Contrib Mineral Petrol (1995) 122: 134–151 C Springer-Verlag 1995

Martine D. Buatier ? Gretchen L. Früh-GreenAnne Marie Karpoff

Mechanisms of Mg-phyllosilicate formation in a hydrothermal systemat a sedimented ridge (Middle Valley, Juan de Fuca)

Received: 9 January 1995y Accepted: 17 June 1995

Abstract We present results of a detailed mineralogicaland geochemical study of the progressive hydrothermalalteration of clastic sediments recovered at ODP Site 858in an area of active hydrothermal venting at the sediment-ed, axial rift valley of Middle Valley (northern Juan deFuca Ridge). These results allow a characterization ofnewly formed phyllosilicates and provide constraints onthe mechanisms of clay formation and controls of miner-al reactions on the chemical and isotopic composition ofhydrothermal fluids. Hydrothermal alteration at Site 858is characterized by a progressive change in phyllosilicateassemblages with depth. In the immediate vent area, atHole 858B, detrital layers are intercalated with pure hy-drothermal precipitates at the top of the section, with apredominance of hydrothermal phases at depth. Sequen-tially downhole inHole 858B, the clay fraction of thepure hydrothermal layers changes from smectite to cor-rensite to swelling chlorite and finally to chlorite. Inthree pure hydrothermal layers in the deepest part ofHole 858B, the clay minerals coexist with neoformedquartz. Neoformed and detrital components are clearlydistinguished on the basis of morphology, asseen bySEM and TEM, and by their chemical and stable isotopecompositions. Corrensite is characterized by a 24 A˚stacking sequence and high Si- and Mg-contents, withFey(Fe1Mg) ratio of©0.08. We propose that corrensiteis a unique, possibly metastable, mineralogical phase andwas precipitated directly from seawater-dominated hy-drothermal fluids. Hydrothermal chlorite inHole 858Bhas a stacking sequence of 14 A˚ with Fey(Fe1Mg) ratios

of ©0.35. The chemistry and structure of swelling chlor-ite suggest that it is a corrensiteychlorite mixed-layerphase. The mineralogical zonation inHole 858B is ac-companied by a systematic decrease ind18O, reflectingboth the high thermal gradients that prevail at Site 858and extensive sediment-fluid interaction. Precipitation ofthe Mg-phyllosilicates in the vent region directly con-trols the chemical and isotopic compositions of the porefluids. This is particularly evident by decreases in Mgand enrichments in deuterium and salinity in the porefluids at depths at which corrensite and chlorite areformed. Structural formulae calculated from TEM-EDXanalyses were used to construct clay-H2O oxygen isotopefractionation curves based on oxygen bond models. Ourresults suggest isotopic disequilibrium conditions forcorrensite-quartz and swelling chlorite-quartz pre-cipitation, but yield an equilibrium temperature of3008 C+308 for chlorite-quartz at 32 m below the sur-face. This estimate is consistent with independent esti-mates and indicates steep thermal gradients of 10–118ymin the vent region.

Introduction

In oceanic hydrothermal systems, newly formed Fe andMg-rich phyllosilicates reflect fluid-rock interaction thatoccurs in the oceanic crust and in the overlying sedi-ments. In order to characterize fluid pathways and toquantify the input and loss of chemical elements throughhydrothermal processes, the mechanisms of phyllosili-cate formation must be well determined. Previous studieshave shown that several factors control the precipitationand evolution of hydrothermal minerals. In oceanicbasalts, hydrothermal phyllosilicates precipitate in voidsand fill vesicles, and can also replace interstitial glassand minerals (Alt et al. 1986; Laverne et al. 1989; Bu-atier and Honnorez 1990). The nature of precipitationand the chemistry of the authigenic phases are highlydependent on permeability, chemistry of the fluids andtemperature. As all of these factors can vary temporally

M.D. Buatier (✉)UniversiteLille I, Laboratoire de Se´dimentologie etGeodynamique, URA 719, F-59655 Villeneuve d’Ascq, France

G. L. Früh-GreenDepartment of Earth Sciences, ETH Zürich, CH-8092 Zürich,Switzerland

A. M. KarpoffUniversiteLouis Pasteur, Centre de Ge´ochimie de la Surface,1, rue Blessig, F-67084 Strasbourg, France

Editioral responsibility: J. Hoefs

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and spatially, significant variations in mineral composi-tions can be observed in the same rock and in the samevesicle (Laverne et al. 1989; Shau and Peacor 1992).When hydrothermal phyllosilicates are formed by directprecipitation at the sea water-oceanic crust interface, theymay form large deposits that can be intercalated with sedi-ments (Buatier et al. 1994; Buatier and Karpoff 1995).

At Middle Valley of the northern Juan de Fuca Ridge,lateral and vertical mineralogical zonations are well de-fined at ODP Site 858, which is located in an area ofactive hydrothermal discharge. The vertical distributionof clay minerals in the sedimentary sequence is charac-terized by the presence of smectite, chlorite and bothregular and irregular Mg-Fe mixed-layered clays, and isvery similar to zonations observed in diagenetic sedi-mentary rocks and low-grade metamorphic rocks (Betti-son and Schiffman 1988; Inoue and Utada 1991; Meunieret al. 1991). In this contribution, we present a mineralog-ical and geochemical study of the hydrothermal phyl-losilicates from the sediments at Site 858. These dataprovide constraints on the mechanisms of formation ofMg-rich phyllosilicates in an active hydrothermal systemwhere fluidyrocks ratios are generally high and allow acomparison with similar sequences in different environ-ments.

Geological setting and description of Site 858

Middle Valley, an axial rift valley of the northern Juan de FucaRidge, was the target of detailed investigations during Leg 139 ofthe Ocean Drilling Program (Davis and Villinger 1992). MiddleValley is a fault-bounded, sediment-covered spreading centre with

an active hydrothermal system detected by heat flow anomalies.Each of the four sites drilled in the eastern part of the valley ischaracterized by a distinct hydrologic environment (Fig. 1). Site855 is an area of fluid recharge at low temperature; whereas Site858 is an area of active fluid discharge at high temperature. Thesetwo areas are considered to be linked by a hydrothermal reservoir(Site 857) where high temperature fluids have been sealed by sed-iments. The fourth site, Site 856, is a former discharge zone, wherean older episode of hydrothermal activity led to extensive alter-ation of the sediments and deposition of a large body of massivesulphides (Davis, Mottl, Fisher et al. 1992). In most of these sites,hydrothermal alteration is characterized by the precipitation ofMg-phyllosilicates, carbonate as cement and in concretions, aswell as local anhydrite and pyrite. An investigation of the petrogra-phy, mineralogy and geochemistry of the sediments at Sites 856,857 and 858 has been presented in two previous contributions; onefocuses on the mineralogical and petrographic record of sediment-fluid interactions (Buatier et al. 1994); the other documents thegeochemical record of convective hydrothermal circulation (Früh-Green et al. 1994a). In this study, we compare the mineralogicaland geochemical characteristics of authigenic silicates recoveredat two selected holes at the site of active hydrothermal venting(Site 858) (Fig. 2).

At Site 858, five holes were drilled across the heat-flow anoma-lies and penetrated the sedimentary cover (Fig. 1).Hole 858A isthe most distal, located 100 m west of the active vent field;Hole858C is approximately 70 m west of the nearest active vent withinthe distal part of the vent field; Hole858B is located within a fewmetres of an active hydrothermal vent; andHoles858D and 858Fare in the centre of the venting area, about 70 m northeast of thenearest active vent (Fig. 2). Fluids currently discharge at tempera-tures between 255 and 2758 C through numerous vents (Davis,Mottl, Fischer et al. 1992). This study focuses onHole 858A (dis-tal area of the active vent region) andHole 858B (directly in thecore of the active upflow zone).

The sedimentary cover at Site 858 consists of turbiditic andhemipelagic sequences, typical of the Middle Valley area, and iscomposed of a mixture of detrital and neoformed minerals in vari-ous proportions (Buatier et al. 1994). In the distalHole 858A, theupper 30 metres of unmodified to weakly altered sediments con-

Fig. 1 Leg 139 location andheat flow map of MiddleValley with the location of thesites of Leg 139

136

Fig. 2 Schematic diagram de-picting lithological variations,bulk mineralogical zonationsand fluid flow at Site 858(modified after Früh-Greenet al., 1994a and Leybourneand Goodfellow, 1994). Nohorizontal scale is inferred

sist of varying proportions of detrital quartz, feldspar, amphibole,mica, clays (smectite, chlorite), and biogenic remains. Quartz,feldspar, and mica persist in hydrothermally altered sediments todepths of 260 mbsf (metres below the seafloor), where they coexistwith newly formed minerals. In the core of the upflow zone, atHole858B, pure hydrothermal layers (sediments with only authigenicphases) are intercalated with unaltered sediments in the upper partof the hole (to depths of 15.5 mbsf); whereas the deeper part of thehole consists of pure hydrothermal sediments.

