Thermal diagenesis of clay minerals within volcanogenic material from the Tonga convergent margin

ELSEVIER Marine Geology 157 (1999) 105–125

Thermal diagenesis of clay minerals within volcanogenic materialfrom the Tonga convergent margin

Frederic Vitali a,Ł, Gerard Blanc b, Philippe Larque a, Joelle Duplay a, Gilles Morvan a

a Centre de Geochimie de la Surface, UMR CNRS 7517, Institut de Geologie, Strasbourg, 1 rue Blessig, 67084 Strasbourg cedex, Franceb Departement de Geologie et d’Oceanographie, UMR CNRS 5805, Universite de Bordeaux I, Avenue des Facultes,

33405 Talence, France

Received 16 September 1997; accepted 15 July 1998

Abstract

Clay minerals are examined in detail in the sediment from the Tonga Trench margin at Site 841 (Leg 135 ODP).The changes in amount and nature of secondary clays with depth provide an alternative explanation for the intensivealteration of volcanogenic material at convergent margins. A characteristic distribution of clay minerals with depth showsfour distinct zones unexplainable by simple burial diagenesis processes. These are named the upper, reactive, lower andrhyolitic zones. The reactive zone is intercalated with numerous sills and is characterized by the dominant iron-richclays such as saponite, corrensite and chlorite associated with analcime. The occurrence of such iron-rich clays, mostlyassociated with a large amount of analcime, yields chemical and mineralogical evidence for thermal diagenesis. Therequired heat for the diagenetic process was transferred from recently intruded basaltic andesite sills. In the vicinity ofthese intrusions, the iron-rich clay minerals may have formed at temperatures up to 200ºC. A zoning with respect to clayand zeolite minerals indicates that the influence of the palaeoheat flow decreased with the distance from the intrusion. Theformation of interlayered I=S, illite, kaolinite and aluminous chlorite, which are recognized as major secondary mineralswithin the rhyolitic complex, was mainly controlled by both early diagenesis at moderately elevated temperatures, andsince the Eocene by burial diagenesis at low temperatures. The occurrence of a steam zone in an early stage of the intrusionis restricted to Miocene tuffs and has overprinted the early alteration of the volcanogenic material within the tuffs and haschanged the originally pristine composition of the pore fluids. 1999 Elsevier Science B.V. All rights reserved.

Keywords: Tonga; volcanogenic material; iron-rich clays; sills; palaeoheat flow

1. Introduction

In the oceanic system, clay minerals are com-monly used as markers of fluid–solid interactions.

Ł Corresponding author. Present address: University of WesternOntario, Department of Earth Sciences, London, Ont. N6A 5B7,Canada. Tel.: C1 (519) 661 3881; Fax: C1 (519) 661 3198;E-mail: [email protected]

Characterization of the mechanism of their forma-tion should permit determination of the history ofthe diagenetic and hydrothermal alteration processesthat have affected the host marine sediments. Duringthe cruise of Leg 135 ODP, the Joides ResolutionSV recovered intensively altered volcanogenic ma-terial from Site 841 in the southwest Pacific Ocean.These materials provide an opportunity for more de-tailed understanding of the alteration processes in

0025-3227/99/$ – see front matter 1999 Elsevier Science B.V. All rights reserved.PII: S 0 0 2 5 - 3 2 2 7 ( 9 8 ) 0 0 1 3 4 - 0

106 F. Vitali et al. / Marine Geology 157 (1999) 105–125

marine sediments. Toward this end, this study wasconducted using transmission and analytical electronmicroscopy (TEM and AEM) to provide efficientcharacterization of submicron-sized clay minerals.The results enable us the following: (1) to proposea clay mineral zonation with depth including di-agnostic paragenesis; (2) to postulate a detrital orauthigenic origin of clays from each zone; (3) tocompare this clay mineral zonation with that of ze-olites previously obtained by Vitali et al. (1995) andwith the distribution of the lithostratigraphic unitsdiscussed in the Scientific Reports of the ODP (Par-son et al., 1992); (4) to determine the chronologicalorder of crystallization and the temperature forma-tion of the clay minerals from Site 841; (5) to definethe causes of the diagenetic processes leading tothe observed clay mineral distribution; and (6) tocompare the clay mineralogy and chemistry fromthe Tonga forearc with those of other drilled sitesfrom the Pacific Ocean where recovered Oligoceneto Pleistocene volcanogenic sediments were recov-ered, including the Izu–Bonin Arc sedimentary basin(Taylor et al., 1990), the Central Mariana Basin(Lancelot et al., 1990), and the North Aoba Basin(Collot et al., 1992). These sediments have beendeposited in different tectonic settings, but, never-theless, show many similarities to those of Site 841.At all sites, they are characterized by the occurrenceof a CaCl2-rich interstitial brine (Blanc et al., 1995)and by an abundance of secondary minerals reachingup to 80 wt% of the sediment. At these sites, pro-cesses related to the formation of clays have beenstated mostly in terms of burial diagenesis associatedto high volcaniclastic sedimentation rates, very lowrates of exchange with deep-sea water, and mod-erately elevated temperature (Egeberg et al., 1990;Egeberg, 1992; France-Lanord et al., 1992; Martin,1994; Gerard and Person, 1994). Our study allows usto propose an alternative explanation for the forma-tion of clay minerals at a tectonically eroded margin.

2. Geological settings

Sediments were recovered from the upper trenchslope at Site 841 (water depth D 4810 m, 23º20.70S,175º17.90W, approximately 55 km west of the axisof the Tonga Trench and 60 km east of the crest

Fig. 1. Location map of Site 841 from the Tonga forearc (fromMcLeod, 1994).

of the Tonga Ridge) (Parson et al., 1992) (Fig. 1).The Tonga Trench is the northern part of the Tonga–Kermadec Trench system that lies between 15º and26ºS latitude, a linear distance of nearly 1200 km.

The sedimentary sequence at Site 841 is 605 mthick. It ranges in age from middle Pleistocene toLate Eocene and consists of clays, vitric siltstones,volcanic conglomerates and breccias, and volcanicsandstones. Basaltic andesite sills are intercalatedwithin the Miocene sediments from 324 m belowseafloor (mbsf) to 497 mbsf (Parson et al., 1992;Fig. 2). These basaltic andesite sills appear to be partof one intrusive episode (Bloomer et al., 1994). Be-low the Late Eocene sediments, the hole penetrateda former subaerial rhyolite volcanic complex, the topof which was intercepted 605 mbsf (Parson et al.,1992). An age of 44.2 Ma for this rhyolitic com-plex is well constrained radiometrically (McDougall,1994).

F. Vitali et al. / Marine Geology 157 (1999) 105–125 107

Fig. 2. Distribution with depth of the lithostratigraphic units from Site 841 (from Blanc et al., 1995).

