STABLE ISOTOPE GEOCHEMISTRY OF CLAY MINERALSminersoc.org/pages/Archive-CM/Volume_31/31-1-1.pdf ·...

Clay Minerals (1996) 31, 1-24

S T A B L E ISOTOPE G E O C H E M I S T R Y OF C L A Y M I N E R A L S " T h e s t o r y o f s l o p p y , s t i c k y , l u m p y a n d t o u g h " C a i r n s - S m i t h ( 1 9 7 1 )

S. M. F. S H E P P A R D AND H. A . G I L G *

Laboratoire de Sciences de la Terre, URA CNRS 726, ENS Lyon, and *UCB Lyon 1, Ecole Normale SupOrieure de Lyon, 69364 Lyon 07, France

(Received 9 February 1995; revised 3 August 1995)

A B S T R A C T : The equilibrium H- and O-isotope fractionations can be approximated by the following equations which are based on experimental, empirical and/or theoretical data:

Hydrogen: 1000 In ~kaolinite-water ~ " -2 .2 x 106 X T - 2 - 7.7 Oxygen: 1000 In ~kaolinite-water = 2.76 x 106 • _ 6.75

1000 In (~smectite-water = 2.55 x 1 0 6 x T -2 - 4.05 1000 In ~illite-water = 2.39 X 106 • T -2 - 3.76

The equilibrium H-isotope fractionation factors v s . 106 • T - 2 for kaolinite and probably smectite and illite are monotonic curves between 350-0~ More complex curves, with a minimum fractionation near 200~ are probably influenced by surface effects and/or disequilibrium fractionations among the different hydrogen sites. The H-isotope fractionations between smectite-water increase by

70% from Fe-poor montmorillonite to nontronite at low temperatures. The pore-interlayer water in smectite H-isotope fractionation at low temperatures is ~20__+ 10%. The presence of organic matter can modify both the ~SD value of the clay analysis and its 'water' content. Clays - - kaolinite, illite, smectite and probably halloysite - - tend to retain their D/H and 180/160 ratios unless subjected to more extreme diagenetic or metamorphic conditions or special local processes. Kinetic information is still only qualitative: for comparable grain sizes, hydrogen exchanges more rapidly than oxygen in the absence of recrystallization. Low-temperature diffusion coefficients cannot be calculated with sufficient precision from the higher temperature exchange data. The H- and O-isotope studies of clays can provide useful information about their conditions of formation.

Studies of the stable isotope ratios D/H (or 2H/1H) and lSO/160 of clay minerals effectively started with Savin's thesis (1967; see Saviu & Epstein, 1970a,b). They established a number of general isotopic systematics of clay minerals from continental and oceanic environments at surface temperatures. Subsequent studies on natural samples and laboratory experiments (see review by Savin & Lee, 1988 and references therein) have developed the stable isotope geochemistry of clay minerals to a wide range of applications such as: geothermometry; provenance of clays: detrital vs. authigenic; diagenetic processes; origin and evolution of upper crustal fluids; hypogene vs. supergene origin; hydrothermal systems; mineral-water interaction; and palaeoclimates.

All of these applications depend directly or indirectly on knowledge of: (1) the H- and O-isotope fractionation factors in clay-water systems (unless specified to the contrary, water is

used to refer to the aque0us-rich fluid phase that is associated with the clay; it may be more or less saline); and (2) the temperature and time when isotopic exchange with and/or in the clay ceased. Either equilibrium or kinetic isotope fractionation with another phase is responsible for variations in the isotopic compositions of clays. Equilibrium isotopic fractionation factors are a function of temperature. However, variations in the chemical composition of the mineral may be of first-order importance for hydrogen (see below). The composition of the aqueous fluid can influence the H- and O-isotope fractionation factors, but this is usually a secondary effect unless salinities are very high and temperatures low (Truesdell, 1974; Graham & Sheppard, 1980; Horita et al., 1993; Kakiuchi, 1994). Kinetic isotope effects are more complicated; they depend on several parameters including temperature, pressure, particle size and form, mineral and fluid composition, fluid/ solid ratio and time of exchange.

�9 1996 The Mineralogical Society

S. M. F. Sheppard and H. A. Gilg

It cannot be over-emphasized that the isotopic composition of clays can provide information about a geological process only if the mineral has retained the isotopic composition it acquired during that process. This implies that the quality of interpretations of the isotopic compositions depend on several factors including the extent of knowledge of the clay-water fractionation factors and the rates and mechanisms of isotopic exchange among the various potential sites in clays for both oxygen and hydrogen: O in tetrahedral, OH and interlayer water; H in OH and interlayer sites.

This paper examines (1) special problems associated with the meaningful isotopic analysis of clays; (2) the present status of clay-water fractionation factors and how they vary with temperature and composition; (3) pore-interlayer water fractionation; and (4) the evidence for or against post- formation isotopic exchange. It will be shown below that, despite a certain number of publications which have questioned the value of H-isotope studies of clay minerals because of apparent late- stage exchange processes, H-isotope studies can yield useful information about the conditions of clay formation or the subsequent evolution of their environment. The clay minerals considered here are principally kaolinite, smectite and illite since there are so many more data for these. Dickite, halloysite, nacrite, glauconite, serpentine and chlorite are mentioned.

D E F I N I T I O N S

The fractionation factor, a, between two phases A and B is:

~A--B = RA/RB = [1000 + 8A]/[1000 + 8B]

where R = D/H or t80/J60.

The per mil fractionation between two phases is: 1000 In ~A-B. A fractionation is considered to be small if it is close to 0 and large if it is numerically quite different from O, such as -30 .

C L A Y M I N E R A L S A N D I S O T O P I C T E C H N I Q U E S

By comparison with many other silicate minerals, isotopic studies of natural clays raise a number of special problems related to their typically much smaller particle size and hence much larger specific surface, the electrical charge that their surface may carry, and the presence of interlayer water in certain

clays. The size of particles may range from the dimensions of their unit-cell, ~ 10 3 /am, to a few gm leading to specific surfaces of up to ~ 7 6 0 m2g -~ (e.g. Newman, 1987; Nadeau, 1987). Isotopic techniques, including the new laser methods (Sharp, 1990), usually require one milli- gram to tens of milligrams for a single analysis. Thus all these techniques measure bulk properties of at least several billions of individual particles. The geometry of associated particles is quite complex. The <0.1 /am size fraction (equivalent spherical diameter) of illite-smectite mixed-layer minerals measured by Whitney & Velde (1993) had a particle thickness (calculated) of 1.00 nm, mean length of 247 nm (range 45-555 nm) and mean width of 139 nm (range 45-330 nm). Mean particle sizes, thicknesses, and aspect ratios of <2 /am fractions of samples with comparable mineralogical composition from one area can vary enormously (Eberl & Srodori, 1988; Inoue et al., 1988; Eberl et al., 1990; Gilg & Frei, 1994). For example, Eberl & Srodofi (1988) showed that mean particle thicknesses of illites from the Silverton Caldera range from 1.6 to 20 nm. This variability contributes to large uncertainties when deriving quantitative kinetic data from experimental isotopic studies where grain size distributions were not determined properly (see below).

Monomineralic clay samples may be composed of a mixture of: (1) detrital and authigenic components; (2) particles of different ages, etc. because the sample has undergone partial recrystallization by a process such as Ostwald ripening (Baronnet, 1982); or (3) particles that have exchanged to varying degrees due to their range of grain sizes. At this fine micron to sub-micron scale, it is often impossible to separate physically these different generations of the same clay mineral. Nevertheless, the analysis of different size-fractions of the same sample or granulometry in combination with careful petrographic analysis using scanning or transmission electron microscopy or other isotopic systems (K/Ar, Rb/Sr, Sm/Nd) can often help to assess the importance of such processes (e.g. Eberl et al., 1987, 1990; Clauer et al., 1990; Gilg & Frei, 1994).

It is often extremely difficult to have a monomineralic clay sample by grain size separation alone. Elimination of other fine-grained phases in the clay fraction, such as hydroxides, oxides, carbonates, other phyllosilicates, or organic matter, is often necessary. Several techniques have been

Stable isotope geochemistry q[ clays

TABLE 1. Pretreatment and isolation methods in isotope studies on clays.

Removal of organic matter 30% H202, 25-100~ 5% NaOCI solution, pH 9.5, 25-100~ Br2 water, 40~ plasma ashing step-heating

Removal of amorphous silica and gibbsite 5% Na2CO3, boiling 15 min.