Systematic variations in temperatures, pore-fluid composi-tions, degree of alteration, and sulphide content occur laterally andwith depth at Site 858. A mineralogical sequence is clearly estab-lished from the sea-waterysediment interface to the deeper parts ofthe holes (Fig. 2). Three general alteration zones have been distin-guished at Site 858: a carbonate-pyrite alteration zone in the upperparts of the holes, grading into a zone of anhydrite-pyrite alter-ation and siliceous high temperature assemblages (quartz-albite-chlorite-pyrite) at the bottom of the holes (Davis, Mottl, Fisheret al. 1992; Leybourne and Goodfellow 1994). The high tempera-ture assemblages may include wairakite and epidote in the coreof upwards fluid flow and intensive fluid-rock interaction atHoles858D and 858F (e.g. Leybourne and Goodfellow 1994). Thecarbonate alteration zone is the thickest in the least altered regionat Hole 858A, occurring from 10 to 80 mbsf, but is essentiallyabsent at the immediate vent region atHole 858B. The mineralogyof the sediments atHole 858B is dominated by authigenic phases.The shallowest hydrothermal unit fromHole 858B is composed ofclays with authigenic pyrite; sequentially downhole, it overliesclay-rich layers with some talc and anhydrite; and in the deepestunit, chlorite is the dominant phyllosilicate coexisting with authi-genic quartz. These pure hydrothermal layers display a mineralog-ical sequence in a sedimentary column of less than 40 metresthick.

Analytical methods

For X-ray diffraction (XRD) and stable isotope analyses, clay-sizefractions (,2-mm) were separated from bulk samples by setting ina water column. Prior to separation, each sample was dispersed indeionized water, disaggregated, decarbonated with a HCl Ny5 so-lution, and washed several times. Clay mineral assemblages weredetermined by routine XRD analysis using a Philips PW1710 dif-fractometer. Samples were run between 2.5 and 32.58 2u, 40 kVy20 mA, using a CuKaradiation, a Ni filter, and a scan speed of18ymin. Scanning electron microscope (SEM) observations usingsecondary electrons or backscattered electrons were made on Au-coated fragments of bulk samples or carbon coated thin sections,respectively, and using a JEOL JSM840 scanning electron micro-scope operating at 80 kV. The microstructures and mineralogies ofseveral samples were further analysed using transmission electronmicroscopy (TEM). Epoxy-impregnated specimens were thin-sec-tioned, mounted on copper grids and ion-milled. Samples wereexamined in a TEM-STEM Philips CM 30 operating at 300 KeV.Lattice fringe images were obtained using only 001 reflectionswith a 10mm diameter objective aperture that included 001 reflec-tions with 1#2. Most images were realised with Sherzer focusconditions. Chemical analyses were obtained in STEM with a Tra-cor X-ray detector for areas that were first characterized by TEMand lattice fringe imaging. Analyses were carried out using asquare raster 1000 A˚ wide to minimise diffusion of ions and sam-ple destruction. Quantitative analyses were obtained following theprocedures of Cliff and Lorimer (1975) and Lorimer and Cliff(1976). Sandard minerals, used to obtain proportionality constants(k values) for each element included: synthetic quartz for Si, natu-ral andalusite for Al, natural jadeite for Na, natural wollastonitefor Ca, and synthetic pure forsterite for Mg and synthetic purefayalite for Fe. A cold-stage holder was used for corrensite-richsamples in order to avoid sample damage.

Authigenic quartz separates were chemically isolated from$2–30mm size fractions using hot, concentrated sulphuric acid.

137

For oxygen isotope analysis, clay separates (,2 mm size fractions)were degassed at room temperature for at least 72 h under highvacuum and then transferred directly to nickel reaction vesselswith minimal exposure to air and heated at 1508 C for 5 h undervacuum. Oxygen was liberated from silicate samples by reactionwith ClF3 at 6008 C (Borthwick and Harmon 1982) and convertedto CO2 by reaction with heated carbon. For hydrogen isotope anal-ysis, clay separates were dried at 1508 C under vacuum overnight,then heated in a vacuum to.11008 C to liberate H2 and H2O.Molecular hydrogen was converted to water by reaction with cop-per oxide. The resulting total water was quantitatively converted tohydrogen by reaction with hot uranium (Bigeleisen et al. 1952).The isotopic ratios of all samples were determined by conventionalmass spectrometric analysis and are reported asd-values in permill (‰) relative to Standard Mean Ocean Water (SMOW). Theoverall reproducibility of oxygen isotope ratios averages+0.1‰for quartz and+0.2‰ for clays, and is+1‰ for hydrogen.

Mineralogical and chemical characterizationof phyllosilicates at Middle Valley

X-ray diffraction data

The interpretation of XRD patterns obtained on orientedpowders of clay fraction (,2 mm) from Holes858A and858B are presented in Table 1. The major difference be-tween the two holes is the persistence downhole of micain the clay fractions of altered sediments inHole 858A.

In both holes, sediments representing the unaltered sedi-mentary cover (Unit I) consist of a mixture of smectite,chlorite and mica, which are probably detrital in origin.In Hole858B, Unit I is interlayered with pure hydrother-mal deposits. Sequentially downhole inHole 858B, theclay fraction of the pure hydrothermal layers changessuccessively from smectite to corrensite to swellingchlorite and chlorite. These minerals are the majorphases of the clay fractions, but talc and allietite (a talc-smectite regular mixed-layered clay) were also detected(Table 1).

The XRD patterns that indicated mixed-layeredphases have been further characterized by comparingexperimental patterns with calculated patterns using theprogramme Newmod (Reynolds1 1985). For example, theXRD pattern of sample B5H3 58–62 displays several(00l) reflections at 31 A˚ , 14.9 A, 9.9 A, 7.4 A, 5.9 A,4.9 A, 3.6 A, 3.3 A and 2.9 A, characteristic of a regularmixed-layered chlorite-smectite mineral, i.e. corrensite.After glycolation, the first three reflections shift to32.9 A, 16 A and 7.9 A. Swelling chlorite was identifiedin samples between 35 and 81 mbsf inHole 858A and atabout 32 mbsf in Hole858B (Table 1). This phase is char-acterized by XRD patterns with reflections at 14.2 A˚ and7.1 A that slightly expand to 14.7 A˚ and 7.9 A. after gly-colation. In the deeper samples in both holes (at about

Table 1 Mineralogical assemblages of unaltered, altered and hy-drothermal sediments fromHoles858A and 858B as determined

by X-ray diffraction analyses. (X very abundant,X abundant,xsmall quantity,A allietite – talc-smectite mixed-layer)

Depth Unit Clay fraction composition

Hole Core, section, interval (mbsf) Mica Smectite Corrensite Swelling chlorite Chlorite Talc

139-858A- 2H-3, 36-40 5.76 I X X . . X .3H-1, 58–60 12.48 I X X . . X .4H-7, 15–17 30.55 I X x . . x .

5H-4, 27-30 35.67 IIA X . . X . .

11X-CC, 12–14 73.59 IIC X . . X x .12X-CC, 14–15 81.94 IIC X . . X . .14X-CC, 11-13 101.11 IIC X . . . X .18X-2, 132–134 142.42 IIC X . . . X .27X-1, 24–26 226.97 IIC x . . . X .31X-2, 3–5 266.17 IIC x . . . X .

139-858B- 1H-1, 124-126 I X X . . X .1H-2, 129-133 2.79 I X X . . X .1H-4, 108–112 5.59 I X X . . X .2H-1, 97–99 8.17 I x X . . X .2H-2, 45–47 9.15 I X x . . X .

2H-3, 75–77 10.95 IV . X . . x .2H-3, 126–129 11.46 IV x . X . x2H-4, 72–74 12.42 IV . . X x . .

2H-5, 113–115 14.33 I x x . x X .2H-6, 89–91 15.59 I X . . . X .

5H-2, 69–73 26.09 IIB . . X . . XA5H-3, 58–62 27.48 IIB . . X . . .5H-4, 55–59 28.95 IIB . . X . . x6H-1, 51–53 32.01 IID . . . X X X6H-1, 67–70 32.17 IID . . . X . .8X-1, 27–31 32.97 IID . . . . X .8X-2, 26–28 34.0 IID . x X . X .8X-4, 22-24 36.3 IID . x X . X .

138

Fig. 3 Backscattering electron microscope images of a weaklyaltered sediment fromHole 858A (a); a partially altered sedimentfrom Hole 858A with detrital minerals (b); and a pure hydrother-mal sediment fromHole 858B showing a very homogeneous tex-ture (c). (Mi mica,Q quartz,Cl clays)

Scanning electron microscope and backscatteredelectron images

The SEM low magnification backscattered electron im-ages allow the various textures of the sediments to beobserved. Samples from Unit I are made up of large de-trital grains of quartz, feldspar and mica surrounded by afine-grained, clay-rich matrix (Fig. 3a). Hydrothermallyaltered sediments fromHole 858A also contain detritalgrains (principally quartz and mica), but the clay matrixis dominant (Fig. 3b). The sediments sampled in purehydrothermal layers fromHole 858B are very homoge-neous and are mainly composed of clays (Fig. 3c). Fromsecondary electron SEM images, each type of clay iden-tified by XRD has a distinct morphology.Smectite lookslike a veil; corrensite morphology resembles corn flakes(Fig. 4a); and chlorite has a typical plate shape (Fig. 4b)(Buatier et al. 1994). Quartz occurs in samples B5H358–62, B6H-1 51–53 and B6H-1 67–70 and, based onmorphology asseen by secondary electron SEM images,is clearly authigenic in origin (Fig. 4).