3. Methods

Morphological features of minerals and qualita-tive elemental compositions were examined with aJSM 840 scanning electron microscope (SEM). The

clay fraction of 45 samples was separated by set-tling (Larque and Weber, 1978; Holtzapffel, 1985)and XRD analyses were done on four types of ori-ented aggregates: untreated, ethylene-glycol treated,hydrazine treated, and heated (4 h at 490ºC). The


diffractometer conditions used were: CoKα radiation,40 kV=20 mA and 1º2θ per minute scan speed. Onthe powder XRD diagrams, the position of the (060)diffraction peak from clay minerals allows dioctahe-dral and trioctahedral clays to be discriminated.

Mineralogical and chemical data of separated clayparticles have been obtained for twelve selected sam-ples on an analytical electron microscope (AEM andTEM) Philips CM12. During the microanalysis, theAEM was operating at 120 kV in raster mode (typ-

Fig. 3. Mineral paragenesis with depth. The weight percent of each mineral phase was determined by XRD and chemical analysis on anhalite-free basis (from Vitali et al., 1995). Corresponding lithostratigraphic units were added on the right side.

ically 300 ð 300 nm), with a beam diameter of 7.5nm, on a single particle.

4. Results

The mineralogy of the samples from the seafloorto 605 mbsf is dominated by clays and zeolites (Vi-tali et al., 1995; Vitali, 1996) (Fig. 3). These hydrousalumino-silicates constituted up to 80% of the sedi-


Fig. 4. Distribution of clay mineral species with depth, Site 841 ODP. Selected clay samples used for SEM, TEM and AEM study areindicated on the right side of the figure.

ments from 0 to 605 mbsf. The underlying highly al-tered rhyolitic volcanic complex is composed mainlyof quartz (30% on average) and plagioclase (20% onaverage). Zeolites were not found in this unit andvarious types of clays were present.

4.1. X-rays clay determination

The identified clay minerals in the settled frac-tions consist of smectite, interlayered illite=smectite(I=S), illite, chlorite, corrensite and kaolinite. Therelative abundance of the clays was approximatedusing the intensities of the (001) reflections on ori-ented lamina treated by ethylene-glycol (Fig. 4).This method has been used for numerous years inthe Centre de Geochimie de la Surface (Larque andWeber, 1978) and in many laboratories; as imperfectas it is, it reasonably illustrates variations of the claymineralogy with depth. The relative percentage ofsmectite sheets in interlayers of illite=smectite wasdetermined using Reynolds (1985).

4.1.1. IlliteIllite represents only a small part of clay mineral

content at Site 841. It occurs dominantly in the Plio–

Pleistocene unit I and in the Eocene rhyolitic tuffunit. Illite represents more than 30% of the clayfraction at the top and at the bottom of the rhyolitictuff unit (Fig. 4).

4.1.2. Interlayered illite=smectite and smectiteExcept for the intervals 317–325 mbsf, 345–352

mbsf, and the bottom part of the rhyolitic tuff unit,interlayered I=S and smectites represent 80 to 100%of the clay fraction in the sediments. On the basis ofthe XRD examinations, two types of interlayered I=Sare distinguished: interlayered I=S with high expand-abilities (S > 70%), and low expandable interlayeredI=S (S < 70%) within which the proportions of illiteand smectite are approximately balanced. In the 0–250 mbsf interval, low expandable I=S is abundant,whereas highly expandable I=S only occur between280 and 460 mbsf.

4.1.3. CorrensiteCorrensite was defined as a 1 : 1 regular inter-

layer of a trioctahedral chlorite with a trioctahe-dral smectite (Bailey, 1982). At Site 841, corrensitewas observed between 317 and 410 mbsf, associ-ated with chlorite (Figs. 4 and 5). In this interval,


Fig. 5. XRD diagram of corrensite, with characteristic peaks at31 A and associated harmonics on oriented powder treated withethylene-glycol. Sample 841B 17R3, 317 mbsf.

corrensite content reaches up to 60% of the clayfraction.

4.1.4. ChloriteChlorite occurs throughout the sedimentary col-

umn. The most important zones of occurrence arebetween 317 and 380 mbsf and within the rhyolitictuff unit. At the bottom of the hole, chlorite accountsfor up to 70 wt% of the clay fraction.

4.1.5. KaoliniteKaolinite represents a secondary mineral in the

rhyolitic tuff unit and accounts for 1=3 of the clay

fraction. It occurs only sporadically within the fiveupper volcano-sedimentary units.

On the basis of preliminary XRD examinationson the clay fraction, twelve representative samplesthrough the sedimentary column were selected fora detailed characterization of clay minerals usingSEM, TEM and AEM methods.

4.2. SEM morphological observations

SEM observations of clay minerals revealedmainly honeycomb structured smectite aggregates(Plate 1a) and large well-crystallized plates of chlo-rite. Honeycomb smectites were commonly observedat 52.5 mbsf and 159.6 mbsf (samples 6H4 and19X1, respectively) (Plate 1a). In the Miocene units,chlorite grows on the surface of other silicate miner-als (zeolites, feldspar) (Plate 1b) (Vitali et al., 1995)and on volcanic glasses. At 296.4 mbsf, where smec-tite minerals are abundant, few and grouped largechlorite minerals have been observed in rare cavities.Such occurrences become more abundant at greaterdepth (samples 21R1 352.4 mbsf and 29R1 430.6mbsf).

4.3. Distribution of clay minerals with depth

The distribution of clay mineral assemblagescould be characterized as a function of depth.

(1) From 0 to 260 mbsf: the clays consist ofdioctahedral I=S and smectites forming honeycombstructures. These clays present a ‘corn-flake’ mor-phology with curved edges (Plate 2a), similar toauthigenic smectites found in marine sediments (Bu-atier, 1989; Buatier et al., 1992; Clauer et al., 1990).

(2) From 260 to 470 mbsf: the clays are repre-sented by large sheets of chlorites (Plate 2b, Fig. 3),by corrensites (Plate 2c,d) and by smectites. Smec-tites are dioctahedral in the 260–310 mbsf intervaland trioctahedral between 310 and 470 mbsf.

(3) From 470 to 605 mbsf: interlayered I=S aredominant phases in the interval.

(4) From 605 m to the bottom: from 605 to 680mbsf, interlayered I=S also constitute the major partof the clay fraction. Our XRD studies could notpermit the identification of vermiculites, althoughSchops and Herzig (1994) reported a small amountof it between 640 and 670 mbsf.


Plate 1. SEM observations. (a) Honeycomb structured smectite aggregates (sample 841A 6H4 52.5 mbsf). (b) Chlorites (a) growing on(b) an altered K-feldspar, and (c) on clinoptilolites (sample 841B 15R1, 296.4 mbsf).

