American Mineralogist, Volume 72, pages 566-579, 1987

Experimental study of hydrogen-isotope exchange betweenaluminous chlorite and water and of hydrogen

diffusion in chlorite

Cor,rN M. Gnq,HA.IvIDepartment of Geology, Edinburgh University, Edinburgh, EH9 3JW, Scotland

J.lNnr A. Vrcr,rNo, Russrr-r. S. HlnwroxDepartment of Geological Sciences, Southern Methodist University, Dallas, Texas 75275,U.5.4.

Ansrucr

Hydrogen-isotope exchange between Mg-Al chlorite of sheridanite composition andwater has been studied experimentally over the temperature range 300-700'C. Problemswere encountered with changes in the bulk D/H ratios of capsule contents during high P-Texperiments. Below 500t, exchange was so slow that equilibrium fractionation factorscould not be determined. Over the temperature range 500-700'C, no temperature depen-dence is observed, with 1000 ln a;n,-"ro around -28. Chlorite-water fractionations are notsimply related to octahedral cation composition; preferential fractionation of the lightisotope into chlorite reflects the presence of hydrogen bonding in the chlorite structure.

A compilation of experimentally and empirically determined chlorite-water hydrogen-isotope fractionation factors over a wide range of temperatures indicates that these are notmeasurably dependent on Fe/Mg in chlorite.

Hydrogen-isotope exchange reactions proceeded by volume diffusion of hydrogen inchlorite, for which activation energies in the range 166-172 kJ/mol H are much higherthan in any other hydrous mineral yet studied. Hydrogen-isotope exchange between chlo-rite and water in slowly cooling metamorphic rocks may nonetheless proceed tosubgreenschist facies temperatures.

Hydrogen-isotope exchange during hydrous metamorphism of the oceanic crust in thetemperature range 200-500"C may be dominated by fractionation factors in the range1000 ln a"D,,r1*",",: -30 to -40.

INrnonuctIoN

Studies ofhydrogen- and oxygen-isotope variations inrocks and minerals provide a powerful method of docu-menting fluid-rock interactions in the Earth's crust (Tay-lor, 1974; Sheppard, 1977). The hydrogen-isotope com-positions of hydrous minerals are particularly sensitiveindicators of the isotope composition of the last fluid withwhich they have equilibrated, because (1) most hydrousminerals contain relatively small amounts of hydrogenrelative to infiltrating hydrous fluids unless fluid/rock ra-tios were very small, (2) waters of different origins in theEarth's crust and hydrosphere have distinctive hydrogen-isotope composition, and (3) hydrogen-isotope exchangebetween hydrous minerals and water should be relativelyfast during mineral-water interactions and crystallizationof hydrous minerals. In order to calculate the isotopecomposition of a fluid phase that has subsequently beenremoved from a water-rock system, it is necessary to knowthe isotope fractionation factors between hydrous min-erals and water as a function of temperature and mineralcomposition.

0003-o04x/87l0506--05 66$02.00

In this paper, we report the results and applications ofan experimental study of hydrogen-isotope fractionationbetween aluminous chlorite and water and of rates forhydrogen diffirsion in chlorite, at high pressures (2-5 kbar)and over a tange oftemperatures (300-700oC) represen-tative of a spectrum of metamorphic conditions and somehydrothermal environments. Aluminous chlorites arecommon and abundant constituents of a wide range ofcrustal rock types crystallizing in these hydrous environ-ments. Chlorites, with about 12 wto/o water, contain themajority of structural water present in many low-grademetamorphic and hydrothermal rocks, most importantlyin hydrated mafic oceanic crust. Therefore the results ofthe study should provide new data applicable to the de-termination of the stable isotope compositions and originsof hydrous fluids in a variety of crustal geologic environ-ments.

Preliminary results of this experimental study have beenreported in abstract form by Viglino et al. (1984) andGraham et al. (1984b). No other experimental data areavailable on hydrogen-isotope exchange between chlorite

566

GRAHAM ET AL.: CHLORITE.WATER

TneLe 1. Chemical analyses of chlorite 9451S

1 2

HYDROGEN-ISOTOPE EXCHANGE

TneLe 2. lsotope compositions of starting ma-terials

6Dt(Va) 1F(a' - 1)

567

0.3432.290.01

5.515.570.01

0.058.90

15.31

Note. Columns are (1) Jones (1981). X-ray fluores-cence and colorimetry analysis. Water content deter-mined by weight loss to 100eC during DrA experi-ments. (2) Shannon and Wherry (1922). Wet-chemicalanalysis.

t HrO content of starting material, determined by in-duction heating in this study, is 12.95%.

and water, although data on hydrogen-isotope exchangebetween serpentine and water (Suzuoki and Epstein, 1 976;Sakai and Tsutsumi, 1978) and brucite and water (Satakeand Matsuo, 1984) have been reported. The serpentineand brucite data are relevant to this study because bothserpentine and chlorite are sheet silicates, with bruciteIayers incorporated in the sheet structure, and serpentineis chemically similar to Al-free chlorite.

Empirical estimates of hydrogen-isotope fractionationfactors for the chlorite-water system as a function of tem-perature and Fe/Mg ratio in chlorite have been madeusing isotope data for chlorites from pelitic schists (Tay-lor, 1974; Kuroda et al., 1976), hydrothermal ore depos-its (Kuroda et al., 1976) and drill-core samples from arecent geothermal field (Marumo et al., 1980). These em-pirical calibrations are usually based on assumptions re-garding the hydrogen-isotope composition of the waterwith which the minerals have equilibrated, the tempera-ture of crystallization of the chlorites, and the absence ofretrograde re-equilibration of the chlorite with hydrousfluid. These assumptions may be poorly founded, espe-cially because many hydrous minerals will readity ex-change hydrogen with water to temperatures well belowthose of primary crystallization during slow cooling (Gra-ham, l98l). Therefore, empirical calibrations of hydro-gen-isotope fractionation factors not derived from directexperimental calibration must be treated with caution.

Experimental measurement of rates of hydrogen-iso-tope exchange between chlorite and water in this studypermits calculation of rates of hydrogen diffusion in chlo-rite. These diffi.rsion rates may be compared with rates of

Note.'ai is the initial (uncorrected) chlorite-water frac.tionation factor.

hydrogen diffusion in other hydrous minerals (Graham,1981; Graham etal.,1984a) and may be used to calculateclosure temperatures for cessation of hydrogen-isotopeexchange between chlorite and water in slow-cooling geo-logic environments.

SranrrNc MATERTALS

The chlorite used in this study is a sheridanite (specimen num-ber 9451 5, Smithsonian Institution Museum of Natural History,Washington, D.C.) whose composition is shown in Table l. Thischlorite is comparably aluminous to typical chlorites in low-grade aluminous metamorphic and hydrothermal rocks (Veldeand Rumble, 1977). Optical, xro, and serr.r examination showedthe sample to be free of contamination.

The specimen was prepared by grinding and sieving to obtainsized starting materials with nominal grain-size fractions of <44pm (fine), 44-63 pm (medium), and 63-105 pm (coarse). Thefinest material was removed from the fine grain-size fraction byelutriation.