Transmission electron microscope data

The textures of both soft and indurated sediments havebeen further observed by TEM. At low magnification (X25 000), corrensite-rich samples (B5H3 58–62 andB5H4 55–59) consist of aggregates of curled crystallites,with each crystallite approximately 25 nm in size(Fig. 5a). At higher magnification, one dimension highresolution images showing lattice fringes with a 24 A˚periodicity, with one layer at 10 A˚ and one layer at 14 A˚ ,can appear by varying the focus. This is characteristic ofthe dehydrated corrensite structure (Figs. 5b, c). Thegrain boundaries between two crystallites sometimeshave a non-topotactic relationship. Edge dislocations arecommonly observed in the crystallites. These defectsgenerally contain a total corrensite layer, i.e. a 24 A˚ thicklayer (Fig. 5c), but do not change the observed periodici-

Fig. 4 Scanning electron microscope images of authigenic quartzcoexisting with corrensite (a) and chlorite (b). Corrensite andchlorite can easily be distinguished on the basis of morphology(flake-shaped for corrensite and plate-shaped for chlorite). (Qquartz,Co corrensite,Ch chlorite)

34 mbsf in Hole858B and 100 mbsf inHole 858A) chlo-rite is the dominant phase. The XRD patterns of thedeepest samples fromHole858A characteristically showlow intensity odd reflections, whereas the even peaks arewell defined and may reflect a high Fe content. The sym-metry of the odd (00l) reflections suggests that Fe is inthe octahedral site of the talc layer and not in the brucitelayer (Moore and Reynolds 1989).

139

Fig. 5a–c Transmission elec-tron microscope (TEM) im-ages of corrensite:a low mag-nification image of a corren-site-rich sediments composedof aggregates of curled crys-tallites; b andc two dimensionhigh resolution image of acorrensite characterized by a24 A periodicity, the narrowareas show the presence of anedge dislocation (D) whichimplies a total layer of corren-site

Fig. 7 Swelling chlorite in sample 12XCC 14–15 observed byTEM: lattice fringe image consists of a majority of chlorite layersintercalated with some 10 A˚ layers

Fig. 6 TEM images of corrensite-chlorite mixed-layer in sample5H3 58–62 (containing pure corrensite according XRD analyses).The diffraction pattern of this area shows the superposition of thechlorite reflection with the corrensite reflection. (Chchlorite layer,Co corrensite layer)

potactic as shown by the image in Fig. 6a and 6b and bythe diffraction pattern showing the superposition of the(00l) reflections of corrensite with the (00l) reflectionsof chlorite (Fig. 6b).

According to XRD patterns, samples B8X2 26–28and A12XCC 14–15 contain swelling chlorite. The onedimension, high resolution TEM lattice fringe imagesshow a predominance of chlorite layers intercalated withsome 10 A˚ layers (Fig. 7). These images can be inter-preted in two ways. One possibility is that swelling chlo-rite consists of mixed layers of chlorite-smectite withapproximately 20 percent smectite layers. The other in-

ty. Diffraction patterns of such crystallites display irra-tional (00l) reflections typical of corrensite.

Sample B5H3 58–62, in which pure corrensite wasindicated by XRD data, contains corrensite mixed withsome 14 A˚ (chlorite) layers. These chlorite intercalationscan occur as one or two layers only, but occasionallypackets of five or six layers can be mixed with corrensite.The relationship between the two phases is always to-

140

Fig. 8a–c Chlorite observed by TEM:a low magnification imagewith straight and rigid aggregates of chlorite crystallites fromsample B6H-1;b low magnification image of sample B8X1;c thesame sample at higher magnification, the thick and coherent pack-ets of chlorite layers are defect free

terpretation is to consider this phase a mixed-layer cor-rensite-chlorite containing approximately 10% corren-site layers. The 10 A˚ layers are always intercalated withchlorite layers. Two 10 A˚ layers were never observededge to edge together between chlorite layers, suggestingthat swelling chlorite is a corrensite-chlorite mixed-lay-ered phase. Selected area electron diffraction (SAED)patterns of this clay mineral display (00l) rational reflec-tions at 14yn A similar to that of chlorite. The number ofcorrensite layers is not sufficiently high to give any re-flection. Large crystals of pure chlorite of a few hundredAngstroms in thickness devoid of any corrensite layerswere also observed in sample B8X2 26–28.

At low magnification, chlorite is easily distinguishedfrom corrensite on the basis of texture. Chlorite formsstraight, rigid aggregates of crystallites. In sample B6H151–53, they are generally composed of stacks of 3 to 10packets of layers of about 100 A˚ in thickness (Fig. 8a).Each packet defines a sequence of coherent layers withidentical crystallographic orientation. At higher magni-fication, the same sample displays (00l) lattice fringeimages with a 14 A˚ periodicity and with local intercalat-ed 10 Aspacings. Deeper samples fromHole 858B dis-play thicker coherent packets of layers that are interpret-ed to be chlorite crystallites of more than 500 A˚(Figs. 8b, c). Large packets of pure chlorite were alsoobserved inHole 858A. They can coexist with mica that,in some samples, appears to be replaced by chlorite.

Analytical electron microscope data

Analytical Electron Microscope (AEM) analyses werecarried out on a selected number of samples that werefirst well characterized by TEM. Typical structural for-mulae are presented in Table 2. In order to use theseanalyses to calculate oxygen isotope fractionation curves(see below), the cation distribution in octahedral andbrucite sheets of corrensite and chlorite has been esti-mated. For chlorite, the amount of Al in the brucite layeris considered to be equal to the Al substituted for Si in thetetrahedral sites. For corrensite, a similar distribution ofcations in all tetrahedral sites was assumed, followingthe recent crystallographic model proposed by Beaufortand Meunier (1994) for the corrensite structure. The

Fig. 9 AEM analyses plotted in diagrams with Fey(Fe1Mg) ver-sus total Al (a) and Siy(Si1Al) versus Fey(Fe1Mg) (b). The cor-rensite composition range is well defined here; this mineral isexceptionally Mg rich compared with saponite and chlorite fromthe same hole. Plotb demonstrates that the range of corrensitecompositions is not intermediate between saponite and chlorite

number of Al in the brucite layer of corrensite was as-sumed to be half of the total aluminium present in thetetrahedral sites.

According to their structural formulae, smectitesfrom sample B2H3 75–77 are saponite and contain ap-proximately 4 cations of Mg in the octahedral site. Thisis consistent with the trioctahedral character of theseclays as indicated by XRD data on the powdered clayfraction (Buatier et al. 1994).

The corrensite-rich sample B5H3 58–62 was investi-gated in detail. According to the ten determinations ofthe structural formulae, calculated on the basis of 25anhydrous oxygens, the chemical composition of corren-site varies significantly. These variations reflect the pres-ence of mixed layers of chlorite-corrensite as observed inhigh resolution images (Table 2 and Fig. 6). All the cal-culated structural formulae display a low Fey(Fe1Mg)ratio, with total Al ranging from 2.5 to 3 (Fig. 9a). Thesecorrensites are exceptionally Mg rich compared withother corrensite formed in pyroclastic rocks during dia-genesis (Inoue 1985) or in hydrothermally altered basaltsfrom DSDPHole 504B (Shau and Peacor 1992), but arevery similar to the composition of corrensite from theTaro Valley, described by Brigatti and Poppi (1984,1985). The number of cations in the octahedral site isbetween 8 and 9, which could be consistent with a combi-nation of one layer of chlorite and one layer of smectitethat respectively contain a maximum of 6 and 3 cations intheir octahedral site.

