An Integrated Chemical and Stable-Isotope Model of the Origin of Midocean Ridge Hot Spring Systems

JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 90, NO. B14, PAGES 12,583-12,606, DECEMBER 10, 1985

TERESASUTER BOWERS 1 AND HUGH P. TAYLOR,JR.

Division of Geologicaland Planetary Sciences,California Institute of Technology,Pasadena

Chemical and isotopic changes accompanying seawater-basalt interaction in axial midocean ridge hydrothermal systemsare modeled with the aid of chemical equilibria and mass transfer computer programs,incorporating provision for addition and subtractionof a wide-range of reactant and product minerals, as well as cation and oxygen and hydrogen isotopic exchangeequilibria. The models involve stepwiseintroduction of fresh basalt into progressivelymodified seawater at discrete temperature intervals from 100ø to 350øC, with an overall water-rock ratio of about 0.5 being constrainedby an assumed •80,2 o at 350øCof +2.0 per mil (H. Craig,personal communication, 1984). This is a realisticmodel because:(1) the grade of hydrothermal metamorphism increasessharply downward in the oceanic crust; (2) the water-rock ratio is high (> 50) at low temperaturesand low (<0.5) at high temperatures;and (3) it allows for back-reaction of earlier-formed minerals during the course of reaction progress.The results closelymatch the major-elementchemistry (Von Damm et al., 1985) and isotopic compositions(Craig et al., 1980) of the hydrothermalsolutions presently emanating from ventsat 21øN on the East Pacific Rise. The calculatedsolution chemistry,for example, correctly predicts completeloss of Mg and SO,• and substantial increases in Si and Fe; however, discrepanciesexist in the predicted pH (5.5 versus 3.5 measured)and state of saturation of the solution with respectto greenschist-faciesminerals. The calculated 6D.2o is +2.6 per mil, in excellentagreement with analyticaldeterminations. The calculated chemical, mineralogic, and isotopic changes in the rocks are also in good accord with observations on altered basaltsdredged from midocean ridges(Humphris and Thompson, 1978; Stakes and O'Neil, 1982), as well as with data from ophiolites(Gregory and Taylor, 1981). Predicted alteration productsinclude anfiydriteand clay minerals at low temperaturesand a typicalalbite-epidote-chlorite-tremolite (greenschist)assemblage at 350øC. The models demand that the major portion of the water-rock interaction occur at temperatures of 300ø-350øC. Interaction at temperatures below approximately 250øC results in negativeg•80.2 o shifts,contrary to theobserved positive g•80 valuesof thefluids exiting at midocean ridge vents.Hydrogen isotope fractionation curvesby Suzuoki and Epstein (1976), Lambert and Epstein (1980),and Liu and Epstein(1984), among others, are compatiblewith the model,and require6D.2 o to increaseat all temperaturesas a result of seawater-basaltinteraction.

INTRODUCTION Two major factors affect the outcome of water-rock interac- The discoveryand subsequentsampling and analysisof sub- tions for a given initial composition of water and rock: tem- marine hot springs(e.g., at 13øN and 21øN on the East Pacific perature of interaction and water-rock ratio. These parameters Rise (EPR)) provide an excellent opportunity for the devel- can vary independently, and it is often difficult to assessthe opment of theoretical models to describethe midocean ridge importance of each from the petrologic record alone. The pri- (MOR) alteration process.This is becausethe chemical and mary goal of this project is to utilize a reaction-path computer stable-isotopecompositions of both the starting solution (sea- model that integrates both the chemical and stable isotope equilibria in an attempt to narrow down the possible paths water) and the final hydrothermal endmember are well con- seawater can follow in simultaneouslyproducing the observed strained. In addition, the observedalteration products in the secondary alteration of basalt and the chemical and isotopic basaltsprovide data on chemicaland isotopic gains and losses to the rock. modification of the submarine hydrothermal fluids. A 'path' is defined as some combination of systematic and sequential Throughout the oceans, and in particular at the MOR spreading centers, the only important rock types are the change in temperature, pressure,and water-rock ratio during characteristic, relatively uniform MOR-type basalts and evolution of the fluids and rocks within the zone of hy- gabbrosmade up of plagioclase+ pyroxene+ olivine + glass. drothermal activity. The following are observational con- Also, the initial water entering the hydrothermal systems is straintswhich shouldbe met by the model. oceanwater, whichhas had a fairly uniform g180, gD, and The mineralogical characteristics of seafloor basalts and chemical composition over the course of geologic time. Thus, their associatedalteration products are describedby lto and the simplicity and uniformity of the starting materials in sub- Anderson[1983], Humphris and Thompson[1978], and Rona marine hydrothermal systemsallow us to make very realistic [1978, 1976], among others. For example, the data of calculations that have broader applicability than is the case Humphris and Thompson[1978] provide a number of impor- for most active or fossilsystems on continents. tant constraints on the predictions of our computer models. They describedtwo seriesof greenschist-facies(albite-chlorite- actinolite-epidote)rocks with minor amounts of quartz and pyrite, one a chlorite-rich, and another an epidote-rich assemblage. Consistent with the results of laboratory experiments • Now at Department of Earth, Atmospheric and Planetary Sci- (see below), the chlorite-rich rocks exhibit a gain of Mg and ences,Massachusetts Institute of Technology, Cambridge. H20 and a loss of Ca and Si, while the epidote-rich rocks Copyright 1985 by the American GeophysicalUnion. exhibit slightly lower Mg contents and slightly higher Ca con- Paper number 4B5398. tents compared to unaltered basalt. Humphris and Thompson 0148-0227/85/004B-5398505.00 propose that the differencesin rock type are a function of

12,583 12,584 BOWERSAND TAYLOR: CHEMICALAND ISOTOPICMODEL OF HOT SPRINGS

temperature, water-rock ratio, composition of the circulating PREVIOUS WORK fluid, and rate of fluid flow relative to reaction rates. Ito and Clayton [1983] measuredg•80 and gD valuesof Laboratory Experiments altered gabbros dredged from the Mid-Cayman Rise in the Seawater-basaltinteractions have been widely studied ex- Caribbean. These samplesfrom the deeper parts of the ocean perimentally. Bischoff and Dickson [1975] reacted seawater crust were altered at high temperatures (>350øC) and low and basalt in a 10 to 1 mass ratio at 200øC and 500 bars. The water-rock ratios (<2). Stakes and O'Neil [1982] found that resulting solutions and alteration mineral assemblagescom- several rocks from the EPR and MAR were altered at 35 ø to pared favorably to those observedin the Icelandic geothermal 350øC and water-rock ratios ranging from > 50 at low tem- fields.Further experimentswere designedto determinewheth- peratures to ,-, 1 at high temperatures.The low-temperature, er the interaction of seawater and basalt could: (1) produce saponite-richrocks are enrichedin •80 comparedto mid- the types of metabasaltscollected from midocean ridges and ocean ridge basalt (MORB), and the high-temperature ophiolite complexes; (2) be responsiblefor fixing the con- chlorite- and epidote-richgreenstones are depletedin •sO. centrationsof some chemical speciesin seawater; and (3) pro- The whole-rockglsO valuesrange from +3.7 (ultramafic)to duce seafloormetalliferous deposits. + 10.1 per mil (saponite-richpillow breccias);gD valuesrange Several additional experimentswere conducted at various from -79 (saponite) to -13 per mil (epidote). Bdihlkeet al. temperatureswith diverserock types (basalt, basalt glass,dia- [1984] report g•80 as high as +24 for smectitesand K- base, peridotite) and over a range of water-rock ratios [Bis- feldspar and q-30 for calcite in MOR basalts from DSDP choff and Rosenbauer, 1983; Seyfried and Mottl, 1982, 1977; cores. Estimated alteration temperatures range from 8ø to Janecky, 1982; Seyfried and Bischoff, 1981, 1979, 1977; Sey- 50øC. fried and Dibble, 1980; Mottl and Seyfried, 1980, 1977; Mottl et An additional source of data is provided by mineralogical al., 1979; Bischoffand Seyfried,1978, 1977; Mottl and Holland, and isotopic studies of ancient oceanic crust, such as the 1978; Mottl, 1976; Hajash, 1975]. A summary of the resultsof Samail ophiolite in Oman [Greqory and Taylor, 1981; Stakes these experiments is given by Mottl [1983]. At water-rock et al., 1984]. Whether they formed at midocean ridges or in ratios < 50, nearly all the dissolvedMg is removed from sea- back-arc spreadingbasins, such ophiolites clearly exhibit the water and taken up by secondaryalteration minerals in the effects of a fossil marine hydrothermal system that involved basalt, with a concomitant drop in pH. At higher water-rock heated ocean water; such systems are probably similar to ratios the removal is less complete, but still dramatic. The present-day MOR systems.Greqory and Taylor [1981] report decreasein pH speedsup the alteration process,particularly a pattern of mineralogical alteration in the Oman ophiolite the leaching of heavy metals. Seyfriedand Mottl [1982] note ranging from amphibole plus some chlorite and epidote in the that the additionalH + ions are consumedduring alterationof high-levelgabbros to common greenschistassemblages includ- the primary silicates,resulting in competing effectson the pH. ing actinolite,chlorite, saussurite, leucoxene and epidotein the Thus, oncethe Mg ++ concentrationdrops to a low value,the aliabase,to zeolites and secondary carbonate in the pillow pH increasesrapidly back to a value near neutrality as a result basalts.This is accompaniedby whole-rockrS•sO values rang- of continued silicatehydrolysis. ing from a high of about + 13 per mil in the pillow basaltsto In addition to Mg removal, various experimentsindicate a minimum of about + 3.8 per mil at depths 1.5 to 2 km below removal of some Na + at low water-rock ratios and addition of the diabase-gabbro contact. Ca + + to the solution(leaching of Ca + + from primarysilicates Chemical and isotopic analysesof MOR hot-spring waters balancesthe removal of Mg ++ to alteration phases[Mottl [Edmond et al., 1979; Edmond, 1980; Welhan and Craig, 1979; and Holland, 1978; $eyfriedand Bischoff,1981]). In theseaque- Craig et al., 1980; Craig, 1981; Michard et al., 1984; l/on ous solutions,dissolved silica contentsare commonly near the Damm et al., 1985] provide a picture of the 350øC hy- level of quartz saturation. drothermal endmember: magnesium and sulfate have been Although the laboratory experiments have worked out quantitatively removed; silica has reached a concentration fairly well in producing aqueous solutions similar to those controlled by quartz or amorphous silica saturation two observed in the Icelandic geothermal fields, the Galapagos orders of magnitude higher than in seawater; iron con- Rift and the East Pacific Rise, poor agreement is observed centrations are approximately six orders of magnitude higher between the rock alteration assemblagesproduced in the lab- than seawater;and pH has dropped off from a value of about oratory and the known mineralogy of altered oceanic rocks 8 for seawaterto about 3.5. In addition, the g•80 and gD of [Mottl, 1983]. Greenschist-faciesbasalts are most often either seawater(which both start at zero per mil relative to standard chlorite- or epidote-rich [Humphris and Thompson, 1978], mean ocean water (SMOW)) have increased in the hy- while laboratory alteration products are most often smectite- drothermal endmember to approximately + 2.0 and + 2.5 per rich. Nucleation difficulties and kinetics may in part explain mil, respectively [H. Craig, personal communication, 1984; thesediscrepancies. Note that a recent experimentby Seyfried Craig et al., 1980]. Becauser5•80 is the one variablethat is and Mottl [1982] resulted in the formation of a mixed-layer most simply and straightforwardlyrelated to water-rock ratio, smectite-chloriteduring a reaction of basalt glass with seawa- its value in the hydrothermal endmember is used to set the ter. endpoint in the calculatedreaction paths discussedbelow. Utilizing these starting conditions and constraints,we de- ThermodynamicCalculations velop an integrated stable-isotope, chemical equilibria and In addition to laboratory experiments,the consequencesof mass transfer computer model based on the EQ3/6 computer water-rock interaction can be calculated with the aid of software package [Wolery, 1978, 1979, 1983]. We then use this chemical equilibria and mass transfer computer programs to calculate chemical and isotopic modifications of seawater using thermodynamic data for minerals and solutions. How- together with mineralogical and isotopic alterations of fresh ever, although fairly realistic calculations have been carried basalt, for comparison with the known petrology and chemis- out for cation speciesin minerals and aqueousfluids [Wolery, try of seafloorhydrothermal systems. 1978; danecky, 1982; Reed, 1983], up to the present no one BOWERSAND TAYLOR: CHEMICAL AND ISOTOPICMODEL OF HOT SPRINGS 12,585 has performedsimilarly detailed calculations for •80/•60 or such alteration products would have on the chemical and D/H. oxygenisotope massbalances. Wolery [-19781revised and extended the mass transfer computer program of Helgeson•19681 and Helgesonet al. [-19701 DESCRIPTION OF COMPUTER PROGRAMS to calculate chemical equilibria among aqueous and mineral AND DATA BASES phasesin basalt-seawatersystems. He conducteda computer experimentin which seawaterwas heated from 0ø to 350øC at Computer Modeling 500 bars, the results of which compared favorably with the The EQ3/6 package used in the present study was devel- experimental results of Bischoffand Seyfried [-1978•. In addi- oped by Wolery [1978, 1979, 1983]. EQ3 computes a tion, the experimentaccurately reproduced changes in compo- distribution-of-speciesin the aqueous solution and generates sition of the solution, including an acidic pH at T > 200øC. concentrationsand activities of ions and complexes;it then Differencesbetween the computer calculations and the results calculates which, if any, of the mineral phasesare saturated of the laboratory experimentswere attributed to lack of equi- with respect to the solution. EQ6 then calculates chemical librium in the experiments,sluggish reaction rates in experi- equilibrium and mass transfer in aqueous solution-mineral ments at T < 150øC [Seyfried, 1976], and inaccuraciesin the systems.These programs are similar to, and derivative from, thermodynamicdata for important speciesin the theoretical those of Helgeson [1968] and Helgeson et al. [1970]. Math- calculations. ematical equations describing mass action, mass balance, Wolery [1978] also calculated solution composition charge balance,and nonideality are solved at each step in the changesand rock alteration accompanyingreaction of sea- reaction progress(for eachincrement of reactantphases added floor basalt with seawaterin isothermal,closed systems at 500 to the solution)with a Newton-Raphson technique. bars and various temperaturesfrom 0 ø to 340øC. Among the EQ6 includes provision for: (1) variations in temperature, alteration mineralspredicted were smectites,chlorites, kaolin- (2) variable rates of titration of reactants,(3) suppressedfor- ite, muscovite,albite, quartz, anhydrite, dolomite, calcite, and mation of undesiredprecipitates (this allows, for example,for- severaliron and coppersulfides. mation of kinetically favored, although thermodynamically More recently, Janecky [1982] used EQ3/6 to calculate metastablephases), and (4) reaction in either a closed-system mineralogical and chemical changes attending seawater- or open-systemmode. periodite interaction at 300øCand 500 bars, employingexperi- The closed-systemmodel is directly applicable to hy- mentally determinedrelative dissolutionrates for the common drothermal bomb experiments.Reactant solid phasesare titra- peridotite minerals. Substantial agreement was found between ted at specifiedrelative rates into solution, and the chemical the experiments and the calculations for both the solution compositionof the solutionis progressivelymodified by disso- composition trends and for the precipitates.Reed [1983] has lution of the reactant phases,as well as by precipitation and modeled deposition of massivesulfides in fluids derived from dissolutionof product phases,which remain in the systemand high-temperature seawater-basalt interaction followed by continue to take part in the equilibria. The reaction progresses cooling of the fluid and subsequentmixing with seawater. until equilibrium is completelyattained. In contrast to the situation described above, dissolution of product phases is not allowed in the "simple" open-system Oxygen Isotope Modeling model. Secondary phases are instantaneouslyremoved from Previous computer models of oxygen isotope mass transfer the system as they precipitate to simulate a packet of fluid among minerals and aqueous solutions include those ad- flowing through unaltered,pristine rock composedonly of the vanced by Cathles [1983], Parmentier [1981], and Norton and reactant phases.In this case, the alteration products do not Taylor [1979], none of which took into account the detailed interact with the systemafter they have formed, and they are variations in product and reactant mineral phases. Cathles assumedto be left behind in the hydrothermal conduit. In this modeled•80, anhydrite,and silica redistributionin a hy- paper, we have not made any model calculationsof this type, drothermal system and compared the results to the Kuroko becausewe feel that a "complex" open-systemmodel is more massive sulfide deposits. Kinetic constraints and formation realistic. The "complex" open-systemis one in which products permeabilitieswere considered,and the calculated whole-rock are periodically removed from the system (for example, at 6•80 anomaliesshowed a patternsimilar to thoseassociated specified temperature intervals), but also one in which the with the Fukazawa massivesulfide deposit in Japan. Norton product minerals remain available between these intervals, and Taylor [1979] analyzed fossilgeothermal systemsassoci- where they can dissolveor otherwisetake part in the equilib- ated with the Skaergaard intrusion, using a numerical ap- ria of the system. proximation of heat and mass transport [Norton and Knight, 1977] and porosity and permeability [Norton and Knapp, Thermodynamic Data Base 1977-1,which permitted simulation of the thermal history and pattern of energy loss during crystallization of this layered The thermodynamicproperties of a large number of min- gabbro. In this system,little or no mineralogical change oc- erals are either well known [Helgeson et al., 1978; Robie et al., curred in the gabbro, and Norton and Taylor [1979] were thus 1978], or can be estimated [Wolery, 1978]. Except for a few able to make a close match between their calculated 5•80 clays,chlorites, and epidoteswhose compositions are listed in valuesand the measured5•80 of Taylor andForester [1979] Figure 10, the chemicalformulas for all mineralsconsidered in for the principal rock types in and around the intrusion. Al- this study are standardend-member compositions; they are all though that study was quite successful,particularly in predict- given in the work of Helgeson et al. [1978] and Wolery ing permeabilitiesof large massesof rock in the earth's crust, [1978]. Wolery's data indicate that clays remain stable at tem- neither it nor the study by Parmentier [1981] attempted to peraturesas high as 350øC,although mineralogicalstudies of account for any secondary alteration minerals, or take into oceanic rocks indicate the occurrence of chlorites rather than account the consequencesthat formation of large amounts of clays at high temperatures.Thus in the present work, forma- 12,586 BOWERSAND TAYLOR: CHEMICALAND ISOTOPICMODEL OF HOT SPRINGS

TEMPERATURE (øC) thermodynamic data base we are using for this work has al- 400 200 I00 50 25 0 ready been tested extensivelyin theoretical calculations, and it 4o typically yieldsremarkably good agreementwith observations. For example, theoretical modeling of porphyry copper ores in Butte, Montana by Brimhall and Ghiorso [1983] and Brimhall [1980, 1979], and of serpentinizationof peridotites by Janecky [1982], and water-sedimentinteraction in Guaymas basin by 30-- Bowers et al. [1985] has provided an excellent match with observedmineral assemblages.