From 680 mbsf to the bottom of the hole, kaoli-nite, chlorite, and long-laths of illite (Plate 3) occur,in the same amount. Chlorite plates are smaller insize (0.5 µm in average) than those from the inter-val 260–470 mbsf. This mineral is also frequentlyfractured and corroded.

4.4. AEM chemical analysis

The elemental analysis and structural formulaeof clay samples are reported in Table 1. The struc-

tural formulae of all the trioctahedral clay mineralswere calculated by the methods of Brigatti and Poppi(1984a,b) and Inoue (1985), which can be sum-marized as follows: (1) the total iron is calculatedas ferrous; (2) the layer charge for corrensite isbalanced on the basis of O10(OH)5; and (3) the oc-tahedral cations in corrensite are evenly distributedbetween three octahedral layers of the interstrati-fied structure. Following the distribution with depthestablished above, AEM chemical analyses of clayminerals from Site 841 show mainly the following.

112F.Vitali

etal./M

arineG

eology157

(1999)105–125

Plate 2. TEM observations. (a) ‘Corn-flake’ smectite (sample 841A 6H4, 52.5 mbsf, ð36,000). (b) Hexagonal chlorite (sample 841B 18R2 325.9 mbsf, ð29,000). (c) Corrensite(Sample 841B 21R1, 352.4 mbsf, ð55,000). (d) Diffraction diagram of the corrensite particle from (c).

F.Vitaliet

al./Marine

Geology

157(1999)

105–125113

Plate 3. TEM observations. (a) Kaolinite (sample 841B 62R1, 748.7 mbsf, ð105,000). (b) Chlorite (sample 841 62R1 748.7 mbsf, ð55,000). (c) General view: chlorites andlaths of illites (sample 841B 62R1, 748.7 mbsf, ð16,500).

114F.Vitali

etal./M

arineG

eology157

(1999)105–125

Table 1AEM analysis and structural formulae from representative clay minerals of Site 841

F.Vitaliet

al./Marine

Geology

157(1999)

105–125115

Table 1 (continued)

(a) Dioctahedral smectites and interlayers I=S; (b) dioctahedral illites; (c) trioctahedral smectites; (d) kaolinites; (e) corrensites; (f) chlorites.


Fig. 6. Distribution of Al, Fe, Mg cations in octahedral layer of TOT clays from the Tonga forearc: (a) 0–260 mbsf, (b) 260–470 mbsf,(c) 470–605 mbsf, (d) 605 mbsf to the bottom of the hole.

(1) Between 0 and 260 mbsf (Fig. 6a): Al, Feand Mg occupy approximately equal proportions inthe octahedral layers of the smectites. Octahedralcharge and tetrahedral content (Fig. 7) suggest thatsmectites are montmorillonitic in composition.

(2) Between 260 and 470 mbsf (Fig. 6b): Si intrioctahedral smectites (21R1 352.4 mbsf and 29R1430.6 mbsf) is extensively replaced by Al. Fe con-stitutes more than 50% of the octahedral cations andAl around 10%. Dioctahedral smectite (15R1 296.4mbsf) contains around 40% Al and 35% Fe withinthe octahedral site. The layer charge diagram (Fig. 7)indicates that the dioctahedral smectites from this260 to 470 mbsf interval consist of nontronite, whilethe trioctahedral smectite, that exhibits an excess ofoctahedral charge is iron-saponite. Corrensite andchlorite from the 260 to 470 mbsf interval containsmainly Fe (40–70%) and Mg (30–50%) in the octa-hedral layers (Fig. 8).

(3) Between 470 and 605 mbsf (Fig. 6c): TOT

contain the same proportion of Al, Fe, and Mg inthe octahedral layer and they were identified in thecharge diagram as interlayer illite=montmorillonite(Fig. 7).

(4) From 605 mbsf to the bottom of the hole(Fig. 6d): TOT clays are characterized by the pre-dominance of Al cations (>60%) in the octahedrallayer. Al represents between 30 and 45% of the oc-tahedral layer of the chlorite minerals (62R1 748.7mbsf and 65RCC 776.2 mbsf) and Fe less than 30%.

5. Discussion

5.1. Iron-rich clay formation into the sedimentaryunits

In the Late Eocene to Pleistocene sedimentaryunits (i.e. from I to V), SEM observations show hon-eycomb structures of smectite and large, well-formed


Fig. 7. Distribution of Site 841 TOT clay in a layer charge diagram from Koster (1982): (A) 0–260 mbsf; (B) 260–470 mbsf; (C)470–605 mbsf.

flakes of clay growing directly on silicate mineralssuch as K-feldspar, zeolites (analcime, clinoptilolite)(Plate 1b) and on glass. These observations suggestthat a major part of clay minerals from units I to Voriginated from diagenetic processes. A diageneticorigin has been already considered for the zeoliteminerals (Vitali et al., 1995; Blanc et al., 1995). Inthe vicinity of the basaltic andesite sills intrudedinto the sediment of units II, III, and IV Miocene inage (Fig. 9), the observations such as clays growingon clinoptilolite and analcime, and analcime forma-tion after clinoptilolite allow us to propose an orderof mineral crystallization: clinoptilolite–analcime–clays. SEM analysis of the clay minerals located

next to the sills and identified as chlorite, corren-site and saponite, shows an elevated amount of iron(25% < FeO < 35%), and, in a lower proportion, ofmagnesium (9% < MgO < 18%) (Table 1, Fig. 6b,Fig. 9). The main source of the Fe (and also Mg) ofthe abundant chlorite, corrensite and saponite formedmainly between 315 mbsf and 470 mbsf could be theintensive dissolution of the glassy matrix of the vol-caniclastic material and also of augite phenocrysts,observed by SEM, next to the sills. Mg dissolvedin the interstitial water must be also an additionalsource for Mg incorporated subsequently into theseclay. The hot iron-rich hydrothermal fluid formedby pore-water–rock interactions in response to the


Fig. 8. Distribution of Al, Fe, Mg cations in chlorite and corren-site octahedral layers from Site 841.

sill intrusions is probably a reducing fluid, allowingthe crystallization of ferrous minerals such as Fechlorite, chlorite, and saponite. The upper limit ofthe influence of this reducing fluid is indicated bythe co-occurrence of ferric nontronite and ferrouschlorite observed at 296.4 mbsf. Variation in thehydrothermal fluid chemistry resulting from the pre-cipitation of the Fe-rich clays below this depth couldhave changed the redox conditions (Andrews, 1980;Grassi et al., 1992) and permitted around 296 mbsfthe growth of abundant ferric nontronites, whereaschlorite minerals were only observed in rare cavities.