141

Tabl

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20

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61

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65

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67

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70

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73

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7.3

67

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37

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76

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37

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6.7

86

.51

6.5

37

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0.8

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.58

0.6

40

.69

0.6

70

.67

0.6

80

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31

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0.9

70

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1.2

21

.49

1.4

70

.86

Alo

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0.4

70

.52

0.4

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1.5

82

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0.3

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1.5

51

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1.0

20

.20

1.4

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.99

0.9

30

.61

0.6

10

.83

0.7

50

.57

0.6

10

.64

0.7

60

.85

0.7

60

.65

0.6

60

.72

0.8

1M

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.90

3.7

74

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4.5

34

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2.9

42

.24

3.1

91

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4.9

53

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23

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3.9

85

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3.0

7

AlB

0.3

30

.34

0.4

20

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0.7

30

.49

0.4

60

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0.7

40

.73

0.4

3M

gB

2.6

72

.66

2.5

82

.98

2.2

72

.51

2.5

42

.39

2.2

62

.27

2.5

7

Ca

0.2

70

.23

0.3

20

.24

0.3

70

.05

0.0

80

.04

0.0

90

.15

0.1

00

.07

0.0

90

.16

0.1

20

.07

Na

1.6

71

.30

1.0

90

.65

0.9

50

.12

0.1

40

.14

0.0

60

.10

0.1

00

.11

0.0

90

.10

0.2

00

.08

Sp

eci

me

nC

hlo

rite

B6

H1

51

–5

3

H3

–1

43

H3

–1

44

H3

–1

45

H3

–1

46

H3

–1

47

H3

–1

48

H3

–1

50

H3

–1

51

H3

–1

52

H3

–1

53

H3

–1

54

H3

–1

57

H3

–1

49

H3

–1

56

Si

7.0

36

.80

6.7

76

.31

6.4

46

.94

7.0

06

.72

6.9

27

.18

7.3

26

.74

7.3

16

.68

AlT

0.9

71

.20

1.2

31

.69

1.5

61

.06

1.0

01

.28

1y0

80

.82

0.6

81

.26

0.6

91

.32

Alo

c1

.93

1.6

71

.71

0.9

80

.86

1.7

91

.89

1.6

21

.73

2.5

12

.32

1.5

42

.46

1.6

1F

eo

c2

.96

3.0

22

.12

2.0

83

.16

2.4

23

.05

2.3

72

.33

2.2

12

.19

3.3

02

.25

2.4

1M

go

c0

.10

0.4

01

.30

2.4

41

.51

0.8

50

.07

1.1

61

.04

0.0

00

.25

0.3

60

.00

1.1

4

AlB

0.9

71

.20

1.2

31

.69

1.5

61

.06

1.0

01

.28

1.0

80

.82

0.6

81

.26

0.6

91

.32

Mg

B5

.03

4.8

04

.77

4.3

14

.44

4.9

45

.00

4.7

24

.92

5.1

15

.32

4.7

44

.79

4.6

8F

eB

0.0

00

.00

0.0

00

.00

0.0

00

.00

0.0

00

.00

0.0

00

.07

0.0

00

.00

0.5

20

.00

Ca

0.0

30

.08

0.0

20

.01

0.0

30

.05

0.0

20

.03

0.0

30

.03

0.0

30

.04

0.0

20

.04

Na

0.0

00

.00

0.0

00

.00

0.0

40

.00

0.0

50

.00

0.0

00

.00

0.0

80

.00

0.0

80

.00

Sp

eci

me

nC

hlo

rite

B8

X2

26

–2

8C

hlo

rite

A2

7X

12

4–

26

S9

-12

2S

9–

12

4S

9–

12

5S

9–

12

5S

9–

12

6S

9–

12

8S

9–

12

9J3

–1

04

J3–

10

5J3

–1

06

J3–

10

7J3

–1

10

J2–

11

1J3

–1

13

Si

5.6

75

.74

5.7

35

.73

5.6

35

.53

5.6

85

.60

5.6

05

.62

5.6

46

.07

6.0

65

.95

AlT

2.3

32

.26

2.2

72

.27

2.3

72

.47

2.3

22

.40

2.4

02

.38

2.3

61

.93

1.9

42

.05

Alo

c0

.52

0.6

50

.57

0.5

70

.63

0.0

50

.58

0.3

20

.53

0.4

30

.64

0.9

00

.93

0.8

7F

eo

c1

.94

4.0

93

.19

3.1

93

.10

3.7

23

.84

5.0

75

.20

5.3

54

.95

4.0

93

.94

3.3

0M

go

c3

.19

0.8

71

.84

1.8

41

.87

2.1

21

.22

0.0

00

.00

0.0

00

.00

0.4

70

.59

1.2

8

AlB

2.3

32

.26

2.2

72

.27

2.3

72

.47

2.3

22

.40

2.4

02

.38

2.3

61

.93

1.9

42

.05

Mg

B3

.67

3.7

43

.73

3.7

33

.63

3.5

33

.68

3.1

83

.03

3.1

53

.07

4.0

74

.06

3.9

5F

eB

0.0

00

.00

0.0

00

.00

0.0

00

.00

0.0

00

.42

0.5

70

.46

0.5

70

.00

0.0

00

.00

Ca

0.0

00

.00

0.0

00

.00

0.0

20

.01

0.0

00

.04

0.0

00

.00

0.0

10

.01

0.0

20

.00

Na

0.1

50

.08

0.1

60

.16

0.1

30

.13

0.1

30

.00

0.0

00

.00

0.1

20

.11

0.0

80

.19

142

A total of 28 analyses was carried out on various chlo-rite-rich samples (Table 2). All the chlorite composi-tions, calculated on the basis of 28 anhydrous oxygens,display a large range of compositions, but are distinct-ly more Fe-rich have lower Siy(Si1Al) ratios than cor-rensite (Fig. 9b). The two chlorite samples fromHole858B are chemically distinct. Sample 6H1 58–62 at32.01 mbsf that contains chlorite with few intercalated10 A layers is less Fe-rich than the large crystals of chlo-rites in sample B8X2 26–28 located at 34 mbsf. The lat-ter displays similar compositions to chlorites from sam-ple A27X1 24–26 located at 226.97 mbsf inHole 858A;both of which are characterized by low Siy(Si1A) ratiosand high Fey(Fe1Mg) ratios (Fig. 9).

In Figure 9, corrensite analyses form a cluster ofpoints with a high Siy(Si1Al) ratio compared with thechlorite analyses. Furthermore, the compositional rangeof corrensite is not intermediate between saponite andchlorite compositions as shown in the diagram Siy(Si1Al) versus Fey(Fe1Mg). If we consider corrensiteto be composed of 50% chlorite and 50% smectite layers,its structural formula should be the sum of the structuralformulae calculated for half cells of both minerals. Asdepicted in Table 2, the structural formulae of corrensiteanalysed in this study establish that this mineral can notbe described chemically as an intermediate phase be-tween the hydrothermal saponite and chlorite depositsfrom Hole 858B.

Stable isotope constraints on clay formation

Oxygen and hydrogen isotope ratios determined on theclay fractions from the distalHoles858A and 858C andthe vent region,Hole858B, are presented in Table 3. Sys-tematic lateral as well as vertical variations in the oxygenand hydrogen isotope ratios of the clay fractions fromvarious samples are observed at Site 858. The shallowestunaltered samples consisting of a mixture of smectite,chlorite and mica (Unit I) haved18O-values of 12 to 14‰and adD-value of284‰ (Fig. 10). These values clearlyreflect the detrital origin of the clays and are distinctfrom the isotope ratios of altered and pure hydrothermalsamples from all three holes. The effects of hydrothermalalteration are reflected by shifts in isotopic compositionsaway from these background values. InHole858B, theseeffects are observed at shallow depths (9.15 mbsf), asevident by sample B2H2 45–47 from Unit I (also consist-ing of smectite, chlorite and mica) withdD- andd18O-values of246‰ and 9.8‰, respectively. A progressivedepletion in18O with depth is characteristic of the clayminerals analysed in this study (Fig. 10). InHole 858B,d18O decreases to values between 2.9‰ and 4.4‰ insamples of authigenic corrensite or chlorite below27.5 mbsf. Similar depletions are observed in the distalHole 858A, but over a thickness of 200 metres of sedi-ments. The XRD analyses from the distalHole 858C in-dicate nearly pure chlorite mineral compositions (Bu-atier et al. 1994). All the samples are depleted in18O

relative to unaltered detrial compositions and haved18O-values similar to those inHole 858A, but over shallowerdepths (Table 3).

The decrease in oxygen isotope composition of theprecipitated phyllosilicates reflects both the high ther-mal gradients that prevail at Site 858 and extensive inter-action with hydrothermal fluids (Davis, Mottl, Fisheret al. 1992; Baker et al. 1994; Früh-Green et al. 1994a;Peter et al. 1994). In the immediate vent areas, tempera-tures of 1978 C at 19.5 mbsf inHole 858B and©2088 Cat 20.9 mbsf inHole 858D were measured during Leg139. These temperatures are similar or slightly lowerthan temperatures determined from fluid inclusion data(Peter et al. 1994) or estimated from oxygen isotope ra-tios of carbonates in the neighbouringHole858D (Bakeret al. 1994; Früh-Green et al. 1994a). InHole 858D, permil fractionations between carbonates and pore watersrange from 18.1‰ at 6.75 mbsf to 11.1‰ at 16.1 mbsf(G. L. Früh-Green, unpublished data) and correspond toan average temperature gradient of 11.48 Cym in the up-per 20 m of the hole. Assuming a similar temperaturegradient for past hydrothermal alteration, temperaturesclose to 300–3508 C would have prevailed at the bottomof Hole 858B during phyllosilicate precipitation. Maxi-mum temperatures of approximately 3208 C have beenestimated from fluid inclusion data of Site 858, whichsuggests that at the time of authigenic mineral precipita-tion, the hydrothermal fluids may have been approxi-mately 408 C hotter than the present-day vent fluidswhich are currently venting at 255–2758 C (Peter et al.1994).

In three pure hydrothermal layers in the deepest partof Hole 858B, the clay minerals coexist with neoformedquartz. According to SEM images on sample B6H1 51–53, the mineralogical assemblage of quartz and chloriteappears to have formed contemporaneously (Fig. 4). Theoxygen isotope ratios determined on different size frac-tions of quartz from these three samples are presented inTable 3. The values of the authigenic quartz coexist-ing with corrensite (at 27.48 mbsf) and chlorite (at32.01 mgsf) are very similar, with an averaged18O-valueof 11.5‰ (+0.17,n54); whereas the quartz from sam-ple B6H1 67–70 (at 32.18 mbsf) has a lowerd18O valueof 10.6‰ (+0.14,n53).