Isotopic Fractionation Factors The integrationof 180/160 and D/H isotopemass distri- butions into EQ6 requires an additional data base composed of the equilibriumfractionation factors for x80/160 and D/H exchangeamong minerals and water as a function of temperature. Mineral-mineral and mineral-H20 fractionation curves for oxygen isotopes are well known for a large number of geologicallyimportant systems.Experimentally determined •J and calculated180/160 fractionationfactors (through 1976) for 180 exchangebetween H20 and calcite,dolomite, anhydrite, quartz, K-feldspar, plagioclase,muscovite, paragonite, biotite, magnetite, pyroxene, garnet, amphibole, and elivine are summarized by Friedman and O'Neil [1977]. More recent papers include Kulla and Anderson[1978] (kaolinite), Chiba et ¸-- al. [1981-] (anhydrite), and Matthews et al. [1983a, b, c] (quartz, albite, anorthite, jadeite, diopside, wollastonite, calcite, and zoisite). Empirically determined180/160 frac- ticnation factors are available over limited temperatureranges for kaolinite, serpentine,chicrite, gibbsite, illite, and smectite [Savin and Epstein, 1970a, b; Wenner and Taylor, 1971; Law- 2 4 6 8 I0 12 rence and Taylor, 1971; Eslin•ler and Savin, 1973; Yeh and 106/T2 (t,<2) Savin, 1977, 1976; O'Neil and Kharaka, 1976; O'Neil and Taylor, 1969]. Oxygen isotope fractionation curves employed Fig. 1. Experimentally and empirically determined equilibrium in this study (Figure 1) can be representedby oxygen isotope fractionation curves employed in our calculations, plotted as a function of temperature. Quartz-H20, Clayton et al. [1972]; anhydrite-H20, Chiba et al. [1981]; magnesite-H20, Perry and Tan [1972]; calcite-H20, O'Neil et al. [1969]; albite-H20, O'Neil 103In 0•min_H20 =a + b(-•) and Taylor [1967]; smectite-H20, Yeh and Savin [1977]; muscovite-H20, O'Neil and Taylor [1969]• zoisite-H20, Matthewset • •min-- •H20'-- Amin- H20 (1) al. [1983c]; anorthite-H20, Matthews et al. [1983b]; kaolinite-H20, where b = 0 for the curves which plot as straight lines. Be- Kulla and Anderson [1978]; serpentine-H20, Wenner and Taylor [1971]; magnetite-H20, Becker[1971]. causea varietyof simplerelationships exist between x80/160 fractionation and mineral structures[e.g., Taylor and Epstein, 1962; O'Neil et al., 1969], minerals for which no experimental tion of smectites has been suppressed at temperatures of 180/160 data are availablecan be assignedthe fractionation 200øC and above. curves of chemically or structurally similar compounds. Al- In addition to the above-mentioned thermodynamic data though the calculationspresented below involve extrapolation for minerals, recent studies have provided solid-solution of severalcurves beyond the temperature range for which they models for some of the important minerals, e.g., epidote [Bird were determined, differencesof only 1 to 3 per mil arise for and Helgeson, 1980] and chlorite rStoessell,1984]. However, eventhe most poorly defined A18Ornin_H20 values; such uncer- the EQ3/6 computer program does not contain adequate pro- tainties make little difference in the outcome of the calcula- vision for solid solutions at the present time. Inclusion of tions. nonideal solid solution models in the calculations would im- D/H fractionation factors of mineral-H20 systemsare less prove the predictiveability of the models. well-definedthan thosefor oxygenisotopes. Several curves are Thermodynamic data for water and aqueous speciesare shown in Figure 2 for micas and amphiboles [Suzuoki and taken from Helgeson and Kirkham r1974a, b-I, Helgeson et al. Epstein,1976], kaolinite, illite, montmorillonite, gibbsite[Lam- [1981], and Helgeson [1969]. Additional data for aqueous bert and Epstein, 1980; O'Neil and Kharaka, 1976; Lawrence complexes have been compiled by Wolery and he supplies and Taylor, 1972, 1971; Savinand Epstein, 1970a], epidote and supporting data files for EQ3/6 which cover a temperature zoisite [Graham et al., 1980], and serpentine[Sakai and Tsut- range of 0ø to 300øC(extended to 350øC where data are avail- sumi, 1978; Wenner and Taylor, 1973]. Some of these curves able or extrapolated if necessary)at a pressurecorresponding and individual data points are incompatible (e.g., to the water-steam saturation curve. All calculations in the kaolinite-H20 and chrysotile-H20). These problemsare a represent work were carried out at 500 bars. Thermodynamic flection of the difficulty of the experiments,the large effects data for the complexeswere generatedat 500 bars by applying that changesin chemicalcomposition of a mineral have on its the isodielectriccorrection suggestedby Helgeson[1969]. The D/H fractionation and the ambiguity associated with hy- BOWERSAND TAYLOR:CHEMICAL AND ISOTOPICMODEL OF HOT SPRINGS 12,587

TEMPERATURE, oc

700 500 300 200 150 I00 50 25

20-•_ o 0

IO -' --

r.,,o-I0 - KAOLINITE I

• KAOLINITE • )_ •,,•-30 o EPIDOTE(GSH) / ",i•KAOLINITE(OK) L -4o - •TE (OK) -- • \% , 0 M-o•T"MORJLLO•TE(OX) g -50 E3 L• (WT --

.• -60 -- --

0 B•OEHMITE(SE) MONTMORILLONITE -7o- • • (SEE)- -80-- • 0 DIASPORE(GSH) -- -9o 0 I 2__ 3 4 5 6 7 8 9 I0 II I:::' IOe/T• (K-•) Fig. 2. Summaryof experimentaland empiricaldeterminations of equilibriumhydrogen isotope fractionation curves as a functionof temperature.Size of the box symbolsrepresents approximate error. SNT, Sheppardet al. [1969]; SaE, Savinand Epstein[1970a-I; LT, Lawrenceand Taylor [1971]; WT, Wennetand Taylor [1973]; OK, O'Neil and Kharaka [1976]; SE, Suzuokiand Epstein[1976]; ST, Sakaiand Tsutsumir1978,1; GSH, Grahamet al. r1980]; LE, Lambertand Epstein [1980].

drogen occurrence in more than one structural OH site in discussapproximate fractionationfactors for kaolinite-H20 someminerals, e.g., chlorites. between25 ø and 400øC obtained empirically from kaolinites Becauseof the above problems,we somewhatarbitrarily in hydrothermally altered rocks of the Vailes Caldera. These developeda consistentset of fractionationfactors for D/H have been substantiatedexperimentally by Liu and Epstein fractionation betweenminerals and H•_O,requiring exclusion [1984]. Below230øC the curveis linear with respectto 1IT 2 of someof the availabledata. The curvesused in thisstudy are (equation(2)). The curve then goesthrough a maximum at a shown in Figure 3, and their positionson the diagram are temperaturearound 230øCand has a positiveslope between basedprimarily on the observationsof Suzuokiand Epstein 230ø and 400øC.This ties in with the positiveslope exhibited [1976], Lambert and Epstein [1980], and Liu and Epstein at temperaturesbelow 400øC for the micas and amphiboles. [1984]. Between 450ø and 800øC these fractionation factors Deuteriumfractionations between minerals and water can ap- are approximatelylinear on a plot of 1IT 2, and the D/H parently be bestrepresented by curvesthat changeslope twice, fractionationsrelative to H20 for two minerals of constant with a negativeslope at both high and low temperaturesand a but differingcomposition plot as parallel lines (i.e., the frac- positive slope at intermediate temperatures. The high- tionationbetween the two mineralsis independentof temper- temperaturereversal in slope occursclose to 374øC, the criti- ature).These relationships can be expressedas cal temperatureof water. We have chosento representthe fractionationfactors of all the OH-bearing mineralsexcept epidote as three straight-line 103In0•min_H20 •---a+ •,'•Tj (2) segmentswith a maximum at 245øC and a minimum at 374øC which is the sameexpression used for most •sO/•60 frac- (Figure 3). This representationresults in largerinaccuracies in tionation curves.At temperaturesbetween 450 ø and 400øC the the fractionationcurves in the temperatureneighborhood of curvesdeviate from the lineardependence on 1IT2 and a pos- the extrema, but it has little effect on the outcome of the siblereversal in slopeis indicated.Lambert and Epstein[1980] calculationspresented herein, because the important feature of 12,588 BOWERSAND TAYLOR: CHEMICALAND ISOTOPICMODEL OF HOT SPRINGS

TEMPERATURE, øC 1976]. We have thereforeassigned Mg-chlorites to the serpen- 500 300 200 I00 50 25 4O tine curve and Fe-chlorites to the annite curve.

2o MATHEMATICS OF INCORPORATION OF OXYGEN AND HYDROGENISOTOPES IN EQ6 In contrast to the simplisticway in which isotopic modeling of hydrothermal systemshas been done in the past, in this -20 study we wish to take into account as many aspects of the formation of secondary mineral phases as possible. EQ3/6 -40 allows us to determineexactly how the production or dissolu- .3 tion of each mineral will affect the overall transfer. For example, if there is an approximately 1:1 replacementof a primary • -60 4 mineral by a secondaryphase for which the A•80 fractionation happensto be zero, there obviously will be no net changein the •xso of the fluid.As water-rockratios are often calculated on the basis of oxygen isotope exchange without -- -I000 2 4 6 8 I0 12 106/T2 (K-2) regard to the effectsof secondarymineral formation, incom- plete or erroneousinformation may result. Fig. 3. Approximated equilibrium hydrogen isotope fractionation Water-rock ratios have often been calculated from oxygen curves employed in our calculations, based on the data shown in isotope analysesof whole-rocks and fluid, using a simplified Figure 2. The curve labeled 1 correspondsto the fractionation between H20 and muscovite, paragonite, beidellites, margarite and material-balance procedure [Taylor, 1974, 1977], where for a prehnite; 2, kaolinite, phlogopite, saponites, tremolite and talc; 3, closedsystem chrysotile,clinochlore, amesite and antigorite; 4, annite, minnesotaite, nontronitesand daphnite; and 5, epidote. •rockf -- •rocki W/R= gH20'--(•rock s-- A) (3) the curves is that the aqueous fluid concentratesdeuterium where A = grockS--gH20s. This model requiresadequate relative to OH-bearing minerals throughout the temperature knowledgeof both the initial (i) and final (f) isotopicstates of range of interest, and the typical fractionations are on the the system. Taylor [1977] suggeststhat it is a reasonable order of 20 to 60 per mil. The epidote curve determined by approximationto set g•SOrockequal to rS•soof plagioclase. Graham et al. [1980] and used for this study showsno temper- This procedurewas followed by the East Pacific Rise Study ature dependenceabove 260øC and a large slope between260 ø Group [1981] in an estimation of water-rock ratios for the and 150øC. Note that the hydroxyl position in the epidote 350øChydrothermal fluids venting at 21øN on the EPR. They structure is considerably different than that of most of the calculatewater/plagioclase of about 1.7 which gives W/R • 1. common silicate minerals. However, at the relatively low temperaturesof the EPR solu- Figure 3 exhibits four parallel curves displaying two ex- tions (< 350øC),secondary alteration mineralsare ubiquitous; trema. The top curve corresponds to muscovite (pure A1- therefore,approximating the alteredbasalt as plagioclasegives endmember of the mica series), the second curve to both a poorrepresentation of 518Orock. phlogopite (pure Mg-endmember) and kaolinite, the third to Possiblemechanisms of isotopicexchange between minerals serpentine and magnesium chlorites, and the lowest curve to and H20 include solution-precipitation(rccrystallization) pro- annite (pure Fe-endmember)and iron-chlorites.In the temper- cessesand diffusion. Solution-precipitation has been observed ature range of interest, the kaolinite curve may be the best attending isotopic exchangeby several workers; recent scan- constrained of the four, based on the work of Liu and Epstein ning electronmicroscope and rate studiesby Matthews et al. [1984] and Lambert and Epstein [1980], and the low- [1983a, b, c] demonstrate that solution-precipitationis the temperature portions of the mica curves are assumed to be major mechanism of isotopic exchange in quartz-, parallel to the kaolinite curve. The phlogopite and kaolinite wollastonite-, diopsidc- and zoisite-watcr reactions in the lab- curves roughly coincide so they are shown as one curve. The oratory. Experimental work by O'Neil and Taylor [1967], serpentine curve correspondsapproximately at low temper- Yund and Anderson[1974], and Matthews et al. [1983a], as atures to empirical fractionation factors determined by well as calculations by Cole et al. [1983], indicate that the Wennet and Taylor [1973]; however, at higher temperatures dominant mechanism of isotopic exchange is by dissolu- the Wenner-Taylor curve has been changed to follow the ex- tion/precipitation reactions when the fluids and minerMs are trema pattern establishedby the mica and kaolinite curves. out of equilibrium, but that a much slower, diffusional mecha- Smectitescontain OH-groups in structurally similar sites to nism takes over once chemicalequilibrium has been achieved. those in mica. Therefore, beidellites have been assignedto the For the purpose of modeling isotopic exchangewith EQ6 muscovite curve, saponitesto the phlogopite curve, and non- we chose to ignore diffusional mechanisms and concentrate tronites to the annite curve. only on solution-precipitation.EQ6 utilizes stepwisedissolu- Chlorites present a problem in that there are no empirically tion of reactantsthat remain out of chemicalequilibrium with or experimentally determined fractionation factors, and they the rest of the system(fluid + product mineral). If reactants show a wide range of A1 contents and contain OH-groups in are also out of isotopic equilibrium with the system, their two different structural sites;this causesdifficulty in predicting dissolutionresults in unequal additions of each isotopeto the the fractionation curves. Because of the two distinct OH sites massbalances of the equilibrium system.Subsequently formed and the resulting differencesin bond strength, chlorites that product minerals precipitate in isotopic equilibrium with the are compositionally similar to micas will probably exhibit fluid. slightly lower •D valuesthan the micas [$uzuoki and Epstein, Conservation of mass of the isotopes of an element is a BOWERSAND TAYLOR: CHEMICAL AND ISOTOPICMODEL OF HOT SPRINGS 12,589 logicaladditional step to theconservation of charge and mass no' refers to moles of oxygen in reactants added in this in- of each element that currently serve as constraints for the cremental step only, non refers to moles of oxygen in product mass transfer calculations performed by EQ6. If the initial phasesprecipitated prior to thisstep (at •ø),and t5•80.• o and fluid is undersaturated with respect to all possible products, t5•80n retaintheir values calculated from (7) and(8). A new after the first increment of reactants has been added to the value of t518Os(•)is computedfrom (9) and used in (7) to system,the r5•80 of the systemat • (reactionprogress vari- redistribute the isotopesamong product phases(including the able) can be expressedas current increment to the products) and fluid. Note that in the casedescribed above the entire product phaseis reequilibrated F•OH20 1 $ t•18OH20(•ø) + • no' 80, (4) with the fluid at each step in the reaction progress,resulting in •• 80s(•) = no nos its uniform isotopic composition.In contrast, products will be isotopically zoned if only the additional increment to the where product phase at a given step of reaction progress is con- nos = F•OH20 + • no' (5) sidered to be in isotopic equilibrium with the fluid. The small chemical affinities of isotopic reactions suggestthis likelihood [Walther and Rye, 1982]. However, for simplicity we have whereno "2ø represents the molesof oxygenin the fluid phase, assumed uniform isotopic composition of products because no' is the molesof oxygenin reactantphase r, and nos denotes molesof oxygenin the entiresystem. •o representsone •-step very little zoning of isotopes would develop in an isothermal reaction. before the present value of •, in this case, the initial state of The underlying assumptionsof this model are that isotopic the systemprior to any dissolutionof reactants. Note that (4) equilibrium is achieved as fast as chemical equilibrium, and and (5) together with the following equations can readily be that thereis no net flux of •80 eitherinto or out of the system, written for hydrogen (or other) isotopes as well, but for the other than that achieved through dissolution of reactants or sake of simplicity the procedure is shown only once, for removal of product phasesfrom the system.This also signifies oxygen. As longas no productsform, t5•80 of the fluid registersonly that isotopic exchangewith undissolvedreactants (in this case, the responseto dissolutionof the reactants,and r5•8Os(•) = mineralogically unaltered basalt) is not considered,i.e., no dif- rSx8OH2o(•). As soonas productsform, rSx80 s must be redis- fusion. In other words, if there is no recrystallization, there is no isotopic exchange; at the relatively low temperatures con- tributed betweenthe fluid and theseproduct phasesaccording to sideredhere, this seemsto be a realistic assumption. A further assumption is that the composition of the fluid 1 has no effect on the degree to which it fractionates isotopes. t518Os(•)= nos nos Taube [1954] and Truesdell [1974] have in fact shown that nøH2øtS'8OH:o(•)+ • • noVtS•8On(•) (6)this is not strictly true. The value of A solution-water for 1 to 4 molal solutionsof KC1, CaC12and MgC12, for example, may where nos has the samevalue as in (4) but noH:ø has been vary from approximately -2 to +2 per mil. However, this modifiedby the precipitationof products,and g•8OH:o and variance shows no clear relationship to temperature and is g•80n at • mustnow be calculatedfrom equilibriumfrac- therefore difficult to incorporate mathematically into the tionation curvesfor the product minerals and H20. Equation model; in any case NaC1 is the most worrisome impurity in (6) can be rewritten as natural hydrothermal fluids, and addition of this salt has a •OH20 1 negligible effect [Taube, 1954; Kendall et al., 1983-1,contra- a18Os(•)= • tlOs tS•8On•o(•)+•F/O s • nøn dicting the observationsof Truesdell [1974]. ß[(1000 + t•18OH20(•)) 0•n_H20 -- 1000] (7) MODEL CALCULATIONS OF SEAWATER-BASALT INTERACTION where Startinq Compositions 1000+a18Op(•) (8) The chemical composition of seawateris well known (Table %,-H20= 1000 + •18OH20(•) 1' see Von Damm et al. [1985]). Seawater has a high pH Equation(7) canbe solvedfor r51sO.2o(•),and rS•8On(0 is (--• 7.8), is supersaturatedwith respectto dissolvedoxygen, and subsequentlycalculated from (8). at 2øC (temperature of deep ocean water) is also super- The next incremental step in reaction progressinvolves ad- saturated with respectto dolomite and quartz. These minerals dition of isotopes through dissolution of reactants to an equi- are not precipitating from seawater at 2øC, presumably be- librium systemalready containing products. If the entirety of cause of kinetic inhibitions. Oxygen and hydrogen isotopic these products is retained in isotopic equilibrium, redistri- 5-values of deep ocean water are very uniform at zero per mil, bution of the isotopesis calculatedin the following manner: relative to SMOW. We have chosento use the same basalt starting composition •OH20 employed by Wolery [1978]. The composition shown in Table •18Os(•) = a18Omo(• o ) •O s 2 correspondsto sampleV25-RD1-T3, a fresh basalt from the Mid-Atlantic Ridge reported by Miyashiro et al. [1969], modi- + - • no•a + • noVa•8Ov(•ø) (9) fied to contain 1100 ppm sulfur. This value is midway in the •O1 s 180••O1 s • range from 680 to 1800 ppm described by Moore and Fabbi [1971] and Mathez [1976] for fresh basalt glasses,and is se- where lected based on a linear relationship between Fe and S concentrations proposed by Mathez [1976]. Typical fresh MOR nos= no"2ø + • no'+ • non (10) basaltshave very uniform 518 0 = + 5.7 _+0.2. r p 12,590 BOWERSAND TAYLOR: CHEMICAL AND ISOTOPIC MODEL OF HOT SPRINGS