The isotope compositions of the chlorite and the waters usedas starting materials in this study are given in Table 2. Theisotope compositions of the waters were chosen so that hydro-gen-isotope equilibrium between chlorite and water would beapproached by exchange from two directions during experiments(cf. Suzuoki and Epstein, 1976; Graham et a1., 1980). All min-eral-water charges were sealed into separate capsules as part ofthe same batch, in order to circumvent problems arising fromvariation of starting water composition with time and to ensureinternal consistency ofdata (cf. Graham et a1., 1980).

ExprnrnrnNTAr, AND ANALyTTCAL METHoDS

Experimental and analytical methods are comparable to thosedescribed by Graham et d. (1980, 1984a). Exchange mns weremade in Au capsules of 5 mm O.D. x 0.5 mm wall x 20 mm(approx.) length. Each run typically contained 125-mg chloriteand 16-mg water, capsules and mineral powders having beenstored and dried at 110'C prior to loading in order to minimizeadsorption of atmospheric moisture. Runs were made in Ni-monic-105 cold-seal pressure vessels (2-kbar runs) and internallyheated gas-pressure vessels (5-kbar runs) using an Ar pressuremedium. After quenching, capsules were weighed to check forleakage, heated at l20t overnight under vacuum to removeadsorbed moisture, and then pierced under vacuum to releasefree water that was collected by freezing into a liquid-N, coldtrap while heating capsules at 120'C for up to 4 h to ensureremoval of all free water. Hydrogen gas formed by passage ofthis water over hot IJ was measured manometrically to furtherensure complete removal and collection of free water.

sio,Al,03Tio,Fe.O"FeOMgoMnOKrONaroHrO.

Total

28.8126.43

0.240.40

31 21

0.350.14

13.89101.47

s.235.66

0.030.068.45

0.080.05

16.84

-33.56.43.4

-26.5-50.8

-39.6-36.8-7.218.2

29.7925.540.06

12.40100 43

Chlorite 94515Water 1Water 2Water 3Water 4

SiAITiFe3*Fe2*MgMnKNaOH

( ' N : t O O c ? N F t SN N O N N O * $ N O

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d'6'vv er9d'd'5 vd'=6'tr6'd' 6'd d'6' i(F N F N O o o O b O @ @ O O @ t s o a 6 O $

F : c d d + c i c i o i R i d @ d c ' j r c i 6 i ; d ! a l d 6 i t r i( O F O I t r N S I t O F S F o F $ F -r r t l t t l

q u ? u ? q q q r q \ \ n u ? \ o ? r a o ? r q \ aN O F N d ) O r ( o 6 @ O N ( 9 N | o n F F < t ( ) r f , $

+ + + + + 1 r + I l + I I r I + I | | |

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r + o N N o o g o ( 9 9 o

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o o o o o o o o o o o o o o o o o 9 0 0 0o o o o o o o o o o o o o o o o a a o o oN t s N N @ @ @ @ @ S r ) o b b o r o \ t * c ) t r ) r )

> > O O | ! L > > O L l ! r ! > > O O u - ( ) L u r

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t

Chlorite run products were washed with distilled water andexamined optically and by xno and snv to check for absence ofmineral alteration and thermal decomposition. For isotope anal-ysis, 20-30 mg ofchlorite were loaded into previously outgassedMo (at Southern Methodist University) or Pt (at Scottish Uni-versities Research and Reactor Centre) crucibles heated undervacuum at 120'C ovemight to remove adsorbed moisture. Struc-tural water was released by induction heating for about I h, andhydrogen yields after conversion by passage over hot U weremeasured manometrically to check water contents of chlorite forconsistency with initial values and to ensure complete removalof structural water.

D/H ratios in waters and chlorites were measured on a Fin-negan MAr 251 mass spectrometer (at Southern Methodist Uni-versity) and a VG-Micromass 602B mass spectrometer (at Scot-tish Universities Research and Reactor Centre). Values for therelative concentration are given relative to standard mean oceanwater (SMOW) in the familiar 6 notation (in 7o) in Tables 2 and3 to a precision of about + l7o where

/ t ^ ' - - ' \

dD*-on: (# : f . . t . - l ) x lo , ( l )1tD/H)*o'o"o

'I

and the equilibrium fractionation factor, a", is related to 6D by

569

ing uncertainty of interpolation of a" at low extents ofisotope exchange has been noted by Graham et al. (l 980).Uncertainties are discussed in detail below.

Graham et al. (1980) suggested that the interpolationmethod may not necessarily be applicable to mineralssuch as chlorite and serpentine that contain hydrogen inmore than one site, depending upon the rate-determiningmechanism in the exchange reaction. This method wasused by Sakai and Tsutsumi (1978) in their experimentalstudy of hydrogen-isotope exchange between serpentineand water, but none of their experiments attained iso-topic equilibrium so that independent verification thatthe interpolated values of a!",o-1rr6 are correct, or that theinterpolation method is valid, is not possible. In this studywe find good agreement between fractionation factors cal-culated from partial-exchange experiments at 600"C us-ing the interpolation method and determined by isotopeequilibrium in longer runs at the same temperature (Ta-ble 3). This observation is consistent with-but not con-clusive proof of-the validity of the interpolation methodfor the chlorites, and by inference possibly also for theserpentlnes.

Applicability of the interpolation method to the chlo-rites with hydrogen in different structural sites may in-dicate that the rate-determining step in the process ofisotope exchange is related to the transport of hydrogenthrough the mineral structure (e.g., Graham, 1981), rath-er than to the breaking and reforming of O-H bonds.

Rnsur,rsResults of valid experiments on hydrogen-isotope ex-

change between chlorite and water are given in Table 3,together with final fractionation factors (ar) and interpo-lated equilibrium fractionation factors (a"). Discussion ofthese results necessitates (l) a consideration ofthe exper-imental problems of hydrogen diffusion through capsulewalls and hydrogen-isotope exchange between capsulecontents and the external experimental environment and(2) a consideration ofthe uncertainties arising from thissource and from analytical uncertainties, in order that thevalidity of any experiment may be assessed and realisticconstraints placed upon the uncertainty in a".

Hydrogen-isotope exchange between capsules andpressure medium

Hydrogen diffusion through capsule walls and D-H ex-change between capsules and pressure medium during highP-Zgas-media experiments may cause significant shiftsin the bulk dD of experimental charges. These changes inDD in experiments where complete exchange was not at-tained may seriously limit the precision with which equi-Iibrium fractionation factors may be constrained by theinterpolation method, and necessitate a consideration ofthe mechanisms and relative rates of hydrogen-isotopeexchange between mineral and water and between themineral-water system and the external experimental en-vironment (pressure medium, pressure vessel, furnace,etc.).