5.2. Clay minerals from the rhyolitic tuff unit

Conversely to the sedimentary units, the TEM ob-servations of samples 62R1, 748.7 mbsf and 65RCC,776.2 mbsf from the rhyolitic tuff unit shows the oc-currences of both small-grained and fractured chlo-rite (Plate 3b) and long-lath illite (Plate 3c) some-times associated with hexagonal kaolinite (Plate 3a).The long-lath morphology of illite and the well-formed kaolinite seem to be an argument in favour ofan authigenic origin for these minerals, while doubtexists concerning the origin of the chlorite whosealtered features could result from hydrothermal al-teration that acted during the cooling of the subaerialEocene rhyolitic complex. This assumption is sup-ported by Schops and Herzig (1994) who proposedan early crystallization of chlorite at high tempera-tures followed by a later illitization=kaolinitizationat lower temperatures. Our observations of small-

grained, corroded and fractured chlorite, long-lathillite, and euhedral kaolinite is an argument with thisorder of crystallization.

5.3. Hydrous alumino-silicate distribution into theTonga forearc

The nature of hydrous alumino-silicates (i.e. ze-olites and clays) and their observed vertical dis-tributions are in contradiction to the present aver-age geothermal gradient of 2.84ºC=100 m. Thus wewill show that the heat of the cooling basaltic an-desite sills was a major control on the formationof authigenic clay and zeolite minerals in Site 841sediments. The vertical distribution of the hydroussilicates in Hole 841 from the Tonga forearc can bedivided in four diagenetic zones (Fig. 9).

From 0 to 260 mbsf, the upper zone (A) con-sists of clays with I=S and montmorillonite, associ-ated with phillipsite. This mineral association can beformed at low temperatures, which is consistent withthe present geothermal gradient.

From 260 to 470 mbsf, the reactive zone (B) ischaracterized by the occurrence of numerous basalticandesite sills. Nontronite, saponite, corrensite, andchlorite constitute the clay fraction. In the first 50m, nontronite constitutes the main clay mineral andassociated zeolites are clinoptilolite and analcime.Nontronite and clinoptilolite disappear below 310mbsf. From 310 to 470 mbsf, Fe-rich saponite,corrensite, and chlorite (Fig. 6b, Fig. 9) are pre-dominantly associated with analcime which formeddirectly from volcanic glasses. Fe-rich clays are alsosometimes associated with other zeolites such asmordenite, and wairakite. The abundant occurrenceof thaumasite is notable in two samples of this zone(Fig. 3). This secondary mineral is stable at tempera-tures lower than 60ºC (Giampaolo, 1986).

The following sequence occurs as a functionof the increasing proximity to the major basalticandesite sills: nontronite=Fe-saponite=Fe-corrensite–chlorite toward the intruded sills. Such mineral para-genesis relative to a thermal zone has been observedin the vicinity of active hydrothermal vents (Tomas-son and Kristmannsdottir, 1972; Stakes and O’Neil,1982; Fruh-Green et al., 1994; Buatier et al., 1994,1995; Buatier and Karpoff, 1995; Marumo et al.,1995), and around intrusions of sills into volcani-


Fig. 9. Mineralogical zonation based on the synthesis of XRD, SEM, AEM studies of hydrous silicates, Site 841 ODP: (A) upper zone(0–260 mbsf); (B) reactive zone (260–470 mbsf); (C) lower zone (470–605 mbsf); (D) rhyolitic zone (605 mbsf to the bottom of thehole). Sill intrusions and lithostratigraphic units are added on the right side of the figure.

clastic sediments (Einsele et al., 1980; Kelts, 1982;Desprairies and Jehanno, 1983; Inoue et al., 1984;Coulon et al., 1985; Inoue, 1987; Inoue and Utada,1991b; Kurnosov et al., 1994; Esposito and Whit-ney, 1995). Fe–Mg corrensite is estimated gener-ally to crystallize at temperatures between 100º and250ºC (Desprairies and Jehanno, 1983; De Kimpe,1988; Chamley, 1989; Mexias et al., 1990; Shauet al., 1990; Harvey and Browne, 1991; Inoue andUtada, 1991a,b; Hillier, 1993; Buatier et al., 1994,1995; Kurnosov et al., 1994), whereas nontronite isfrequently found in environments where the tem-perature is below 100ºC (Stakes and O’Neil, 1982;McMurtry et al., 1983; Singer et al., 1984; Cole,

1985; Stoffers et al., 1990; Schops et al., 1993;Buatier and Karpoff, 1995). An intermediate tem-perature between 80º and 200ºC generally favoursFe-saponite formation (Kohyama et al., 1973; Sudoand Shimoda, 1978; Badaut et al., 1985, 1990, 1992;Chang et al., 1986; Duplay et al., 1989; Decarreau etal., 1990; Tardy and Duplay, 1994). With increasingtemperature, iron saponite has been described as apotential precursor of corrensite, and subsequentlyof chlorite in volcanic-rich sediments (Inoue, 1987;Harvey and Browne, 1991; Meunier et al., 1991) andis often associated with zeolites (Sudo and Shimoda,1978; Besse et al., 1981; Schops et al., 1993; Ger-ard and Person, 1994). From the previous data, we


propose that the vertical distribution of clays withinthe reactive zone (B) at Site 841 resulted from a highthermal gradient created by the cooling of basalticandesite sills. The present thermal gradient impliesthat the clays and zeolites are not in thermodynamicequilibrium with the present temperature environ-ment but they persist because of their slow rates ofdissolution. The formation of thaumasite in this in-terval could be considered as a late stage formationduring thermal diagenetic processes in the coolingreactive zone.

From 470 to 605 mbsf, the lower zone (C), canbe defined by the predominance of I=S and by theoccurrence of various zeolites (phillipsite, clinoptilo-lite, analcime, erionite, chabazite and heulandite).This secondary paragenesis suggests that zone C hasbeen weakly influenced by the thermal pulse than theoverlying zone B.

For the three intervals described above, the zoneboundaries do not conform to the lithologic bound-aries. This discrepancy can be best understood interms of the recent and local heat flow due to theintrusion of sills into the Miocene volcaniclastic sed-iments. Consequently, the recent thermal pulse effectcould have obliterated a mineral assemblage thatcrystallized in the coarse-grained Miocene tuffs dur-ing early burial diagenesis. The present secondaryminerals zonation appears to have been establishedas a function of the distance from the intruded sillsand changes in the sediment grain-size. Compared tosiltstones, the higher permeability of coarser-grainedtuffs could facilitate leaching of elements and result-ing alteration of the sediment by hot fluids (Parson etal., 1992). Thus, the large proportion of siltstones atthe top of the lower zone (C) have probably limitedthe influence of the alteration processes below 470mbsf.