TABLE 1. Chemicaland Isotopic Composition of Seawater,of theCalculated Hydrothermal Fluids at Various Temperatures,and of the 350øCEast Pacific Rise (21øN) Hot SpringFluids

Elements Concentration,mmol/kg or Species Seawater (2øC) 100øC 150øC 200øC 250øC 300øC 350øC EPR Fluid

Na 463. 476. 484. 503. 527. 542. 607. 430-510 C1 540. 541. 541. 545. 550. 553. 574. 490-580 K 9.8 0.743 2.20 4.96 9.22 15.1 29.1 23-26 AI 2.0x 10-5 2.1x 10-3 0.134 0.716 2.63 5.00 0.179 0.004-0.005 Si 0.18 0.471 2.11 3.85 6.14 7.27 11.6 15-20 Mg 52.6 2.43 5.2 x 10-2 4.8 x 10-2 2.1 x 10-2 1.1 x 10-2 9.8 x 10-2 0 Fe 1.5x 10-6 4.4 x 10-3 4.3x 10-• 2.6x 10-3 2.2x 10-3 0.191 63.2 0.7-2.5 Ca 10.2 48.3 33.4 18.9 8.11 1.04 0.774 11-20 SO d 28. 18.5 5.69 1.5 x 10-9 0 0 0 0 H2S 0 1.7 x 10-• 6.9 x 10+ 3 0.419 2.31 5.72 14.2 6-9 pH 7.8 6.38 6.31 5.99 5.80 5.98 5.67 3.3-3.8 logfo2 - 1.24 - 51.5 -45.0 -44.4 -40.2 - 36.3 - 32.1 - 30.9 •80 0.0 -0.126 -0.202 -0.107 0.145 0.508 2.02 2 rSD 0.0 0.021 0.022 0.200 0.406 0.721 2.67 2.5

Pathlinesin Open VersusClosed Systems unstable,all under isothermal(100øC) conditions.Once the In nature, the interaction of basalt with heated seawater specifiedamount of freshbasalt and newly-formedalteration takes place as neithera perfectlyopen systemnor a perfectly mineralshave completely reacted, all solidphases are removed closed system in the sensedefined above in the discussionof from the systemand the packetof exchangedseawater (which the EQ6 model. A packet of seawaterbeing heated and react- is now somewhatmodified both chemicallyand isotopically, ing with fresh basalt at some temperaturecan perhapsbe and whichmay be lessthan one kilogramas a resultof hy- envisionedin this way: as temperatureincreases with depth dration reactions)is heated up to 150øC.No reaction with and proximity to the ridge axis, this packet of seawaterwill basaltis consideredduring the heatingstage. However, note presumably move along a pathline downward and inward that simplyheating the solutionfrom 100ø to 150øCmay toward the heat source(MOR), and along the way it should result in supersaturationsof some minerals; in our model encounter some relatively unaltered basaltic material. On the theseare precipitatedat the beginningof reactionwith basalt other hand, it will also encounterbasalts altered to varying at 150øC,implying that the concentrationsof certainspecies in degrees by earlier packets of seawater that have passed solutionat the initiationof the 150øCstep may be slightly through the system.The assumptionis made here that EPR- differentfrom their valuesat the end of the 100øCstep. At type hydrothermal fluids result h'om interaction of seawater 150øC,after the chemicalspecies in the fluid are redistributed with basalt along a steep temperaturegradient where fresh accordingto the new equilibriumconditions, the packet of basaltis beingcontinually supplied to the system(i.e., by mag- (modified)seawater is againallowed to reactisothermally with matism at the ridge axis). In contrast,off-axis hydrothermal freshbasalt in a closedsystem. circulation probably entails low-temperaturealteration of ba- In the above scenario,seawater reacts isothermally with saltspreviously altered at higher temperatures.This latter sce- fresh basalt at each incrementin temperature.We continue nario is not consideredin the presentstudy. this procedurein stepsof 50øC to the final temperatureof Whatever the valuesof parameterssuch as temperatureor 350øC. Back-reactionof the exchangedfluid with alteration water-rockratio, two things seemcertain. (1) The alteration productsoccurs during eachisothermal stage of the reaction; processshould not be modeled simply as a closedsystem however,at the conclusionof reactionat each temperature undergoinga stepwisetemperature increase where all newly step,all secondaryminerals are removedfrom the equilibrium formed alteration products are available for back-reactionat system.This correspondsto a progressivelydecreasing water- higher temperatures.(2) It should also not be modeled as a rock ratio with increasingtemperature (more precisely, it is an "simple,"or perfectlyopen systemwhere all secondaryminerals formed are instantaneouslyremoved from the system, TABLE 2. CalculatedChemical Compositions (wt %) of the with no back-reactionwhatsoever being allowed. Utilizing the PredictedHydrothermally Altered Basaltsat the End-Points above constraints,together with the capabilitiesand limi- of the VariousTemperature Steps Described in the Text tations of the EQ6 model, the following path of seawater- basalt interaction is used as the basis for the calculations. Oxide 100øC 150øC 200øC 250øC 300øC 350øC V25-RD1-T3 Interaction of unaltered seawater and fresh basalt is as- SiO 2 47.57 47.25 49.21 49.25 49.54 49.80 49.87 sumedto begin at 100øC,and an arbitrary amount of basaltis A1203 15.00 14.95 15.50 15.49 15.56 15.71 15.70 dissolvedisothermally into onekilogram of seawaterusing the FeO 0.62 0.63 6.54 7.75 7.38 7.93 8.24 EQ6 closed system model. Note that this solution- Fe203 9.59 9.57 3.31 2.04 2.49 1.51 1.40 MgO 15.09 8.25 8.27 8.28 8.32 8.36 8.37 redepositionprocedure requires that the basalt reacts homo- CaO 2.71 12.56 10.90 10.81 10.78 10.65 10.66 geneouslywith seawater.At present,no provisionis made for Na20 1.42 2.14 2.68 2.63 2.76 2.81 2.93 the probability that somecomponents of the basalt will alter K20 1.69 0.0 0.15 0.12 0.08 0.15 0.21 more rapidly than others,or that somecomponents will un- H20 4.16 2.40 3.26 3.55 3.02 2.99 1.00 dergo isotopic exchangemore easily. Product minerals form S 1.18 1.27 0.18 0.08 0.07 0.09 0.11 as a resultof the 100øCinteraction, and they may back-react SampleV25-RD1-Te [Miyashiroet al., 1969] representsthe unal- at any given stagein the reaction progressthat they become tered basalt. BOWERSAND TAYLOR' CHEMICALAND ISOTOPICMODEL OF HOT SPRINGS 12,591

increasing rock-water ratio, in the way the titration in the I00 øC • computer calculation is actually carried out, as more and more freshrock is added to the initial kg of seawater).Predic- tions of the model are dependenton the amount of rock inter- IøøI %7 acting with fluid at each temperature,and this parameter can ao be varied in different models.Also, different paths of temperature and water-rock ratio can be tested for agreementwith the analytically determined chemical and isotopic compo- - sitionsof the EPR hydrothermalsystem. The computationalprocedures outlined above representa realistic compromise between the constraints of the EQ6 models and actual MOR hydrothermal systems.Our calcula- • 2o tions produce a vertically stratified oceaniccrust in which the grade of hydrothermalmetamorphism increases downward, as i o it should.The model also has the advantagethat at the start- -7 -$ -5 -4 -2 0 ing point of each successivetemperature step, the only reactant material that has changedits chemical and isotopic composition is the fluid. This means that the reader can more Fig. 4. Calculatedvolume percent of alterationphases predicted readily visualizewhat is going on at each temperature interval to form as the result of seawater-basaltinteraction at I•øC, plotted as a function of log reactionprogress where log • = 0 correspondsto than if we had substituted some arbitrary, path-dependent, 216.0 grams of basalt dissolvedin 1 kg of seawater. PY, pyrite; altered basaltic material at the start of each step. For example, NONT, nontronite; PA, paragonite. by the time one had traveled through several temperature steps in any of several more complex types of temperature- composition paths that one might imagine, it would be diffi- quently, at least at 21øN it appears that, even at depth, the cult not to lose sight of which parametersare most important fluids were never much hotter than 350øC. in producing specificcause-and-effect relationships. There are Further support for our assumedrange of temperatures so many variables and so many possiblepaths that even if a comesfrom the estimatesof Lister [1982] and Mottl [1983]. good isotopic and chemical match were to be made between Also, Norton [1984] has shown that the transportproperties calculation and observation,it would be difficult to singleout of pure H20 (buoyancy, viscosity,etc.) go through extrema the critical features of the models. near the critical point (374øC); this leads to relatively rapid Temperaturesof Reaction convectionat suchtemperatures, explaining why this is a commonly observedupper limit in exploredhydrothermal systems We have made a somewhatarbitrary selectionof 100øCas a (e.g.,EPR, Salton Sea, Larderello). starting temperature for our calculations, even though data Although we are assumingin our model that the MOR such as those of Bb'hlkeet al. [1984] and Stakes and O'Neil fluidshave beenat maximumtemperatures on the order of [1982] document formation of alteration phases at temper- 350ø-365øC, this does not apply to all submarine hy- atures below 100øC. We can do this because the effect of drothermal systems,and indeed there is evidence that much low-temperature interaction on a given packet of seawater is higher-temperature, sub-seafloor hydrothermal circulation small as a result of kinetic factors and the high water-rock must exist. For example, Gregory and Taylor [1981] estimate ratios involved at the lower temperatures.Only minor errors that hydrothermalcirculation locally occurredat a depth of are introduced by neglecting interaction at temperatures more than 5 km in the Samail Ophiolite, at temperatures below 100øC in our calculations. greater than 500ø-600øC,similar to valuesestimated by Stakes An upper limit of approximately 350øC is imposed on our et al. [1984] for gabbrosfrom the Indian Ocean. Gregory and model becauseof present constraintson the thermodynamic Taylor [1981] and Bischoffand Rosenbauer[1984] suggest data base, but this limit seemsappropriate for several other that the deeper penetration and higher temperaturesoccur reasonsas well. The exiting solutionsat 21øN on the EPR are predominantlywithin the layered gabbrosand along the sides the highest-temperaturefluids (350øC _+5 ø) yet measured at and beneath the axis magma chamber, and therefore are not a midocean ridges.Recently, higher temperatureshave been re- part of the venting axial systemsuch as that from which the ported by Delaney et al. [1984] and Kim et al. [1984], but 21øN fluids are being collected. these measurementshave yet to be substantiated. Von Damm et al. [1985] calculatetemperatures at depth for the EPR vent RESULTS OF THE COMPUTER MODEL fluids assuming adiabatic cooling and using silica concentration in the fluids as a geobarometer.This procedure General Statement indicates a depth of reaction of 0.5 to 2.0 km beneath the There exist essentially an infinite number of possible seafloor,which correspondsto maximum temperaturesbefore temperature/water-rock ratio paths, only a few of which are adiabatic cooling of approximately 355ø-365øC. Von Damm et consideredin the presentstudy. However, most of thesepossi- al. [1985] note that ocean bottom seismicdata at 21øN [Rie- ble paths can be rejected becausethey do not result in the desel et al., 1982] indicate that most seismic events occur at simultaneous occurrence of the necessary mineralogical, about these same depths, and this is interpreted to be the chemical,and isotopic characteristicsactually observedin the maximum depth of large-scalehydrothermal circulation. The basalts and hydrothermal fluids. Thus, after severaliterations rapid exit velocitiesproposed by MacDonald et al. [1980] for and adjustments,we settled on a particular path and final vent fluids suggest that the fluids could cover the distance end-point that fits most of the available observational data; from depth in a few minutes, making it highly unlikely that the resultsare shown in Figures4 to 20. The path was initiat- substantial heat would be lost during the ascent. Conse- ed with one kilogram of seawaterheated to 100øC.26.6 grams 12,592 BOWERSAND TAYLOR: CHEMICAL AND ISOTOPICMODEL OF HOT SPRINGS

150 øC 250øC IOO I00 i • I /' I i r • MAGNETITE/ • CHLORITE • 8¸ - -

• 60-

o I,-.. I i I i i I a-• I , -7 -6 -5 -4 -5 -2 -I 0 o-7 -6 -5 -4 -5 -2 o log log • Fig. 5. Calculated volume percent of alteration phasespredicted Fig. 7. Calculated volume percent of alteration phasespredicted to form at 150øC from interaction of basalt with a fluid of compo- to form at 250øC from interaction of basalt with a fluid of composition correspondingto the results of the 100øCcalculation shown in sition correspondingto the results of the 200øC calculation shown in Figure 4, plotted as a function of log reaction progress.PY, pyrite; Figure 6, plotted as a function of log reaction progress.The dashed HM, hematite; PARAG, paragonite. curve representsvolume of alteration minerals at any stage as a frac- tion of the total volume of alteration minerals produced during the entire isothermal reaction interval. of fresh basalt were dissolved into the seawater at 100øC, 32.5 grams at 150øC, 216.0 grams at each of 200ø, 250ø, and 300øC, and finally, 1080.2 grams were added at 350øC. A total of the oceanic crust, if there is continued flushing of the bulk 1787.4 grams of basalt were thus added to the kilogram of oceanic crust by such fluids, up to the characteristicwater- seawaterover the entire temperature range, resultingin a final rock ratio derived for each temperature-depthstep. (overall) water-rock ratio of approximately0.56 (weight units). Results of the model calculationsare dependent on the end- It is important to understand that although we are con- point of each temperature step, and selection of these end- cerned specificallywith calculating a plausible pathline for a points thereforeinvolves the use of severalcriteria. At 350øC, particular fluid packet that circulates downward close to the the model reactionswere stoppedwhen 6•80 of the fluid roof of a MOR magma chamber and which then exits at reached the observedvalue of + 2.0 per mil, and if the predic- 350øC on the seafloor, the end-points of the calculations at ted 6D of the fluid, the solution chemistry,or the mineralogy, each intermediate temperature step also provide insight into was not substantially in agreement with MOR systems,the the overall alteration of the oceanic crust. For example, a entire temperature-pathwas consideredto be inadequate.The significant proportion of the sea water that is heated to 150øC input parameterswere then modified and the calculationsre- or 200øC will migrate either upward or laterally, possiblyexit- peated until substantialagreement was obtained. However, no ing at the sea floor in off-axis effluent;only a small quantity isotopic constraintsexist for any of the intermediate temper- will circulate to still greater depths and ultimately exit at ature steps,so other criteria that were used included relative 350øC from a MOR vent. The whole-rock isotopic compo- changesin isotopicvalues of the fluid, direction of changein sitions and the mineral assemblagesat the end-points de- the pH, and reasonablenessof predicted mineral assemblages. scribedbelow will adequatelydescribe each depth-segmentof For example,at 100ø and 150øC•180.20 is decreasing,and

