isotope fractionation studies of solute-water interactions.Stable Isotope Geochem.:A Tribute Samuel Epstein3,17. Orlowski,N.,Frede,H.,Brüggemann,N.,Breuer,L.,2013.Validation and application of a cryogenic vacuum extraction system for soil and plant water extraction for isotope analysis.J.Sensors Sensor Syst.2,179–193.
Phillips,F.M.,Bentley,H.W.,1987.Isotopic fractionation during ion?ltration:I.
Theory.Geochim.et Cosmochim.Acta51(3),683–695.Rothfuss,Y.,Vereecken,H.,Brüggemann,N.,2013.Monitoring water stable isotopic composition in soils using gas-permeable tubing and infrared laser absorption spectroscopy.Water Resour.Res.49,3747–3755.
Schoonheydt,R.A.,Johnston,C.T.,2013.Surface and interface chemistry of clay minerals.In:Fa?za,B.,Gerhard,L.(Eds.),Developments in Clay Science.Elsevier, pp.139–172.
Scrimgeour,C.M.,1995.Measurement of plant and soil water isotope composition by direct equilibration methods.J.Hydrol.172,264–274.
Soderberg,K.,Good,S.P.,Wang,L.,Caylor,K.,2012.Stable isotopes of water vapor in the vadose zone:a review of measurement and modeling techniques.Vadose Zone J.11(3).
Sofer,Z.,Gat,J.R.,1972.Activities and concentrations of18O in concentrated aqueous salt solutions–analytical and geophysical implications.Earth Planet.
Sci.Lett.15(3),232–238.
Sofer,Z.,Gat,J.,1975.The isotope composition of evaporating brines:effect of the isotopic activity ratio in saline solutions.Earth Planet.Sci.Lett.26(2),179–186. Soil Survey Staff,1999.Soil taxonomy:a basic system of soil classi?cation for making and interpreting soil surveys.Natural Resources Conservation Service, United States Department of Agriculture.
Soil Survey Staff,2010.Keys to Soil Taxonomy.United States Department of Agriculture,Natural Resources Conservation Service.
Sposito,G.,2008.The Chemistry of Soils.Oxford University Press,329pp.
Stern,L.,Baisden,W.T.,Amundson,R.,1999.Processes controlling the oxygen isotope ratio of soil CO2:analytic and numerical modeling.Geochim.
Cosmochim.Acta63(6),799–814.
Stern,L.A.,Amundson,R.,Baisden,W.T.,2001.In?uence of soils on oxygen isotope ratio of atmospheric CO2.Global Biogeochem.Cycles15(3),753–759.
Tans,P.P.,1998.Oxygen isotopic equilibrium between carbon dioxide and water in soils.Tellus B50(2),163–178.
Taube,H.,https://www.wendangku.net/doc/644504904.html,e of oxygen isotope effects in study of hydration of ions.J.Phys.
Chem.58(7),523–528.
Turton,D.A.,Hunger,J.,Stoppa,A.,Thoman,A.,Candelaresi,M.,Hefter,G.,Walther, M.,Buchner,R.,Wynne,K.,2011.Rattling the cage:micro-to mesoscopic structure in liquids as simple as argon and as complicated as water.J.Mol.Liq.
159(1),2–8.
Wassenaar,L.I.,Hendry,M.J.,Chostner,V.L.,Lis,G.P.,2008.High resolution pore water2H and18O measurements by H2O((liquid))-H2O((vapor))equilibration laser spectroscopy.Environ.Sci.Technol.42(24),9262–9267.
Werner,C.et al.,2012.Progress and challenges in using stable isotopes to trace plant carbon and water relations across scales.Biogeosciences9,3083–3111. Xu,W.,Johnston,C.T.,Parker,P.,Agnew,S.F.,2000.Infrared study of water sorption on Na-,Li-,Ca-,and Mg-exchanged(SWy-1and SAz-1)montmorillonite.Clays Clay Miner.48(1),120–131.
Yakir,D.,Sternberg,L.,2000.The use of stable isotopes to study ecosystem gas exchange.Oecologia123(3),297–311.
E.Oerter et al./Journal of Hydrology515(2014)1–99