From 605 mbsf to the bottom of the hole, thelowermost zone (D) corresponds to the rhyolitic tuffunit. Zeolites are absent in this unit (Vitali et al.,1995). Aluminous interlayered I=S (Fig. 6d) is pre-ponderant in the upper part of this zone, whereasillite, kaolinite and aluminous chlorite are character-istic of the lower part. The early alteration leadingto the Al-chlorite formation is probably due to theheat transferred from the cooling Eocene rhyoliticcomplex. The other secondary minerals were mostlyformed by burial diagenesis during the 40 million

years of subsidence along the tectonically erodedTonga Margin.

5.4. Relationship with pore-water concentrations

In Site 841 sediments the depth variations in themajor chemical components dissolved in pore wa-ters are much more pronounced than those usuallyobserved in deep-sea sediments. The extensive alter-ation of volcanic Miocene deposits to zeolites andclay minerals mostly controls the elemental pore-water compositions with depth, and results in theformation of an interstitial CaCl2-rich brine below257 mbsf and a decrease in Na content from that inseawater concentration to a minimum of 347 mmol=lat 325.9 mbsf (Blanc et al., 1995). Volcanic materialsacted as the major source of Ca and Na. Crystalliza-tion of analcime with an average chemical formula ofNa14.5Si33.5Al14.5Ca0.1O96Ð16H2O (Vitali et al., 1995)seems to be a main process leading to the decreasein Na concentration in the interstitial fluid, whereasCa has accumulated in the pore water. As for Na,the concentration of Cl shows an unusual noncon-servative behaviour with an increase from seawaterconcentration to 612 mmol=l at 257 mbsf (Blanc etal., 1995). Processes supporting such enhanced Clconcentrations include evaporite dissolution, gas hy-drate formation, clay membrane filtration and uptakeof water by secondary hydrous minerals (Gieskeset al., 1990; Martin et al., 1995). Evaporites areabsent and concentrations of organic carbon andmethane are extremely low in the Miocene section(Fowler, 1994); thus, the first two processes are prob-ably not important. If the elevated Cl concentrationswere caused by membrane filtration, a correspond-ing Cl minimum should also have been observed(Gieskes et al., 1990; Blanc et al., 1991) but wasnot. Crystallization of zeolite, mainly analcime, andclay minerals seems to be the main process leadingto the increase in Cl concentration, as both processesremove H2O from the interstitial fluid (Blanc et al.,1995). However, because of its structural formula,analcime formation cannot be the reason for the de-crease of major dissolved chemical components suchas Mg. Instead, clay mineral formation seems to beresponsible for inducing a depletion of Mg in porewater from 50 mmol=l at the seafloor to values closeto 0 mmol=l in the interval 260–605 mbsf (Fig. 10).


Fig. 10. Mg concentrations with depth in pore waters, Site 841.

In the rhyolitic tuffs, the Mg concentration slightlyincreases from 0 mmol=l at 605 mbsf to 19 mmol=lat the bottom of the hole (Blanc et al., 1995). Asno hydrothermal fluid was detected at the bottom ofthe hole (Blanc et al., 1995; Vitali, 1996), this Mgincrease seems mainly to reveal that formation of theclay minerals from the rhyolitic zone occurred ear-lier than the clays found between 260 and 605 mbsf.The recent formation of clay minerals between 260mbsf and 605 mbsf could have strongly depleted thedissolved Mg concentration, in which case the slightincrease below 605 mbsf represents a progressivere-equilibration profile of Mg in pore waters.

5.5. Comparisons with other sites

The large amount of volcaniclastic materials, Ter-tiary to Quaternary in age, the intensive alterationof volcaniclastic sediment to clays and zeolites, andthe occurrence of CaCl2-interstitial brines, have beendescribed in five other sites from the Pacific Ocean:Sites 792 and 793 in the Izu–Bonin forearc (Tay-lor et al., 1990); Site 802 in the Central MarianaBasin (Lancelot et al., 1990); and Sites 832 and833 in the Aoba Basin (Collot et al., 1992). For allthese sites, there is little doubt that the anomalouscomposition of the pore waters is related to the al-

teration of volcanic materials. However, the genesisof numerous hydrous high-temperature minerals ap-pears to be obscure. Thermodynamic calculations atlow temperatures (e.g. 30º to 80ºC) showed that thepore waters are close to equilibrium with respect tomost of the phases recognized in the Izu–Bonin andAoba sites (Egeberg, 1992; Martin et al., 1995). Onthe other hand, the stability of chlorite and high-temperature zeolites such as wairakite cannot beexplained using the present temperature gradients,which do not exceed 67ºC km�1 at these sites. Ege-berg (1992) suggested that these high-temperatureminerals are not diagenetic in origin sensu stricto,but were eroded from volcanic highs. Conversely,Marsaglia and Tazaki (1992) suggested that ther-mal pulses brought about by intrusive events couldhave accelerated the diagenetic reactions in the olderOligocene sections of the Izu–Bonin forearc holes.Drilling at Site 833 in the North Aoba Basin indeedpenetrated a sill complex that was intruded into theLower and middle Pliocene sediments. Gerard andPerson (1994) submitted that the observed alterationof volcanic materials to saponite, nontronite and ze-olites resulted from this sill intrusion. However, sim-ple diffusivity would have dissipated heat from thesills within 104 years. Thus, the altered Late Pliocenesediments could not have been affected by this ther-mal pulse (Martin, 1994). Furthermore, sill intrusionwas not encountered at Site 832. At Site 802 fromthe Mariana Basin, siliceous cementation has prob-ably favoured a post-depositional ‘closed system’alteration of the Miocene tuffs to clays and zeolites(France-Lanord et al., 1992). However, the associ-ation of saponite with analcime into the Miocenetuffs at 50 mbsf suggests an unknown additional heatsource (Karpoff et al., 1992).

At all sites including 841, the present geothermalgradients are not high enough to explain the for-mation of the high-temperature authigenic mineralssuch as corrensite, chlorite and wairakite by simpleburial diagenesis. The presence of systematic verti-cal zonation in the authigenic mineral assemblagesat Site 841 is well related to the presence of sillintrusions. These intrusions supplied locally and re-cently an additional heat flow responsible for theformation of the authigenic minerals. Such mineralspersist up to the present because of their slow rate ofdissolution.


6. Conclusions

Mineralogical and chemical study of the clayfraction from the Tonga forearc sediments permitsthe following conclusions to be drawn.

Clays from Site 841 have an authigenic origin.Their formation was recent for clay from the LateEocene to Pleistocene sediments. The crystallizationof clays in the rhyolitic tuffs probably proceededduring two distinct time periods.

A characteristic distribution of clay minerals ex-ists with depth, showing four distinct zones. Thedistribution is unexplainable by simple burial diage-nesis processes. Among these zones named, upper,reactive, lower and rhyolitic, the reactive zone hostsnumerous sills, and is dominated by iron-rich clayssuch as saponite, corrensite, and chlorite associatedwith analcime.