The hydrogen isotope ratios of the authigenic phasesin Hole 858B are relatively constant, withdD between244‰ and239‰ (Table 3). Although hydrogen isotopefractionation factors for most phyllosilicates are poorlydefined at temperatures below 4008 C, experimental andnatural data suggest that clay minerals such as illite,kaolinite and smectite are relatively insensitive to tem-perature variations. However, H-isotope fractionations inthese minerals as well as in chlorite and serpentine maybe strongly dependent on mineral chemistries (e.g. Savinand Lee 1988; Marumo et al. 1980; Sheppard and Gilg,in press). In spite of the uncertainties in fractionationcalibrations, most experimental and empirical data sug-gest an H-isotope fractionation of approximately 30–50‰ in Mg-rich smectite and chlorite at temperatures

143

Tabl

e3

Oxy

ge

na

ndhy

dro

ge

nis

oto

pe

ratio

so

fcl

ayfr

act

ion

sa

nda

uth

ige

nic

qua

rtz

inO

DP

Hol

es8

58

Aa

nd8

58

B,w

ith

tem

pe

ratu

ree

stim

ate

sb

ase

do

nb

ond

-typ

eca

lcu

latio

ns

and

em

pir

ica

lca

libra

tion

s.In

de

pe

nda

ntte

mp

era

ture

de

term

ina

tion

sa

rein

clu

de

dfo

r

com

pa

riso

n(Sm

sme

ctit

e,S

ap

sap

on

ite

,Ch

lch

lori

te,I

lill

ite

,sw

.C

hls

we

llin

gch

lori

te,

Co

rrco

rre

nsi

te,Q

tzqu

art

z,T

lkta

lc;

min

era

lsin

pa

rent

he

ses

rep

rese

ntle

ssa

bu

nda

nto

rm

ino

rp

hyllo

silic

ate

ph

ase

sa

sd

ete

rmin

ed

by

XR

D,

Bu

atie

re

tal.

19

94

)

Oxy

ge

nis

oto

pe

tem

pe

ratu

res

Ind

ep

end

ent

tem

pe

ratu

res

OD

PH

oley

De

pth

Min

era

log

yG

rain

size

dD

‰d

18O

‰C

lay-

wa

tera

Qu

art

z-w

ate

rbQ

tz-C

hlc

De

pth

Me

asu

redd

Flu

idsa

mp

le(m

bsf

)(X

RD

resu

lts)

(mm

)(S

MO

W)

(SM

OW

)d

18O

(H2

O)

d1

8O

(H2O

)d

18O

H2

Od

18O

(H2

O)

(mb

sf)

incl

usi

one

50

‰5

2.5

‰5

0‰

52

.5‰

85

8A

21

.40

39

A2

H3

36

–4

05

.76

Ch

l1S

m1

Il,

22

84

13

.73

0.9

05

5A

5H

42

7–

30

35

.67

sw.

Ch

l1

Il,

28

.79

2.4

01

19

A1

2X

CC

14

–1

58

1.9

4Il1

sw.

Ch

l,

27

.11

11

.70

15

3A

18

X2

13

2–

13

41

42

.42

Ch

l(Il

),

27

.71

30

17

02

36

.24

25

1–

31

1A

31

X2

3–

52

66

.17

Ch

l(Il

),

22

44

5.0

17

02

20

24

6.2

02

30

–2

65

85

8B

B1

H2

12

9–

13

32

.79

Sm1

Ch

l1Il

#2

12

.5B

2H

24

5–

47

9.1

5C

hl1Il

(Sm

),

22

46

9.8

B2

H3

75

–7

71

0.9

5S

ap

(Ch

l),

22

40

8.5

14

01

80

11

.31

13

7–

16

7B

2H

47

2–

74

12

.42

Co

rr(s

w.

Ch

l),

24

.42

05

26

0B

2H

51

13

–1

15

14

.33

Ch

l(sw

.C

hl

+S

m+

Il)

,2

23

95

.81

70

19

.50

19

7B

2H

68

9–

91

15

.59

Ch

l(Il

),

26

.01

60

21

0B

5H

35

8–

62

27

.48

Co

rr,

22

42

2.9

24

03

20

19

0B

5H

35

8–

62

27

.48

Au

thig

en

icQ

tz1

0#

x#2

01

1.7

23

02

75

B5

H3

58

–6

22

7.4

8A

uth

ige

nic

Qtz

$2

01

1.4

23

02

85

B5

H4

55

–5

92

8.9

5C

orr

(Tlk

),

23

.42

30

30

0B

6H

15

1–

53

32

.01

Ch

l(S

w.

Ch

l1

Tlk

),

22

44

3.6

21

02

80

30

0B

6H

15

1–

53

32

.01

Au

thig

en

icQ

tz2#x#

10

11

.42

30

28

5B

6H

15

1–

53

32

.01

Au

thig

en

icQ

tz$

10

11

.52

30

28

0B

6H

16

7–

70

32

.18

Ch

l#

24

.32

00

26

0(–

)B

6H

16

7–

70

32

.18

Au

thig

en

icQ

tz2#x#

10

10

.62

46

30

0B

6H

16

7–

70

32

.18

Au

thig

en

icQ

tz,

20

10

.62

47

30

0

85

8C

7H

37

4–

77

44

.24

Ch

l,

25

.91

60

26

01

4.4

29

0–

13

87

H4

3–

64

6.0

3C

hl

,2

6.2

15

02

50

14

.60

11

2–

15

08

H1

17

–1

94

6.6

7C

hl(

Il)

,2

8.4

12

51

40

16

.24

12

6–

19

21

1X

15

2–

54

55

.02

Ch

l,

26

.81

40

21

02

2.5

06

71

3X

13

8–

40

74

.08

Ch

l,

24

.71

75

23

04

2.6

01

22

aB

ond

-typ

eca

lcu

latio

ns

ofS

avin

and

Le

e(1

98

8)

and

usi

ng

min

era

lco

mp

osi

tion

sd

ete

r-m

ine

db

yA

EM

an

aly

ses,

as

dis

cuss

ed

inth

ete

xtb

Sh

arp

and

Kir

sch

ne

r(1

99

4)

cC

om

bin

ed

fra

ctio

na

tion

so

f(1

)a

nd(2

)a

bov

ed

Tem

pe

ratu

res

me

asu

red

du

rin

gL

eg

13

9(D

avis

,M

ott

l,F

ish

er

eta

l.,

19

92

)e

Flu

idin

clu

sio

nd

ata

fro

mP

ete

re

tal.

19

94

144

Fig. 10 Distribution of clayfraction assemblages,d18O anddD as a function of depth (me-tres below the seafloor)

between 100 and 4008 C (e.g. Wenner and Taylor 1973;Heaton and Sheppard 1977; Suzuoki and Epstein 1976;Sheppard and Gilg, in press). Thus, thedD-values ofthe authigenic phases inHole 858B are consistent withprecipitation from seawater-dominated hydrothermalfluids.

Pore-waters collected from the sediments at Site 858haved18O-values from21.3 to 2.9‰ anddD values up to7‰ (Früh-Green et al. 1994b; Früh-Green et al., inpreparation). InHoles858A and 858C, the isotope ratiosremain relatively constant withd18O-values of 0+0.2‰anddD-values 0+1‰, which are essentially identical tothe isotopic compositions of the bottom water in theMiddle Valley region (G. L. Früh-Green, unpublisheddata). In contrast, thed18O-values of the pore-watersin Hole 858B vary from21.3‰ to 0.2‰, with a weaktrend towards more positive values with depth. The hy-drogen isotope ratios are slightly enriched in deuteriumrelative to seawater. ThedD-values range from 0‰ atshallow depths to 2‰ at 27 mbsf, where corrensite isfound (Früh-Green et al., in preparation). Although thechanges in isotope ratios provide important constraintson the present-day hydrothermal circulation pattern andinteraction between fluids and sediment at the site ofactive venting, it can not necessarily be assumed that theisotopic compositions of the present-day pore-waters areidentical with the compositions of the fluids from whichthe Mg-phyllosilicates inHole 858B were deposited.Modification of the isotopic and chemical compositionof the pore-waters through fluid-sediment interactions isparticularly evident in the upper 20 m of the neighbour-ing Hole858D. The oxygen isotope compositions of thepore-waters inHole858D vary linearly and reach a max-imum of 2.9‰ at shallow depths (Früh-Green et al.1994b). Increases in thed18O-values correspond directlyto increases in salinity, iron and manganese, and provideevidence for present-day lateral advection of hydrother-

mal fluids at shallow depths below the sedimentyseafloorinterface (Davis, Mottl, Fisher et al. 1992; Baker et al.1994; Früh-Green et al. 1994 a, b).

Formation temperature estimates

The combination of oxygen isotope ratios of coexistingquartz-clay pairs, pore-water isotope data and indepen-dent temperature estimates provides an ideal opportuni-ty to test available fractionation calibrations and to deter-mine the degree of isotopic equilibrium during hy-drothermal alteration. Four approaches are commonlyused in the calibration of oxygen isotope equilibriumfractionation factors: laboratory equilibrium experi-ments; statistical mechanical calculations; empirical es-timates based on the isotopic compositions of naturallyoccurring samples; and calculations based on empiricalbond-type models and increment methods. Large dis-crepancies exist among the various calibrations and, con-sequently, the different approaches yield very differentapparent temperatures. Furthermore, because of slow ex-change rates and isotopic disequilibrium effects, the ex-perimental calibration of fractionations at low tempera-tures and fractionations of hydroxyl-bearing silicates re-mains problematic. In the absence of experimental datafor Mg-phyllosilicates, estimates of oxygen isotope frac-tionations can be made by empirical calculations basedon bond strength and crystal chemical models (e.g.Schütze 1980; Savin and Lee, 1988; Richter and Hoernes1988; Zheng 1991, 1993a, b; Hoffbauer et al. 1994). It isbeyond the scope of this study and is not our intentionhere to discuss the limitations and discrepancies of thevarious calibrations. Critical evaluations and discussionsof empirical and experimental calibrations of fractiona-tion factors for a number of minerals are presented else-where (e.g. Savin and Lee 1988; Kyser 1987; Clayton

145

et al. 1989; Sharp and Kirschner 1994; Sheppard andGilg, in press). In this study, estimates of the tempera-tures of formation of the pure hydrothermal layers inHole858B have been made in three ways: (1) from calcu-lated clay-water fractionation curves; (2) from empiricalquartz-water fractionation data; (3) by combining theclay-water calibrations with empirical quartz-waterfractionations.