GRAHAM ET AL. : CHLORITE-WATER HYDROGEN-ISOTOPE EXCHANGE

_ (D/H).hr.d,"o6nr_u2o:

(D/Ht*

I + (dD"hr"d,"/1000)

I + (dD_.,",/1000)

INrEnpolarroN oF de

Complete isotope exchange at any temperature is dem-onstrated by attainment of the same final fractionationfactor (ar) either in experiments run at the same temper-ature with starting waters of different isotope composi-tion, or in experiments run with waters of the same iso-tope composition run for different times (a, + a,, wherea, is the initial mineral-water isotope fractionation fac-tor). For most experiments, hydrogen-isotope exchangebetween chlorite and water was so slow that equilibriumwas not attained, and the interpolation method of North-rop and Clayton (1966), modified by Suzuoki and Epstein(1976), was used to calculate a" from the relationship

(o, - l) : (a" - l) - A(a, - a1), (3)

(where a, and ai are the final and initial fractionationfactors andAis a constant related to the extent ofisotopeexchange) for two or more experiments of the same du-ration run with different starting waters. In practice, armay be obtaired graphically, algebraically, or by least-squares analysis from the intercept 10,(d. - l), in plotsof 103(a, - l) vs. 103(a, * a,), where I is the slope of theline fitted through the data for any group of experimentswith different starting waters (a,) at constant Z and time.The theoretical kinetic basis of this approach is wellfounded (Northrop and Clayton, 1966), and the internalconsistency of experimental hydrogen-isotope exchangedata observed using this partial-exchange interpolationmethod has been noted in several studies (Suzuoki andEpstein, 1976; Graham et al., 1980, 1984a). The increas-

(2)

570 GRAHAM ET AL.: CHLORITE-WATER HYDROGEN-ISOTOPE EXCHANGE

1 0 + 1 0 +20 +30-20A (mass balance value)

Fig. l. Plot ofmass-balance value (A) against run time (hours) to illustrate shifts in bulk 6D ofchlorite-water hydrogen-isotope

exchange experiments. A values calculated using Eq. 5. Filled circles indicate runs listed in Table 3. Open circles indicate runs

rejected on basis ofanomalous hydrogen yields or inconsistent kinetics. See discussion in text.

a a a

o

a a

a a

a o a

a a

o a aa

l 0 4

1 0 3

1 0 2

oE

- 1 0

Graham et al. (1980) suggested that changes in dD ofexperimental charges could be minimized at tempera-tures below about 650"C by use of thick-walled (0.5 mm)Au capsules. However, the results of hydrogen-isotopeexchange experiments between amphibole and water(Graham et al., 1984a) indicated that Fe-rich amphibolecharges underwent significant shifts in 6D at temperaturesless than 650"C owing to oxy-hornblende reactions driv-en by the buffering effect of bomb walls on the charges,while Fe-poor charges showed little or no significantchange in DD in experiments up to 850"C. Also, experi-ments in internally heated gas vessels were found to bemuch less susceptible to change in bulk 6D than those inNimonic cold-seal.pressure vessels (Graham et al., 1984a,Table 3). In all these studies (see also Graham et a1., 1980,App. I), changes in bulk 6D of charges were always in thedirection of D enrichment.

Mass-balance calculations for isotope exchange exper-iments in this study were calculated from the relationship

xDD;n, + /DDhro : xDDln, * ydDfrro, (4)

where x is the number of moles of H in chlorite and y isthe number of moles of H in the water and i and f referto hydrogen-isotope compositions before and after ex-periments, respectively. From Equation 4, mass-balancevalues were calculated as

a: (xdDLr + /6Dl,o) - (xDD;n, + y6D;,"). (5)

A values for all experiments are plotted against run timein Figure L Clearly, numerous runs suffered significantchanges in 6D down to temperatures as low as 400"C,despite the absence of significant Fe in the chlorite anddespite the use of thick-walled Au capsules, which clearly

did not inhibit D-H exchange between charge and pres-

sure medium. In addition, unlike previous experimentalhydrogen-isotope studies in the high-pressure laborato-ries at Edinburgh, significant changes in bulk DD were inthe direction of both D enrichment and D depletion inexperiments in both cold-seal and internally heated pres-

sure vessels (Fig. l). In addition to hydrogen diffirsionthrough capsule walls during high P-Texperiments, shiftsin bulk 6D of charges may result from gain of atmosphericmoisture or loss of water during loading and welding ofcapsules, analytical error, and incomplete removal of freewater from quenched capsules or ofstructural water fromchlorite. These several alternative sources of error mayoften be identified by yields ofhydrogen that are incon-sistent with starting water or chlorite contents' Indeed,some runs showing large DD shifts in Figure 1 gave anom-alous hydrogen yields. However, numerous runs grvrng

satisfactory hydrogen yields showed shifts in DD that ex-ceed reasonable isotopic analytical error.

The related problems of controlling the 6D of chargesduring high P-Z hydrogen-isotope exchange experimentsand of interpreting and understanding the mechanismsand processes causing DD shifts have been discussed byGraham et al. (1980, 1984a) and recently by Richet et al'(1986). Whatever the sources of hydrogen in the Ar pres-

sure medium in these studies (bomb walls, filler rods,moisture from furnace cement or atmospheric sources,impurities in bottle Ar, etc.), tentative explanations todate have focused on the existence oflarge/", gradients

across capsule walls (Graham et al., 1980; Richet et al.,l936). A detailed discussion ofthese processes and oftheconclusions ofRichet et al. (1986) will be presented else-where. However, here we make the following observa-

GRAHAM ET AL.: CHLORITE-WATER HYDROGEN-ISOTOPE EXCHANGE 571

tions and conclusions regarding DD shifts from experi-mental data of this and previous studies (cited above):(l) The occurrence of both D enrichment and D depletionduring chlorite-water exchange experiments in both in-ternally heated and cold-seal pressure vessels (Table 3)negates the suggestion of Richet et al. (1986) that thepressure medium in the internally heated pressure vesselsin their D-H exchange on melt-vapor systems was muchmore reducing than in our exchange experiments. (2) Wenote that A in Figure I shows no time dependence. (3)We note that the data on diffusivity of protium and deu-terium through Pt (Ebisuzaki et al., 1968), if applicableto other metals of comparable structure such as Au, wouldcause the rapid elimination of initial, transient/", gradi-ents between capsules and pressure media well within thetime scale of our shortest exchange experiments. This lat-ter conclusion is quite consistent with the observed ab-sence of a time dependence for A in our runs. Thereforewe believe that our chlorite + HrO charges approach asituation of hydrogen-isotope exchange equilibrium withhydrogen-bearing contaminants of widely variable D/Hin the pressure medium well within the time scale of ourshortest experiments. Apparent absence of DD shifts inany experiment may be due to coincidental similarity of6D in charge and pressure medium.