200øC 300øC IOO I001MArGArinE 80 ol-

CLINOZOISITE - 60

40-

20-

0 0 I -7 -5 -4 -2 o -7 -6 log( Fig. 6. Calculated volume percent of alteration phasespredicted Fig. 8. Calculated volume percent of alteration phasespredicted to form at 200øCfrom interactionof basalt with a fluid of compo- to form at 300øC from interaction of basalt with a fluid of composition correspondingto the results of the 150øC calculation shown in sition correspondingto the resultsof the 250øC calculation shown in Figure 5, plotted as a function of log reaction progress.PY, pyrite; Figure 7, plotted as a function of log reaction progress. TREM, HM, hematite; TREM, tremolite; QTZ, quartz; AB, albite. tremolite. BOWERSAND TAYLOR: CHEMICAL AND ISOTOPIC MODEL OF HOT SPRINGS 12,593

IOO 25.0 I I ALBITE--•_I • I I Ill ANALCIME TE

z 20.0 I

Ld•I•• 80-60- ARAGONIT• .J --- • l"¾'40 -- %1/,.-,..,-•-• N CHLORITE -- _ 150ø • • • 20- c.•ozo•s•c _ - 200ø 150ø ANORTHITE STARTINGMATERIAL• -5 -4 -5 -2 -I 0 log ( 3500 250•, o FiB. 9. Calculatedvolume perce•t oFalteratJo• •bases predicted 250ø 00ø• to Format 350øCFrom J•teractJo• oF basalt wJtNa •uJd oFcom•o- 0.0-- sJtJo•corres•o•dJ• 8 to tNe resultsoF tNe 300øC calculatio• sbow• i• FiBure 8, plotted as a Fu•ctio• oF lo8 reactio• •rosr•ss. phlo8opite;CNL, cblorite;(A), amesite;(D), dapN•ite;PAR, paras- o•ite; •TZ, quartz; EP, epidote. --6.0 -5.0 -4.0 -:3.0 -2.0 -I.0 0.0 1.0 ß LOG (' Fig. 11. Calculated6t sO whole-rockof the predictedmineral as- pH at high valuesof reactionprogress begins to increaserap- semblagesshown in Figures 4 to 9 as a function of log reaction idly; neithertrend is desirable.In addition,at 100øCepidote is progress.Contours are temperature in øC. The horizontal dashed line predictedat high log • values,and becauseepidote does not at 6•80 = + 5.8is theassumed tsO/t60 ratioof theinitial basalt. typicallyform at temperaturesas low as 100øC,the end-point can be justifiably assignedto a lower value of log •. The end- crease pH. Reaction at 350øC, as noted above, is allowed to point at 150øCis set right beforethe onset of a very rapid increasein pH. proceed only as far as the log • value correspondingto Reaction at 200ø, 250ø, and 300øCis uniformly stoppedat 6•80.,.o- + 2.0per mil. It is importantto notethat in addition to the major set of calculationspresented in Figures4 to log • = 0, after the developmentof epidote-tremolite-chlorite- 20, many "dead-end"paths with different end-pointsat each albite assemblages.There is little point in advancingthese temperature step were testedand rejectedduring the course of temperaturesteps to higher valuesof log • becausethe rate of this work. changein 6•80.2o is too slightfor it to reachits ultimate desired value; in addition, further reaction continues to in-

Mg- SAPONITE 250 ø Ca - SAPONITE I r _ ] o Mg-NONTRONITE I i Ca - NONTRONITE I KAOLINITE I / AMESITE -io DAPHNITE 100 ø HEMATITE PYRITE MAGNETITE PYRRHOTITE DOLOMITE 350 ø I r \ ?. CALCITE ...... [...... l..•.'.}'"-4--•5øø---x- MUSCOVITE PARAGONITE -30 PHLOGOPITE ANHYDRITE LIJ EPtDOTE .J CLINOZOISITE o -40 PREHNITE 300 ø ALBITE ANORTHITE TALC (x:) -50 - TREMOLITE QUARTZ MARGARITE

ANALCIME I,•,, -60

Fig. 10. Summaryof predictedalteration phases as a functionof -70 temperature,corresponding to those shown in Figures 4 to 9. Min- -6.0 -5.0 -4.0 -3,0 -2,0 -I.0 0.0 1.0 eralsincluded here but not shownin Figures4 to 9 compriseless than LOG ( 1-2 volume percent of the alteration assemblage.Mg-saponite, Mg.165 Mg3(AI.33Si3.67010)(OH)2; Ca-saponite, Ca.•65Mg3 Fig. 12. Calculatedwhole rock 5D of the predictedmineral as- (A1.33Si3.67Olo)(OH)2;Mg-nontronite, Mg. x6sFe 2 (AI.33Si3.67010) semblagesshown in Figures 4 to 9 as a function of log reaction (OH)2; Ca-nontronite,Ca.•65Fe2(A1.33Si3,670•0 ) (OH)2; amesite, progress.Contours are temperaturein øC. The 200øC curve beginsat Mg,•AI2(A12Si20•o)(OH)8;daphnite, F%AI(A1Si3Oxo)(OH)8; clino- log • •,- -4.6 becauseno H-bearing alteration productsare predicted zoisite,Ca2A13Si3Ot2(OH); epidote, Ca2FeA12Si30•2(OH ). to form earlierin the reactionprogress. 12,594 BOWERSAND TAYLOR: CHEMICAL AND ISOTOPICMODEL OF HOT SPRINGS

I I I I I I I I I field of stability in thesediagrams, Si concentrationsin solution a•c•often at valuesclose to quartzsaturation. Epidote and ,of EPIDOTE - albite are major precipitateslate in reaction progressfrom 150ø to 350øC. Note that the computer program containsno provisionfor solid solutionbetween clinozoisite and epidote, -20-I0 _ and as a resulteither or both phasescan appearin the calcula- • CHLORITES -30 AMPHIBOLES - tions. Talc and its replacementtremolite (servingas a proxy for actinolite)are predictedfrom 200ø to 350øC. -40 The relativesize of the area occupiedby a phasefrom left to right acrosseach of the Figures4 to 9 doesnot representthe -,50 ;:• SAPON ITES - relativeamounts of thosephases formed (because volume per- -60 cent rather than total volumeis plotted on the ordinate).The total volume of alteration phasesincreases approximately -70 •-350o logarithmicallyfrom left to right. The dashedlogarithmic curveshown in Figure 7 displaysthis, whereif 100 equalsthe

i i total volumeof alterationminerals at log • = 0, the volumeof 0 2 4 6 8 •0 12 14 •6 18 20 mineralsformed at any valueof log • < -2 is too smallto be 8180 shown on the scale.For example,the large field occupiedby tremolite at 250øC and low values of log • is somewhatmis- Fig. 13. 6D versus6x80, modifiedafter Figure 1 of Stakesand O'Neil [1982]. Saponite,amphibole, epidote, and chloritefields from leading;it is largemerely because no otheralteration products oceanicdredge hauls are from Stakeset al. [1984] and Stakesand are formingand the smallamount of tremolitetherefore repre- O'Neil [1982]. Rectangularsymbols represent the rangeof calculated sents 100% of the alteration. 6D and 6•80 valuesfor alteration assemblagesshown in Figures4 to Our model allows dissolutionof products during a given 9 for a giventemperature in øC. The high calculated6D valuesrela- tive to observation at 100ø and 150øC result from the lack of con- temperaturestep (see above).This means that a predicted siderationof an Fe-saponiteend memberin the thermodynamicdata base and the sensitivityof 6D of OH-bearing mineralsto changesin Fe/Mg (seetext). TEMPERATURE, øC 350 300 250 200 150 I00

Mineralogy

The calculated modal mineralogy of the alteration assem- 25 blagesproduced at the varioustemperature steps is shownas a function of log reaction progress(log •) at 100ø, 150ø, 200ø, 250ø, 300ø, and 350øC (Figures 4 to 9). A summary of all mineral phasescalculated for the entire temperaturerange is givenin Figure 10. Mineralsshown in Figure 10 but not in the correspondingFigures 4 to 9 were predicted to appear in quantitiesless than 1 or 2 volumepercent and thuscould not be shown with clarity on thesefigures. Seawater is supersaturatedwith respect to dolomite at 100øC,just as it is at 2øC, and as a result dolomite forms a large proportion of the early alteration at 100øC. A small amount of anhydriteforms toward the end of the 100øCcalcu- lation, and it is a major precipitateat low values of reaction progressat both 150ø and 200øC. As reaction progressincreases, the Ca-uptake in the anhydrite and dolomite is switched over to Ca-smectitesand epidote, while SO4 is reducedand pyrite forms.Abundant smectite forms at 100ø and 150øC; at 200øC and above, precipitation of smectitesis (purpoSely)suppressed in the computer model, and chloritestake their place.At low valuesof reaction progress(equivalent to high water-rock ratio) Mg-rich sapon- itc is the primary smectite.As the Mg concentrationdecreases with further reaction, Mg-rich saponiteis replacedby Ca-rich saponite, and this is joined at 150øC by Ca-rich nontronite. Fig. 14. Equilibrium oxygen isotope fractionations as a function Chlorite forms over the entire temperaturerange; Mg-rich of temperaturefor various minerals.The vertical bars representthe chlorite(amesitc) forms at 100ø and 150øC,as well as at low rangeof valuesof whole-rockt5•80 minus 6•8OH2o calculated for the predicted alteration assemblagesshown in Figures 4 to 9 at various values of reaction progressat 300ø and 350øC, and Fe-rich temperature steps.The long-dashedline is drawn through the set of chlorite (daphnite) forms above 200øC, at higher values of vertical bars, and approximately correspondsto the equilibrium reaction progressthan for amesitc. oxygenisotope fractionation between "altered basalt" and H20. Note Muscovite and paragonitc (sericite) both occur over the that the vertical bars at 150ø and 200øC are above the dashed line, a result of the large amounts of anhydrite predicted to form at these entire temperature range, but muscovitedominates at 100øC temperatures.Solid curvesfor anhydrite, smectite,muscovite, epidote and paragonitc at 150ø to 350øC. Quartz is predicted as a and chloriteare taken from Figure 1. The short-dashedcurve refers to precipitateoccasionally, and although it doesnot have a large a smectite-H20 curvedetermined by Cole [1980]. BOWERSAND TAYLOR'CHEMICAL AND ISOTOPICMODEL OF HOT SPRINGS 12,595

pHC• mined for the alteration assemblagespredicted by the model Neutrolity above to facilitate comparison with the chemical changesob- servedin altered basalts,as summarizedby Mottl [1983] and

50 ø Humphris and Thompson [1978]. Table 2 shows the oxide- 6.5 distribution of a rock representingthe alteration assemblage present at the conclusion of each of the isothermal calculations from 100ø to 350øC. The composition of the fresh basalt starting material from Miyashiro et al. [1969] is also shown. The differencein compositionbetween the fresh basalt and the 6.0 altered rock is most dramatic at the 100øC step, whereas at I 350øC the fresh and altered basalts are chemically nearly in- distinguishable.The reason for this effect is that by the time •150 ø that the high-temperature step is reached, the water-rock /--- / ratios are much lower and the aqueous fluid has thoroughly exchangedwith the basalt; thus, most of the chemical changes 5.5 J 1200ø 250 ø in the basaltsoccur at 100ø-150øCand high water-rock ratios. 300 ø At 100øC the altered basalt shows a substantial gain in Mg, considerableloss of Ca, and a larger proportion of oxidized Fe. In addition, the altered rock at 100øC shows some loss of Na, gain of K, and gain of H20. These computer-generated 5.0 I compositions agree quite well with the rocks described by 50 I I i i I Mottl [1983] as chlorite-quartz rich, which he concluded 40 •5oo probably formed at high water-rock ratios. The altered rocks predicted at high temperature (200ø-350øC) are chemically ,-• 30 200ø very similar to the fresh basalt, but show a slight loss of Fe ' ' 20 250ø and a gain in H20 content, in agreementwith the chlorite-

I0 Iøøø 300 ø , quartz poor rocks described by Mottl. Mottl [1983] reports that these rocks also typically show either small gains or -6.0 -5.0 -4.0 -3.0 -2.0 -I.0 0.0 1.0 LOG •' Fig. 15. Ca concentration(mmol/kg) and pH of the evolvingsolution in equilibriumwith mineralassemblages shown in Figures4 to 9, plotted as a function of log reactionprogress. Contours are temperature in øC. The pH of a neutral solution at various temperaturesis shownalong the right-hand-sideof the diagram; note that neutral pH 3 3ooo ,,• steadilydecreases at higher temperaturesup to about 250øC,so that at their end-pointsall of the calculatedfluids are mildly alkaline. I F :"••250 ø ,500_• ,5.oo.•ooo. ol , •• .,,- •,oo alteration product can back-reactand may be destroyedand 60 replacedby anotherphase once it becomesunstable (e.g., anhydrite and amesite at 100øC and 150øC; talc at 200øC to 350øC).However, in reality a given phase is probably only • 20 partially replacedand any of the alterationproducts predicted 0 at a particular temperatureare likely to persist. The secondaryphases shown in Figures 4 to 9 agree well with observationsof alteredbasalts sampled from the seafloor. 25 For example, saponite-richpillow brecciasare common at low temperatures(<200øC) in seawater-dominatedsystems (wa- 3020 ter/rock > 50; Stakesand O'Neil [1982]), and both chlorite- 15 350ø rich and epidote-rich pillow basalts were described by I0 I0øø-"-• Humphrisand Thompson[1978]. Anhydrite is frequentlyde- 300o..-• 5 25øø scribedin samplesfrom activechimneys on the EPR [Styrt et 200 ø al., 1981; Oudin,1983; Haymon, 1983]; however,because of its O• retrogradesolubility it is presumedto have been redissolved by cold seawater, and is thus not commonly observedin 6ooj , , , , • •,/, dredgedsamples. Figure 4 of Mottl [1983] showsan overview z of the mostcommon alteration phases and their relativepro- ' ' 500 portions as a function of water-rock ratio. Note that the rela- 450 tive proportions of chlorite, epidote, actinolite and albite at -6.0 -5.0 -4.0 -3.0 -2.0 -I.0 0.0 1.0 low water-rock ratios shown by Mottl [1983] correspond almost exactly to the relative amounts of the same phases predictedby our computermodel at 350øC,shown in Figure Fig. 16. Na, K, Mg and A1 concentrations(mmol/kg) of the 9. evolvingsolution in equilibriumwith mineralassemblages shown in Figures4 to 9, plottedas a functionof log reactionprogress. Con- Bulk chemistry in terms of the oxides can also be deter- tours are temperature in øC. 12,596 BOWERSAND TAYLOR' CHEMICAL AND ISOTOPIC MODEL OF HOT SPRINGS

20 high at low valuesof reactionprogress, corresponding to the

15 presenceof dolomiteas the major alterationphase (Figure 4). This is attributable to the large equilibrium value of Adolomite_.2o (Figure 1). When significant amounts of saponite and chloriteform, f•80• dropsby severalper mil reflecting

O' the lowerAmin-.2o values of chloriteand smectite.The minimum in the 100øCf•80 curveat log • • -3.2 correspondsto 40 I I I I the point in Figure4 wherethe abundanceof chloritereaches 30- a maximum; this is causedby the fact that the equilibrium 350 c 20 Achlorite_H20value is lowerthan any of the other A values (Figure 1), exceptfor hematite-H20and magnetite-H20(and I0 300 ø these oxide minerals do not form abundant alteration prod- 0 ucts). The valueof f•80• generallydecreases with increasingtemperature,as is expectedbecause of the accompanyingdecrease in Amin_.2o. However, at low valuesof reactionprogress be- I05/L 3oo350o•o •- .._ tween250 ø and 350øC,f•80• showsthe oppositetrend. This ol I I 12•øø ', •-•-. I'•2øøø is againattributable to theidentity of thealteration products; tremolite-H20at 250øChas a lowerA thandoes epidote-H20 at 300øC.The f•80 (and fD) valuescalculated for higher 30/ I i I i • •oi I valuesof reactionprogress (log • > • -3) may be taken as most representativeof alteredbasalts at the corresponding • 201Io •5oo - temperatures,because the alterationis more pervasiveand 01 2øøø .. thusbetter represents an averageof true alteredbasalts than do assemblagesconsisting of onlyone product. f•80• curves ru -3o! , , , • , ,;ooo. 250ø ' -;•ool_._j at 250ø-350øCat log • > • -1 show a moderatelyincreas- ing trendwith increasingreaction progress, and the highest o -50 valueof f•8ORat 350øCis higherthan any value at 300øCand __1-60 I I I I / almostequivalent to thehighest value at 250øC.This trend of -6.0 -5.0 -4.0 -3.0 -2.0 -I.0 0.0 1.0 increasingf•80, withreaction progress along each isothermal LOO ( temperaturestep is attributedto a correspondingincrease in Fig. 17. Log oxygenfugacity and SO4, H2S,Fe and Si con- 6•80.=o, because by materialbalance at 250ø-350øCboth the centrationsof the evolvingsolution in equilibriumwith mineralas- H20 andthe altered rock can increase in f•80 asa resultof semblagesshown in Figures4 to 9, plottedas a functionof log the continuedcontribution to the systemof freshrock that is reactionprogress. Contours are temperature in øC. 5.8per mil higherin x80 thanthe initial seawater. The fid curvesshown in Figure 12 are even lessregular lossesof Mg and small lossesof Ca. Humphris and Thompson than the f•80 curves;again, the variationsmay be attributed [1978] presentsimilar results,including evidencefor takeup of Mg and H20 into altered basaltsand loss of Ca and Si, all of which are typical of the low-temperature assemblagespredicted by our computer model. In general, the agreementin bulk 2.0- 21 øN MOR • -

chemistrybetween the predictionsmade in this study and the 1.8- observationsof midoceanridge samplesis quite good. 1.6-