Removal of calcite 2m NaOAc-HOAc, pH 4.8 HCI, pH 4.5

Removal of Fe, A! and Mn oxides Na-citrate-bicarbonate-dithionite solution, 80~

Removal of chlorite 10% HCI, 80C, 1 h + 3% Na2CO3

Removal of kaolinite heating to 550~ 2 h + 0.5 N NaOH, 4 h

Removal of mica fusion with Na2S207 + HC1 + NaOH

Removal of feldspars H2SiF6

Jackson, 1979; Yeh, 1980 Anderson, 1963; Bird et al., 1992 Mitchell & Smith, 1974 Taieb, 1990; Hogg et al., 1993 Gilg, Sheppard & Kahr (unpublished data)

Jackson, 1979; Bird et al., 1992

Jackson, 1979; Yeh, 1980

Jackson, 1979; Yeh, 1980

Longstaffe, 1986; Ayalon & Longstaffe, 1988

Hashimoto & Jackson, 1960

Syers et al., 1986

Syers et al., 1986

developed (Jackson, 1979) and employed as pretreatments for stable isotope analyses. They are compiled in Table 1. Organic matter can pose considerable problems especially in hydrogen isotope analyses: (1) it is ubiquitous in certain clays and especially in the finer fractions either absorbed on the surfaces and edges or accommo- dated in interlayer spaces; (2) it is not detected by most of the widely used mineral identification techniques; (3) it can contribute measurable amounts of hydrogen with its own D/tt ratio (Fig. 1) to many isotopic analyses; and (4) it can resist conventional cleaning treatments (e.g. Farmer & Mitchell, 1963; Jackson, 1979). Step-heating extraction techniques can sometimes be used to separate organic hydrogen (and carbon), which typically is liberated between 250 ~ and 350~ (Fig. 1), from the hydroxyl hydrogen released at temperatures >400~ (Gilg, Sheppard & Kahr unpublished data). Although hydrogen yields are measured as molecular hydrogen after heating the mineral at temperatures less than 120 ~ to 200~ to

remove adsorbed and interlayer water, they are conventionally presented as H20 § contents. Part of this so-called H20 + may in fact have been released as hydrogen or methane gas from the organic constituents. This can account for some of the H20 + values that are higher than the maximum theoretical value for the mineral.

Surfaces of clays, especially expandable ones like smectites, are invariably associated with 1- or 2-layer thick adsorbed water. Mourn & Rosenqvist (1958) and Savin & Epstein (1970a) demonstrated that adsorbed and interlayer water exchange isotopically with atmospheric water vapour in hours. Therefore, they must be separated from the framework [SiO4] and [OH] oxygen and hydrogen. Pre-heating the sample under vacuum at temperatures up to 200~ removes most of this water without major isotopic exchange and satisfactory isotopic data can be obtained (Savin & Epstein, 1970a). However, complete removal of interlayer water with the total absence of isotopic exchange between it and the OH group may not be possible in

6 D (%0)

613C

(%4

S. M. F. Sheppard

exothe~1r

I I I I [ I I I I 100 20o 30o 400 500 60o 700 800 900

Temperature (~

TG

DTG

M18 (H20) M44 (C02)

FIG. 1. Differential thermal (DTA), thermogravimetric (TG) and differential thermogravimetric (DTG) analyses of a hydrothermal illite (B77) and mass spectro- metric analysis of the evolved gases (M18 = H20; M44 = CO2) in air. The ion current intensities of the two masses do not have the same scale. The 6D and 613C values of the gases evolved under vacuum between 180 ~ and 400~ and >400~ are marked on the peaks. The peak at 300~ that marks an exothermic reaction with release of CO2, indicates an oxidation reaction of the organic matter associated

with the illite (modified after Gilg, 1993).

all cases (Lawrence & Taylor, 1971; Newman, 1987; Cuadros et al., 1994). The precision of H- and O-isotope analyses of clays is thus often a little lower than that of most other minerals. A laboratory study on halloysite detected isotopic exchange be tween in te r layer and hydroxyl hydrogen (Lawrence & Taylor, 1971). Thus meaningful D/H measurements may not be obtained from this mineral. This may be attributed to the immediate contact and hydrogen bonding between interlayer water and surface hydroxyls in halloysite (Giese, 1988) in contrast to smectites, where interlayer water is separated by a tetrahedral sheet from the hydroxyl groups. However, many natural halloysites occur in a partially or completely dehydrated 7 A form (metahalloysite). The reconstitution of interlayer water in halloysite is, in contrast to smectites, not reversible. As none of the existing isotope studies mentions the hydration state of the analysed halloysites (10 ,~ vs. 7 ,~), a rigorous evaluation of

and H. A. Gilg

the data cannot be given. However, a compilation of isotope data on halloysite from weathering environments (see Fig. 13) suggests that the above-mentioned problem may not be that serious (see below).

Pre-fluorination techniques should be omitted before starting the oxygen extraction procedures because of fractionation effects (Hamza & Epstein, 1980; Hogg et al., 1993).

F R A C T I O N A T I O N F A C T O R S

Ideally, fractionation factors for minerals should be firmly based on simple isotopic exchange experiments, where phases are neither created nor destroyed, equilibrium is approached from both directions and surface fractionation effects are negligible. Despite a number of such laboratory studies, the results have only been partially successful because the percentage of exchange was often much less than 100%, unless the temperature was quite close to the stability limit of the clay (O'Neil & Kharaka, 1976; Liu & Epstein, 1984). In practice, a multiple approach is used with additional results from synthesis experiments, and empirical and theoretical determinations. Each method is associated with its uncertainties and limitations.

For the partial exchange data, extrapolation techniques are applied to derive an estimate of the equilibrium fractionation factor. They are usually based on the Northrop & Clayton (1966) model or a model where part of the mineral is assumed to have completely re-equilibrated but the rest has not exchanged at all (Liu & Epstein, 1984). For clay minerals, however, the Northrop & Clayton (1966) extrapolation technique may not give results that are similar to empirically derived values when the amount of exchange is < ~75%. Such models implicitly require that all the atoms of the element of interest are equivalent or exchange with the same rate. This is not the case for clays (or serpentine) (Sheppard, t980). Taking kaolinite as an example, although hydrogen only occurs as hydroxyl in the octahedral sheet (Fig. 2), they can be divided into two major types: (1) the inner hydroxyls that are within the same layer as the non-bridging oxygens of the tetrahedral sheet and represent a quarter of the hydroxyls; and (2) the outer or 'inner-surface' hydroxyls. Hydrogens in these two sites neither exchange at the same rate (e.g. Ledoux & White, 1963, 1964) nor necessarily have identical isotopic

t_ i T M

Stable isotope geochemistty of clays

~) outer OH

0 inner OH

� 9 0 AI

�9 Si

FIG. 2. A diagrammatic projection down the b-axis of the structure of kaolinite showing the outer and inner

hydroxyls (after Giese, 1988).

properties, complicating the derivation of equilibrium fract ionat ion factors from incomplete exchange reaction data.

In experiments where exchange with the surface layer of the mineral dominates the system, differences in the fractionation factor between the surface layer and the bulk mineral may be important (Hamza & Broecker, 1974; Stolper & Epstein, 1991; Matthews et al., 1994). For example, Hamza & Broecker (1974) have shown that for the system calci te-CO2 at 200~ the O-isotope fractionation factor between CO2 and the surface layer is 5-6%0 larger than that between CO2 and the bulk calcite. This phenomenon may be important for the finer or high surface/volume ratio fraction of clay minerals or in systems where the degree of exchange is very low.

During a synthesis experiment, kinetic isotope effects are possible and rigorous proof of equilibrium is usually lacking. However, the study of O'Neil & Taylor (1967) on O-isotope fractionations between water and feldspar have demonstrated that simple exchange and synthesis experiments involving cation exchange can give concordant fractionations. Similarly, they showed that cation exchange reactions and synthesis of muscovite from gels also gave concordant results (O'Neil & Taylor, 1969).

Empirical methods may measure the isotopic composition of a mineral and its associated water (or another mineral) and assume that they are in equilibrium at the measured or estimated temperature. However, clays from weathering environments are not necessarily in equilibrium with actual meteoric waters and in active geothermal systems

the clay mineral may not have equilibrated at the measured temperature or with the present associated fluid.

Statistical mechanical calculations of fractionation factors are possible but for mineral systems they are usually limited by the quality of the spectroscopic data and the approximations that are made in order to carry out the calculations (e.g. Kieffer, 1982; Clayton & Kieffer, 1991). No calculations have been published for H-isotope fractionations involving minerals.

The O-isotope fractionations for phyllosilicates in particular have also been estimated from bond-type calculations by assuming that the fractionation between two phases is the weighted sum of the fractionations of the different bond types present (Savin & Lee, 1988). Estimated fractionations for the different bond types are based on the limited experimental and empirically determined fractionations of minerals. Because the approach is not independent of the above methods, the bond-type calculations cannot be used as a check on the other methods. Yet other bond-type calculations have been presented (Sch~tze, 1980; Zheng, 1993). Such calculations are particularly useful for minerals for which no experimental or empirical data exist. They are not discussed further.

The approach here will be to examine critically both the experimental and empirical data on which the fractionation factors have been based. Several significant revisions are presented. The new curves are given as linear functions of T -2, because the combined experimental and empirical data are still too limited to test the justification of other expressions.