The rate of this capsule-pressure medium exchangeprocess is much faster than the rate of exchange of hy-drogen isotopes between chlorite and water in this study,as is evident from an analysis ofthe kinetics ofthe latterexchange (Table 3; see discussion below). Therefore, afirst-order correction to the experimental data for the ef-fects of changing bulk DD during high P-T experimentsmay be made by assuming that the entire shift in 6D,calculated as A (the mass-balance values derived from Eq.5 and listed in Table 3), may be attributed to a shift in6Di{ro (the DD of the starting water) before significant ex-change between chlorite and water has been achieved.Accordingly, in experiments in which 1000/o exchange be-tween chlorite and water has not been attained, modifiedvalues of 6D;r" (DD';r" : DDi,ro - A) have been used tocalculate the effective initial chlorite-water fractionationfactor (oi) used in interpolating the equilibrium fraction-ation factors (Table 3). It is interesting to note that theequilibrium factors thus interpolated are, with one excep-tion (runs 26 and 27), within 5-6Tm of the interpolatedvalues of l03ln a" calculated using uncorrected values ofa, from Table 2.

Partial-exchange experiments listed in Table 3 and usedto interpolate equilibrium fractionation factors have mass-balance values (A) up to over l0 (positive or negative).Many experiments with A values greater than + I l, anda few with smaller A values, gave anomalous hydrogenyields for water or chlorite and were rejected. Beyondthis, internal consistency of results is taken as a primarycriterion of data validity; kinetic analysis of the experi-mental data in Table 3 provides a useful approach todemonstrating such internal consistency (Graham, 1981)and is considered below (Fig. 5).

Consideration of errors and uncertainties

A rigorous quantitative treatment of uncertainties inthe experimental and analytical data is problematical, asis calculation ofrealistic uncertainties in the equilibriumfractionation factors. Several steps are involved in thederivation of each valid datum, each with some attacheduncertainty or error. Some sources of uncertainty (e.g.,analytical error in hydrogen-isotope analysis) are nor-mally distributed, whereas others (shifts in dD of chargesby exchange with the pressure medium) are not. Correc-tion of the experimental data for the latter effect has beenconsidered in detail above. The analytical uncertainty,estimated to be + l7-, is the most tractable, and probablythe largest, source of normally distributed uncertainty.This + l7m uncertainty gives rise to an uncertainty of+l.4lvn in l03ln a, (= dDln, - DDf'r") for any chlorite-water pair. This will be the uncertainty in l03ln a" onlywhen complete exchange is attained (e.g., runs 2 and 3,Table 3). However, it is evident from a consideration ofthe graphical representation of the interpolation methodfor deriving equilibrium fractionation factors from par-tial-exchange experiments (e.9., Northrop and Clayton,1966; Suzuoki and Epstein,1976) that the uncertainty inl03ln a" will be greater than + I .41 and will increase withdecreasing extent of exchange (/in Table 3). Taking lim-iting values of DD for chlorite and water consistent withanalytical uncertainty in each run used to interpolate val-ues ofa", uncertainties in interpolated values ofequilib-rium fractionation factors (l03ln a") were calculated (Ta-

ble 3), ranging from + 7.7 at 800/o exchange (runs 22 and23) to +3.7 at 38o/o exchange (runs 26 and' 27).

Relationship between fractionation factor(a") and temperature

Exchange rates (given by the extent of exchange, /inTable 3) in the system chlorite-Hro were sufficiently slowthat complete exchange (f : l) was attained in only onepair of experiments, run at 600'C for 46 d with the fine-grained starting material; although both charges (runs 2and 3 in Table 3) gained D, attainment of l00o/o ex-change, as demonstrated by the same value of ar, obviatesthe need for interpolation of a'. Runs 16 and 17 with themedium-grained starting material at 600"C for only 7l h,attained only 52o/o exchange, but the interpolation meth-od gave a value of d' comparable within experimentalerror. These results would seem tojustify the applicationof the interpolation method to the chlorites, as discussedabove.

Owing to slow exchange and decreasing extent of ex-change (/), interpolated values of a" are less well con-strained at lower temperature, despite the long run times.Thus at 500'C, runs attained only 38-47o/o exchange (26-43 d). At 300 and 400"C, exchange was insufficient toconstrain interpolated values of a" (Table 3).

Equilibrium fractionation factors a1700,600, and 500"Care given in Table 3 and plotted in Figure 2. Al 700 and600"C, where there is 50-1000/o exchange, equilibrium

572

T " c700 500 300 100

S e r p e n t i n e - E x p

Chlor i te -Mode l

C h l o r i t e- E x p

{ \

.4

Bruc i t e -E x p( 3 )

S e r p e n l i n e - E m p

C h l o r i t e - E m P

1 2 3 4 5 6 7

106/T2(K)

Fig. 2. Relationships between hydrogen-isotope fraction-ation factor (1000 ln ai,n*ur_*u,".) and temperature [plotted as 106/P (K)l for minerals in the system MgO-FeO-AlO3-SiOr-HrO.Chlorite-Exp.: this study; Chlorite-Model: calculated for clino-chlore composition using the model ofSuzuoki and Epstein (1976)(see text). Chlorite-Emp: empirical curve of Taylor (1974); Bruc-ite-Exp.: experimental data of Satake and Matsuo (1984); Ser-pentine-Exp.: experimental data of Sakai and Tsutsumi (1978);Serpentine-Emp.: empirical curve of Wenner and Taylor (1973).Tentative extrapolations (dashed lines) of experimental data (thisstudy) for chlorite-HrO based on fractionation factors given inTable 4 and derived from (1) Kuroda et at. (1976); (2) Heatonand Sheppard (1977); (3) Marumo et al. (1980). Symbols forvarying Fe/(Fe + Mg) in chlorite: filled circle : 0 to 0.09; cross :0.20 to 0.29; diamond : 0.30 to 0.39; triangle : 0.40 to 0.49;square : 0.50 to 0.69. See also Marumo et al. (1980, Fig. 8).

fractionation factors derived from different pairs of runsat each temperature are in reasonable agreement. At500"C, equilibrium fractionation factors calculated fromruns using the fine- and medium-grained starting mate-rials (0.4 <

"f < 0.5) are in close agreement (1031n a' *

-28), but are significantly more negative than those cal-culated from runs using the coarse-grained starting ma-terial (l03ln a' x -21;f :0.38). We place more relianceon the runs with the higher extents ofexchange and con-sider -28 to be the best estimate of the chlorite-waterfractionation at 500'C. There is no evidence in all thesedata for any measurable temperature dependence for thechlorite-water hydrogen-isotope fractionation over thetemperature range 500-700t, the fractionation factor re-maining close to l03ln a' : -28.

DrscussrouControls on fractionation factors

Suzuoki and Epstein (1976) proposed a general rela-tionship between hydrogen-isotope fractionation factor

GRAHAM ET AL.: CHLORITE-WATER HYDROGEN-ISOTOPE EXCHANGE

o@

I -1 06o

o E - 2 0

8

o -3c-30

-40

- 5

and octahedral cation composition (Mg:Fe:Al) applicableto sheet silicates, in which 1000 ln di,i...ur-szo : -22.4(106/

T2) + (2X^t - 4Xve - 68X.") + 28.2, where X' is themole fraction of cation M in octahedral sites, but pre-dicted that chlorite might be deuterium depleted relativeto fractionation factors predicted by the model becauseof hydrogen bonding in chlorite. Comparison of modelversus experimentally measured fractionation factors forthe chlorite used in this study shows this to be the case(Fig. 2), the experimentally measured fractionation fac-tors being 15 to 30V* more negative than predicted bythe model ofSuzuoki and Epstein (1976).