Isotopic Systematicsin the Altered Rock 1.4- Oxygen and hydrogen isotopic values are generated as a 1,2- function of temperature and reaction progressfor each predic- o ted mineralproduct. f•80 and fD can be calculatedfrom the modal abundances of the product minerals in the altered % o.8- o 0.6 - 350 ø - whole-rock, R (assumedto include no residualfresh basalt), by E / the following relationship' oo0.4-0.2 - 300 ø / - • noVf•80•, f•8OR = v (11) o.o 250ø - E nov -o.2 1øøø/ - p 2øøøf15øø1/ I I -6 -5 -4 -3 -2 -i o and a correspondingexpression for fiD. The results of this log • calculation for f•80 and fD of the altered whole-rock are shown in Figures 11 and 12. The mineral assemblagescorre- Fig. 18. fixsoof theevolving solution in equilibriumwith mineral sponding to thesecurves can be determined by comparisonof assemblagesformed at the varioustemperature steps shown in Fig- Figures 11 and 12 with Figures 4 to 9 at the appropriate ures4 to 9, plottedas a functionof log reactionprogress; the coexistingmineral assemblages have the isotopiccompositions given temperature and value of reaction progress.The whole-rock in Figure11. Contours are temperaturein øC.The starcorresponds isotope curves are quite sensitiveto the identity of the min- to 5x80 of the 21øN MOR fluid as determinedby H. Craig (personal erals which they represent.For example,at 100øCf•8OR is communication, 1984). BOWERSAND TAYLOR: CHEMICALAND ISOTOPICMODEL OF HOT SPRINGS 12,597

2.8 I I have been summarizedby Taylor [1983]. Ophiolite pillow

2.6- 21øN MOR F lavashave typical 6•80 valuesfrom + 10 to + 16, closely 2.4- correspondingto 6•80 valuescalculated from our modelat

2.2- 100ø-150øC.Sheeted diabasesand high-level gabbros range from about + 3 to + 10, correspondingto the modelcalcula- 2.0- tions at > 200øC.Chlorites and amphibolesin the ophiolite • 1.8- complexescover a 6D rangefrom about -65 to -35 permil, •-• 1,6-- o correspondingvery well with the 250ø-350øCrange of 6D -r 1.4- valuesgenerated in our model. 0o 1.2 An averagebasalt-H20 oxygenisotope fractionation curve 1.0 canbe approximated from our model. The range of A•soR_.:o

0,8- 350 ø values(excluding some from low valuesof reaction progress

0.6- where alteration had not advancedsignificantly) are plotted 300ø •/// on Figure 14 as a functionof temperature.Equilibrium frac- 0.4 tionationcurves for An5oplagioclase, anhydrite, smectite, epi- 0.2 dote, muscoviteand chlorite from Figure 1 are includedfor 150o - 200 ø _ 0.0 250ø• I00 ø reference.A dashedline is drawn approximatelythrough the

-0.2 , I I 1 I I -6 -5 -4 -5 -2 middleof the calculatedAalteredbasalt_H20 values,and thisrepre- log (' sentsthe alteredbasalt-H20 •80/•60 fractionation.Compari- son of this curve with the curve for muscovite shows their Fig. 19. gD of the evolving solution in equilibrium with mineral similarity, particularly at low temperatures. The assemblagesformed at the various temperature steps shown in Fig- ures 4 to 9, plotted as a function of log reaction progress; the muscovite-H20fractionation was suggestedby Spooneret al. coexistingmineral assemblageshave the isotopic compositionsgiven [1977] to be a plausibleapproximation of the basalt-H20 in Figure 12. Contours are temperature in øC. The star corresponds fractionation. Note that the dashed curve in Figure 14 lies to gD of the 21øN MOR fluid as determined by Craig et al. [1980]. closer to the smectitecurve at low temperaturesand to the chlorite curve at high temperatures,reflecting the relative contributions of smectite and chlorite to the alteration mineral primarily to the identity of the mineralsmaking up the alter- assemblageat low and high temperatures,respectively. The ation assemblage.At high valuesof reaction progressthere is rangeof calculated6•80 valuesat 150ø and 200øCare some- a generallydecreasing trend of •iD with temperaturefrom 150ø what above the dashedcurve. The high valuesare a result of to 350øC, correspondingroughly to the similar trend in frac- the large amounts of anhydritepredicted to form at these tionation factorsshown in Figure 3. temperaturesand thus are not truly representativeof altered Figure 13 showsa comparisonof •i•80 and •iD valuesfor basaltscontaining no anhydrite. altered basalts calculated from our computer model with For comparison, the smectite-H20 fractionation curve analysesof altered basalts from the EPR and MAR [Stakes derivedby Cole[1980] from 6•80 analysesof seawater-basalt and O'Neil, 1982] and the Indian Ocean [Stakes et al., 1984]. This figure is after Figure 1 of Stakesand O'Neil [1982]. The high-temperaturecalculations agree reasonably well with observationson greenschist-faciesbasalts; however the 100ø and 150øCcalculated fields are considerablyricher in •iD than the saponite-richpillow brecciasanalyzed by Stakesand O'Neil. There are two major reasonsfor the •iD discrepancy.The •iD of OH-bearing minerals is quite sensitiveto changesin Fe/Mg and Fe/A1 (seeabove), and even small amounts of Fe substitutioninto the octahedral sites in the clays will signifi- cantly reduce the •iD signature of the clay. The computer model correctly predicts the occurrenceof saponitesat low temperatures,but there is no provision for clay solid solution 2 in the models,nor is there an Fe-saponiteend member in the thermodynamicdata base.As a result,the predictedsaponites 350 are Mg- and Ca-rich and their calculated•iD valuesare higher than those observedfor natural oceanicsaponites containing 250ø300• / Z... •øø• •5øøI 2øøø•' - someiron substitution.Even though the low-temperature•iD •'•-• •00o calculationsare lessthan perfect,this causesnegligible change in the overallresults of the calculationsfor •iD.2o (seebelow), I00 50 20 I0 5 2 I 0.5 becausemost of themodifications in •i•80.2 o and•iD•o take WATER/ROCK(wt. units) place at higher temperatures at relatively low water-rock Fig. 20. As a function of water-rock ratio in weight units (solid ratios. curves),this figureshows the •80 and •iD of the evolvingsolution in In addition to isotopic analysesof midocean ridge altered equilibriumwith the mineral assemblagesshown in Figures4 to 9' basalts,ophiolites provide an added source of isotopic data the latter have the isotopiccompositions given in Figures 11 and 12. Dot-dash curve for 6•80 represents a second calculated for comparisonwith the calculationspresented here. Oxygen H20 temperature/water-rockpath wheremore interactionis consideredat and hydrogen isotopic studiesof ophiolites by Javoy [1970], lower temperatures(see text). The dashedcurves represent approxi- Magaritz and Taylor [1974, 1976], Spooner et al. [1974], mateextrapolations of 6•80.2oand 6Dn2oif reactionis limitedto Heaton and Sheppard[1977] and Gregory and Taylor [1981] maximum temperaturesof 250ø or 300øC. 12,598 BOWERSAND TAYLOR: CHEMICALAND ISOTOPICMODEL OF HOT SPRINGS experimentsis also shownin Figure 14. Cole'scurve was used shownin Figure 17. The increasingconcentration of sulfidein to representthe •80/•60 fractionationduring basalt-seawater solution from 200 ø to 350øC is a result of contributions of interaction by Cathles [1983] in his isotopic modeling of the sulfur from the fresh basalt, dampened by removal of sulfide hydrothermalsystems that formed massivesulfide deposits in into pyrite and pyrrhotite.Fugacity of oxygenis initially quite the Hokuroku Basin in Japan. Note that the curve of Cole high because seawater is supersaturatedwith respect to [1980] is severalper mil higherthan the curvefor basalt-water oxygen;however, reaction with only a small amount of basalt approximatedfrom our model, and that its slopeis lesssteep overcomesthis supersaturationand log oxygenfugacity drops than nearly all of the experimentallydetermined fractionation markedly, from • -0.8 to -,- - 51.0 at 100øC. curvesfor minerals.Cole [1980] givesno reasonwhy his curve Na and K concentrationsincrease in solution largely as a is so drasticallyout of step with all other experimentaldata, as result of their removal from basalt; some Na is taken up in the well as with data and inferences from natural occurrences. It is development of albite as an alteration phase, and even less K well known that A18Oquartz_H20exhibits the largestfrac- is consumedby the precipitation of minor amounts of K-mica. tionation observedin any silicate-H20 system [Taylor and Dissolved Si (Figure 17) in the hydrothermal fluid increasesat Epstein, 1962; Matthews et al., 19836; Friedman and O'Neil, every temperature step. The fluid attains equilibrium with 1977],and yet the Cole [1980] A•aO .... tite-H20values are even quartz at 200ø and 350øC during a portion of the reaction biggerthan any of the publishedA18Oquartz_H20 values. We progresspath (hatchured on the diagram), and it is close to thus feeljustified in rejectingthe Cole [1980] curve. quartz saturation at intermediatetemperatures as well. The Fe concentration (Figure 17) increases dramatically Solution Chemistry over the course of reaction, from an initial value of • 1.5 Calculated dissolved concentrations of several elements, to- X 10-6 mmol/kg in unalteredseawater to as high as 60 gether with SO4 and H2S concentration,pH, and log fugacity mmol/kg at the conclusionof the 350øC calculation.This re- of oxygen as a function of reaction progressand temperature quires substantialleaching of Fe from basalt, although some are shown in Figures 15, 16, and 17. Many of the trends agree Fe is consumedin the production of products such as non- with those exhibited in laboratory experiments.A large pro- tronite at low temperatureand Fe-rich chlorite (daphnite) at portion of these curves, for example potassium (Figure 16), high temperature. A1 concentration (Figure 16) varies con- show no change in concentration from the end of a given siderably over the course of reaction, showing an increasing temperature step to the beginning of the next higher temper- trend at temperaturesup to 300øC, and then dropping dra- ature step. However, the concentrationsof certain speciesdrop matically at 350øC. over the course of the heating interval, when supersaturations The influence of a particular process(dissolution, precipi- occur. Examples include the drop in SO4 concentration tation, hydrolysis,electrical balancing of Mg uptake) on pH is (Figure 17) from the 100øCtemperature step to the start of the difficult to ascertain, because these processesare occurring 150øC temperature step as a result of anhydrite precipitation, simultaneouslywith possiblyopposite effects. Calculated vari- and the drop in A1 concentration (Figure 16) from the 300øC ations in pH as a function of temperature and reaction step to the start of the 350øC step as a result of precipitation progressare shown in Figure 15. At 100øC,the initial pH of of clinozoisiteand paragonite. ---6.0 correspondsto that calculated from a distribution-of- Mg concentration (Figure 16) decreasesfrom 52.6 to 2.4 speciesin the solution at 100øC,disallowing any precipitation mmol/kg at 100øC as a result of its uptake into Mg-rich clay from supersaturationsat that or lower temperatures. Predic- and chlorite alteration phases.This processis approximately tions based on laboratory experiments summarized by Mottl charge-balanced by an increase in Ca concentration in solu- [1983] suggesta drop in pH during the stage at which Mg is tion from 10.2 to 48.3 mmol/kg (Figure 15) over the same being taken up by a Mg(OH)2 component of the alteration range of reaction progressat 100øC. Reaction at 150øC com- assemblage,followed by a rebound to near neutrality by sub- pletes Mg removal, resulting in a final Mg concentration of sequentconsumption of H + in silicatehydrolysis reactions. •0.005 mmol/kg. Charge balancing of Mg removal by Ca The calculationspresented here show pH decreasingat 100øC replacement in solution was also observed in laboratory ex- over the range of reaction progresswhere Mg concentration periments by Mottl and Holland [1978] and Seyfriedand Bis- remains high (log • < -2.2); however the decreasein pH is choff [19813. slight compared to that observed in laboratory experiments Ca concentration increases slightly with Mg removal at [Mottl and Seyfried, 1980]. Nevertheless,the outcome of the 150øC, and subsequentlydecreases due to precipitation of an- calculationsdoes agree with laboratory experimentalresults in hydrite, a phenomenon also observed in several laboratory that both show near-neutral pH solutionsat the conclusionof experiments summarized by Mottl [1983]. Ca concentration the experiments.Except at 300øC, the calculated curves at continues to drop with increasingtemperature, predominantly each temperature step show an initial decreasein pH at low as a result of its uptake into epidote and tremolite. The initial values of reaction progressfollowed by a sharp increaseat increase of Ca in solution, followed by its decrease, corre- high values of reaction progress.The lowest pH reached is sponds well to observations by Humphris and Thompson •5.25 at 350øC, corresponding to the value of reaction [1978]. They found that the chlorite-rich assemblagesshow progress where the fluid is also saturated with respect to significant depletions in Ca relative to unaltered basalt, while quartz. In all of the above discussion,it should be remembered the epidote-rich assemblages(which presumably formed at that the pH of a neutral solution changesfrom 7.0 at 25øC to higher temperatures and lower water-rock ratios) show little 6.0 at 100øC and 5.5 at 350øC. change. Table 1 shows the predicted values of element or species Seawater sulfate is removed by anhydrite precipitation concentration in solution at the conclusion of each isothermal below 200øC (Figure 17). At 200øC the remaining sulfate is calculation, for comparison with the starting solution and the reduced to sulfide, which remains the dominant sulfur species range of concentrations observed in analyzed samples of hy- in solution at higher temperatures.The reduction corresponds drothermal fluids from 21øN on the EPR [Von Damm et al., to a drop in oxygenfugacity of approximately 5 log units, also 1985]. Several of the predicted concentrationsin the modeled BOWERSAND TAYLOR:CHEMICAL AND ISOTOPICMODEL OF HOT SPRINGS 12,599