As a first approximation, the fractionation factors ofchlorite-water, corrensite-water and saponite-water havebeen calculated by applying the oxygen bond model ofSavin and Lee (1988) for the measured chemical compo-sitions. This method assumes that the isotopic behaviourof oxygen in a chemical bond is similar regardless of themineral in which the bond occurs. By assuming that oxy-gen isotope fractionation is simply the weighted sum ofthe fractionation of oxygen bonds, the relative propor-tions of different oxygen bonds in a mineral can be deter-mined with help of a crystal model. This method thusrequires that the type – and fractionation equations – ofthe different cation-oxygen and cation-OH bonds beknown, and consequently the nature of cation distribu-tion in the brucite layer must be determined. Further-more, the accuracy of the fractionation curves calculatedin this manner will clearly be limited by the accuracy ofthe individual bond fractionations. Structural formulae,calculated from AEM analyses presented in Table 2,were used for the determination of the relative proportionof the various oxygen-containing bonds (Savin and Lee1988).

Because oxygen isotope fractionation factors betweenquartz and water are poorly constrained or controversialat low temperatures (e.g. Matsuhisa et al. 1979; Claytonet al. 1989; Sharp and Kirschner 1994), available experi-mental data should not be extrapolated to temperaturesbelow 400–6008 (e.g. see discussion in Clayton et al.1989). In this study, we have chosen the empirical cali-bration of quartz-water recently published by Sharpand Kirschner (1994): 1000 lna(Qtz-H2O)5 3.653106yT2

22.9. This calibration was derived from natural datafrom various low – to medium-grade metamorphic rocks(100–6008 C) in which isotopic equilibrium could beshown. We prefer this calibration because the use of nat-ural equilibrium fractionations avoids uncertainties dueto kinetic effects during recrystallisation in direct-ex-change experiments and yields more geologically rea-sonable temperatures, particularly at low temperatures(see discussion in Sharp and Kirschner, 1994).

Calculated fractionation curves

The fractionation curves obtained for saponite-water,chlorite-water and corrensite-water from five sampleswith different chemistries are compared in Fig. 11.These calculations predict that Mg-rich saponite (sam-ple B2H3 75–77) and corrensite (sample 8B5H3 58–62)have larger fractionation factors than Mg-Fe chlorite atany given temperature. As the bond-type model yields

Fig. 11 Oxygen isotope fractionation curves calculated for claycompositions given in Table 2 and using the bond-type approach ofSavin and Lee (1988). The estimated fractionations correspond tothe following equations for 1000 lnaclay-water:

Saponite-water (B2H3)27.213103T2116.943106T2221.1533109T2310.10331012T2421.92Corrensite-water (B5H3)20.103103T2114.283106T2220.6983109T2310.06331012T2428.78Chlorite-water (A27X) 1.483103T2113.923106T2220.7693109T2310.06931012T24211.25Chlorite-water (B6H1) 3.413103T2113.103106T2220.5583109T2310.05031012T24212.63Chlorite-water (B8X2) 1.583103T2113.873106T2220.7563109T2310.06831012T24211.31

the same fractionation coefficient between Fe-O andMg-O bonds, the curves obtained for chlorite with vary-ing Fe- and Mg-contents are very similar. This results inoverlapping curves for samples B8X2 26–28 and A27X24–26 as seen in Fig. 11. Almost identical curves werealso obtained for the Mg-rich corrensite from this studyand the Fe-rich corrensite analysed by Shau and Peacor(1992) from altered basalt inHole 504B. The fractiona-tion curve calculated for the more Si-rich chemical com-position of the chlorite sample B6H1 51–53 (compareTable 2) is intermediate between the corrensite and thepure Fe-Mg chlorite curves and reflects the small amountof corrensite layers in this sample (see TEM section andTable 2).

The bond-type chlorite-water fractionation curve cal-culated for sample B6H1 51–53 is further compared withother experimental and empirical chlorite-water calibra-tions in Fig. 12a. At temperatures between approximate-ly 130 and 4008 C, the calculated chlorite-H2O curve

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agrees well with the experimental data of Cole (1985),but predicts larger fractionations than the incrementmethod of Zheng (1993b) and recalculated empiricalcalibrations of Wenner and Taylor (1971). If more Si-poor, Fe- or Mg-rich compositions of chlorite were as-sumed (as in sample A27X1 24–26), the calculatedcurves would show greater agreement with the incrementmethod of Zheng (1993). The original chlorite-H2O cali-bration of Wenner and Taylor (1971) was calculated fromnatural and experimental fractionations and was in partbased on preliminary experimental data of the quartz-water oxygen isotope fractionation, which has long sincebeen revised. In this study, we have used the fractionationrelationships discussed in the original work of Wennerand Taylor, but have recalculated the curves in two ways.Curve (4) in Fig. 12a is based on the published quartz-water fractionation factors of Clayton et al. (1972), as-suming a CO2-H2O fractionation factor of 1.0412 at258 C (Friedman and O’Neil 1977). Curve (5) was calcu-lated from the empirical calibration of quartz-water ofSharp and Kirschner (1994), as discussed above. Thiscurve yields similar fractionations as the original curvepublished by Wenner and Taylor (1971). The quartz-chlorite fractionation curve for sample B6H1 51–53, de-termined by combining the bond-type chlorite-waterfractionation curve with the quartz-water fractionationcurve of Sharp and Kirschner (1994), is compared withother experimental and empirical quartz-chlorite frac-tionation curves in Fig. 12b.

Temperatures of precipitation

Because of the variability in chlorite and corrensite min-eral chemistries as well as the discrepancies in experi-mental and empirical determinations of the quartz-waterfractionation factors, only a range of possible tempera-tures can be determined for authigenic mineral precipita-tion at Site 858 (compare Fig. 12). These data, however,provide constraints on the degree of isotopic equilibriumduring precipitation of the authigenic phases and, to-gether with independent temperature estimates, allowthe applicability of the oxygen bond method to be evalu-ated. Two approaches have been taken. The first ap-proach takes into acount the range of oxygen isotoperatios in the pore-waters (21.3 to 2.9‰) at Site 858 andcalculates equilibrium temperatures for phyllosilicateprecipitation from unaltered seawater with ad18O-valueof 0‰ and from an altered hydrothermal fluid with ad18O-value of 2.5‰. In samples that coexist with quartz,these estimates are then compared with temperatures es-timated from the empirical curve of Sharp and Kirschner(1994). The second approach estimates temperaturesfrom the calculated quartz-chlorite fractionations (asdiscussed above and shown in Fig. 12b) and then calcu-lates the oxygen isotope ratio of water in equilibriumwith quartz at that temperature (using the calibration ofSharp and Kirschner, 1994). The calculated oxygen iso-tope composition of the fluid can then be compared with

Fig. 12a Comparison of oxygen isotope fractionation factors be-tween chlorite and water. The curve labelled1 represents bond-type calculations for chlorite sample B6H1 51–53 (see Table 2 forcompositions and Fig. 11 for equation). Curve2 was calculated byZheng (1993b) using the increment method. Curve3 (Cole 1985)is based on experimental formation of chlorite by hydrothermalalteration of biotite. Curve4 is expressed by the equation:1031naChl2H2O 5 1.693106 y T22 24.57, modified after Wennerand Taylor (1971) to take into account corrections in the experi-mental quartz-water fractionation data of Clayton et al. 1972 (seediscussion in text). Curve5 is expressed by the equation:1031naChl2H2O 5 1.513106 y T2224.57, recalculated after Wen-ner and Taylor (1971) to include the empirical quartz-water frac-tionation factors of Sharp and Kirschner (1994). A temperaturerange of 508 C is predicted from the five curves for a chlorite witha value of 3.6‰ in equilibrium with unaltered seawater, as shownby thehorizontal finely stippled line.If an altered fluid composi-tion of 2.5‰ is assumed, higher temperatures with a range of 408 Care indicated by thehorizontal dotted line.b Estimates of theoxygen isotope fractionation between quartz and chlorite. Curve1combines the empirical quartz-water fractionation curve of Sharpand Kirschner (1994) with the bond-type fractionation curve ofchlorite-water for sample B6H1 51–53. Curve2 has been calculat-ed by Zheng (1993b) using the increment method. Curve3 com-bines the empirical quartz-water fractionation curve of Sharp andKirschner (1994) with the experimental data of Cole (1985) forchlorite-water. Curve4 combines the quartz-water fractionationdata of Clayton et al. 1972 (in Friedman and O’Neil, 1977) withCurve 4 in Fig. a. Curve 5 combines the empirical quartz-waterfractionation curve of Sharp and Kirschner (1994) with curve5 inFig. a. The fractionation of 7.9‰ measured for sample B6H1 51–53 is shown by thehorizontal stippled line

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measured pore-water compositions. A summary of theresults is given in Table 3.