Hydrogen isotope f rac t iona t ions

Kaolinite. Figure 3 presents a summary of the experimental and empirical data on kaolinite-water fract ionations over the temperature interval 15-350~ The experimental data of Sheppard et al. (1969), O'Neil & Kharaka (1976) and Liu & Epstein (1984) are based on the simple exchange reaction method, whilst Kulla (1979) used a synthesis technique. In these latter experiments, a calcinated gel with the chemical composition of the clay was recrystallized to the clay mineral in the presence of waters with different isotopic compositions and 'equi l ibr ium' was approached and bracketed from opposite directions. Water-kaolinite ratios were >1 except in the Liu & Epstein (1984)


Temperature (~ 700 400300 Z00 150 100 75 50 2__5 10 0

10 L , , , J , ? , , . . . . .

~ t ? ~ IN ? ?ONM

-1o ? { ~ ? N?

f -20 ?

-40

-50 - 1000 Inc~ = -2.2"106/T2-7.7

f -7o D k a o l i n i t e

0 2 4 6 8 1 12

106/T 2 14

�9 Kulla (1979) exp: synthesis, high W/R

�9 Liu & Epstein (1984): exp: exchange, low W/R

�9 O'Neil & Kharaka (1976) recalculated [] Sheppard et al. (1969)

exp: exchange, high W/R

[] Maruino et al. (1980) [] Lambert & Epstein (1980) [] Sheppard et al. (1969)

emp: geothermal

O Savin & Epstein (1970a) (I) Lawrence & Taylor (1972)

crop: weathering

FIG. 3. H-isotope fractionations between kaolinite and water determined experimentally and empirically. NM = New Mexico; ? = experiment with low percentage of exchange or uncertainty in 8D value of meteoric water (see text); exp = experimental;

emp = empirical; W/R = water-rock ratio.

runs where they were so small ( ~ 1/86) that surface exchange processes dominated for small percen- tages of exchange (T<275~ Data above 330~ are not considered because kaolinite becomes unstable. The kaolinite data should be the simplest because variations in its chemistry are too minor to influence significantly the fractionation factors. Nevertheless, between 100 and 300~ both the experimentally and empirically derived fractionations are quite dispersed. The experimental data around 300~ are considered to be satisfactory

because the extent of exchange is >75% and the different techniques give concordant results. All the ~ 2 0 ~ data are based on combining the measured isotopic composition of the clay from a weathering environment with an estimated value for present- day meteoric water (Savin & Epstein, 1970a; Lawrence & Taylor , 1972). F rom the f ive kaolinites, only one is recent (Lawrence & Taylor, 1972). Their value ( 0 ~ k ao l_ w a te r = 0.968) has been used to derive the fractionation curve (Fig. 3). The precision of this fractionation is not particularly high, because there is no isotopic analysis of the directly associated meteoric waters. Three of the other factionations scatter about this value, since the isotopic compositions of the local meteoric waters have probably been rather similar to actual meteoric waters for the past tens of millions of years because of minor changes in latitude and altitude, and hence climate (Hassanipak & Eslinger, 1985). The New Mexico kaolinite (NM, Fig. 3), with a fractionation of +3 (see Savin & Epstein, 1970a, Table 2), plots on the kaolinite line of Savin & Epstein (1970a), indicating that it formed in equil ibrium with meteoric waters which were ~35%o enriched in deuterium relative to present waters. It comes from the post-depositional alteration of a Cretaceous sandstone (Savin & Epstein, 1970a). Because of the major change in altitude of New Mexico since the Cretaceous, gD of meteoric water has decreased by about 30 to 40%, (compare Fig. 3a with 3b in Sheppard, 1986). Its H- and O-isotope composit ions are consis tent with a Cretaceous-Tertiary age. Much of the dispersion of the data at ~ 2 0 ~ is thus considered to be due to uncertainties in the 6D value of the water; variations in the temperature of formation also contribute, but are minor.

Between 300 and 20~ there are both experimental and empirical data suggesting that either the fractionation factor changes essentially monotoni- cally (Sheppard et al., 1969; Kulla, 1979; Marumo et al., 1980) or the curve is more complex with a max imum (min imum frac t ionat ion) at 200~ (Lambert & Epstein, 1980; Liu & Epstein, 1984). We propose that the best e s t ima te of the fract ionation relat ion is the monotonic curve shown on Fig. 3, for the following reasons.

In the experiments of Liu & Epstein (1984) below 300~ the calculated percentage of kaolinite reacted was always <55%. In addition to the problems associated with the extrapolation technique for low degrees of exchange mentioned

Stable isotope geochemistry of clays

above , surface f rac t iona t ion effects may be important . At 200 and 250~ the calculated kaolinite-water fractionation becomes either less positive or more negative with increase in reaction time (Liu & Epstein, 1984; Table 1). This implies that the fractionations after complete exchange would be even more negative than the estimated values, decreasing the difference between our curves. The comparable empirical data of Lambert & Epstein (1980) are based on whole-rock drill cuttings samples of hydrothermally altered tufts with only 3 - 9 % kaolinite coming from two wells in the Valles Caldera, New Mexico. The 230 and 280~ data are on blown-out fragments. Subsequent detailed mineralogical and geochemical studies on this well have shown that kaolinite is restricted to temperatures <70~ (Hulen & Nielson, 1986) and that the actual well waters are not in chemical equilibrium with the kaolinitic alteration (White, 1986). The isotopic data at these temperatures cannot be used directly to derive fractionation factors. The three kaolinites with temperatures of ~ 6 5 ~ 150 ~ and ~ 1 8 0 ~ come from a well in another area of the caldera. Although no isotopic data on waters from this well are reported, waters from neighbouring wells are isotopically variable, due to boiling and mixing processes, with values of 15-40%o heavier than local meteor ic waters (Vuataz & Goff, 1986). The hot waters are acid and could be in chemica l equ i l ib r ium with kaolinite. However, the fractionations for these kaolinites estimated by Lambert & Epstein (1980) could readily be too small by ~ 15%o because they used the local meteoric water value. Their data become consistent with the proposed curve if such a correction is applied.

The other empirical data on pure kaolinite have been used to derive our proposed curve. The kaolinite and/or dickite separates are stable in these acid hot springs waters. They come from three geothermal systems with very different 6D water values in the USA (Sheppard et al., 1969) and Japan (Marumo et al., 1980). Al though O-isotope data have not been published in detail on either the waters or clays from the two Japanese areas, preliminary data of Marumo et al. (1982) indicate that in the Ohnuma well the O-isotope fractionations for two of three clays are a little different from the values expected for the estimated temperatures if, as assumed, the ~ilSo values of the waters are constant throughout the well. Empirical fract ionat ion data estimated from alunite and

kaolinite/dickite data from acid sulphate alteration at 100-300~ given in Rye et al. (1992) also support the proposed fractionation curve. There is no ev idence for e i the r a H- or O- i so tope fractionation between dickite, nacrite and kaolinite (Lombardi & Sheppard, 1977; Marumo et aL, 1980; Marumo, 1989).

Smectite. Smectites are expected to be more compl ica ted than kaol ini te because of major variations in their chemistry and certain authigenic oceanic smectites could have formed at temperatures down to 0~ Figure 4 summarizes the experimental and empirical results. Kulla (1979)

r

o o o

lO

o

-lO

-20

-30

-40

-50

-60

-70

-80 o

T e m p e r a t u r e ( *C)

7oo4oo3oo 2oolso loo 7s so zs ao o I [ I I I I i I I I

�9 oo o%

~ ,A,. o %

? m*

D/H smectite

i I ~ I I

2 4 6 8

1 0 6 / T 2

i l

1o

|

i12 I 14

Kulla (1979): �9 montmorillonite (Fe-free) "A" Mg-saponite

exp: synthesis, high W/R

�9 O'Neil & Kh.araka (1976)-recalculated: montmorillonite (XFe = 0.1) exp: exchange, high W/R

[] Capuano (1992) illite-smectite mixed--layer minerals (XFe = 0.1 to0.2) emp: sedimentary basin

@ Savin & Epstein (1970a) emp: oceanic no chemical data, probably Fe-rich

@ Lawrence & Taylor (1971) emp: weathering montmorillonite Fe-poor (XFe<0.2)

| Yeh & Epstein (1978) emp: oceanic no chemical data

O France-Lanord & Sheppard (1992) emp: oceanic Mg-Fe-smectite (XFe = 0.35)

FIG. 4. H-isotope fractionations between smectite or illite-smectite and water determined experimentally and empirically. Xve refers to the fraction of the octahedral site occupied by iron (largely as Fe 3+) and

see Fig. 3.

8

D/H 0

- 2 0

S. M. F, Sheppard and H. A. Gilg

Fractionation Smectite-Water , , , , T , , , ,

0.4 0.6 0.8 1.0

Fe3+/(Fe3++Mg+AI)

- 4 0 r

- - - 6 0 Q O O - 8 0

- I 0 0

-120 0.0 0.2

octahedral

FIG. 5. Empirical determination of H-isotope fractionation between smectite and water as a function of iron content at low temperatures. Data from Gilg &

Sheppard (1995).

synthesized Al-rich Fe-free montmorillonites and Mg-saponites. Exchange experiments by O'Neil & Kharaka (1976) on Fe-poor montmorillonites had a high percentage of exchange at 350~ (86%) for the longest run, but the formation of mixed-layer illite- montmorillonite was observed. A value of 1000 In of - 1 9 %o is calculated from these experiments. Surprisingly few data are available on chemically well characterized smectites. The experimental data for Fe-poor montmorillonites are essentially concordant at T ~300~ and this fractionation (1000 In = -15- t -5 %o) is taken as the best estimate. The saponite fractionation is systematically ~15%0 larger, presumably reflecting the influence of the chemical composition of the smectite on the fractionation. This difference is, however, substan- tially larger than that predicted by the Suzuoki & Epstein (1976) atomic mass/charge relationship for the octahedrally coordinated cation. The empirical values below 25~ are also considered to be reasonable with the large variations predominantly reflecting variations in the chemistry of the clay. The fractionation increases with increase in Fe 3+ content (Fig. 5) but the data (Gilg & Sheppard, 1995) do not follow the atomic mass/charge relationship of Suzuoki & Epstein (1976) that was principally based on data for di-and tri-octahedral micas where Fe is principally divalent.