The efect ofhydrogen bonding on D-H fractionationin mineral-water systems was studied in detail by Gra-ham et al. (1980), who confirmed the prediction of Su-zuoki and Epstein (1976) that hydrogen-bonded mineralsshould strongly concentrate protium relative to non-hy-drogen-bonded minerals. A qualitative model relatinghydrogen-bond structure and geometry to fractionationfactor and exchange rate (Graham et al., 1980) was sub-sequently shown not to be supported by more recent neu-tron-diffraction studies of hydrogen in epidote.

Recent studies of the chlorite structure show that hy-droxyls occupy two structurally different sites in the bruc-ite and talc layers. The structure permits hydrogen bond-ing between hydroxyls of the brucite layer and oxygensof the talc layer, whereas hydrogens within the talc layerare not hydrogen bonded (Phillips et al., 1980; Joswig etal., 1980). Long hydrogen bonds between the brucite andtalc layers stabilize the chlorite structure and contributesignificantly to the interlayer bond energy (Bish and Giese,l98l). Therefore, we might expect the hydrogen-bondedhydroxyls associated with the brucite layer to concentrateprotium relative to the talc layer; that is, hydrogen iso-topes will fractionate differently between the brucite layerand water than between the talc layer and water. Dehy-dration of chlorites occurs in two stages, with dehydra-tion of the brucite layer occurring at lower temperaturesthan dehydration of the talc layer, and isotopic analysisof water released during step-heating of natural chloriteshows that water in the brucite layer is lower in AD thanwater in the talc layer (Suzuoki and Epstein, 1976), aswould be expected if hydrogen associated with the brucitelayer in chlorite is hydrogen bonded.

The experimentally determined chlorite-water frac-tionations (Fig. 2) are sigrrificantly more negative thanfractionations in the systems brucite-water (Satake andMatsuo, 1984) and serpentine-water (Sakai and Tsutsu-m| 1978), reflecting the hydrogen bonding in chlorite thatis absent in brucite and serpentine (Suzuoki and Epstein,1976; Satake and Matsuo, 1984). Comparison of thechlorite-water and brucite-water curves (Fig. 2) is alsocomplicated by the distribution of Al in chlorites, whichis strongly partitioned into the brucite layer (Phillips etal., 1980; Joswig et al., 1980), such that in the chloriteused in this study, between one-third and one-halfofthebrucite-layer octahedra are occupied by Al.

GRAHAM ET AL.: CHLORITE-WATER HYDROGEN-ISOTOPE EXCHANGE 573

TneLe 4. Chlorite-water fractionation factorsestimated from chlorite D/H analvses

f ChloritefC) 103 In c&1,",* Fel(Fe + Mg)

From Kuroda et al. (1976)200-300 -42.1 0.06200-300 -34.7 0.07

From Heaton and Sheppard (1977)

130140160170170180200220230230240250250250250250250250250250250

-31.7-31.7-36.9-32.8-34.8-33.8

From Marumo et al. (1980)_13.3_ 16 .6_16.6-35.8-42.2-28.2-39.1-33.4-40.3-36.0_40.9-37.5-36.4-37.7-31.7-35.3-33.5-46.7_47.8-37.4-39.1

0.51o.420.480.400.630.28

0.280.230.360.350.430.430.430.430.430.440.510.360.390.360.440.36U.JC

0.520.480.340.36

Comparison with empirical chlorite-waterfractionations and effect of FelMg

Taylor (1974) calculated an empirical curve for D-Hfractionation in the system chlorite-water (Fig. 2) usingdata for chlorites from pelitic schists. The compositionsand origins of these chlorites are not reported, but it maybe surmised that they were comparably or slightly lessaluminous than the chlorite used in this study (Velde andRumble, 1977)but with extensive and variable replace-ment of Mg by Fe2+. The extrapolation of this curve tohigher temperature lies close to the experimentally deter-mined fractionation factors (Fig. 2). On the other hand,a more securely based empirical calibration of chlorite-water hydrogen-isotope fractionation discussed belowsuggests diferent fractionation factors from Taylor's cal-ibration while remaining consistent with the experimen-tal data.

A limited amount of data on the D/H of natural chlo-rites from metamorphic (ocean floor), geothermal, andhydrothermal environments in which temperature, 6D ofcoexisting water, and Fe/Mg (chlorite) are well-known(Kuroda et al.,1976; Heaton and Sheppard,1977;Mar-umo et al., 1980) allow us to assess the effects of Fe/Mgand temperature on chlorite-water fractionations over arange of temperatures down to 100'C (Table 4). Althoughthe empirical calibrations must be treated with caution

o.2 0 .4 0 6F%"*ue

Fig. 3. Plot ofhydrogen-isotope fractionation factor (1000 lnaEn,-*,*J vs. Fe/(Fe + Mg) (chlorite). Data from this study (ex-perimental data), and Kuroda et al. (1976), Heaton and Shep-pard (1977), and Marumo et al. (1980). Empirical data listed inTable 4. Data of Marumo et al. (1980) are underlined.

because of uncertainty regarding crystallization temper-atures, 6D of coexisting water, and continued equilibra-tion of chlorites with water to temperatures well belowthose of crystallization of chlorite during slow cooling(Graham, l98l), the data in Table 4 represent low-tem-perature chlorites from geologic situations where coolingrates are relatively fast by geologic standards and DD (HrO)and temperature are quite well constrained, particularlythe geothermal drill-core data of Marumo et al. (1980) inwhich Z and DD (H'O) may be measured directly.

Marumo et al. (1980) have asserted that there is a cor-relation between a"n1-*1r6 and Fe/Mg (chlorite), and thata"n,."ro "does not depend on the temperature explicitly."However, the numerically smallest and largest fraction-ations (1000 ln a) in their data set, on which their sup-posed Fe/Mg effect is based, are from their lowest- andhighest-temperature samples, respectively (Fig. 3). Theeffects of temperature and Fe/Mg (chlorite) on chlorite-water fractionation cannot be unambiguously separatedin the data of Marumo et al. (1980), and, taking the dataof Table 4 together with the experimental data (Figs. 2,3), the above assertion of Marumo et al. is not supported.The data in Table 4 do not include information on theoctahedral Al contents of the chlorites, which could alsoinfluence chlorite-water D-H fractionations. However,from these data we may construct a tentative and empir-ical calibration of chlorite-water D-H fractionations forZ < 400'C that is broadly consistent with the 500-700"C

20

400400400370370370

;'I

Ic -30

o

o r o o - r a 9 " cr t s o - t 9 9 o co 200-2490C. 2 5 0 - 2 9 9 o cx goo-aooocg soo-zoooc

E x o e r i m e n t a l d a t a

xI

.!x

- !