350øC hydrothermal endmember agree well with actual sam- in solution. Lack of consideration of solid solution between ples of the 350øC hydrothermal fluid. Mg and SO,• con- clinozoisite and epidote, of Fe and A1 substitution in am- centrationsdrop to essentiallyzero, in agreementwith obser- phiboles,and of chlorite and smectitesolid solutionsmay con- vations. Silica increases to a value in accord with observation tribute substantially to the minor problems in the predicted during a portion of the 350øC calculation (Figure 17), al- Fe, Ca and A1 concentrationsof the aqueousfluids. though it is somewhat lower at the end-point of the 350øC The final oxygen fugacity calculatedis approximately 1 log calculation. The EPR hydrothermal fluid is apparently satu- unit below the value estimated for the EPR fluids, based on rated with respectto quartz [Von Dammet al., 1985]. assumed equilibrium between dissolved CO2 and CH,• con- Prediction of C1 concentration is in line with observation; centrationsreported by Welhan and Craig] [1983]. Dissolved dissolvedC1 increasesthroughout the calculations only as a H 2 concentrations[Welhan and Crai•], 1983] yield a similar result of the concentrating effect of H20 loss from solution log oxygenfugacity and becauseof the inherent uncertainties and uptake by the altered portions of the rock. In reality, in the measurementsand the assumptionsof equilibrium, we some C1 is also consumedby uptake into amphiboles,a pro- considerthe calculationsto representgood agreement. cess which has not been considered in these calculations and which may accountfor the lower end of the C1 range observed Isotopic Compositionsof the Fluids in the EPR fluids. Predicted Na concentration is slightly Valuesof g•so and gD of the evolvinghydrothermal solu- higher than that actually observedin the EPR fluids. Calcu- tion as a function of reaction progress and temperature are lated K concentration initially decreasesat low temperature shown in Figures 18 and 19. At low temperaturethe alteration and subsequentlyincreases to a valuewhich agrees well with productsare generally rich in •so (seeabove) and g•80 of the EPR observations.Calculated H2S concentrationis somewhat solution decreases,dropping most sharply at the lowest tem- higher than observed,and can be related to the amount of peratureconsidered, in thiscase, 100øC (Figure 18). g•sO.2 o sulfur entering the solution from fresh basalt and whether or continues to decline at 150øC, reaching a low value of about not saturation of the solution with respectto sulfide phasesis -0.2 per mil. Not until midway through the reaction at 250øC reached. When no sulfur is added from the basalt, calculated doesg•so.2 o increaseback to its initialvalue of 0 per mil. H2S concentrationsare considerablylower than observed.A Further reaction with basalt at 300 ø and 350øC results in a basalt with a lower sulfur content than the one used in these more rapid increasein g•so.•o until the 350øCreaction is calculationsshould yield a closeragreement between predicted terminatedat thepoint where g•so.2 o ,• + 2.0per mil, which and observedH2S concentrations. is the value reported by H. Craig (personal communication, Less agreementis seen between calculated and observed 1984) from measurementson the EPR fluids. concentrations of Fe, A1 and Ca, and pH (see below). Al- In contrast to g•so, the gD of the solution increasesat all though the calculated Fe concentrationsat lower temper- temperatures,because there is negligible,if any, input of H atures are reasonable,at the end of the 350øC step the predic- from the fresh basalt, and the ADmin_H2o fractionations for ted values are about 30 times those observed. Fe is also in- hydrogenare nearly all lessthan 0 exceptat very high temper- creasingvery rapidly with log • near the end of the 350øCstep atures outside the range of this study. The least rapid increase (Figure 17).This is probablyrelated to the inadequaciesof the in 6D.2o (Figure 19) occursat 250øCbecause the ADmin_H20 thermodynamicdata basein not allowing for the formation of fractionations for some of the minerals forming the alteration Fe-rich solid solutionsin the computer model. It might also be assemblageare closeto 0 (Figure 3). 5D.•o increasesfairly accountedfor by someprecipitation of Fe-rich mineralsin the rapidly at 350øCto a value of ,• + 2.6 per mil at the termina- hydrothermalconduit during ascentof the fluid, just prior to tion of the 350øC reaction. This is in excellent agreement with reachingthe seafloor. We know that abundantFe sulfidesare the value of +2.5 measured by Craig et al. [1980] for EPR indeed precipitatedthere. In any case, this is not such bad hydrothermal fluids. agreementconsidering that Fe concentrationsmust increase Although5 •sO.2o is controlledprimarily by temperatureof approximately 6 orders of magnitude from the initial cold interaction,gD.•o is chieflyaffected by the identity of the seawaterto the observedhydrothermal endmember,and our alteration products.For example,over 50% of the hydrogenis calculations show an Fe increase of about 7 orders of mag- taken up by clays during low-temperaturealteration and by nitude. Fe-rich chlorite during high-temperaturealteration. This ac- A1 concentration in solution initially increases,then de- countsfor the sharpincrease in c•D.•o at the highertemper- creases,but reachesa final value which is high compared to atures becauseof the large negative AD fractionations em- the observed concentrations. Much of the excess A1 in solution ployed for Fe-rich chlorites. could be taken up by A1 substitutioninto amphiboles.As little Variations in g•so and gD of solution as a function of as 1 weight percent substitutionof A1 into the amount of water-rock ratio over the temperature range employed are tremolite predictedat 350øC would require more A1 than is shown in Figure 20. High water-rock ratios correspondto low presentin the solution.Note that at 250ø and 350øC,the A1 temperature, where only a small amount of rock has inter- concentrationsare enormous and are increasingvery rapidly acted with a given packet of fluid. At high temperaturesthis with log •, extensivereaction at 350øCis requiredto drop the samepacket of fluid has traveledthrough and interactedwith A1 valuesanywhere close to the observedEPR value (Figure considerably more rock, correspondingto low water-rock 16). ratios. The lower diagram in Figure 20 showsthe decreaseand Ca concentrationalso presentsa problem, initially increas- subsequentincrease in g•so of the solutionas a functionof ing, then decreasingin accordwith laboratory experimentsas temperature to a final value of +2 per mil at water/ well as whole-rock bulk chemical alteration trends. However, rock ,• 0.56. Below this curve a second curve is shown for a the final Ca concentrationpredicted is an order of magnitude similar series of seawater-basalt interactions from 2øC up to below that observed in the EPR fluids. Na substitution for Ca 350øC. Note that the final g•so of this latter solution at wa- in amphibolewould have the effectof simultaneouslyreducing ter/rock • 3.2 is still <0 per mil. This givessome indication of the quantity of Na predictedand increasingthe amount of Ca the proportion of low- and high-temperatureinteraction that 12,600 BOWERSAND TAYLOR:CHEMICAL AND ISOTOPICMODEL OF HOT SPRINGS

pH• water-rock ratio fluids at shallow levels in the pillow basalts Neutrality : 25 ø or in off-axiscirculation systems. DissolvedMg and SO4 and pH are shownin Figure 21 as a

I00 ø / function of reaction progress and temperature from 100ø- 6- --, i00 o 350øC. The results of the calculations at 100øC are identical to 200 ø 300 ø those shown in Figures 4 and 15, 16, and 17. However, the

_ solution composition used to initiate the 150øC reaction in 200 ø / Figure 21 is taken from log • •-, -2.2 at 100øC rather than 250ø log • •-, -0.9 as in the formerset of calculations.The termina- • _.•- •' / 300 ø / / _ tion of eachisothermal calculation shown in Figure 21 is rep- 350 ø resentedby the end of the solid curve. The dotted portion of the curve at higher values of reaction progressshows the re- I I i I i i sults if the calculation had been carried further. At each tem- 60 i I I perature, reaction is terminated at the point where the lowest 100 ø 50- 150 ø value of pH is reached. Therefore, the pH of •-,3.5 for the 350øC solution is as low a pH as could be calculatedwith this 40- procedure.However, note that the low pH at each temper-

30- 350 ø ature correspondsto high Mg and SO• concentrationsin the 20- solution.When Mg beginsto drop (dottedportion of curvesin Figure 21), SO• also drops and pH risesdramatically. This is I0- in accordwith observationsby Seyfriedand Bischoff[1977], i I i 0 summarizedby Mottl [1983], who showedin laboratory ex- perimentsthat pH risesfrom acidic levelsto near-neutrality once Mg concentrationdrops. 30 too ø Alteration productsprecipitated along this reaction path 150ø 200ø (J3 20 250ø '""'• are summarizedas a function of temperature in Figure 22. 3ooø/ ;5oo/ Predominant products are Mg-saponite, chlorite, and muscovite plus small amounts of hematite and pyrite at 100ø and - . -6.0 -5.0 -4.0 -3.0 -2.0 -I.0 0.0 150øC, and chlorite, talc, anhydrite, quartz and hematite LOG ( above 200øC. These results correspondwell to the chlorite- Fig. 21. Mg and SO4 concentration(mmol/kg) and pH of an quartz-rich rocks describedby Mottl [1983] as typical alter- evolvingsolution corresponding to a high water-rockratio, low-pH ation productsat high (> 50) water-rock ratios. path, plotted as a functionof log reactionprogress. Contours are temperaturein øC.Reaction at eachtemperature except 100øC begins SUMMARY AND CONCLUSIONS with a solutioncomposition corresponding to that at the termination of the solidportion of the curveat the next lower temperature.The We have attemptedto model the chemical,mineralogical, dashedportion of the curvesexhibits the changesin concentrations, and isotopiceffects of hydrothermal circulation of seawaterat had weconsidered continued reaction at eachtemperature. a midocean ridge, where seawater presumably encounters fresh basalt, diabase,and gabbro at decreasingwater-rock ratios with depth and increasingtemperature up to a maximum temperature of approximately 350øC. Our results are a givenpacket of fluid must undergoto exhibita + 2 per mil summarizedin schematicfashion in Figure 23, which showsa 6•80 signature,as measured for exitinghydrothermal fluids at hypotheticalpath taken by a packet of fluid that will eventual- 21øN on the EPR. ly exit at a MOR vent. Also shown are schematicflow lines for lower-temperature, off-axis fluids. The model is realistic be- Calculations at High Water-Rock Ratios causeit providesfor sharplyincreasing temperatures and de- Observationssummarized by Mottl [1983] that the pH of creasingwater-rock ratios as the fluid approachesthe top of hydrothermal solutions produced in laboratory experiments the magmachamber. The effectivewater-rock ratio, definedby initially becomesquite acidic and remains so while Mg con- the amount of rock that chemicallyinteracts with a given centration is high and water-rock ratio is greater than 50 are amountof initial fluid, decreasesdownward because: (1) much further supported here by a set of computer calculationsat fluid that circulatesat shallowdepths in the low-temperature high water-rock ratio. This seriesof calculatedreaction paths from 100ø-350øC is the outgrowth of an attempt to find a possiblepath resulting in an acidic fluid similar to those actu- Mg- SAPOI'•ITE ally sampled at 21øN. DOLOMITE' ANTIGORITE The solution composition generatedby this seriesof calcu- CLINOCHLORE lations at a water/rock of about 150 does indeed exhibit an AMESITE TALC acidicpH of around 3.5 at 350øC.However, low pH and satu- MUSCOVITE QUARTZ ration of the fluid with respectto quartz are the only simi- ANHYDRITE laritiesbetween the calculatedfluid compositionand the 21øN PYRITE analyzed fluid compositions.Mg and SO4 concentrations HEMATITE remainhigh, and 6D and g•80 of the fluid remainclose to 0. Althoughthis seriesof calculationsin no way representsthe Fig. 22. Summary of predicted alteration phasesas a function of hydrothermalendmember that is currentlyexiting from vents temperature, correspondingto a calculated high water-rock ratio, on the EPR, it may correspondto the low-temperature,high low-pH path, part of whichis shownin Figure 21. ]lOWERSAND TAYLOR: CHEMICALAND ISOTOPICMODEL OF HOT SPRINGS 12,601

18 8 0=0.0 Bi80.l•2.0 0.25KB

SAPONITE- BEARING / ,• .• ,,,..•, !

BASALTS ...... + PREH•ITE-BEARI•G GREENSCHIST - DIKES ASSEMBLAGES •o GREENSCHIST ASSEMBLAGES o o o 0.5 KB

GABB RO S ;••••••••• ;?'" GABBROS

Fig. 23. Schematicillustration of the model describedin this paper, showinga hypotheticalpath taken by the fluid from its initial penetrationof seafloorbasalts at low temperature,its circulation downward to the magma-chamber contact,and finally to its eventualexit at 350øCfrom MOR vents.Cumulative water-rock ratio, percentreaction progress, and6x8OR2 o are shown along the two symmetrical path-lines at each50øC temperature step at whichthe calculations were carried out. The temperaturesincrease and the effectivewater-rock ratios decreasevery rapidly with depth. The calculations lead to a sharplyincreasing grade of hydrothermalmetamorphism downward and they providea good match with the assemblagesactually observedin nature. In the near-surfaceenvironment, saponite-bearing assemblages form at high water-rockratios as a result of low-temperatureinteraction of the fluids with submarinebasalts. With increasingdepth, theseassemblages are replacedby prehnite-bearingtransitional greenschist assemblages, and eventuallyby true greenschist assemblagesformed at 300ø-350øC.Pressure ranges from approximately250 bars at the seafloorto approximately500 bars at the top of the magma chamber [Von Darnrnet al., 1985]. All of the model calculationswere performed at a constantpressure of 500 bars,but this is a reasonableassumption because variations over this limited pressurerange have little or no effecton the calculations.It is alsoassumed that no significantalteration of the fluid occursduring its trip from depth to the seabedinterface. Details concerningthe calculatedmineral assemblagesand calculatedchemical and isotopic compositionsof the fluidsat eachtemperature step are shownin Figures4 to 20.

regionsdoes not reachthe high-temperatureregions at depth, agreement with observationson dredged samples from the and (2) reaction rates are more rapid at high temperature, oceansand in ophiolite complexes;these include saponitesat implying that more rock will interact with a given packet of low temperatures,Mg-rich and chloritic assemblagesat high water per unit time in the high-temperatureregions of the water-rock ratios, and epidote-chlorite-actinolite (tremolite)- hydrothermal system.Note that the effectivewater-rock ratio albite assemblagesat high temperaturesand low water-rock is a material-balanceparameter that is meaningfulonly for a ratios. The calculatedconcentrations of the major elementsin systemof sufficientsize (e.g., a significantportion of the entire the evolved fluid also agree substantiallywith analysesmade MOR hydrothermalcirculation system). This parameteris not on actual samplesof the 350øC hydrothermal endmember. relatedin any simpleway to the actual amount of H20 that There are two important discrepanciesbetween the calcula- may interactwith a cubicmeter of rock' for example,in the tions (or the laboratory experiments)and the sampled hy- active upflow region beneatha MOR hydrothermalvent, the drothermal solutions:pH and saturation state of the solution integratedmass flux of H,•O can easilybe as high as 6 x 106 with respectto certain minerals. kg/m2, whichin weightunits translatesinto an integrated The lowest pH reached in the main set of reaction-path water-rock ratio of 2000 for the cubic meter of basalt [Norton calculationsis 5.25,and most laboratory experimentsconduct- and Taylor, 1979' Cathles,1983]. ed to date also exhibit similar, near-neutralpH valuesexcept Although appreciablewater clearlymust penetrateinto re- (1) in the initial stagesof the experimentswhere Mg con- gionswith temperaturesabove 350øC,we are assumingthat centration is still high [Mottl, 1983], or (2) at 400ø to 500øC that componentof the fluid doesnot contributesignificantly [Mottl et al., 1979]. The sampled21øN-EPR solutionshave a to the upflow zone at the ridge crest, at least in the active pH in the range 3.3 to 3.8 (25øCmeasurements) and are satu- venting situation presentlyobserved on the East Pacific Rise. rated only with respect to quartz [Von Datum et al., 1985]. Water-rockratios of about0.5 are necessaryto obtainthe x80 Solutions sampled from 13øN on the EPR also have low pH and D enrichmentsobserved in the EPR hot-springfluids values: approximately 3.8 to 3.9 for the samples with the (Figure 23). Our calculationsindicate that the major propor- smallestcomponent of seawatermixing [Michard et al., 1984]. tion of the water/rock interactionhad to take place at at least The pH and mineral saturation problems are related. In- 300ø-350øCin order to producethe appropriateisotopic sig- creasingthe pH from 3.5 to 5.5 will result in many silicates natures in the fluids. being closerto saturation,even with no other changeto the Calculatedalteration assemblages in this work are in good solutionbut an adjustedNa value to maintain chargebalance. 12,602 BOWERSAND TAYLOR'CHEMICAL AND ISOTOPICMODEL OF HOT SPRINGS

For example, if the 21øN EPR fluid composition given in water-rockratio is lessthan 10, fx8Oa•o and fDa2o also Table 1 is consideredto have a pH of 5, calculatedaffinities of become sensitive to the value of the water-rock ratio. This is the minerals show that the resulting solution would be close best illustratedby the curvesshown in Figure 20, which repre- to saturation with respect to albite, epidote, prehnite and sentthe dependencyof fx8Oa•o and fDa•o on temperature anorthite, among others. Slightly higher A1 contents would and water-rock ratio for the particular "path of interaction" result in saturation of the fluid with respectto many of these utilized in the present study, and describedin detail above. phases,and this problem is most likely attributable to poor The factthat calculatedgx8Oa• o and fDa•o simultaneously thermodynamicdata on aqueousAl-complexes, or to a lack of attain values substantiallyin agreementwith those reported consideration of solid solutions. Note that the state of satu- for midocean ridge hot springs at a cumulative water-rock ration of the fluid with respectto quartz is not substantially ratio of 0.56 arguesstrongly that the particular combination affectedby changesin pH alone. of amount of interaction at the chosen temperatures is an A contrast to midocean ridge solutionsis found in geother- adequate descriptionof how these hydrothermal fluids could mal solutions from Iceland, which also formed from seawater have evolved. The proportion of high-temperatureto low- flowing through basalt. Solutions from Svartsengi temperatureinteraction must be large in order to producethe [Ragnarsddttir et al., 1984] more closely resemblethe results positivef•80 anomaliesactually observed in the hot-spring of calculationspresented here in that they exhibit high pH fluids_,quhqtantia! amounts of interaction at low temperatures values (• 5.4) and are saturated with respectto several min- will resultin negativefx8Oa2 o valuesand indeed, any interac- erals observed in drill cuttings from the hydrothermally al- tion at all between seawater and basalt below approximately tered zone in the tholeiitic basalt, including quartz, calcite, 200ø-250øCwill decreasef x8Oa• o, as demonstratedin this albite, chlorite and epidote. Our calculations predict study and that by Stakes and O'Neil [1982]. Although it is greenschist-faciesalteration of basalt, including formation of quite likely that off-axis hydrothermal alteration at low tem- chlorite, epidote,actinolite, and quartz at 350øC.The solution peraturesis occurringand doesresult in negativefx8Oa• o is, therefore,obviously saturated with respectto the alteration values, it is clear that the fluid sampled at the MOR vents productsthat have precipitatedfrom it. It appearson the basis cannot have undergone any substantial stage of low- of laboratory experimentsand analysesof geothermal water, temperature, low-water/rock interaction with newly-formed such as those at Svartsengi,that this processoccurs at neutral oceanic crust. If it had, it would be virtually impossible to pH, between 5 and 6, much the same as the pH calculated producethe observed f x80}h o valueof + 2 permil. here. The lightly dashedlines in Figure 20 illustrate approximate Several possibilitiesremain which could account for this water-rock ratios for both oxygenand hydrogen that must be discrepancyin pH. These include pressureeffects occasioned attained if the observed isotopic ratios are to result from a by the fluid's ascent up the conduit, inaccuraciesin the maximum temperature of interaction of 300ø' or 250ø rather measurementsat 21øN (and at 13øN), possibleselective pre- than 350øC as assumedin this study. At 300øC, the water-rock cipitation of some mineralsfrom the fluid or degassingof the ratiorequired to attainf x80}ho = + 2 permil is •0.3, andto fluid before the measurementsare made, inaccuraciesor omis- attain a fDH20-- q-2.5 per mil, it is •0.2. At 250øCthese sionsin the thermodynamicdata base used to calculate the in water-rockratios are 0.04 and 0.015, respectively.A maximum situ pH values from the 25øC values,or the possibilitythat we temperatureof alteration above 350øC would result in attain- have overlooked a major, pH-affecting processwhich occurs mentof theappropriate c5t80 and fD valuesof thevent fluids before the fluids vent. Also, there may have been some addi- at somewhathigher wa•ter-rockratios. Note, however,that tion of high-temperaturemagmatic H20 at depth, or possibly onlyat 350øCare the appropriatefx80 and fD valuesfor the water/rock interaction occurred at higher temperatures than fluid of + 2.0 and + 2.5, respectively,achieved simultaneously considered here (400ø-450øC), and then during ascent there at approximatelythe samevalue of water-rock ratio. was more extensivecooling than we have envisioned(mixing Muehlenbachs and Clayton [1976] proposed that hy- with cooler fluids?). We are left with a puzzle that will require drothermal circulation within the oceanic crust "buffers" the further work to sort out. f x80 of oceanwater as a resultof a near-perfectcancellation The results of the calculationspresented here seemto vali- of the x80 enrichmentof seawaterdue to high-temperature date the selectedhydrogen isotope fractionation curves used alterationby a correspondingxsO depletionresulting from in this work, at least over the temperature range for which low-temperaturealteration. Support for their hypothesiswas theyhave been employed (Figure 3). Calculationsof fDa2o for obtainedby a material-balanceintegration of a fx80 profile the same temperature/water-rock interaction paths shown in through a sectionof the Samail ophiolite in Oman [Gregory Figures 4 to 9, but utilizing different D/H fractionationcurves, and Taylor, 1981]. In the presentstudy, we have emphasized give valuesthat are markedly differentfrom the observeddeu- the need for a higher proportion of high-temperature(300 ø- terium enrichmentsof + 2.5 per mil in midocean ridge hot 350øC)alteration products to producethe +2 per mil f•80 springs.For example,assuming a constantD/H fractionation value observedin the EPR exiting hot springsby H. Craig curve of Amin_H20-'--20 for all mineralsresults in a final (personalcommunication, 1984). Obviously,if this were the fDa2o = q-1.2 per mil, a valuetoo low by half. If Amin_H20-' only processaffecting the isotopic compositionof seawater, --60, the calculatedfDI•2O = q-3.5, which is too high.Clearly, the effectwould be to producean increasingf x80 of ocean the averagevalue of Ami•_•o over the reactionpath con- water with time as a result of hydrothermal alteration of the sideredin this work must be about -40, which is in fact very oceaniccrust. However, the processproducing ridge-cresthot closeto the averagemeasured difference between seawater and springs,which is responsiblefor thepositive fx80,2 o anoma- hydrothermal OH-bearing minerals obtained from marine ly, is apparentlyonly a relativelyshallow, axial process,affect- dredgehauls and ophiolite complexes. ing perhapsthe upper 2 to 3 km of the oceaniccrust; it clearly At high water-rockratios, f Oa•o and fD•o are affected cannot representthe whole story. primarily by temperatureof interaction (and compositionof Activehot springshave not, to date, beenobserved off-axis. the mineral phasesin the case of fiD). However, when the Off-axis circulationcells probably involve larger quantitiesof BOWERSAND TAYLOR:CHEMICAL AND ISOTOPICMODEL OF HOT SPRINGS 12,603