The empirical values between 100 and 150~ (Fig. 4) are for mixed-layer illite-smectite, where

the illite content increases from 25 to 80%, from the Gulf Coast sedimentary basin given in Capuano (1992). She combined a single 8D value for illite- smectite because of its relatively constant value in the geopressured zone in one well (Yeh, 1980) with 8D analyses of waters, which vary systematically with temperature, from the geopressured zones from other wells in the same field. Taking into account the effect of the chemistry of these brines on the fractionations would only decrease the values by about 1-2%o (Horita et al., 1993). The Fe203 content of the <0.1 gm illite-smectite fraction decreases from 6.5 to 3.5 wt% with increase in illite content (Hower et al., 1976). The O-isotope compositions of these clays are quite variable (Yeh & Savin, 1977) but they cannot be simply related to the variations in the water analyses from the other wells at the measured temperatures (Capuano, 1992). Assuming that the H-isotope fractionation factors for montmorillonites and illite are similar at the same temperature (compare Figs. 4 and 6 ), then these empirical fractionations could be approximately correct.

At cont inenta l surface tempera tures , the montmorillonite-water fractionation is indistinguish- able from that of kaolinite (Lawrence & Taylor, 1971). Similarly, fractionations between hydrothermal montmorillonite and kaolinite must be close to zero (Sheppard et al., 1969). For all of these reasons, the Iow-Fe montmorillonite-water fractionation as a function of temperature could be very similar to that shown on Fig. 3 for kaolinite.

lllite. The experimental and empirical fractionation data for muscovite and illite are shown on Fig. 6. The experimental data are for muscovite above 400~ (Suzuoki & Epstein, 1976) and illites with variable extent of exchange (26% at 200~ and 70% at 350~ - - recalculated values) by O'Neil & Kharaka (1976). Most of the empirical and experimental data suggest that the fractionation is not very sensitive to temperature variations in the range 400 to 120~ (1000 In cx = -25__+5 %0). The ~280~ fractionation value from the Geysers system of Lambert & Epstein (1980) is smaller than all the other illite values except the mixed- layer illite-smectites with high illite contents shown on Fig. 4. Since no details of the Geysers system have been published, further comment is not warranted. The large fractionation reported for glauconites (~gia .... . ~,t~r = 0.927) by Savin & Epstein (1970a) is related to its high Fe content, largely present as Fe 3..


T e m p e r a t u r e ( ~

700 4130300 200 150 100 75 50 25 10 0 1 0 i i ~ i i i i 1 t I I

0 | -10 ?l~lg [ ]

-20 7 , - - ? j [ ] - - [ 3 [ 3 0 -30 m

0 0 -40

-50

-60

- 7 o D / H i l l i t e / m u s c o v i t e O

I I I I U ~ I 112 - 8 0 0 2 4 6 8 1 4

1 0 6 / T 2

�9 Suzuoki & Epstein (1976): muscovite �9 O'Neil & Kharaka (! 976)-recalculated: illite

exP: exchange, high W/R

[] Marumo et al. (1980): illite Lambert & Epstein (1980): illite emp: geothermal

m Blamart eta/ . (1992): muscovite [] Landis & Rye (1974): muscovite [] Gilg eta/ . (unpubl.): muscovite

emp: fluid inclusion

O Savin & Epstein (1970a): glauconite emp: oceanic (XFe - 0.5-0.6)

FIG. 6. H-isotope fractionations between muscovite, illite or glauconite and water determined experimen-

tally and empirically. Symbols: see Figs 3 and 4.

Oxygen isotope fractionations

T e m p e r a t u r e ( *C)

zoo 4oo3oo zoo~so ~ooTs so zs ~o o 3~ I I I I I I I I I I

30 1 8 0 / 1 6 0 k a o l i n i t e 7o NM /

~ 2 0

o 15

2 lo /

1000 lmx = 2.76" 106/T 2 - 6.75

- 5 I , i i , i , I 0 , l t2 , 2 4 6 8 1

106/T 2

�9 Kulla & Anderson (1978) exp: synthesis

A Sheppard et al. (1969) bond-type model

[] Sheppard et al. (1969) (Yellowstone, USA)

[] Eslinger (1971) (Broadlands, N. Zealand)

[] Marumo eta/. (1982) (Ohnuma, Japan)

[] Marumo etal. (1982) (Otake, Japan) emp: geothermal

O Savin & Epstein (1970a) (I) Lawrence & Taylor (1972)

emp: weathering

1 4

FIG. 7. O-isotope fractionations between kaolinite and water determined experimentally and empirically.

Symbols: see Fig. 3.

Kaolinite. The experimental and empirical fractionation factors for kaolinite are shown on Fig. 7. All the experimental data are based on synthesis reactions from gels by Kulla & Anderson (1978). The empirically derived fractionations at surface temperatures by Savin & Epstein (1970a) and Lawrence & Taylor (1972) are probably more tightly grouped than those presented on Fig. 7, after corrected 8180 w a t e r values are used. Similar to that for hydrogen (see above), the fractionations derived by Savin & Epstein (1970a) for the New Mexico (NM) and Arkansas kaolinites (both marked by "?" in Fig. 7) are probably too large because they

did not re-equilibrate with present meteoric waters. For example, the fractionation for the New Mexico kaolinite could easily be 3-5%o smaller than that marked on Fig. 7, if it equilibrated with Cretaceous or Tertiary waters. The recent kaolinite from Elberton, Georgia (Lawrence & Taylor, 1972) i s considered to give the best estimate of the fractionation: ~kaolinite-water = 1.026 + 0.001. The curve on Fig. 7 has been drawn through this point and the experimental data for T > 170~ Empirical data for geothermal systems plot on both sides of this curve. Certain of these points are unlikely to represent equilibrium values, perhaps because either

10 s. M. F. Sheppard and 14. A. Gilg

TABLE 2. Comparison of the O-isotope fractionation between kaolinite and water given by different fractionation equations.

Temperature (oc)

1000 In 0co• This study Kulla & Anderson Land & Dutton Savin & Lee Zheng

(1978) (1978) (1988) (1993) 1 2 3 4 5

20 25.4 20.0 26.2 25.7 30.0 50 19.7 15.8 21.1 21.5 23.2 100 13.1 10.9 15.1 16.1 15.6 150 8.7 7.6 11.1 12.1 10.8 200 5.6 5.3 8.3 8.9 7.6 300 1.7 2.4 4.7 4.4 3.9

1:1000 In ~ = 2.76x 106/7'2-6.75 2 :1000 In c~ = 2.05 x 106/7e-3.85 3 :1000 In cx = 2.50 x 106/TZ-2.87 4 :1000 In ~ = 0.42x 106/7e+10.6 x 103/T-15.34 5 :1000 In ~ = 4.29 x 106/Tz-6.44 x 103/T+2.03

the measured temperature or the 5180 value of the water was not identical to that during formation of the kaolinite. The equation proposed by Land & Dutton (1978) was based on a single kaolinite from the Broadlands geothermal field analysed by Eslinger (1971) and the surface kaolinite value of Savin & Epstein (1970a). It is significantly different from the equation proposed here (Table 2). A comparison of the fractionations as a function of temperature are given in Table 2 for the principal equations which have been proposed in order to emphasize the importance of some of the differences.

Smectite. Figure 8 presents the experimental synthesis data of Kulla (1979) and the low- temperature empirical data. The smectites are montmorillonites except, possibly, for some of the oceanic samples. Fractionations tend to decrease with increase in Fe content (Lawrence & Taylor, 1972). However , too few iso topic data on chemically characterized smectites exist to quantify the effect.

lllite. No experimental data which can be used to derive equilibrium fractionations for illite have been published. Exchange rates are exceedingly slow even at 350~ (O'Neil & Kharaka, 1976). At high temperatures (>400~ fractionations derived from synthesis and cation exchange experiments on muscovite by O'Neil & Taylor (1969) are shown on Fig. 9. The low-temperature (22~ experimental

datum derived from limited partial exchange of chemically treated (Na-EDTA) and modified illite by James & Baker (1976) is not considered here because it does not fulfill the necessary criteria to apply the Northrop & Clayton (1966) extrapolation technique. The empirical data were recalculated from the illite data on the Broadlands geothermal system of Esl inger & Savin (1973) using a measured value of -3%o for the water. Although illite may not be stable at surface temperatures, the curve is drawn using the low-temperature value derived by theoretical calculat ions of Kieffer (1982). A fractionation factor derived from an oceanic glauconite (Savin & Epstein, 1970a) is included on Fig. 9. It plots slightly below the curve. Taking into account the role of Fe, this empirical fractionation is consistent with the proposed curve.