The lower limit of the reactive zone is situatedat 470 mbsf. This accurately locates the zonationobtained by the study of zeolite minerals (Vitali etal., 1995; Blanc et al., 1995) and confirms that thedistribution of the secondary minerals does not co-incide with the lithostratigraphic limit, but is insteadderived from an intense alteration of the Miocenetuffs next to the basaltic andesite sills.

The SEM observations have revealed that theorder of crystallization in the reactive zone isclinoptilolite–analcime–clay minerals and that theformation temperatures of iron-clays may reach upto approximately 200ºC in the vicinity of the sills.

In the middle part of the Miocene volcaniclas-tic tuffs, the occurrence of iron-rich chlorite, cor-rensite, and saponite, associated with analcime andwairakite, is a mineralogical and chemical indica-tor for a high-temperature diagenetic process as-sociated with cooling basaltic andesite sills. Thisprocess is interconnected with the formation ofCaCl2-interstitial brines. Both intensive dissolutionof the glassy matrix of the volcaniclastic materialand of augite phenocrystals could act as the majorsource of iron for the authigenic clays. The magne-sium for the formation of these clays was derivedfrom the volcanic materials and from pore waters.The thermal flux was less efficient far from the in-trusions; consequently the clay minerals formed inthe interval 470–605 mbsf are interlayered I=S. Withrespect to the other strata at the site, the rhyolitic

zone shows a distinct secondary mineral assemblagewhose formation is mainly attributed to burial dia-genesis, ongoing in this subsiding igneous complexsince the Eocene. Additional heat brought by the in-trusion of the recent basaltic andesite sills apparentlydid not have an effect on the alteration of the basalrhyolitic zone.

Site 841 shows numerous similarities to other Pa-cific drilled sites in the Izu–Bonin arc sedimentarybasin, in the Central Mariana Basin, and in the NorthAoba Basin in terms of the amounts of volcaniclasticmaterial, the content of secondary clays and zeo-lites, and the occurrence of CaCl2-interstitial brines.Despite these similarities, the systematic zonationof the authigenic minerals and the relationships be-tween this diagenetic zonation and the occurrenceof sill intrusions can be well defined at Site 841only. Thus, the data obtained in this study are uniquein establishing intimate relationships among the ap-pearance of authigenic minerals, the dissolution ofthe volcanic materials, and the changes in pore-waterchemistry.

Acknowledgements

We acknowledge thoughtful reviews of themanuscript by Atsuyuki Inoue and an anonymousreviewer. Efficient editorial handling by M.A. Arthurand H. Chamley is also gratefully acknowledged.Thomas Pedersen is thanked for his English correc-tions. We would like to thank the ODP for technicalassistance during Site 841 sediment sampling on-board the Joides Resolution, and on shore. Thisresearch was supported by the French Committee ofMarine Geosciences from INSU-CNRS.

References

Andrews, A.J., 1980. Saponite and celadonite in layers 2 basalts,DSDP leg 37. Contrib. Mineral. Petrol. 73, 323–340.

Badaut, D., Besson, G., Decarreau, A., Rautureau, R., 1985. Oc-currence of a ferrous trioctahedral smectite in recent sedimentsof Atlantis II deep, Red Sea. Clay Miner. 20, 389–404.

Badaut, D., Blanc, G., Decarreau, A., 1990. Variation desmineraux ferriferes, en fonction du temps et de l’espace, dansles depots metalliferes de la fosse Atlantis II en Mer Rouge.C. R. Acad. Sci. Paris 310 (2), 1069–1075.


Badaut, D., Decarreau, A., Besson, G., 1992. Ferripophyllite andrelated Fe3C-rich 2 : 1 clays in recent deposits of Atlantis IIdeep, Red Sea. Clay Miner. 27, 227–244.

Bailey, S.W., 1982. Nomenclature for regular interstratification.Clay Miner. 17, 243–248.

Besse, D., Desprairies, A., Jehanno, C., Kola, V., 1981. Lesparageneses de smectites et de zeolites dans une serie pyro-clastique d’age eocene moyen de l’Ocean Indien (D.S.D.P.,Leg 26, Site 253). Bull. Mineral. 104, 56–63.

Blanc, G., Boulegue, J., Gieskes, J.R., 1991. Chemical evidencefor advection of interstitial fluid in the sedimentary series ofthe Barbados accretionary complex (Leg 110). Oceanol. Acta14, 33–49.

Blanc, G., Vitali, F., Stille, P., 1995. Unusual diagenetic alter-ation of volcanoclastic sediments in the Tonga fore-arc: evi-dence from chemical and Sr isotopic composition of interstitialwaters. Geochim. Cosmochim. Acta 59, 4633–4644.

Bloomer, S.H., Ewart, A., Hergt, J.M., Bryan, W.B., 1994. Geo-chemistry and origin of igneous rocks from the outer forearc(Site 841). Proc. ODP, Sci. Results 135, 625–646.

Brigatti, M.F., Poppi, L., 1984a. Crystal chemistry of corrensite:a review. Clays Clay Miner. 32, 391–399.

Brigatti, M.F., Poppi, L., 1984b. ‘Corrensite-like minerals’ in theTaro and Ceno Valleys, Italy. Clay Miner. 19, 59–66.

Buatier, M., 1989. Genese et evolution des argiles vertes hy-drothermales oceaniques: les monts du rift des Galapagos(Pacifique equatorial). These de doctorat de l’Universite LouisPasteur, 175 pp.

Buatier, M., Karpoff, A.-M., 1995. Authigenese et evolutiond’argiles hydrothermales oceaniques: exemples des monts dela ride des Galapagos et des sediments de la ride de Juan deFuca. Bull. Soc. Geol. Fr. 166 (2), 123–136.

Buatier, M., Peacor, D.R., O’Neil, J.R., 1992. Smectite–illitetransition in Barbados accretionary wedge sediments: TEMand AEM evidence for dissolution=crystallisation at low tem-perature. Clays Clay Miner. 40 (1), 65–80.

Buatier, M., Karpoff, A.-M., Boni, M., Fruh-Green, G.L.,McKenzie, J.A., 1994. Mineralogical and petrographicalrecords of sediment–fluid interactions in the sedimentary se-quence of Middle Valley, Juan de Fuca Ridge, Leg 139. Proc.ODP, Sci. Results 139, 133–154.

Buatier, M., Fruh-Green, G.L., Karpoff, A.-M., 1995. Mecha-nisms of Mg-phyllosilicate formation in a hydrothermal sys-tem at a sedimented ridge (Middle Valley, Juan de Fuca).Contrib. Mineral. Petrol. 122, 134–151.

Chamley, H., 1989. Clay Sedimentology. Springer, Berlin, 623pp.