! l!

tt' a

x l

o

- x

o9 oa

574

experimental data for Fe-poor aluminous chlorite. Thesedata (Fig. 2) are only sufficient to constrain 1000 lna:hr-H2o to +5Vm at best at any temperature, but allow usto make the following tentative conclusions: (l) Any Fe/Mg (chlorite) dependence is within the scatter of the datafor 0 < FelMg < 0.6 and must await experimental ver-ification. (2) The experimental data for Fe-poor sheridan-ite indicate that 1000 ln o;n,_rro is about -28 over thetemperature range 500-700'C, with no temperature de-pendence. (3) Over the temperature range 500-200t,fractionations show only a small temperature dependenceat most, with 1000 ln a;n,.rro becoming more negativewith decreasing temperature, but lying in the range -30to -40. (4) Below 200t, fractionations may become morepositive with decreasing temperature, as found in the ex-perimentally studied systems serpentine-HrO (Sakai andTsutsumi, 1978), epidote-H,O, and boehmite-HrO (Gra-ham et al., 1980). (5) The empirical calibration in Figure2 is shifted by 5 to l}Vm to more positive fractionationsthan those deduced by Taylor (1974) using chlorites frompelitic schists.

Irnplications for experimental actinolite-HrOD-H fractionations

Graham et al. (1984a) experimentally determined theD-H fractionation for actinolite-water at 400"C to be 1000ln a;.,6o,,,.-"ro : -29, using an actinolite from a meta-morphosed basaltic rock. In these experiments, 55-600/oexchange was attained in 59 d. However, the actinolitestarting material contained a small amount of chloritethat could not be separated, and therefore the measuredfractionation is only approximate. In this study, muchlonger experiments (105 and 157 d) produced negligibleD-H exchange between chlorite and water at 400"C.Therefore, unless the chlorite intergrown with actinolitehas a particularly small ffictive grain size, this chloritein all probability underwent very limited exchange duringthe actinolite-water exchange experiments. Therefore, themeasured actinolite-water D-H fractionation at 400"C maybe a good approximation to the equilibrium fractionationfactor.

KrNnrrcs oF D-H ExcHANGE

The kinetics of hydrogen-isotope exchange betweenchlorite and water may be quantified by application ofreaction-kinetics theory (Graham, 198 l). Here we as-sume, followingGraham (1981) andGraham etal. (1984a),that exchange reactions may be approximated (empiri-cally) by second-order kinetics, so that

f/(r - J): K,t, (6)

where/is the fractional approach to equilibrium (Table3), / is the time of exchange, and K. is the second-orderrate constant. Activation energies (O) for the rate-deter-mining step may be derived from the Arrhenius relation-ship

log K,: log a - Q/2'303RT, (7)

where c is a constant (the pre-exponential factor).

GRAHAM ET AL.: CHLORITE-WATER HYDROGEN-ISOTOPE EXCHANGE

Satake and Matsuo (1984) chose to assume that hydro-gen-isotope exchange between brucite and water followedfirst-order kinetics, although they were unable to dem-onstrate this unequivocally (cf. Graham, l98l). Second-order kinetics give an equally good linear fit of the brucitedata to the Arrhenius relationship as do first-order ki-netics. Nonetheless, it is quite likely that the brucite-waterexperimental data fit a dffirent rate law than other ex-perimental OH mineral-water exchange data, because Sa-take and Matsuo (1984) observed that dissolution-repre-cipitation was very likely the major process of D-Hexchange rather than volume diffusion ofhydrogen in thesolid (cf. Graham, 1981).

Minerals may change in grain size during exchange ex-periments owing to mechanical breakdown or dissolu-tion-reprecipitation; ifthis occurs, exchange rate may varywith time. Chlorite-water exchange experiments usedsmall amounts of water (cf. Graham et al., 1980) in orderto minimize grain dissolution-reprecipitation. Chlorite runproducts of flve exchange experiments were examined byseu in order to assess the extent of either of these pro-cesses. Comparison of run products with starting mate-rials (Fig. 4) indicates that no significant changes in grainsize or morphology have occurred and that chlorite D-Hexchange must proceed dominantly by volume diffusionofhydrogen in the chlorite.

An Arrhenius plot of second-order constants (log Kr)vs. 103/Z for the chlorite-water exchange data (Fig. 5)gives good and consistent linear fits for the medium andcoarse grain-size fractions and is a useful test and dem-onstration of the internal consistency of the experimentaldata set. Activation energies for hydrogen transport inchlorite are around 150 kJ (36 kcal)/mol H for the me-dium and coarse grain sizes (Table 5). Rate constants forthe coarse grain size are displaced to more negative val-ues of log K, than those for the medium and fine grainsizes (Fig. 5), consistent with the conclusion that the ef-fective chlorite grain size has not varied during exchangeexperiments, that exchange has proceeded by volume dif-fusion of hydrogen in chlorite, and that the measuredgrain sizes are the effective grain sizes for diffusion. Cal-culated activation energies are much higher than for hy-drogen exchange between other hydrous minerals andwater (cf. Gt'aham, l98l; Graham et al., 1984a), as wouldbe predicted from the very slow exchange rates at Z s400"C.

DrrrusroN oF HYDRocEN IN cHLoRrrE:CLOSURE TEMPERATURES

Calculation of diffusion coefficients

Closure temperatures for cessation of hydrogen-isotopeexchange between chlorite and hydrous fluid in coolingmetamorphic and hydrothermal systems may be calcu-lated from estimated diffusion parameters for hydrogendiffusion in chlorite. Methods of calculating hydrogen-difrrsion coefficients from experirnental data on hydro-gen-isotope exchange are described in detail by Graham(1981), and their application to the chlorites is brieflyoutlined here.

GRAHAM ET AL.: CHLORITE-WATER HYDROGEN-ISOTOPE EXCHANGE 575

Fig. 4. Scanning electron microscope (snv) photographs of fine-grained chlorite in starting material (b) and in run product (a)of exchange experiment at 600t for 217 h. Scale bar is in micrometers. Photographs indicate absence of solution-reprecipitationor mechanical breakdown during exchange experiments.

The calculation of diffusivities involves a knowledge ofgrain geometries in order to account for diffusional an-isotropies, especially for minerals with well-developedcleavages such as the sheet silicates. The mean effectivegrain diameters, plate thicknesses, and aspect ratios forthe three chlorite grain-size fractions were measured op-tically and are given in Table 6. sru observations showedthat measured grain thicknesses were greater than truethicknesses owing to the difficulty of viewing the platygrains edge-on in optical mounts, and grain thicknesseshave been adjusted accordingly to half the optically es-timated thicknesses to take account of this problem andto assess the effect on calculated diffusion coefficients ofvarying the grain thickness. Two geometrical models areapplicable to hydrogen diffusion in platy minerals: (l) theinfinite cylinder model in which hydrogen diffuses par-allel to the layers and (2) the infinite plate model in whichhydrogen diffirses perpendicular to the layers. For a stirredsolution of limited volume, Equations 5.36 and 4.43 re-spectively of Crank (1975) are applicable (see Graham,1981). Corresponding values of log D calculated usingthese equations are given in Table 3. For the plate model,both measured and adjusted plate thicknesses were used;the effect ofreducing plate thickness by halfis to reduceD by about half an order of magnitude.