Single-mineral O-isotope fractionations

Minerals with oxygen atoms in different structural sites, such as tetrahedral [SiO4] and octahedral hydroxyl, are potentially the basis of single-mineral thermometers . Fol lowing the deve lopmen t of analytical techniques to separate the two types of oxygen in micas and kaolinite by Hamza & Epstein (1980), Bechtel & Hoernes (1990) have refined the technique further and applied it to Fe-free illite. They presented the following preliminary empirical

T e m p e r a t u r e (~

7oo 4o03oo zoo 1so too 7s so zs lo 3 5 i i i i i i i i i i i .

3o 180/160 smectite / ~

25

20

c) 15 Q

0 1 0

5

0 1000 lnot = 2 . 5 5 " 106IT 2 - 4 . 0 5

- 5 - ~ I t i i J i_ i ; i i 0 2 4 6 8 1 1 2 1 4

1 0 6 / T 2

�9 Kulla (1979): montmorillonite (Fe-free) / Mg-saponite exp: synthesis

& Sheppard et al. (1969): montmorillonite bond-type model


T e m p e r a t u r e (~

7oo4oosoo ~o~so roo ts so zs ~o o 3 5 - - i , , , i ~ ~ ~ , i , .

30 l

25

20

0 15 0 0 tO

5

0 ~ 1000 lnot = 2 . 3 9 " 1 0 6 / r - 3 . 76

i i i ] , I ; i I i

- 5 2 4 6 8 1 1 P

1 0 6 / T 2 �9 O'Neil & Taylor (1969): muscovite

exp: synthesis, cation exchange

(1) Savin & Epstein (1970a) emp: oceanic no chemical data, probably Fe-rich

@ Lawrence & Taylor (1972) emp: weathering Fe-rich montmorillonite (X Fe=0.2)

@ Lawrence & Taylor (1972) emp: weathering Fe-poor montmorillonite (XFe=0. l)

@ Yeh & Savin (1976) emp: oceanic no chemical data

C) Hein et al. (1979) emp: oceanic Fe-rich montmorillonite (XFe=0.27)

[] recalculated from Eslinger & Savin ( 1973): illite emp: geothermal

/x Kieffer (1982): muscovite theoretical calculation

C) Savin & Epstein (1970a): glauconite emp: oceanic (XFe = 0.5-0.6)

FIG. 8. O-isotope fractionations between smectite and water determined experimentally and empirically.

Symbols: see Figs. 3 and 4.

11

1 4

FIG. 9. O-isotope fractionations between muscovite, illite or glauconite and water determined experimentally, empirically and theoretically. Symbols: see Figs.

3 and 4.

calibration of the fractionation between the total mineral oxygen and the hydroxyl oxygen in the temperature range 200-300~

1000 In ~ ( i l l i t e - O I t ) "= -0.076t + 30.42 (t in ~

These results are very encouraging since the temperature dependence of the fractionation is large. Further studies and applications are merited.

Summary

The preferred estimated H- and O-isotope fractionation equations for kaolinite, smectite and illite are given in Table 3. Equations for H-isotope

fractionations have not been derived for smectite or illite because of the influence of chemical composition on the former and lack of data for the latter; the fractionation equations could be rather temperature insensitive for both illite and Fe-poor montmorillonite, like that for kaolinite. Data for related minerals like chlorite and serpentine have not been presented. Experimental and/or empirical data are either non-existent or limited for these minerals (see review by Savin & Lee, 1988). Differences between the fractionations for T < 300~ derived from the H-isotope experimental data for chlorite (Graham et al., 1987) and serpentine (Sakai & Tsutsumi, 1978)

12 S. M. F. Sheppard and H. A. Gilg

and the empirical expressions (Wenner & Taylor, 1973) increase with decrease in temperature (or decrease in percent of exchange). Perhaps surface related fractionations and/or disequilibrium fractionations among the different sites are important for the low percent of exchange experiments and the calculated fractionations do not represent the bulk mineral. Further advances and refinements must await new experimental, empirical and theoretical d a t a on c h e m i c a l l y a n d p h y s i c a l l y well-characterized minerals.

I N T E R L A Y E R W A T E R

Routine isotopic analyses of smectites are carried out after removal of the adsorbed and interlayer water. The measured 8D value is thus that of the hydroxyl hydrogen whilst the 81SO value measures the combined tetrahedral and hydroxyl oxygen (depending on the clay, hydroxyl oxygen accounts for 17-44% of the framework oxygen).

Since isotope geochemists routinely 'throw away' the interlayer water, its role during the evolution of a smectite-rich rock has perhaps been neglected. However, a smectite contains ~ 15 wt% interlayer water and only 8 wt% water as hydroxyl. Therefore, in a smectite-rich sediment, interlayer water can represent 15-50 mol % of the total rock (minerals + pore-water) hydrogen. Knowledge of the pore- interlayer water H-isotope fractionation is necessary in order to understand the evolution of the H-isotope composition of the total rock system. This fractionation factor has been estimated using two very different metbods.

Ta~'eb & Sheppard (unpublished data, 1986) exchanged sorbed, i.e. interlayer plus adsorbed, water of Wyoming smectite (montmorillonite) with four different water vapours of known isotopic composition for several days at room temperature. Although the experimental set up was not perfect because the transfer of the smectite from its controlled water vapour atmosphere to the vacuum extraction line took several minutes, seven independent measurements of the liquid water-sorbed water H-isotope fractionation gave

(~fi-ee wal t . . . . . . . bed water = 1.012 __. 0.006. Sorbed water yields calculated a s n H 2 0 gave values of n from 4 to 17 with no obvious systematics between a and either the 8D value of the vapour or the value of n.

During Leg 129 of the Ocean Drilling Program in the Central Western Pacific, C. France-Lanord took

6 mm whole-rock microcore samples (minerals + pore-water) and sealed them in pre-weighed quartz tubes and sampled pore-waters from an identical sample by squeezing (France-Lanord & Sheppard, 1992). Three quartz-tube samples, two of which contained ~90% smectite (Fe-Mg montmorillonite (France-Lanord et al., 1992)), were broken under vacuum and heated to 140~ for 20 h to liberate both the pore and sorbed waters (a very small fraction of the interlayer water may not be removed under these conditions). The mass and isotopic composition of the sorbed water was measured. From these and the 8D value of the pore-water, the H-isotope composition of the sorbed water can be estimated from the mass balance equation:

T~Dtota I = P ~ D p o r e + S~Dsorbed

where T, P and S are the hydrogen contents of the total bulk sediment, the pore-water and sorbed water, respectively. Sorbed water was assumed to be dominated by interlayer water because zeolites are essentially absent and adsorbed water is probably minor. From the interlayer water content, I, the 8D value of the interlayer water was estimated from:

8Dinter layer = {T~Dto ta I - PSDpore}/1 and T = P + 1

Following Bird (1984), the interlayer water contents of the smectites were estimated to be between 15 and 20% of the dry weight, leading to estimated pore-interlayer water fractionations of about 20 to 30%0 (Fig. 10). The agreement with the ~ values of Tareb & Sheppard mentioned above is considered to be very reasonable taking into account all the uncertainties on the various values.

Lawrence & Taviani (1988) proposed that interlayer water is enriched in D by about 20%o relative to pore-water because interlayer water has an ice-like structure and the ice-liquid water fractionation is ~ 2 0 (Friedman & O'Neil, 1977). This is in the opposite sense of the results from both of the above methods. It emphasizes again the lack of even qualitative theoretical models for predicting H-isotope fractionations in mineral systems.

Interlayer water of smectites thus represents a D- depleted reservoir of hydrogen relative to the associated pore-waters. The influence of this reservoir will be more important the more closed the behaviour of the system. As both interlayer and hydroxyl hydrogen are depleted in D relative to

Stable isotope geochemistry of clays 13

| 40

3o

o ~ 20 ~ ~

O I I I t I

3o

~ ~ 2O - ~ 40

o 54 l o

I0 IS 20 25 30 35 40 "i" interlayer-water content of smectite

(% dry wt.)

FIG. 10. Variation of (A) H-isotope fractionation between pore and interlayer water, and (B) pore-water content of sediment (P), as a function of the interlayer water content of the smectite (i). The calculations take into account the 8D values of pore and total waters and the amount of total water (T) and of smectite in each sample. The shaded area on A and the thicker lines on B correspond to interlayer water content of smectite between 15 and 20% dry weight which are plausible values for the P-T conditions. The labels 37, 49 and 54 are the sample numbers (after France-Lanord &

Sheppard, 1992).

pore-waters, alteration of volcanic rocks to smectites will tend to enrich the pore-water in D unless the water/smectite ratios are large. Such increases in 8D of pore-waters (and breaks in the cation concentration profiles (France-Lanord et al., 1992)) have been observed, for example, at sites 800 and 801 of Leg 129 (France-Lanord & Sheppard, 1992), implying that pore-waters in the upper layers of the oceanic crust are not necessarily in diffusive communica t ion with seawater. At these sites, overlying chert and radiolarite units form efficient lithologic barriers to diffusion and the smectite-rich rock plus pore-water behaves approximately as a closed system.