The bond calculation of the saponite sample B2H375–77 (Fig. 11) predicts a temperature of 1408 C forequilibrium with seawater and a temperature of 1808 Cfor equilibrium with an altered fluid of 2.5‰. This tem-perature range is essentially identical to the fluid inclu-sion temperatures (137–1678 C) measured on anhydritejust 37 cm deeper in the hole (Peter et al. 1994). Highertemperatures of 1658 C (in equilibrium with seawater)and 2308 C (in equilibrium with an altered fluid) are pre-dicted by the smectite-water fractionation factor (1000lna5 2.553106T22 24.05) recently proposed by Shep-pard and Gilg (in press). The bond calculations of equi-librium temperatures for nearly all the clays below12 mbsf in Hole 858B yield consistent temperatureswithin a range of 200–2308 C for precipitation from sea-water and 260–3008 C for precipitation from an alteredfluid of 2.5‰. These temperature ranges are similar tothose calculated from the three samples of coexistingquartz (Table 3). The+0.2‰ uncertainty in the isotopicanalyses corresponds to an uncertainty in the tempera-ture determination (in this temperature range) of approx-imately +58 C. Comparison of the calculated chlorite-water curve for sample B6H1 51–53 with the other cali-brations shown in Fig. 12 indicates a temperature rangeof 508 C for chlorite with a value of 3.6‰ in equilibriumwith unaltered seawater and a range of 408 C for equi-librium with an altered fluid.

The bond-type calculation of the quartz-chlorite frac-tionation of 7.9+0.4‰ for sample B6H1 51–53, shownas curve (1) in Fig. 12b, yields temperatures of300+158 C and indicates equilibrium with a fluid of3.2+0.5‰. This estimate is in good agreement with tem-peratures predicted from curves (3) and (5) in Fig. 12b,both of which have been combined with the quartz-wateroxygen fractionations of Sharp and Kirschner (1994).The other two curves, which are based on the quartz-wa-ter calibration of Clayton et al. (1972) consistently yieldlower temperatures. A temperature of 300+158 C at32 mbsf corresponds to a temperature gradient of ap-proximately 108 Cym, in agreement with the temperaturegradient predicted from carbonate data inHole 858D(Baker et al. 1994; Früh-Green et al. 1994a) and mea-sured inHole 858B (David, Mottl, Fisher et al. 1992).Considering the uncertainties in the isotope determina-tions and in the fractionation calibrations, we considerthis temperature to be a reasonable estimate within atemperature range of+308 C. The calculated fluid com-position of 3.2+9.5‰ further agrees well with pore-wa-ter compositions measured in the neighbouringHole858D where present-day lateral advection of hydrother-mal fluids is considered to be occurring. These data thusprovide evidence that the oxygen bond method of Savinand Lee (1988) yield reasonable estimates of Si-richchlorite-water oxygen isotope fractionations, at least inthe temperature range of approximately 100 to 3508 C.

For the corrensite-quartz pair from sample B5H3 58–62 at 27.5 mbsf, the bond-type calculation yields a tem-

perature of approximately 2408 C for corrensite and2308 C for quartz in equilibrium with unaltered seawater.If we assume a composition of 2.5‰ for the hydrother-mal fluid, formation temperatures would be 3008 for cor-rensite and 2808 for quartz. Bond-type calculations ofquartz-corrensite oxygen isotope fractionation give atemperature of 1908 C with a value ford18O of 22.4‰for the coexisting water, which is approximately 1‰more negative than any of the pore-waters measured atMiddle Valley (Früh-Green et al., in preparation). Dise-quilibrium conditions are indicated for the third quartz-clay mineral pair in sample 6H1 67–70. The chlorite-wa-ter and quartz-water fractionations yield temperatures of200 and 2458 C, respectively, for deposition from unal-tered seawater (see Table 3) and temperatures of 260 and3008 C, respectively, for deposition from a fluid of 2.5‰.As the calculated quartz-chlorite fractionation curvesbecome relatively constant at fractionations of 6.4‰(Fig. 12b), no temperature estimates for this quartz-chlorite pair were possible using the bond-type methoddescribed above.

Application of the oxygen bond calculations furtherallows an evaluation of temperatures and isotopic re-equilibration in the distal holes at Site 858 (Table 3). Thepresence of small quantities of mica in altered samplesfrom Hole 858A cannot unequivocally be distinguishedby the oxygen isotope data; however, they are probablydetrital in origin. Their occurrence is always correlatedwith the presence of detrital feldspar in the bulk rock.The TEM analyses on these micas indicate that they arelarge crystals of muscovite. Their presence inHole858Areflects its distal position in the vent area. Hydrothermalalteration is probably more diffuse at this hole and per-mits the preservation of detrital phases. The preservationof 18O-rich detrital mica would explain the relativelypositive d18O-values of the deepest samples inHoles858A and sample 8H1 17–19 inHole 858C. In Table 3,calculated equilibrium temperatures of relatively purechlorite samples inHoles858A and 858C are comparedwith temperatures measured during Leg 139 and thosedetermined from fluid inclusions (Peter et al. 1994). Ingeneral, the calculated temperatures for equilibriumwith unaltered seawater are lower than those determinedfrom fluid inclusion data from samples at similar depthsor those predicted by borehole temperatures extrapolat-ed to depth. If a fluid compositions of 2.5‰ is assumed,the calculated temperatures approach those of the fluidinclusion data, which may provide evidence that, in thepast, the hydrothermal fluids were more enriched in18Othan the present day pore-waters. The discrepancy be-tween predicted and measured pore-water isotopic com-positions and temperatures in the two distal holes, as wellas the variable isotopic compositions of the clays inHole858C, may reflect fluctuating conditions of fluid flow inthe distal regions and incomplete oxygen isotope ex-change during hydrothermal alteration.

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Mechanisms of Mg-phyllosilicate formation

The mineralogical zonation determined by XRD analy-ses is clearly established in bothHoles858A and 858B.The clay fractions show a successive change from smec-tite, to corrensite, to swelling chlorite and to chlorite inHole 858B. InHole 858A, this zonation is present over adeeper sedimentary sequence, although no corrensitewas identified but only swelling chlorite. Smectite-richchlorite-smectite mixed-layered clays were not ob-served; however, the occurrence of swelling chlorite in-dicates that chlorite-rich, chlorite-smectite mixed layerscan be present. The HRTEM analyses have shown thatswelling chlorite could be interpreted as a corrensiteychlorite mixed-layered clay. Most studies on trioctahe-dral clay minerals report similar sequences with the ab-sence of smectite-rich, chloriteysmectite mixed layers.For example, Inoue and Utada (1991) and Inoue et al.(1984) have studied the conversion of smectite to chloritein pyroclastic and volcanoclastic sediments and haveshown that the expandability of the clay fraction decreas-es from 100–80% to 50–40% to 15–10% in the morealtered samples. These authors propose two steps, with atransition of smectite to corrensite and a transition ofcorrensite to chlorite, and suggest that corrensite is aunique phase (Inoue and Utada 1991). Other smectite-chlorite sequences have been observed in active geother-mal fields and in metavolcanic rocks (Kristmannsdottir1975, 1979; Bettison and Schiffman 1988; Bettison-Var-ga et al. 1991). According to these studies, the observedsequences are the result of a progressive conversion ofsmectite to chlorite, mainly induced by a temperaturegradient. The HRTEM and spectroscopic studies providefurther constraints on the nature and crystal structure ofcorrensite. For example, infrared spectroscopy studies ofcorrensite from the Taro Valley (Bergaya et al. 1985)suggest that the non-swelling component is a chloritelayer that is physically distinct from the neighbouringlayer that contains exchangeable cations. However, morerecent studies that combine HRTEM and chemical analy-ses indicate that corrensite can be considered to be aunique phase (Shau and Peacor 1992; Beaufort and Meu-nier 1993; Beaufort et al., in preparation). According toShau and Peacor (1992), the distribution of Si and Al inthe tetrahedral sites is asymmetric, with a higher Al sub-stitution in tetrahedral sites adjacent to the brucite sheet,suggesting that the expandable layer cannot be inheritedfrom a precursor smectite. Beaufort and Meunier (1994)analysed corrensite formed in fractures of metamorphicrocks by microprobe and Mossbauer spectroscopy andproposed a different model for the corrensite structure inwhich the expandable layers are characterized by a highcharge with similar cation distributions in tetrahedralsheets in the structure.

The present study confirms that corrensite is a uniquephase: dislocations probably formed during corrensitegrowth, always disrupt a complete layer of 24 A˚ , i.e. acorrensite layer suggesting that the 24 A˚ layer is the base

unit repeated along c* to form the corrensite crystal. Fur-thermore, in swelling chlorite, the 10 A˚ stackings, thatcould represent dehydrated smectite layers, were alwaysobserved on HRTEM images as isolated layers between14 A layers suggesting that they are a part of the corren-site structure. The structural formulae calculated fromAEM analyses show that corrensite forming at MiddleValley is more Mg-rich than corrensite previously de-scribed in other environments. It is not a high chargecorrensite as suggested by Beaufort and Meunier (1993)for corrensite formed in fractures of metamorphic rocks.The unusual composition of corrensite from the MiddleValley hydrothermal sediments may reflect its formationin a sea-water dominated system. The mechanism of for-mation of these corrensite-rich deposits can be explainedby a smectite-chlorite transition or by direct precipita-tion from hydrothermal fluids.