The correct difusion model should yield similar dif-fusion coefrcients for the different grain sizes at any tem-

perature. However, although chlorite is a platy mineral,the aspect ratios of the starting materials are not largeenough to allow a clear-cut choice to be made betweenthe two models (cf. muscovite; Giletti and Anderson,1975; Graham, 1981), because for both models all data,irrespective of grain size, may be fitted to a single linearregression on an Arrhenius plot (Fig. 5), with a correla-tion coefficient (r'z) better than 0.97 (Table 5). Althoughthe data may be interpreted to mean that hydrogen dif-fusion occurs at comparable rates parallel and perpendic-ular to the layers, it is more likely, by analogy with mus-covite (Graham, l98l) that hydrogen diffusion will occurfaster parallel to the layers (cylinder model). For this case,diffusivities are one to two orders of magnitude fasterthan for the plate model (depending on the adopted platethickness).

Calculated activation energies for both models arecomparable at 172 kJ (41 kcal)/mol H and 167 kJ (40kcal)/mol H for cylinder and plate models, respectively(Table 5). These activation energies are of considerableinterest, being significantly the largest for any hydrousmineral studied to date. By comparison the activationenergies for hydrogen diffusion in muscovite, hornblende,tremolite, and zoisite are l2l (29),84 (20),71 (17), and103 (24.5) kJ (kcal)/mol H, respectively (cylinder model;Graham, 198 I ; Graham et al., 1984a) (Fie. 5). It is pos-sible that the large interlayer site in muscovite, for ex-

A f ine gra in s izeO med ium gra in s izeO coarse gra in s iz€

576

1 01 o 3/T(K )

ample, allows faster transport pathways parallel to thelayers than are available in chlorite, where interlayer sitesare lacking. Diffusion coefficients for hydrogen diffusionin chlorite, adopting the cylinder model for comparison,are up to 3 to 4 orders of magnitude slower over the rangeof temperatures of crystallization and cooling of meta-morphic and hydrothermal rocks than diffi.rsion coeffi-cients for hydrogen diffirsion in other hydrous minerals,excluding epidote (Graham, l98l; Graham et al., 1984a).The implications of these results for calculated closuretemperatures for the cessation of hydrogen diffusion in

GRAHAM ET AL. : CHLORITE-WATER HYDROGEN-ISOTOPE EXCHANGE

r cc)600

N

Y

o -

700

r o3rr(r)

r ("c)600

A l i n e g r a i n s i z e

O O m e d i u m g r a i n s i z e

O ! c o a r s e g r a i n s a z e

Fig. 5a. Arrhenius plot ofsecond-order rate con-stants (log Kr) vs. reciprocal temperature for hy-drogen-isotope exchange experiments betweenchlorite and water. Data given in Table 3. See textfor discussion.

Fig. 5b. Arrhenius plot of diffusion coefrcientsvs. reciprocal temperature for diffusion ofhydrogenin chlorite, using both cylinder and plate models.Plate (l) refers to optically estimated grain thick-nesses; plate (2) refers to adjusted grain thicknesses(see Table 6). Muscovite data (cylinder model) fromGraham (1981) shown for comparison. Symbols asin Fig. 5(a). Data given in Table 3. See text fordiscussion.

chlorites relative to other hydrous minerals are consid-ered below.

Exchange mechanisms

Matthews et al. (1983) conducted one oxygen-isotopeexchange experiment using clinochlore and water at 500tfor 3 d. This experiment achieved only ll0/o exchange.By comparison, the medium-grained chlorite used in thisstudy attained 45o/o exchange of hydrogen isotopes at500"C in 26 d. The grain size of the starting material inthe oxygen-isotope exchange study was not stated by

1 Q1 . 0

-o

* - r z

o

- 1 8

GRAHAM ET AL.: CHLORITE-WATER HYDROGEN-ISOTOPE EXCHANGE 577

400

Tc( 'C)

300

coolins Rate too $f ("clmv)

Fig. 6. Plot of calculated closure (blocking) temperature (71) vs. logarithm of cooling rate dT/dt (dee/m.y.) for cessation ofhydrogen-isotope exchange between chlorite and water, using both plate and cylinder models. Grain sizes (in mm) refer to diameterof chlorite grain (cylinder model) or thickness of chlorite plate (plate model). Calculation method described in text.

Matthews et al. (1983), but was most likely very muchfiner grained than the chlorite used in this study (Mat-thews, pers. comm.), indicating that hydrogen diffusionin chlorites is more rapid than oxygen diffusion. This isconsistent with the observed absence of solution-repre-cipitation in our experiments and with the observationsof Graham (1981) that hydrogen diffusion in muscoviteproceeded fastest parallel to the layers and was about frve

TABLE 5. Activation energies (O) derived from the Arrhenius re-lationshio

(kJ/mol) (kcal)

Chlorite-water hydrogen-isotope exchange at 500-700'C.

orders of magnitude faster than oxygen diffusion, whichoccured fastest perpendicular to the layers (Giletti andAnderson, 1975). Similarly, O'Neil and Kharaka (1976)observed negligible oxygen-isotope exchange between clayminerals and water in low-temperature (=350"C) ex-change experiments in which significant hydrogen-iso-tope exchange occurred. These authors concluded thathydrogen and oxygen isotopes exchanged by diferentmechanisms in the absence of solution-reprecipitation.

Closure temperatures

Assuming that the diffusion coefficients for hydrogendiftrsion in chlorite determined at 500-700t in this studymay be extrapolated to lower temperatures (i.e., that dif-fusion mechanisms are not temperature dependent), clo-

TneLe 6. Chlorite grain geometries

Nominal grain size (pm)Average grain diameter (pm)Measured grain thickness (pm)Adjusted grain thickness (pm)Asoect ratio

Note.' Average grain diameters and thicknesses measured optically.Measured grain thicknesses are too large due to difficulty of viewing platygrains edge-on. Grain thicknesses were adiusted to half the measuredvalues, and both sets of thicknesses were used in diffusion-coetficientcalculations to assess the effect of varying grain thickness.

Medium grain sizeCoarse grain size

Cylinder modelPlate model

(measuredthickness) -6.70Plate model

(adlustedthickness) -7.32

Hydrogen diffusion in chlorite at 500-70eC.'-5.21 -8.97 171.7 41.0 0.97

3.70 -7.89 151.1 36.1 0.993.56 -7.95 152.2 36.4 0.99

-8.72 166.9 39.9 0.98

8.70 166.5 39 I 0.98

63-105 44-63 <44167.6 1 09.9 34.520.6 16.7 5.010.3 8.4 2.51 6.3 13.1 13.8Note: b : 412.303R; r, is the correlation coefficient; f is in kelvins

. O calculated by least-squares fitting of calculated second-order rateconstants (Kr) to the equation log K,: a + q103/7). Grain geometries aregiven in Table 6.

'- O calculated for both the infinite-cylinder and plate models by least-squares fitting of calculated diffusion coefficients to the equation log D:a + 410317).