The transformation of smectite to illite will tend to deplete the pore-waters in D, unless water-rock

ratios are large. In a closed system, H-isotope depletions of the pore-waters of up to 10 to 15%o could be produced during this transformation. A part of the ~SD variations of formation waters from a given sedimentary basins could be a result of such reactions, with meteoric water being not necessarily the only source of waters depleted in D relative to connate seawaters.

K I N E T I C S O F I S O T O P E E X C H A N G E

Quantitative isotopic exchange rate data on clay mineral-water systems can be derived from well designed and controlled experimental studies (see review and Table 1 by Cole & Ohmoto, 1986). The available experimental data on clays, however, are rather limited and are essentially non-existent for temperatures below 100~ because of the extremely slow exchange rates, even for hydrogen. In fact, few experiments have been specifically designed to derive information on the kinetics, i.e. the rates and mechanisms, of isotopic exchange in clay minerals.

O 'Nei l & Kharaka (1976) exchanged illite, kaolinite and montmorillonite that was <44 lam (stating that 30% of the kaolinite, 75% of the montmorillonite and 10% of the illite was <2 ~tm) with 0.04 N NaC1 solutions at temperatures between 350 and 100~ for times up to almost 2 yr. Unless the temperature was close to the stability limit of the mineral, the degree of exchange was between 5 and 80% for hydrogen and 0 and 19% for oxygen. The rates of isotopic exchange were in all cases much larger for hydrogen than for oxygen for a given time; O'Neil & Kharaka (1976) thus proposed a proton exchange mechanism. A similar relationship can be deduced from studies on micas at higher temperatures (Suzuoki & Epstein, 1976; Fortier & Giletti, 1991). In the latter investigations, volume diffusion is most probably the mechanism of isotopic exchange. Oxygen diffusion coefficients for transport perpendicular to the layers are three to four orders of magnitude smaller than parallel to the layers (Fortier & Giletti, 1991). Similarly, hydrogen transport occurs dominantly parallel to the layers (Graham, 1981). Calculated activation energies using an infinite cylinder model are of the order of 30 kcal/mol for hydrogen and 3 5 - 4 0 kcal/mol for oxygen. Cole & Ohmoto (1986) calculated diffusion constants from the experiments of O'Neil & Kharaka (1976) assuming that diffusion is the principal mechanism of exchange. The resulting activation energies ( 8 - 1 7 kcal/mol) are much

14

- 1 5

IogD

- 1 7

- 1 9

- 2 1

- 2 3


Temperature (~ 4 0 0 2 0 0 1 O0 25 0

\ \

- 2 5 , I n 2 3 4

1 O 0 0 / T (K)

Fro. 11. Arrhenius plot of hydrogen diffusivity for illite, based on the data of O'Neil & Kharaka (1976) and assuming a particle diameter of 2 gm. The 95%

confidence interval is shown as a dashed line.

smaller than those of micas and are similar to activation energies of surface-controlled reactions in sheet silicates, cation exchange or recrystallization. However, reliable quantitative kinetic data cannot be extracted from the experiments of O'Neil & Kharaka (1976). Grain-size distributions of the starting material were not well constrained and run products were not examined carefully for recrystallization. The exchange mechanisms in these experiments are, therefore, not known. Furthermore, extrapolations to ambient temperatures, using the constants in the Arrhenius relationship calculated by Cole & Ohmoto (1986), yield unreasonably high hydrogen diffusivities for clay m i n e r a l s ( 1 O g D k a o l i n i t e = - 1 8 . 5 8 cm2/s , logDin i te = - - 1 7 . 1 2 c m 2 / s a n d logDmonunorillonit e =

-16.59 cm2/s). These diffusion coefficients imply that 95% D/H exchange between a clay particle of 2 gm diameter and water at 25 ~ is reached after only 220, 20 and 6 yr for kaolinite, illite and montmor i l lon i te , respect ively . Changing the assumed average particle size of the clay also has a large influence on the calculated diffusion coefficients. For example, Cole & Ohmoto used a grain radius of 17 gm for the illite data of O'Neil & Kharaka (1976) and ca lcula ted a logD of -15 .84 cm2/s at 200~ whereas a radius of 2 I.tm would yield a value of only -17.10 cm2/s. The

uncertainties in the regression of the kinetic data in an Arrhenius plot are significant when extrapolating to low temperatures. For example , a 95% confidence interval on the fit of the illite data (Fig. 11), yields an error interval for the diffusion coeff icient at 0~ covering three orders of magnitude! High-temperature isotope exchange data cannot be extrapolated to ambient temperatures with any precision. In the case of montmorillonite, extrapolations of Cole & Ohmoto's regression data down to 25~ as presented by Kyser & Kerrich (1991), yield higher diffusivities for oxygen than for hydrogen. Thus, meaningful model calculations of kinetic data of isotope exchange in clay minerals must await better constrained or better designed experiments.

P R E S E R V A T I O N IN N A T U R A L S Y S T E M S

Do natural clay minerals retain their initial isotopic compositions? Evidence concerning the extent of isotopic exchange for natural systems is contra- dictory. Early studies (e.g. Savin & Epstein, 1970a,b; Sheppard et al., 1969; Lawrence & Taylor, 1971, 1972; Sheppard, 1977) addressed the problem and with a number of subsequent studies concluded that many clay minerals such as kaolinite, smectite and illite are often out of equilibrium with their present-day local waters. This is not to imply, however, that these clay minerals never undergo any post-formational or retrograde O- and/or H-isotope exchange.

Several workers have interpreted the H- and/or O-isotope compositions of their clays in terms of retrograde alteration with more recent waters without the intervention of special processes to provoke or facilitate exchange (Yeh & Savin, 1976; Wilson et al., 1987; Bird & Chivas, 1988; Longstaffe & Ayalon, 1990; Kotzer & Kyser, 1991). Others have invoked special processes such as a high radiation flux to catalyse exchange (Halter et al., 1987) or the presence of strongly D-depleted species in the fluid like CH4 (Bray et al., 1988).

Before examining some of these examples in more detail, we suggest that convincing evidence for complete O- and/or H-isotope exchange without recrystallization is usually lacking, unless the clay has subsequently been subjected to either higher temperatures or a special process. This implies that some information of past environments is usually recorded.

Stable isotope geochemistry of clays 15

~ 1 0 . . . . i . . . . I ' ' ' ' l l r ' ' l ' ' i / ' ' ' 1 . . . . . r . . . . I . . . . I . . . . I . . . .

I ~ 0

- - _ 4 o o 3 0 o 2 0 < _ - ..~- ~t~ -

- 4 0 ~ / . . . ! ~ _ ~ �9 e l t 1 0 ~

-60. o"~ / /~J~r 'KI 8D -80

- 1 O 0 ~ ~ ~ ~ ~

- 1 2 0 .

- 1 6 0 . . . . ' . . . . ' . . . . ' . . . . ' . . . . ~ , , r , , . . . . , . . . . J . . . . ) .... - 2 - l S - l O - 5 0 5 1 0 1 5 2 0 2 5 3 0

81~3

FIG. 12. 8D vs. 8a80 plot of the meteoric water and kaolinite lines. Kaolinites K1 and K2 are in equilibrium with waters W1 and W2. Trajectories A and A' are possible exchange paths between K1 and K2 and B between K2 and K1 (see text for discussion). Kaolinites in equilibrium with meteoric water W3 at

10, 20, 30 and 40~ are shown for reference.

Weathering environments

A thought exercise is considered to help understand the consequences of post-formational isotopic exchange in the weathering environment. In Fig. 12, kaolinites which are in equilibrium with meteoric waters plot along the kaolinite line (SD = 7.558180 - 219; modified using the fractionation factors proposed above) of Savin & Epstein (1970a). Kaolinites K1 and K2 are in equilibrium with waters W1 and W2, respectively. For K1, the composition of the water changes to W2, because, for example, the climate becomes cooler due to an increase in lati tude and/or altitude, giving a kaolinite that is out of equilibrium with its local m e t e o r i c waters . I f K1 u n d e r g o e s i so top ic exchange, a number of trajectories are possible depending on the relative rates of H- and O-isotope exchange: curve A if hydrogen re-equilibrates essentially completely before oxygen, or along the kaolinite line for equal rates of exchange or curve A' for an intermediate situation with hydrogen exchang ing faster than oxygen. Similar ly, a kaolinite K2 that migrates to an environment with meteoric waters of composition W1 will follow trajectories such as B or the line K2 to KI depend ing on the re la t ive rates of H- and O-isotope exchange. Based on structural arguments and the limited experimental data of O'Neil &

Kharaka (1976), differential isotopic exchange is probable with hydrogen exchanging more rapidly than oxygen in systems where phases are neither destroyed nor created. Oxygen isotope exchange may be more complicated in detail because of the two different structural sites within the framework. Thus, unless essentially complete H- and O-isotope exchange is the rule rather than the exception, considerable dispersion about the kaolinite line is to be expected and, for a given locality, a 'J ' (curve A) or inverted 'J ' (curve B) trend will be given. The effect of temperature on a kaol ini te in equilibrium with meteoric water (W3) is also given in Fig. 12.