water, and they certainly involve lower temperaturesthan do though less information is typically available for the high- those in the immediate neighborhood of the ridge-crest temperature parts of active systemson continents, the possi- (Figure 23). Seawater-basaltinteraction at temperaturesless bility of imposing constraints on temperatures,temperature thanapproximately 200øC results in negative6•8OH2 o values; gradients,pressures, and water-rock ratios suggestthat similar thus, extensiveoff-axis circulation at low temperaturescould computer models that combine a chemical and stable isotope serveto cancelthe effectsof the positive6180 fluidsventing approachmay prove to be very useful. fromthe ridgecrests. The 618OH2o results of two calculated paths shown in Figure 20 indicate that low-temperaturealter- Acknowledgments. Financial support for this researchwas provid- ation possiblylowers 6180 of seawaterto -0.2 to -0.5. ed by the National Science Foundation, grants OCE-8019021 and OCE-8315280, and by The ResourceGeology Research Fund of the These valueswould require that approximately4 to 10 times Division of Geological and Planetary Sciences,California Institute of the volume of water circulate through the off-axis circulation Technology. We are indebted to John Edmond, Sam Epstein, Hal systemswith a concomitantreduction in 6180 in order to Helgeson, Dave Janecky, Peter Larson, Denis Norton, Debra Stakes, cancel the ridge-crestenrichment of 180. If the low- and Karen Von Damm for their helpful suggestions,assistance, and encouragementduring the course of this research, and we wish to temperatureeffects on 618OH20 are not asnegative as suggest- thank Harmon Craig for allowing us to quote his unpublished iso- ed here, an even larger volume of water would be required to topic analyses of MOR hydrothermal fluids. We would also like to circulate through the off-axis systemsin order to achieve the expressour appreciation to J. K. B6hlke, Dave Janecky, Mike Mottl, "buffering"effect on seawaterof 6180 proposedby Muehlen- K. Valla Ragnarsd6ttir, and Karen Von Damm for their constructive bachs and Clayton [1976] and Gregory and Taylor [1981]. reviews of the manuscript and helpful suggestionsfor improvement. Contribution 4146 of the Division of Geological and Planetary Sci- The problem is compoundedby the fact that in this work we ences,California Institute of Technology. have also ignored the truly high-temperature,very deep (> 3 km) circulation of exchangedsea water down along the flanks REFERENCES (and underneathT) the MOR axial magma chamber.Although Becker, R. H., Carbon and oxygen isotope ratios in iron-formation Gregory and Taylor [1981] showed that this circulation in- and associatedrocks from the Hamersley Range of western Aus- volvesa very low water/rock ratio (• 0.2), in the Oman ophio- tralia and their implications, Ph.D. thesis, Univ. of Chicago, lite its isotopic effects are significant,because they are ob- Chicago, Ill., 1971. served throughout a large part of the lower oceanic crust, at Bird, D. K., and H. C. Helgeson, Chemical interaction of aqueous solutionswith epidote-feldsparmineral assemblagesin geologicsys- least down to the MOHO. To even a greater degreethan the tems, I, Thermodynamic analysis of phase relations in the system shallow,axial systemconsidered extensively in this study, the CaO-FeO-Fe203-SiO2-H20-CO2, Am. J. Sci.,280, 907-941, 1980. deep systemalso producesmarked 180 depletionsin the Bischoff, J. L., and F. W. Dickson, Seawater-basalt interaction at rocks; thus, its effectsmust also be counter-balancedby an 200øC and 500 bars: Implications for origin of seafloorheavy-metal deposits and regulation of seawater chemistry, Earth Planet. $ci. appropriate amount of low-temperature, off-axis hy- Lett., 25, 385-397, 1975. drothermal activity and submarine weathering. In order to Bischoff, J. L., and R. J. Rosenbauer,A note on the chemistry of obtain a completepicture of the isotopic,chemical, and miner- seawater in the range 350ø-500øC, Geochim.Cosmochim. Acta, 47, alogicalalteration of the entire oceaniccrust, future computer- 139-144, 1983. modeling studiesshould include provision for the effectsof the Bischoff,J. L., and R. J. Rosenbauer,The critical point and two-phase boundary of seawater, 200-500øC, Earth Planet. $ci. Lett., 68, 172- deep hydrothermal circulation, as well as the off-axis hy- 180, 1984. drothermal alteration. Unfortunately, at the present time, ap- Bischoff,J. L., and W. E. Seyfried,Jr., Seawateras a geothermalfluid: propriate high-temperature thermodynamic data (>400øC) Chemical behavior from 25ø to 350øC, in Proceedingsof the 2nd are almost nonexistent,and most of the constrainingparame- International Symposiumon Water-Rock Interaction, pp. IV165- IV172, International Association for Geochemistry and Cosmo- ters for the low-temperaturecirculation are also very poorly chemistry,Strasbourg, France, 1977. known. Bischoff, J. L., and W. E. Seyfried, Jr., Hydrothermal chemistry of In contrastto the 6180 effectsdescribed above, the hy- seawater from 25ø-350øC, Am. d. Sci., 278, 838-860, 1978. drothermal alteration of oceanic crust has a much different B6hlke, J. K., J. C. Alt, and K. Muehlenbachs,Oxygen isotope-water effect on the 6D of seawater. The 6D of seawater should stead- relations in altered deep-sea basalts: Low-temperature mineralogical controls, Can. J. Earth $ci., 21, 67-77, 1984. ily increaseat all temperaturesof hydrothermalinteraction up Bowers, T. S., K. L. Von Damm, and J. M. Edmond, Chemical evolu- to approximately 700øC. This is attributable to the fact that tion of mid-ocean ridge hot springs,Geochim. Cosmochim. Acta, in freshbasalt contributesessentially no H to the system,and all press, 1985. hydroxyl-bearingminerals formed as a result of hydrothermal Brimhall, G. H., Lithologic determination of mass transfer mechanisms of multiple-stage porphyry copper mineralization at Butte, alteration exhibit negative 6D values compared to H20 Montana: Vein formation by hypogeneleaching and enrichment of (Figure 3). Becauseocean water is a substantiallylarger reser- potassium-silicateprotore, Econ. Geol., 74, 556-589, 1979. voir of hydrogen relative to the whole earth than it is for Brimhall, G. H., Deep hypogene oxidation of porphyry copper oxygen, the steady increasein 6D of ocean water would take potassium-silicateprotore at Butte, Montana: A theoretical evalu- place very slowly. It is also probably counterbalancedby recy- ation of the copper remobilizationhypothesis, Econ. Geol., 75, 384- 409, 1980. cled H20 from subductedhydrous minerals in the form of Brimhall, G. H., and M. S. Ghiorso, Origin and ore-forming conse- emanationsof low-D magmatic H20 elsewherein the world quencesof the advanced argillic alteration processin hypogene [Taylor, 1974]. Even though there are still many gaps in the environments by magmatic gas contamination of meteoric fluids, required data-base,it would be usefulto carry out a complete Econ. Geol., 78, 73-90, 1983. Cathles, L. M., An analysisof the hydrothermal systemresponsible material-balancecomputer modeling study of both the D/H for massive sulfide deposition in the Hokuroku Basin of Japan, and the '-80/'-60 in the oceans,taking into accountthese Econ. Geol. Monogr., 5, 439-487, 1983. other factors,as well as the predictionsof the presentwork. Chiba, H., M. Kusakabe, S. Hirano, S. Matsuo, and S. Somiya, Integrated chemical and stable-isotopemodels such as the Oxygen isotope fractionation factors between anhydrite and water from 100 to 550øC, Earth Planet. $ci. Lett., 53, 55-62, 1981. one presentedhere can also be used to place important con- Clayton, R. N., J. R. O'Neil, and T. K. Mayeda, Oxygen isotope straints on other types of (on-land) hydrothermal systems,in exchangebetween quartz and water, J. Geophys.Res., 77, 3057- addition to seawater-basalt hydrothermal circulation. A1- 3067, 1972. 12,604 BOWERSAND TAYLOR: CHEMICAL AND ISOTOPIC MODEL OF HOT SPRINGS

Cole, D. R., Mechanism and rates of stable isotope exchangein hy- from the Mid-Cayman Rise: Petrographic and mineralogic con- drothermal rock-water systems,Ph.D. thesis,Pa. State Univ., Uni- straints on hydrothermal processesat slow-spreadingridges, Con- versity Park, 1980. trib. Mineral. Petrol., 82, 371-388, 1983. Cole, D. R., H. Ohmoto, and A. C. Lasaga, Isotopic exchange in Ito, E., and R. N. Clayton, Submarine metamorphism of gabbros mineral-fluid systems,I, Theoretical evaluation of oxygen isotopic from the Mid-Cayman Rise: An oxygen isotopic study, Geochim. exchangeaccompanying surface reactions and diffusion, Geochim. Cosmochim.Acta, 47, 535-546, 1983. Cosmochim.Acta, 47, 1681-1693, 1983. Janecky, D. R., Serpentinization of peridotitc within the oceanic Craig, H., Hydrothermal plumesand tracer circulation along the East crust: Experimental and theoretical investigations of seawater- Pacific Rise: 20øN to 20øS (abstract), Eos Trans. AGU, 62, 893, peridotitc interaction at 200øC and 300øC, 500 bars, Ph.D. thesis, 1981. Univ. of Minn., Minneapolis, 1982. Craig, H., J. A. Welhan, K. Kim, R. Poreda, and J. E. Lupton, Geo- Javoy, M., Ultisation des isotopes de l'oxygene en magmatologie, chemical studies of the 21øN EPR hydrothermal fluids (abstract), th•se de doctorat d'etat des sci. phys., Facult• des Sci. de Paris, Eos Trans. AGU, 61,992, 1980. Paris, 1970. Delaney, J. R., R. E. McDuff, and J. E. Lupton, Hydrothermalfluid Kendall, C., I.-M. Chou, and T. B. Coplen, Salt effect on oxygen temperaturesof 400øC on the Endeavor segment,northern Juan de isotopeequilibria (abstract),Eos Trans. AGU, 64, 334-335, 1983. Fuca (abstract), Eos Trans. AGU, 65, 973, 1984. Kim, K.-R., J. A. Welhan, and H. Craig, The hydrothermal vent fields East Pacific Rise Study Group, Crustal processesof the mid-ocean at 13øN and liøN on the East Pacific Rise: ALVIN 1984 results ridge, Science,213, 31-40, 1981. (abstract),Eos Trans. AGU, 65, 973, 1984. Edmond, J. M., The chemistry of the 350øC hot springs at 21øN on Kulla, J. B., and T. F. Anderson, Experimental oxygen isotope frac- the East Pacific Rise (abstract), Eos Trans. AGU, 61, 992, 1980. tionation between kaolinitc and water, U.S. Geol. Surv. Open File Edmond, J. M., C. Measures, R. E. McDuff, L. H. Chan, R. Collier, B. Rep., 78-701,234-235, 1978. Grant, L. I. Gordon, and J. B. Corliss, Ridge crest hydrothermal Lambert, S. J., and S. Epstein, Stable isotope investigationsof an activity and the balancesof the major and minor elementsin the active geothermal systemin Valles Caldera, Jemez Mountains, New ocean:The Galapagosdata, Earth Planet. Sci. Lett., 46, 1-18, 1979. Mexico, J. Volcanol. Geotherm.Res., 8, 111-129, 1980. Eslinger, E. V., and S. M. Savin, Mineralogy and oxygen isotope Lawrence, J. R., and H. P. Taylor, Jr., Deuterium and oxygen-18 geochemistry of the hydrothermally altered rocks of the Ohaki- correlation: Clay minerals and hydroxides in Quaternary soils Broadlands,New Zealand geothermal area, Am. J. Sci., 273, 240- compared to meteoric waters, Geochim.Cosmochim. Acta, 35, 993- 267, 1973. 1003, 1971. Friedman, I., and J. R. O'Neil, Compilation of stable isotope frac- Lawrence, J. R., and H. P. Taylor, Jr., Hydrogen and oxygen isotope tionation factorsof geochemicalinterest, U.S. Geol.Surv. Prof. Pap., systematicsin weathering profiles, Geochim.Cosmochim. Acta, 36, 440-KK, 1977. 1377-1393, 1972. Graham, C. M., S. M. F. Sheppard, and T. H. E. Heaton, Experi- Lister, C. R. B., "Active" and "passive"hydrothermal systemsin the mental hydrogen isotope studies, I, Systematicsof hydrogen iso- oceaniccrust: Predicted physicalconditions, in The DynamicEnvi- tope fractionationin the systemsepidote-H20, zoisite-H20 and ronmentof the Ocean Floor, edited by K. A. Fanning and F. T. A10(OH)-H:O, Geochim.Cosmochim. Acta, 44, 353-364, 1980. Manheim, pp. 441-470, D.C. Heath, Lexington, Mass., 1982. Gregory, R. T., and H. P. Taylor, Jr., An oxygen isotope profile in a Liu, K.-K., and S. Epstein, The hydrogen isotope fractionation be- section of Cretaceous oceanic crust, Samail ophiolite, Oman: Evi- tween kaolinitc and water, Isot. Geosci.,2, 335-350, 1984. dencefor •5•80 bufferingof the oceansby deep(> 5 km) seawater- MacDonald, K. C., K. Becker, F. N. Speiss,and R. D. Ballard, Hy- hydrothermal circulation at midocean ridges,J. Geophys.Res., 86, drothermal heat flux of the "black smoker" vents on the East Pacif- 2737-2755, 1981. ic Rise, Earth Planet. Sci. Lett., 48, 1-7, 1980. Hajash, A., Hydrothermal processesalong mid-ocean ridges: An ex- Magaritz, M., and H. P. Taylor, Jr., Oxygen and hydrogen isotope perimental investigation, Contrib. Mineral. Petrol., 53, 205-226, studies of serpentinization in the Troodos ophiolite complex, 1975. Cyprus, Earth Planet. Sci. Lett., 23, 8-14, 1974. Haymon, R. M., Growth history of hydrothermal black smoker Magaritz, M., and H. P. Taylor, Jr., Oxygen, hydrogen, and carbon chimneys,Nature, 301, 695-698, 1983. isotope studiesof the Franciscanformation, Coast Ranges,Califor- Heaton, T. H. E., and S. M. F. Sheppard, Hydrogen and oxygen nia, Geochim. Cosmochim.Acta, 40, 215-234, 1976. isotope evidence for sea-water hydrothermal alteration and ore Mathez, E. A., Sulfur solubility and magmatic sulfidesin submarine deposition, Troodos complex, Cyprus, in Volcanic Processesin Ore basalt glass,J. Geophys.Res., 81, 4269-4276, 1976. Genesis,Spec. Pap. 7, pp. 42-57, Geological Society of London, Matthews, A., J. R. Goldsmith, and R. N. Clayton, On the mecha- London, 1977. nisms and kinetics of oxygen isotope exchangein quartz and feld- Helgeson, H. C., Evaluation of irreversible reactionsin geochemical spars at elevated temperaturesand pressures,Geol. Soc. Am. Bull., processesinvolving minerals and aqueous solutions,I, Thermody- 94, 396-412, 1983a. namic relations, Geochim.Cosmochim. Acta, 32, 853-877, 1968. Matthews, A., J. R. Goldsmith, and R. N. Clayton, Oxygen isotope Helgeson,H. C., Thermodynamicsof hydrothermalsystems at elevat- fractionationsinvolving pyroxenes:The calibration of mineral-pair ed temperaturesand pressures,Am. J. Sci., 267, 729-804, 1969. geothermometers,Geochim. Cosmochim. Acta, 47, 631-644, 1983b. Helgeson, H. C., and D. H. Kirkham, Theoretical prediction of the Matthews, A., J. R. Goldsmith, and R. N. Clayton, Oxygen isotope thermodynamic behavior of aqueous electrolytesat high pressures fractionation between zoisite and water, Geochim.Cosmochim. Acta, and temperatures,I, Summary of the thermodynamic/electrostatic 47, 645-654, 1983c. properties of the solvent,Am. J. Sci., 274, 1089-1198, 1974a. Michard, G., F. Albarede, A. Michard, J.-F. Minster, J.-L. Charlou, Helgeson, H. C., and D. H. Kirkham, Theoretical prediction of the and N. Tan, Chemistry of solutionsfrom the 13øN East Pacific Rise thermodynamic behavior of aqueouselectrolytes at high pressures hydrothermal site, Earth Planet. Sci. Lett., 67, 297-307, 1984. and temperatures,II, Debye-Huckel parameters for activity coef- Miyashiro, A., S. Fumiko, and M. Ewing, Diversity and origin of ficientsand relative partial molal properties,Am. J. Sci., 274, 1199- abyssaltholeiite from the Mid-Atlantic Ridge near 24øN and 30øN 1261, 1974b. latitude, Contrib. Mineral. Petrol., 23, 38-52, 1969. Helgeson,H. C., T. H. Brown, A. Nigrini, and T. A. Jones,Calcula- Moore, J. G., and B. P. Fabbi, An estimate of the juvenile sulfur tion of mass transfer in geochemicalprocesses involving aqueous content of basalt, Contrib. Mineral. Petrol., 33, 118-127, 1971. solutions, Geochim.Cosmochim. Acta, 34, 569-592, 1970. Mottl, M. J., Chemical exchangebetween seawater and basalt during Helgeson, H. C., J. M. Delaney, H. W. Nesbitt, and D. K. Bird, hydrothermalalteration of the oceaniccrust, Ph.D. thesis,Harvard Summary and critique of the thermodynamic properties of the Univ., Cambridge, Mass., 1976. rock-forming minerals, Am. J. Sci., 278A, 1-228, 1978. Mottl, M. Jo, Metabasalts, axial hot springs, and the structure of Helgeson, H. C., D. H. Kirkham, and G. C. Flowers, Theoretical hydrothermalsystems at mid-oceanridges, Geol. Soc. Am. Bull., 94, prediction of the thermodynamicbehavior of aqueouselectrolytes 161-180, 1983. at high pressuresand temperatures,IV, Calculation of activity coef- Mottl, M. J., and H. D. Holland, Chemical exchange during hy- ficients,osmotic coefficients, and apparent molal and standardand drothermal alteration of basalt by seawater,I, Experimental results relative partial molal propertiesto 600øC and 5 kb, Am. J. Sci., 281, for major and minor componentsof seawater,Geochim. Cosmochim. 1249-1516, 1981. Acta, 42, 1103-1115, 1978. Humphris, S. E., and G. Thompson, Hydrothermal alteration of Mottl, M. J., and W. E. Seyfried, Jr., Experimental basalt-seawater oceanic basalts by seawater, Geochim. Cosmochim.Acta, 42, 107- interaction: Rock- vs. seawater-dominatedsystems with the origin 125, 1978. of submarine hydrothermal deposits (abstract), Geol. Soc. Am. Ito, E., and A. T. Anderson,Jr., Submarinemetamorphism of gabbros Abstr. Programs,9, 1104-1105, 1977. BOWERSAND TAYLOR: CHEMICAL AND ISOTOPIC MODEL OF HOT SPRINGS 12,605