Chang, H.K., Mackenzie, F.T., Schoonmaker, J., 1986. Compar-isons between the diagenesis of dioctahedral and trioctahedralsmectites, Brazilian offshore Basins. Clays Clay Miner. 34 (4),407–423.

Clauer, N., O’Neil, J.R., Bonnot-Courtois, C., Holtzapffel, T.,1990. Morphological, chemical, and isotopic evidence for andearly diagenetic evolution of detrital smectite in marine sedi-ments. Clays Clay Miner. 38, 33–46.

Cole, T.G., 1985. Composition, oxygen isotope geochemistry,and origin of smectite in the metalliferous sediments of the

Bauer Deep, southeast Pacific. Geochim. Cosmochim. Acta49, 221–235.

Collot, J.Y., Greene, H.G., Stokking, L.B., the Shipboard Scien-tific Party (Eds.), 1992. Vanuatu. Proc. ODP, Init. Rep. 134.

Coulon, H., Debrabant, P., Lefevre, C., 1985. Donnees petro-graphiques, mineralogiques et geochimiques sur la transitionbasaltes–sediments dans l’Atlantique Nord. Ann. Soc. Geol.Nord CIV, pp. 219–233.

Decarreau, A., Badaut, D., Blanc, G., 1990. Origin and tempera-ture formation of Fe-rich clays from Atlantis II deep deposits(Red Sea). An oxygen isotopic geochemistry approach. Geo-chemistry of the Earth’s surface and of mineral formation. 2ndInt. Symp., July 2–8 1990, Aix en Provence, France. Chem.Geol. 84, 363–364.

De Kimpe, C.R., 1988. Geographic distribution of corrensite andassociated minerals in southeastern Ontario. Can. Mineral. 26,957–964.

Desprairies, A., Jehanno, C., 1983. Parageneses minerales lieesa des interactions basalte–sediment–eau de mer (Sites 465 et456 des Legs 65 et 60 du D.S.D.P.) Sci. Geol. Bull. 36 (2–3),93–110.

Duplay, J., Paquet, H., Kossovskaya, A., Tardy, Y., 1989. Estima-tion de la temperature de formation des parageneses saponite–celadonite et glauconite–nontronite dans les alterations sousmarines de basalte, par la methode des correlations entre ele-ments au sein de populations monominerales. C. R. Acad. Sci.Paris 309 (2), 53–58.

Egeberg, P.K., 1992. Thermodynamic aspects of Leg 126 inter-stitial waters. Proc. ODP, Sci. Results 126, 519–529.

Egeberg, P.K., the Leg 126 Shipboard Scientific Party, 1990.Unusual composition of pore waters found in the Izu–Boninfore-arc sedimentary basin. Nature 344, 215–218.

Einsele, G., Gieskes, J.M., Curray, J., Moore, D.M., Aguayo, E.,Aubry, M.-P., Fornari, D., Guerrero, J., Kastner, M., Kels, K.,Lyle, M., Matoba, Y., Molina-Cruz, A., Niemitz, J., Rueda, J.,Saunders, A., Schrader, H., Simoneit, B., Vacquier, V., 1980.Intrusion of basaltic sills into highly porous sediments andresulting hydrothermal activity. Nature 283, 441–445.

Esposito, K.J., Whitney, G., 1995. Thermal effects of thin ig-neous intrusions on diagenetic reaction in a Tertiary Basin ofsouthwestern Washington. Geol. Surv. Bull. 2085-C, 40 pp.

Fowler, M.G., 1994. Organic geochemistry of sediments fromLeg 135 (Lau Basin) and of an oil seep (Tonga). Proc. ODP,Sci. Results 135, 667–676.

France-Lanord, C., Michard, A., Karpoff, A.-M., 1992. Major el-ement and Sr isotope studies in interstitial waters of sedimentsfrom Leg 129: the role of diagenetic reactions. Proc. ODP, Sci.Results 129, 267–281.

Fruh-Green, G.L., McKenzie, J.A., Boni, M., Karpoff, A.-M.,Buatier, M., 1994. Stable isotope and geochemical record ofconvective hydrothermal circulation in the sedimentary se-quence of Middle Valley, Juan de Fuca Ridge, Leg 139. Proc.ODP, Sci. Results 139, 291–305.

Gerard, M., Person, A., 1994. Low hydrothermal impact involcaniclastic sediments of the North Aoba Basin: Sites 832and 833. Proc. ODP, Sci. Results 134, 131–176.


Giampaolo, C., 1986. Dehydration kinetics of thaumasite atambient pressure. Neues Jahrb. Mineral. Monatsh. 3, 126–134.

Gieskes, J.M., Blanc, G., Vrolijk, P.J., Elderfield, H., Barnes, R.,1990. Interstitial water chemistry — major constituents. Proc.ODP, Sci. Results 110, 155–177.

Grassi, V.J., Ulmer, G.C., Grandstaff, D.E., 1992. Water–rockhydrothermal experiments: influence of rock type on solutionpH and oxygen fugacity. In: Kharaka, Maest (Eds.), Water–Rock Interaction. Balkema, Rotterdam, pp. 1415–1418.

Harvey, C.C., Browne, P.R.L., 1991. Mixed-layer clay geother-mometry in the Wairakei geothermal Field, New Zealand.Clays Clay Miner. 39 (6), 614–621.

Hillier, S., 1993. Origin, diagenesis, and mineralogy of chloriteminerals in Devonian lacustrine mudrocks, Orcadian Basin,Scotland. Clays Clay Miner. 41 (2), 240–259.

Holtzapffel, T., 1985. Les mineraux argileux: preparation, anal-yse diffractometrique et determination. Soc. Geol. Nord Publ.12, 136 pp.

Inoue, A., 1985. Chemistry of corrensite: a trend in composi-tion of trioctahedral chlorite=smectite during diagenesis. ChibaUniversity, Journal of the College of Arts and Sciences B-18,69–82.

Inoue, A., 1987. Conversion of smectite to chlorite by hydrother-mal and diagenetic alterations, Hokuroku Kuroko minerali-sation area, Northeast Japan. In: Schultz, L.G., van Olphen,H., Mumpton, F.A. (Eds.), Proceeding of the InternationalClay Conference, Denver, 1985. The Clay Minerals Society,Bloomington, IN, pp. 158–164.

Inoue, A., Utada, M., 1991a. Hydrothermal alteration in theKamikita Kuroko Mineralisation Area, northern Honshu,Japan. Min. Geol. 41 (4), 203–218.

Inoue, A., Utada, M., 1991b. Smectite-to-chlorite transforma-tion in thermally metamorphosed volcanoclastic rocks in theKamikita area, northern Honshu, Japan. Am. Mineral. 76,628–640.

Inoue, A., Utada, M., Nagata, H., Watanabe, T., 1984. Con-version of trioctahedral smectite to interstratified chlorite=smectite in Pliocene acidic pyroclastic sediments of the Ohyudistrict, Akita Prefecture, Japan. Clay Sci. Soc. Jpn. 6 (3),103–116.