Figure 13 presents the ava i lab le data on kaolinites from weathering environments which have not subsequently had a burial history. Many of these samples are not in equilibrium with present meteoric waters. Taking into account that some of the samples are not absolutely pure kaolinites, the temperature of exchange is not the same for all and some meteoric waters plot 1-2%o 8180 from the meteoric water line, the dispersion about the kaolinite line is considered to be very minor. This implies that differential isotopic exchange is unimportant for kaolinite, at least in the weathering environment.

Bird & Chivas (1988) have argued for major post-formational H-isotope exchange, without O- isotope exchange, for Permian kaolinites of weathering origin from the Gunnedah Basin, Australia. The kaolinites come from cores at 11 -1276 m depth and are all depleted in 8180 (6.1-9.4%o) and 8D ( - 1 0 7 to -94%0) relative to most kaolinites of weathering origin (Fig. 13). They are not included on Fig. 13 because they have had at least a mild burial history; they would plot to the left of the S/H line. Bird & Chivas (1988) estimated that burial depths did not exceed ~ 1 5 0 0 m but were ~ 5 0 0 - 7 0 0 m for some of the kaolinites and that the maximum temperature for these shallowest kaolinites was ~45~ The case for at least H-isotope exchange is well made. This implies that relatively mild heating is sufficient to reset the D/H ratios, if the estimates of the maximum t e m p e r a t u r e s are correc t . I f e s sen t i a l l y no O-isotope exchange occurred, the initial 8D values were ~ - 1 7 0 % 0 for kaolinite and -140%o for meteoric water. All other known kaolinites of weathering origin, or their associated meteoric waters, have 8D values at least 30%0 heavier (91 samples in Fig. 13). Such heavier values are more


8D

20

0

-20

-40

-60

- 80

-100

-120

-140

-160

. . . . ' . . . . . . . . . ' / . . . . . . . . . . . . . . . . . . ' . . . . . i / i SMOW / / - ~/

~ / . % / /

" --~ O �9

- 2 0 - 1 5 - 1 0 - 5 O 5 10 15 20 P5 30

8 1 8 0

weathering :

�9 Kaolinite + Halloysite

hydrothermal :

�9 Kaolinite [ ] Dickite A Nacrite

water was ~-70%0. Such a value is at least comparable with a few other values in the world (Fig. 13). From analyses of nearby coals, Smith e t

al. (1983) inferred that Permian meteoric waters had values of 6D = ~-120%o, supporting an intermediate value. Both interpretations raise a number of problems and they hinge critically on the minimum value for the temperature of burial.

Data for halloysites from weathering environments are plotted on Fig. 13. They also display a systematic variation along a line parallel to, but slightly to the left of, the kaolinite line. These data suggest that many halloysites have preserved their natural isotopic compositions and have undergone neither major partial nor complete O- and/or H-isotope exchange with the local laboratory atmosphere. Two implications of these systematics for halloysite are (1) the H- and/or O-isotope fractionation factors for halloysite-water are slightly smaller than those for kaolinite-water, assuming that their temperatures of formation are similar to those for kaolinite, and (2) palaeoclimatic information can be derived from such analyses.

O c e a n i c e n v i r o n m e n t s

F~G. 13. A compilation plot of 8D vs. 8180 for kaolinites (k), dickites (d), nacrites (n) and halloysites (h) from weathering and hydrothermal environments. There are 129 samples with the following distribution: (1) soil and weathering zones: k = 71; h = 13; (2) supergene oxidation zones over ore deposits: k = 20; h = 10; (3) hydrothermal systems: h = 46; d = 9; n = 3. The meteoric water, kaolinite weathering and supergene/hypogene lines are given for reference. Data from Sheppard et al. (1969); Savin & Epstein (1970a); Lawrence & Taylor (1971, 1972); Hall et al. (1974); Sheppard & Taylor (1974); Sheppard & Gustafson (1976); Lombardi & Sheppard (1977); Sheppard (1977); Jackson et al. (1982); Marumo et

al. (1982); Chivas et al. (1984); Hassanipak & Eslinger (1985); Bernard (1987); Bird & Chivas (1989); Marumo (1989); Arehart et al. (1992); Rye et al. (1992); Pruett & Murray (1993); Vennemann et al.

(1993); Hedenquist et al. (1994).

reasonable because the total amount of precipitation tends to increase with increase in 8D value and the formation of kaolinite requires relatively high rainfall. If both the D/H and JSo/~60 ratios re-equilibrated with meteoric water during burial, requiring T > ~ 75~ the 8D value of the meteoric

Detrital clays in oceanic sediments usually retain a continental-type H- and O-isotope signature (Savin & Epstein, 1970b; Yeh & Savin, 1976; Yeh & Epstein, 1978; Eslinger & Yeh, 1981; Yeh & Eslinger, 1986). Different size-fractions were analysed in a number of cases. Yeh & Savin (1976) showed that the <0.1 lam size-fractions of illite- smectite-rich oceanic sediments become more 180-rich if they are older than several 104 yr. Although this can be interpreted in terms of exchange with sea water (Yeh & Savin, 1976), the growth of authigenic smectite could also account for the l go trend (O 'Nei l , 1987). Similarly, Yeh & Epstein (1978) suggested that the only detectable H-isotope exchange (8--28%) in a 2 - 3 Ma old smec t i t e - i l l i t e - r i ch deep sea sediment was in the minor ( ~ 5 % ) <0.1 lam fraction. Additional studies on Gulf of Mexico and DSDP clays including smectites by Yeh & Eslinger (1986) have found no evidence for O-isotope exchange for periods of several 106 years. There is thus no compelling evidence for significant H- or O-isotope exchange in the absence of diagenetic processes, implying that isotopic studies provide a powerful tool for distinguishing detrital from authigenic clays.

Stable isotope geochemistry of clays /17

Diagenet ic environments

Convincing evidence for H-isotope exchange without major O-isotope exchange has been reported by Longstaffe & Ayalon (1990) for a few kaolinites from the Cretaceous sandstones (Viking Formation) in a sedimentary basin in Alberta, Canada. The very early diagenetic kaolinite, which formed at 10-30~ about 100 Ma, exchanged its H-isotopes with later (40-1-10 Ma) to present formation waters at 120-40~ and without important O-isotope exchange. However, these early kaolinites have had a complex history (see Fig. 2 in Longstaffe & Ayalon, 1990) and have been heated to 120~ during the thermal evolution of the basin. Equally, Longstaffe & Ayalon (1990) suggested that the late diagenetic kaolinite, illite and illite-smectite could also have re-equilibrated their D/H ratios with present formation waters. This interpretation is debatable. They showed that these minerals formed after the maximum of temperature (~120~ and burial about 50-60 Ma ago, with illite-smectite forming at T ~50-1-10~ and later than the second generation of kaolinite. Although the H-isotope compositions of these clays could be in equilibrium with present formation waters, no data were presented on how the D/H ratio of the formation waters evolved over the past 50 Ma. If the ~iD values of the waters were not very variable, a realistic possibility, the small range of 8D values but relatively large range of 8180 values of the clays may just be reflecting formation over a range of temperatures because the H-isotope fractionations factors, unlike those for oxygen, are not very sensitive to temperature in this range (Figs. 3 and 6).

Hydrothermal environments

Many kaolinites, dickites, nacrites and smectites from hydrothermal systems plot to the left of the supergene-hypogene line, S/H, (equivalent to kaolinite in equilibrium with meteoric waters at temperatures of about 35~ shown on Fig. 13. They are consistent with the preservation of a hydrothermal signature. A thought exercise for hydrothermal kaolinites in equilibrium with meteoric waters which undergo exchange at surface temperatures can be developed like that above. Although some exchange cannot be completely discounted, major differential exchange is not supported by the data. These clays are not in hydrogen isotope equilibrium with recent meteoric

waters. Based on these observations, Sheppard et al. (1969) proposed that the combined H- and O-isotope compositions of clays can be used to discriminate hypogene from supergene clays (Fig. 13). This proposition is still considered to be generally viable.

Radiat ion- induced retrograde exchange?

Chlorite, illite and kaolinite occur in both barren and uranium mineralized rocks within the Proterozoic Athabasca Basin, Saskatchewan, Canada. These unconformity-related deposits have very high grades (e.g. the main ore body at Cedar Lake averages 14% U308 (Bray et al., 1988)). Some deposits are within a few metres of the surface whilst others are at depths >100 m. Illites associated with mineralization from all six deposits studied are both strongly depleted in D ( -90 > 8D > -170%o) and have young apparent K/Ar ages (<750 Ma) relative to the surrounding unminer- alized rocks (-48 > 8D > --62%o; 1265 Ma) (Halter et al., 1987, 1988; Wilson et al., 1987; Bray et al., 1988; Kotzer & Kyser, 1991). Chlorites and kaolinites from the ore zones are also systematically D-depleted. Additionally, the H20 + content of illite from mineralized samples is higher than the theoretical value as well as that of illite from the barren zones. Halter et al.(1987) and Bray et a/.(1988) also observed that both 8D and K/Ar 'ages' decrease systematically with increase in U content (Fig. 14). A continuous variation of the 8D values from the high group associated with the ancient barren regoliths to the low values intimately associated with U mineralization is not observed; the geologically subdivided groups have their own characteristic 8D values.