The AEM analyses on the various clay particles docu-ment compositional differences between chlorite andcorrensite. The structural formulae of corrensite, chlor-ite and saponite presented in Table 2 show that corrensitecannot be considered to be a phase intermediate betweenchlorite and saponite because the amount of Fe in theoctahedral site of corrensite is too low. Corrensite ismore Mg-rich than both saponite and chlorite, and nocompositional trend between saponite, corrensite andchlorite minerals is observed (see Fig. 9). This suggeststhat corrensite is a unique phase and that in the Juan deFuca sediments, it is formed by direct precipitation fromheated seawater. In other sedimentary sequences where asmectite-chlorite conversion has been described, a pro-gressive evolution of the mineralogy and chemistry withdepth is generally described (Meunier et al. 1991). Theoccurrence of pure hydrothermal layers of well-differen-tiated chemical and mineralogical composition (Buatieret al. 1994), and the fact that some hydrothermal de-posits are intercalated with unmodified sediments sub-stantiates the hypothesis of direct precipitation for thehydrothermal sequence observed atHole858B. However,corrensite has never been observed as the sole phase ofthe pure hydrothermal sediments of Middle Valley. Al-though XRD data of sample B5H3 58–62 indicated onlycorrensite, the HRTEM images show that it is intercalat-ed with chlorite layers forming a corrensite-rich chlor-ite-corrensite mixed-layered phase (Fig. 6). This sug-gests the presence of metastable products (Shau andPeacor 1992).

Controls on fluid-sediment interactionat a sedimented hydrothermal system

The lateral and vertical distribution of mineralogically,chemically and isotopically distinct alteration zones atSite 858 reflect the overall temporally and spatially fluc-tuating hydrothermal system, with convective fluid cir-culation, high geothermal gradients and extensive fluid-sediment interactions prevailing in this area. Based onintegrated stable isotope, mineralogical, bulk chemical

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and pore-water data, Früh-Green et al. (1994a) suggestthat large-scale convective circulation drives smaller-scale fluid advection and causes near-surface seawaterrecharge at Site 858. The hydrothermal fluids are strong-ly reducing directly at the vent regions and precipitateMg-silicates, sulphides and minor sulphates. Away fromthe vents, the reducing fluids migrate outwards from thedischarge conduits and are mixed with more oxidizingfluids, producing the more extensive, thicker carbonateand sulphate mineral zones at the distal regions.

The superposition of hydrothermal units with unal-tered sediments in the upper part of theHole 858B sug-gests that permeability heterogeneities in the sedimenta-ry cover may facilitate lateral circulation of fluids. Thenewly formed phases are Mg- and Fe-rich, whereas thedetrital clay fractions are characterized by high Al-con-tents (Buatier et al. 1994). Comparison of pore-waterand clay chemistries suggests that part of the chemicalelements required for hydrothermal clay formation areprovided by the fluids. The pore-water profiles and theauthigenic mineral chemistries at Site 858 indicate acomplex history of dissolution and precipitation reac-tions, diffusion, and advection (see also Davis, Mottl,Fischer et al. 1992). The control of fluid-sediment inter-action on pore-water chemistry is particularly evident byCa and Mg concentrations in the pore-waters. At the dis-tal regions, increasing Ca-contents below approximately75 mbsf in Hole858A is associated with decreasing Mg-contents, and most likely reflect an upwards flux of Careleased during albitization of feldspar and concurrentMg-depletion during chloritization. The decrease in Mgcorresponds directly with a change from swelling chlor-ite to chlorite in the clay fractions and the presence ofchlorite as the dominant phyllosilicate phase (compareTable 1). The role of clay precipitation in controlling iso-topic and chemical compositions of the pore-waters isparticularly well observed at the immediate vent region.At Hole 858B, pore-water Mg-contents decrease drasti-cally from seawater concentrations (52.95 mmolykg) to12.72 mmolykg at 13.15 mbsf and coincide directly withincreases in chlorinity, silica,d18O anddD in the pore-waters (Früh-Green et al., in preparation) and to the firstoccurrence of Mg-rich corrensite.

The progressive decrease ind18O-values of the clayfractions at Site 858 is most likely the combined result ofincreasing temperatures, changing fluid compositionsandyor varying fluidyrock ratios with depth. An oxygenisotope equilibrium temperature of 3008 C+308, indicat-ed by the quartz-chlorite pair in sample B6H1 51–53, isonly slightly lower than that predicted by a temperaturegradient of approximately 11.48 Cym, as estimated fromoxygen isotope analyses on carbonate concretions andpore-waters inHole858D, but is higher than present-dayvent fluid temperatures. Furthermore, the calculatedtemperature for sample B6H1 51–53 implies equilibriumwith an altered seawater composition of approximately3.2‰. Although this estimate is slightly more positivethan the oxygen isotope ratios measured on the present-day pore-waters inHole 858B, it is consistent with en-

richments in 18O observed in the neighbouringHole858D, where new formation of clays may presently beoccurring (Früh-Green et al., in preparation). These re-sults thus suggest that both the temperatures of clay for-mation and the oxygen isotope composition of the pore-fluids have changed since the chlorite was deposited. Incontrast to the chlorite sample B6H1 51–53, the isotopedata of the corrensite-quartz pair in sample B5H3 58–62at 27.5 mbsf yield lower calculated temperatures thanthose predicted by a steady-state temperature gradient ofapproximately 11.48 Cym and more negatived18O-valuesof the coexisting fluids than those measured at MiddleValley. This discrepancy suggests that corrensite may nothave precipitated in equilibrium with quartz and the flu-ids. Disequilibrium conditions are consistent with theTEM and AEM data which indicate that this sample iscomposed of corrensiteychlorite mixed-layered (corren-site-rich) clays that are generally considered metastableproducts.

The distinct variations in isotopic and chemical com-positions observed over small vertical distances in thepore-water profiles at the immediate vent region reflectthe strongly fluctuating nature of the hydrothermal sys-tem and emphasise the strong control of fluid-sedimentreactions and precipitation of authigenic phases on thechemistry of the pore fluids. In order to alter the chemi-cal and isotopic compositions of the hydrothermal fluids,reaction rates must be high relative to fluid flux rates.

Conclusions

Detailed mineralogical and geochemical data on hy-drothermally altered sediments and pure hydrothermaldeposits from ODP Site 858 at Middle Valley provideconstraints on the mechanism of Mg-Fe phyllosilicateformation and indicate that mineral reactions directlycontrol the chemical and isotopic composition of hy-drothermal fluids.

Hydrothermal alteration at Site 858 is characterizedby a progressive change in phyllosilicate assemblageswith depth. In the distal regions atHole 858A, the clayfractions show a gradation from a detrital assemblage ofmica, smectite and chlorite at the top of the hole, pro-gressing to swelling chlorite and finally chlorite withdepth and increasing degree of alteration. In the immedi-ate vent area, atHole858B, detrital layers are intercalat-ed with pure hydrothermal precipitates in the upper15 m, whereas at depth hydrothermal phases are pre-dominant. Hydrothermal Mg-smectite is the major phyl-losilicate at the top ofHole 858B; with increasing depthcorrensite becomes dominant, followed by swellingchlorite and finally chlorite at the bottom of the hole. Inthree pure hydrothermal layers in the deepest part ofHole 858B, the clay minerals coexist with neoformedquartz. According to HRTEM and AEM analyses, cor-rensite, chlorite and swelling chlorite are well distin-guished by their chemistry and stacking sequence. Ourresults suggest that corrensite is a unique mineralogical

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phase and was precipitated directly from seawater-domi-nated hydrothermal fluids.

The mineralogical zonation inHole858B is accompa-nied by a systematic decrease ind18O, reflecting both thehigh thermal gradients that prevail at Site 858 and exten-sive sediment-fluid interaction. Comparison with pore-fluid chemical and isotopic compositions suggests thatthe precipitation of the Mg-phyllosilicates directly con-trols Mg-content of the fluids and, in the vent region,results in an enrichment in deuterium in the hydrother-mal fluids.

Structural formulae calculated from AEM analyseswere used to construct clay-H2O oxygen isotope frac-tionation curves, based on oxygen bond models. The cal-culated curves allow an evaluation of isotopic equilibri-um and approximations of precipitation temperature tobe made. The results suggest isotopic disequilibriumconditions for corrensite-quartz and swelling chlorite-quartz precipitation, but yield an equilibrium tempera-ture of 300+308 for chlorite-quartz at 32 mbsf. This esti-mate is consistent with temperatures measured duringLeg 139 and calculated from oxygen isotope ratios incarbonates in an adjacent hole, indicating steep thermalgradients of 10–118ym in the vent region. Disequilibriumprecipitation of corrensite-rich deposits and swellingchlorite is consistent with TEM data, which suggest thatmixed-layered corrensiteychlorite minerals are meta-stable products of hydrothermal alteration. Non-equi-librium conditions during the direct precipitation of cor-rensite-chlorite mixed layers from seawater may be aconsequence of high reaction rates in this active hy-drothermal system.

Acknowledgements We would like to thank M. Boni at the Uni-versity of Naples for providing us with samples, A. Fallick and T.Donnelley at the Scottish Universities Reactor and Research Cen-tre for hydrogen analyses and Philippe Recourt from the URA 719,University of Lille for his technical assistance on XRD analyses.Many thanks to S. Bernasconi for his scientific input and to J. C.Doukan and A. Baronnet for scientific discussions on AEM analy-ses and HRTEM images. This study was supported by INSU-ISTgrant No. 91GEO2y3.06 to M.B. and A.M.K. and ETH grant No.0–20–710–93 to G.F.G.

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