Karpoff, A.M., France-Lanord, C., Lothe, F., Karcher, P., 1992.Miocene tuffs from Mariana basin, ODP. Leg 129, Site 802:a first deep-sea occurrence of Thaumasite. Proc. ODP, Sci.Results 129, 119–135.

Kelts, K., 1982. Petrology of hydrothermally metamorphosedsediments at Deep Sea Drilling Site 477, Southern GuaymasBasin Rift, Gulf of California. Init. Rep. DSDP 64, 1123–1136.

Kohyama, N., Shimoda, S., Sudo, T., 1973. Iron-rich saponite(ferrous and ferric forms). Clays Clay Miner. 21, 229–237.

Koster, H.M., 1982. The crystal structure of 2 : 1 layer sili-cates. In: Development in Sedimentology 35, InternationalClay Conference 1981, pp. 41–71.

Kurnosov, V., Murdmaa, I., Rosanova, T., Kashintzev, G.,Eroshchev-Shak, V., Krasnov, S., 1994. Mineralogy of hy-drothermally altered sediments and igneous rocks at Sites

856–858, Middle Valley, Juan de Fuca Ridge, Leg 139. Proc.ODP, Sci. Results 139, 397–413.

Lancelot, Y., Larson, R.L., Fisher, A., and the Shipboard Sci-entific Party (Eds.), 1990. Old Pacific Crust. Proc. ODP, Init.Rep. 129.

Larque, Ph., Weber, F., 1978. Techniques de preparation desmineraux argileux en vue de l’analyse par diffraction desrayons X. Note technique de l’Institut de Geologie, Strasbourg(4th ed.), 33 pp.

Marsaglia, K.M., Tazaki, K., 1992. Diagenetic trends in Leg 126sandstones. Proc. ODP, Sci. Results 126, 125–138.

Martin, J.B., 1994. Diagenesis and hydrology at the New He-brides Forearc and intra-arc Aboa Basin. Proc. ODP, Sci.Results 134, 92–104.

Martin, J.B., Kastner, M., Egeberg, P.K., 1995. Origin of salinefluids at convergent margins. In: Active Margins and MarginalBasins of the Western Pacific. Geophys. Monogr. 88, 219–239.

Marumo, K., Longstaffe, F.J., Matsubaya, O., 1995. Stable iso-tope geochemistry of clay minerals from fossil and active hy-drothermal systems, Southwestern Hokkaido, Japan. Geochim.Cosmochim. Acta 59 (12), 2545–2559.

McDougall, I., 1994. Data Report: dating of rhyolitic glass in theTonga Forearc (Hole 841B). Proc. ODP, Sci. Results 135, 923.

McLeod, C.J., 1994. Structure of the outer Tonga Forearc at Site841. Proc. ODP, Sci. Results 135, 313–329.

McMurtry, G.M., Wang, C.H., Yeh, H.W., 1983. Chemical andisotopic investigation into the origin of clay minerals from theGalapagos hydrothermal mount field. Geochim. Cosmochim.Acta 47, 475–489.

Meunier, A., Inoue, A., Beaufort, D., 1991. Chemiographicanalysis of trioctahedral smectite-to-chlorite conversion se-ries from the Ohyu Caldera, Japan. Clays Clay Miner. 39 (4),409–415.

Mexias, A., Formoso, M., Meunier, A., Beaufort, D., 1990.Composition and crystallisation of corrensite in volcanic andpyroclastic rocks of Hilario Formation (RS) Brazil. SciencesGeologiques. Proc. 9th Int. Clay Conf., Strasbourg, 1989,Mem. 88 (4), 135–145.

Parson, L., Hawkins, J., Allan, J., Shipboard Scientific Party(Eds.), 1992. Proc. ODP, Init. Rep. 135.

Reynolds, R.C., 1985. NEWMOD: a computer program for thecalculation of one-dimensional diffraction pattern of mixed-layered clays. R.C. Reynolds, 8 Broke Rd, Hanover, NH03755.

Schops, D., Herzig, P.M., 1994. Hydrothermally altered felsicrocks of the Tonga Forearc. Proc. ODP, Sci. Results 135, 653–663.

Schops, D., Herzig, P.M., Halbach, P., Friedrich, G., Blum, N.,1993. Mineralogy, chemistry and oxygen isotope thermome-try of nontronitic smectites from Central Pacific seamounts.Chem. Geol. 106, 331–343.

Shau, Y.H., Peacor, D.R., Essene, E.J., 1990. Corrensite andmixed-layer chlorite=corrensite in metabasalt from northernTaiwan: TEM=AEM, EMPA, XRD, and optical studies. Con-trib. Mineral. Petrol. 105, 123–142.

Singer, A., Stoffers, P., Heller-Kallai, L., Szafranek, D., 1984.


Nontronite in a deep-sea core from the south Pacific. ClayMiner. 32, 375–383.

Stakes, D.S., O’Neil, J.R., 1982. Mineralogy and stable isotopegeochemistry of hydrothermally altered oceanic rocks. EarthPlanet. Sci. Lett. 57, 285–304.

Stoffers, P., Singer, A., McMurtry, G., Arquit, A., Yeh, H.-W.,1990. Geochemistry of a hydrothermal nontronite deposit fromthe Lau Basin, Southwest Pacific. Geol. Jahrb. D 92, 615–628.

Sudo, T., Shimoda, S., 1978. Clays and clay minerals of Japan.Developments in Sedimentology 26, 326 pp.

Tardy, Y., Duplay, J., 1994. Stability fields of smectites and il-lites including glauconites as a function of temperature andchemical composition. Diagenesis IV. Developments in Sedi-

mentology 5, 95–132.Taylor, B., Fujoka, K., Janecek, T.R., the Shipboard Scientific

Party (Eds.), 1990. Proc. ODP, Init. Rep. 129.Tomasson, J., Kristmannsdottir, H., 1972. High temperature alter-

ation minerals and thermal brines, Reykjane, Iceland. Contrib.Mineral. Petrol. 36, 123–134.

Vitali, F., 1996. Geochimie des series volcano-sedimentaires etdes eaux interstitielles du sud-ouest Pacifique (Exemple dela marge active Tonga). These de doctorat de l’Universite deStrasbourg I, 148 pp.

Vitali, F., Blanc, G., Larque, Ph., 1995. Zeolite distributionin volcaniclastic deep-sea sediments from the Tonga Trenchmargin (SW Pacific). Clays Clay Miner. 43 (1), 92–104.

Thermal diagenesis of clay minerals within volcanogenic material from the Tonga convergent margin

Documents

Transcript of Thermal diagenesis of clay minerals within volcanogenic material from the Tonga convergent margin