Halter et al. (1987) interpreted these relations in terms of radiation-catalysed retrograde H-isotope and K-Ar exchange of the clays at low temperatures with post-Cretaceous meteoric waters because they, like Bray et al. (1988), observed such a good correlation with the U content. Radiolysis phenomena have been documented in the fluid inclusions from these deposits by Dubessy et al. (1988). Wilson et al. (1987) and Kotzer & Kyser (1991) related the decrease in ~iD and K/Ar 'ages' to simple low-temperature exchange with recent meteoric waters that entered along reactivated shear zones, with a few modified illites coming from U-free samples near mineralization. They proposed that the extent of retrograde exchange is related to


" -50

�9 - I 0 0

-1 SO

<

o ~ v~o 0 0 0

o o $

barren halo ~e kaolinite "~ illite 0 0 �9 �9 chlorite 0

-200 ........ ' ........ ' ........ ' ...... 1500" 10 100 1000 10000

. . . . . . . . i . . . . . . . . , . . . . . . . . r . . . . . . . .

0 0 0 0 0

1 0 0 0 0

4~

S O 0

0 . . . . . . . . i . . . . . . . . i . . . . . . . . i . . . . . .

1 0 1 0 0 1 0 0 0 1 0 0 0 0

Uwr (ppm)

FIG. 14. Variation of the uranium content of the whole rock (wr) with (A) the 8D value of the mineral, and (B) the K/Ar age of the chlorite or illite (after Halter et

al., 1987).

the amount of reactivation and late fracturing. Bray et al. (1988) argued against simple retrograde exchange with later meteoric waters in the ore zone and proposed that the low 8D values are a resul t of se lec t ive modi f ica t ion by s t rongly D-depleted CH4 -F HzS • H2 fluids which had p r e v i o u s l y i n t e r a c t e d w i t h the g r a p h i t i c metasediments.

Several arguments do not support the simple but extremely selective retrograde alteration hypothesis of Kyser and his co-workers. Although they do not observe a systematic relation between U content and 8D value, all of their samples with >1000 ppm U have low 8D values and those samples with a really low 6D but little or no U come from within a metre of ore. Since many of their samples come from near-surface deposits, in contrast to the relatively deep Cluff Lake deposit studied by Halter et al. (1987), U may have been leached from some of their samples by oxidizing meteoric waters. Their model is thus not considered to

explain why chlorites, illites and kaolinites from above and around the ore zones and their alteration halos have not been isotopically modified. In fact, the isotopic compositions of these clay minerals are consistent with exchange with the low latitude Proterozoic formation waters during the diagenetic evolution of the basin before mineralization at temperatures of 150-220~ and their subsequent preservation (Pagel, 1975; Halter et al., 1988).

The high H20 + illites, which are invariably D-depleted, are also restricted to the mineralized zones. Although Kotzer & Kyser (1991) showed that these illites display no conspicuous effects of retrograde alteration at the scale of the scanning electron microscope and proposed that the water is present as interlayer water, the few recent studies on illites with above normal H20 + contents have shown that the extra water (or hydrogen) is present as either extremely fine intergrowths of chlorite, kaolinite and other high water content minerals or organic matter and not as interlayer water (Peacor, 1993; Jiang et aI., 1994; Gilg, Sheppard & Kahr, unpublished data). The H-isotope depletion is thus related to either chemical and/or mineralogical changes or contamination.

C O N C L U S I O N S

A new set of equilibrium H- and/or O-isotope fractionation factors between kaolinite, smectite, illite and water are derived by re-examining critically the experimental, empirical and theoretical values (Table 3). The level of confidence is considered to be high for temperatures around 300~ The fractionation factors at surface and in te rmedia te t empera tu res are domina t ed by empirical results, where available. Although the uncertainties in these values are difficult to assess, criteria are presented for selecting or rejecting data. Taking either the hydrogen or oxygen results as a whole, a general coherence in both the magnitude and temperature dependence of the fractionation is observed. The H-isotope fractionations between bulk kaolinite and water, and probably for illite and smecti te too, vary monotonical ly in the temperature range 350 -0~ and are rather insensitive to temperature variations. The more complex fractionation curves that have been proposed, for example the kaolinite expression of Liu & Epstein (1984), are considered to be influenced by surface and/or inter-site fractionation effects. The importance of such effects has probably been under-


TABLE 3. Estimated H- and O-isotope fractionations between mineral and water.

19

Element Mineral Equation: 1000 In ~lnineral-water

H K a o l i n i t e - 2 . 2 • 106 • T -2 - - 7.7 O Kaolinite 2.76 • 106 • T -2 - 6.75 O Smectite 2.55 X 106 • T - 2 - 4.05 O Illite 2.39 x 106 • T -2 - 3.76

estimated in the past. All experimental partial exchange studies with only low to moderate degrees of exchange must be closely examined for such fractionation effects. Estimation of equilibrium fractionations from partial exchange experiments is complicated for phyllosilicate minerals because both hydrogen and oxygen are distributed between two or more structurally different sites with different exchange and isotopic properties. These remarks may also apply to the serpentine-water and chlorite-water systems and could equally account for some of the differences observed between experimentally and empirically derived fractionations. As shown for smectites, H-isotope fractionations are quite sensitive to chemical variations with the smectite-water fractionation increasing by ~ 70%o from Fe-poor montmorillonite to nontronite. Although some of the revisions to the O-isotope fractionation expressions are small, the new equation for kaolinite-water, which is based on a much wider data set, gives temperatures that are ~ 2 5 - 4 0 ~ lower in the temperature range 100-150~ than many expressions used previously (Table 2).

The preliminary calibration of the fractionation of oxygen between tetrahedral and hydroxyl sites of the same phyllosilicate (Hamza & Epstein, 1980; Bechtel & Hoernes, 1990), indicates that such single-mineral thermometers are rather temperature sensitive. More rigorous calibration and testing on natural samples is required to develop these promising thermometers.

For smectites, the estimated pore-water/interlayer water fractionation at oceanic to surface temperatures is about 20-t-10%o. Since interlayer water can represent 15-50 mol % of the total rock (minerals + pore-water) hydrogen of a smectite-rich sediment, interlayer water can be a significant reservoir of D-depleted hydrogen if the system behaves as a closed system.

The meaningful isotopic analysis of clays requires a number of pretreatment procedures to remove interfering constituents (Table 1). Organic matter is often intimately associated with certain clays. As it contains hydrogen, neglecting its presence or its incomplete removal can modify the liD value of the clay analysis and its hydrogen content, which is conventionally presented as H20 +, may thus be higher than the theoretical upper limit.

Clay minerals - - kaolinite, dickite, nacrite, illite, smectite as well as halloysite - - tend to retain both their original H- and O-isotope compositions unless they have been subjected to more extreme post- formational diagenetic or metamorphic conditions. All examples presented as evidence for complete or partial isotopic exchange in the complete absence of the intervention of higher temperature conditions or special processes such as an exceptionally high radiation field lack conviction. The state of the data is such that only qualitative statements can be made about the exchange characteristics of clays with comparable grain sizes. In the absence of recrystallization processes and for comparable grain sizes, hydrogen exchange is more rapid than oxygen exchange. The uncertainties in the regression of the limited experimentally derived kinetic data in an Arrhenius plot are significant when extrapolated to surface temperatures and prevent useful application of diffusion data to natural systems. However, there is an urgent necessity to develop criteria to recognize the importance of exchange processes. The isotopic analysis of different size-fractions of the same clay is important but not sufficient by itself because variations in the liD and/or 8180 values could be due to (1) differential exchange, (2) the presence of varying mixtures of detrital and authigenic or neoformed clays, or (3) evolution of the distribution of grain sizes (and hence ages or isotopic compositions) by a process such as Ostwald ripening. Other


techniques and approaches are needed to assess the importance of these different possibilities.

Both H- and O-isotope studies of clay minerals can provide useful information about the conditions of formation or subsequent thermal evolution of their environment. Further advances in the stable isotope geochemistry of clay minerals will require: (1) the careful application of multiple techniques such as SEM, TEM, other isotopic systems (K/At, Rb/Sr, Sm/Nd) and particle distribution studies on natural systems; and (2) well conceived experimental studies, with careful characterization of materials before and after the experiment, that are designed to yield equilibrium fractionation factors, and/or kinetic information (rates and mechanisms).

ACKNOWLEDGMENTS

Andrew Parker is thanked for encouraging the authors to look at the stable isotope geochemistry of clays more closely. We are grateful to our many co-authors and colleagues for discussions over the years. HAG is grateful for a Human Capital and Mobility Fellowship of the EU. Tony Fallick is thanked for his careful review and correction of the "franglais'.

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