578

sure temperatures for the cessation of hydrogen-isotopeexchange between chlorite and water during cooling ofmetamorphic or hydrothermal rocks may be modeled asa function of grain size and cooling rates. The closuretemperatures (2.) were calculated using the relationship

T " : Q/R (8)

GRAHAM ET AL.: CHLORITE-WATER HYDROGEN-ISOTOPE EXCHANGE

after Dodson (1973, Eq. 23), which is readily solved byiterating values of T.; Q is the activation energy for dif-fusion of hydrogen in chlorite (Table 5), I is the diffu-sional anisotropy parameter (: 27 for cylinder; : 8.65plate), R is the gas constant (: 8.314 J.mol '.deg-'), Dois the pre-exponential factor in the Arrhenius relation-ship, a is the grain-size term (radius of cylinder, half-thickness of plate), and dT/dt is the cooling rate (deg.s '), where cooling rate is assumed to be linear with l/Znear 2". Calculated closure temperatures will show smallvariations with varying fluid/mineral ratio, as noted byCole et al. (1983) for experimental oxygen exchange data(e.g., for cylinder model, l0-20 deg variation for two or-ders of magnitude variation in water-chlorite ratio), butincorporation ofthis factor into the calculation of 7" re-quires a more elaborate analysis than has been attemptedto date. Figure 6 shows the calculated relationship be-tween I. and cooling rate for various chlorite grain sizes,for both plate and cylinder models.

The following general conclusions may be drawn fromFigure 6. First, in slow-cooling geologic environments suchas those of regional metamorphism [e.g., I < log(dT/dl) <21, even the largest chlorite crystals in metamorphic rockswill be susceptible to continued hydrogen-isotope ex-change with coexisting hydrous fluid to temperatures be-low those of greenschist-facies metamorphism. Loss ofhydrous fluid will of course permit the chlorite to pre-serve a record of higher-temperature fluid interactionsduring slow cooling (e.g., Graham, l98l).

The much faster cooling rates likely to characterize hy-drothermal greenschist metamorphism of the oceanic crustor contact metamorphism of mafic rocks may be suffi-cient to lead to closure temperatures for cessation ofhy-drogen-isotope exchange between chlorite and water thatare within the greenschist facies and not far removed fromtemperatures of chlorite crystallization. This conclusionwill, however, be dependent on chlorite grain size and onthe diffusion model adopted (Fig. 6).

Finally, computed closure temperatures for chloritesare signiflcantly higher than those for other hydrous min-erals (cf. Graham, 1981; Graham et al., 1984a), as wouldbe expected from the high activation energies for hydro-gen diffusion in chlorites (Fig. 5). The variations in clo-sure temperatures and activation energies among difer-ent hydrous minerals will mean that mineral-mineralhydrogen-isotope equilibrium is unlikely to be preservedif a hydrogen-bearing fluid is present even in small

amounts during slow cooling (see also Graham, l98l);preservation of hydrogen-isotope equtlibrium, on the otherhand, might indicate loss of such a fluid at an early stageofthe cooling process.

Pnrnolocrc APPLICATIoNS

Metamorphism of basic rocks to chloritic "green-stones" in the oceanic crust is thought to have occurredin environments dominated by relatively high fluid/rockratios, with fast cooling rates being indicated by the veryyoung ages of many studied samples (e.g., Stakes and O'-Neil, 1982). For these circumstances, closure tempera-tures for cessation ofhydrogen-isotope exchange betweenchlorite and water are in excess of 250'C (Fig. 6). Hydro-gen-isotope fractionation factors between chlorite andwater at temperatures above 250"C should lie in the ap-proximate range 1000 ln a!,,1-*u,". : -30 to -40vm (Fig.2). Inasmuch as chlorite is a major reservoir of hydrogenin the oceanic crust, the above fractionation factor shouldrepresent a dominant component of the overall seawater-hydrated oceanic crust hydrogen-isotope fractionation forthat portion of Earth's history during which sea-floorspreading has operated.

There are currently limited data available on the dD ofchlorites from ocean-floor, metamorphic, or hydrother-mal-geothermal environments. Some of these data (Ku-roda et al., 1976:, Marumo et al., 1980; Heaton and Shep-pard, 1977) have been utilized and discussed above inassessing the effect of FelMg on chlorite-HrO D-H frac-tionation and the fractionation factors at low tempera-tures.

Surprisingly few data are available on the DD ofchloritein hydrated mafic rocks of the ocean crust and of ophiolit-ic rocks. Satake and Matsuda (1979) analyzed the whole-rock 6D of hydrated metabasalt from the Mid-AtlanticRidge in which chlorite (typically 24 modal percent) is thedominant hydrous mineral, and therefore DD** t 6D"r,r :-38 to -30V*. Separated chlorites from Mid-AtlanticRidge greenstones (-317m, -34Vm; Stakes and O'Neil,1982) lie within this range of values. At the likely closuretemperatures for cessation of D-H exchange during cool-ing of hydrated ocean crust, water in equilibrium withthese rocks should be close to dDro : 07- (using fraction-ation factors estimated from Fig. 2), consistent with aseawater ong1n.

Similarly, little data exist on the dD of chlorites andcoexisting hydrous minerals in metamorphic rocks. Ryeet al. (1976) studied the stable isotope composition ofcoexisting hydrous minerals in high-pressure metamor-phic rocks of Naxos (Greece). In a low-grade schist crys-tallized at about 480'C, chlorite of dD : -55V* coexistswith white mica, consistent with available experimentaland empirical fractionation data (this study Suzuoki andEpstein, 1976), and the calculated dD'ro for the meta-morphic water is in the range -25 to -15Vm.

Magaritz and Taylor (1976) reported dD.hr and DD**for chlorite-rich samples in a wide range of metamorphicrocks of the Franciscan Formation of California. Except

for samples for the vicinity of the San Luis Obispoophiolite (see discussion in Sheppard, 1980), chloritesgivecalculated DD"ro for metamorphic fluids within the range- 36 to - 37m. These estimates are compatible with DD"roof metamorphic pore fluids calculated by Magaritz andTaylor (1976).

Acrlqowr-rncMENTS

This work was supported by NATO Grant 0551/82 (to CM.G andR.S.H.) and NSF Grant EAR 82-18380 (o R.SH). The ExperimentalPetrology Unit in Edinburgh and isotope geology facilities at the ScottishUniversities Research and Reactor Centre are supponed by NERC. Dis-cussions with Simon Sheppard, Alan Matthews, and Gawen Jenkin andhelpful reviews from John Valley and J. Rice are acknowledged.

RrrnnnNcnsBish, D.L., and Giese, R.F. (1981) Interlayer bonding in IIb chlorite.

American Mineralogist, 66, 12 | 6-1220Cole, DR., Ohmoto, H., and Lasaga, A.C. (1983) Isotopic exchange in

mineral-fluid systems I. Theoretical evaluation ofoxygen isotopic ex-change accompanying surface reactions and diffusion. Geochimica etCosmochimica Acra. 47 . 168 l-1693

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MeNuscnrpr RECETVED NowuseR. 5, 1985MnNuscnryr AccEprED Fennunv 13, 1987

GRAHAM ET AL. : CHLORITE-WATER HYDROGEN-ISOTOPE EXCHANGE