Mottl, M. J., and W. E. Seyfried,Jr., Sub-seafloorhydrothermal sys- mineral formation and implications for the transport of heavy tems: Rock- vs. seawater-dominated,in SeafloorSpreading Centers: metals, Geochim. Cosmochim.Acta, 45, 135-147, 1981. HydrothermalSystems, edited by P. A. Rona and R. P. Lowell, pp. Seyfried,W. E., Jr., and W. E. Dibble, Jr., Seawater-peridotiteinterac- 66-82, Dowden, Hutchinson,and Ross,Straudsburg, Pa., 1980. tion at 300øC and 500 bars: Implications for the origin of oceanic Mottl, M. J., H. D. Holland, and R. F. Carr, Chemical exchange serpentinites,Geochim. Cosmochim. Acta, 44, 309-321, 1980. during hydrothermal alteration of basalt by seawater,II, Experi- Seyfried, W. E., Jr., and M. J. Mottl, Origin of submarine metal-rich mental resultsfor Fe, Mn, and sulfur species,Geochim. Cosmochim. hydrothermal solutions: Experimental basalt-seawaterinteraction Acta, 43, 869-884, 1979. in a seawater-dominated system at 300øC, 500 bars, in Proceedings Muehlenbachs,K., and R. N. Clayton, Oxygen isotope composition of the 2nd International Symposiumon Water-Rock Interaction, pp. of the oceaniccrust and its bearing on seawater,J. Geophys.Res., IV173-IV180, International Association for Geochemistry and 81, 4365-4369, 1976. Cosmochemistry,Strasbourg, France, 1977. Norton, D., Theory of hydrothermal systems, Annu. Rev. Earth Seyfried, W. E., Jr., and M. J. Mottl, Hydrothermal alteration of Planet. Sci., 12, 155-177, 1984. basalt by seawater under seawater-dominated conditions, Geochim. Norton, D., and R. Knapp, Transport phenomenain hydrothermal Cosmochim.Acta, 46, 985-1002, 1982. systems:The nature of porosity, Am. J. Sci., 277, 913-936, 1977. Sheppard, S. M. F., R. L. Nielson, and H. P. Taylor, Jr., Oxygen and Norton, D., and J. Knight, Transport phenomenain hydrothermal hydrogen isotope ratios of clay minerals from porphyry copper systems:Cooling plutons, Am. J. Sci.,277, 937-981, 1977. deposits, Econ. Geol., 64, 755-777, 1969. Norton, D., and H. P. Taylor, Jr., Quantitative simulation of the Spooner, E. T. C., R. D. Beckinsale,W. S. Fyfe, and J. D. Smewing, hydrothermalsystems of crystallizingmagmas on the basisof trans- •80 enrichedophiolitic metabasic rocks from E. Liquria (Italy), port theory and oxygen isotope data: An analysis of the Skaer- Pindos (Greece), and Troodos (Cyprus), Contrib. Mineral. Petrol., gaard Intrusion, J. Petrol., 20, 421-486, 1979. 47, 41-62, 1974. O'Neil, J. R., and J. K. Kharaka, Hydrogen and oxygen isotope ex- Spooner, E. T. C., R. D. Beckinsale,P. C. England, and A. Senior, change reactions between clay minerals and water, Geochim.Cos- Hydration,•80 enrichmentand oxidationduring ocean floor hy- mochim.Acta, 40, 241-246, 1976. drothermal metamorphism of ophiolitic metabasic rocks from E. O'Neil, J. R., and H. P. Taylor, Jr., The oxygenisotope and cation Liguria, Italy, Geochim.Cosmochim. Acta, 41, 857-872, 1977. exchangechemistry of feldspars,Am. Mineral., 52, 1414-1437, 1967. Stakes, D. S., and J. R. O'Neil, Mineralogy and stable isotope geo- O'Neil, J. R., and H. P. Taylor, Jr., Oxygen isotope equilibrium be- chemistry of hydrothermally altered oceanic rocks, Earth Planet. tween muscoviteand water, J. Geophys.Res., 74, 6012-6022, 1969. Sci. Lett., 57, 285-304, 1982. O'Neil, J. R., R. N. Clayton, and T. K. Mayeda, Oxygen isotope Stakes,D. S., H. P. Taylor, Jr., and R. L. Fisher, Oxygen isotope and fractionationin divalent metal carbonates,J. Chem. Phys., 51, geochemicalcharacterization of hydrothermal alteration in ophio- 5547-5558, 1969. lite complexesand modern oceaniccrust, in Ophiolitesand Oceanic Oudin, E., Hydrothermal sulfide depositsof the East Pacific Rise Lithosphere,edited by I. G. Gass, S. J. Lippard, and A. W. Shelton, (21øN), I, Descriptivemineralogy, Mar. Min., 4, 39-72, 1983. pp. 199-214, Blackwell Scientific Publications, Oxford, England, Parmentier,E. M., Numericalexperiments on •80 depletionin ig- 1984. neous intrusionscooling by groundwaterconvection, J. Geophys. Stoessell, R. K., Regular solution site-mixing model for chlorites, Res., 86, 7131-7144, 1981. Clays Clay Miner., 32, 205-212, 1984. Perry, E. C., and F. C. Tan, Significanceof oxygen and carbon iso- Styrt, M. M., A. J. Brackmann, H. D. Holland, B.C. Clark, V. tope variations in early Precambrian cherts and carbonate rocks of Pisutha-Arnond, C. S. Eldridge, and H. Ohmoto, The mineralogy southern Africa, Geol. Soc. Am. Bull., 83, 647-664, 1972. and the isotopic composition of sulfur in hydrothermal sul- Ragnarsd6ttir,K. V., J. V. Walther, and S. Arn6rsson,Description fide/sulfate depositson the East Pacific Rise, 21øN latitude, Earth and interpretation of the compositionof fluid and alteration miner- Planet. Sci. Lett., 53, 382-390, 1981. alogy in the geothermal system,at Svartsengi,Iceland, Geochim. Suzuoki, T., and S. Epstein, Hydrogen isotope fractionation between Cosmochim.Acta, 48, 1535-1553, 1984. OH-bearing minerals and water, Geochim.Cosmochim. Acta, 40, Reed, M. H., Seawater-basaltreaction and the origin of greenstones 1229-1240, 1976. and related ore deposits,Econ. Geol., 78, 446-485, 1983. Taube, H., Use of oxygen isotope effectsin the study of hydration of Riedesel, M., J. A. Orcutt, K. C. MacDonald, and J. S. McClain, ions, J. Phys. Chem.,58, 523-528, 1954. Microearthquakes in the black smoker hydrothermal field, East Taylor, H. P., Jr., The application of oxygen and hydrogen isotope Pacific Rise at 21øN, J. Geophys.Res., 87, 10,613-10,623, 1982. studiesto problemsof hydrothermal alteration and ore deposition, Robie, R. A., B. S. Hemingway,and J. R. Fisher, Thermodynamic Econ. Geol., 69, 843-883, 1974. propertiesof minerals and related substancesat 298.15øK and 1 bar Taylor, H. P., Jr., Water/rockinteractions and the origin of H20 in (105pascals) pressure and at highertemperatures, U.S. Geol.Surv. granitic batholiths,J. Geol.Soc. London, 133, 509-558, 1977. Bull., 1452, 456 pp., 1978. Taylor, H. P., Jr., Oxygen and hydrogen isotope studies of hy- Rona, P. A., Pattern of hydrothermal mineral deposition: Mid- drothermal interactionsat submarineand subaerialspreading cen- Atlantic Ridge crest at latitude 26øN, Mar. Geol., 21, M59-M66, ters, in NATO SymposiumVolume on Hydrothermal Processesat 1976. Seafloor Spreading Centers, edited by P. A. Rona, K. Bostr6m, L. Rona,P. A., Criteriafor recognitionof hydrothermalmineral deposits ' Laubier, and K. Smith, Jr., pp. 83-139, Plenum, New York, 1983. in oceanic crust, Econ. Geol., 73, 135-160, 1978. Taylor, H. P., Jr., and S. Epstein,Relationship between O•8/O •6 Sakai, H., and M. Tsutsumi,Experimental determination of D/H frac- ratios in coexistingminerals of igneousand metamorphicrocks, I, tionation factors between serpentineand water at 100ø to 500øC Principles and experimental results, Geol. Soc. Am. Bull., 73, 461- under 2000 bars water pressure: Some implication to the D/H 480, 1962. ratios of natural serpentinites,U.S. Geol. Surv. Open File Rep., 78- 701, 370-372, 1978. Taylor, H. P., Jr., and R. W. Forester, An oxygen and hydrogen isotopestudy of the Skaergaardintrusion and its country rocks: A Savin, S. M., and S. Epstein, The oxygen and hydrogen isotope geo- descriptionof a 55-m.y. old fossilhydrothermal system, J. Petrol., chemistry of clay minerals, Geochim.Cosmochim. Acta, 34, 25-42, 20, 355-419, 1979. 1970a. Savin, S. M., and S. Epstein,The oxygenand hydrogen isotope geo- Truesdell, A. H., Oxygen isotope activities and concentrations in chemistry of ocean sediments and shales, Geochim. Cosmochim. aqueoussalt solutionsat elevated temperatures--Consequencesfor Acta, 34, 43-63, 1970b. isotopegeochemistry, Earth Planet. Sci. Lett., 23, 387-396, 1974. Seyfried, W. E., Jr., Seawater-basalt interaction from 25ø-300ø and Von Damm, K. L., J. M. Edmond, B. Grant, C. I. Measures, B. 1-500 bars: Implications for the origin of submarinemetal bearing Walden, and R. F. Weiss, Chemistry of submarine hydrothermal hydrothermal solutionsand regulation of ocean chemistry,Ph.D. solutions at 21øN East Pacific Rise, Geochim. Cosmochim.Acta, in thesis,Univ. oI ;b.CallI., Los Angeles,1976. press,!985. Seyfried,W. E., Jr., and J. L. Bischoff,Hydrothermal transport of Walther, J. V., and D. M. Rye, Possiblestable isotopicsystematics of heavy metals by seawater:The role of seawater/basaltratio, Earth progressivemetamorphism (abstract), Geol. Soc. Am. Abstr. Pro- Planet. Sci. Lett., 34, 71-77, 1977. grams,14, 641, 1982. Seyfried,W. E., Jr., and J. L. Bischoff,Low temperaturebasalt alter- Welhan, J. A., and H. Craig, Methane and hydrogen in East Pacific ation by seawater:An experimentalstudy at 70øC and 150øC, Ge- Risehydrothermal fluids, Geophys.Res. Lett., 6, 829-831, 1979. ochim.Cosmochim. Acta, 43, 1937-1947, 1979. Welhan, J. A., and H. Craig, Methane, hydrogen and helium in hy- Seyfried,W. E., Jr., and J. L. Bischoff,Experimental seawater-basalt drothermal fluids at 21øN, East Pacific Rise, in NATO Symposium interaction at 300øCand 500 bars: Chemical exchange,secondary Volume on Hydrothermal Processesat Seafloor SpreadingCenters, 12,606 BOWERSAND TAYLOR: CHEMICAL AND ISOTOPICMODEL OF HOT SPRINGS

edited by P. A. Rona, K. Bostr6m, L. Laubier, and K. L. Smith, Jr., Yeh, H., and S. M. Savin, The extent of oxygen isotope exchange pp. 391-409, Plenum, New York, 1983. between clay minerals and sea water, Geochim.Cosmochim. Acta, Wenner, D. B., and H. P. Taylor, Jr., Temperatures of serpen- 40, 743-748, 1976. tinization of ultramaficrocks basedon •80/•60 fractionationbe- Yeh, H., and S. M. Savin, Mechanism of burial metamorphism of tween coexisting serpentine and magnetite, Contrib. Mineral. argillaceoussediments, 3, O-isotope evidence, Geol. Soc. Am. Bull., Petrol., 32, 165-185, 1971. 88, 1321-1330, 1977. Wenner, D. B., and H. P. Taylor, Jr., Oxygen and hydrogen isotope Yund, R. A., and T. F. Anderson, Oxygen isotope exchangebetween studiesof serpentinizationof ultramafic rocks in oceanicenviron- potassiumfeldspar and KC1 solution, in GeochemicalTransport and ments and continental ophiolite complexes,Am. J. Sci., 273, 207- Kinetics,edited by A. Hoffman, B. J. Giletti, H. S. Yoder, and R. A. 239, 1973. Yund, pp. 99-105, Carnegie Institution of Washington, Wash- Wolery, T. J., Some chemical aspectsof hydrothermal processesat ington, D.C., 1974. mid-oceanicridges--A theoreticalstudy, I, Basalt-seawaterreaction and chemical cycling between the oceanic crust and the oceans,II, T. S. Bowers, Department of Earth, Atmospheric and Planetary Calculation of chemicalequilibrium betweenaqueous solutions and Sciences, MassachusettsInstitute of Technology, Cambridge, MA minerals, Ph.D. thesis, Northwestern Univ., Evanston, Ill., 1978. 02139. Wolery, T. J., Calculation of chemical equilibrium between aqueous H. P. Taylor, Jr., Division of Geological and Planetary Sciences, solutionsand minerals: The EQ3/6 software package,Rep. UCRL California Institute of Technology,Pasadena, CA 91125. 52658, Lawrence Livermore Natl. Lab., Livermore, Calif., 1979. Wolery, T. J., EQ3NR, a computerprogram for geochemicalaqueous (ReceivedDecember 14, 1984 speciation-solubilitycalculations, Rep. UCRL 5, Distrib.Category revisedJuly 26, 1985; UC-70, Lawrence Livermore Natl. Lab., Livermore, Calif., 1983. acceptedJuly 30, 1985.)