textures of the hellyer volcanic-hosted massive sulfide...

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IntroductionTHE HELLYER orebody, now largely mined out, lies within theMount Read massive sulfide province of Cambrian age inwestern Tasmania (Fig. 1). It consists of a single, mildly de-formed, elongate lens of massive sulfide partially overlain bybarite-sulfide and silica-pyrite rocks (McArthur, 1989;McArthur and Dronseika, 1990). Unfolding of the orebody byDowns (1993, see Solomon and Khin Zaw, 1997) indicatedthat the orebody occupied a fault-controlled basin on theCambrian sea floor. It was originally thought to be a Kuroko-like deposit (Large, 1992; McArthur, 1996) that developed byzone refining of a capped mound on the sea floor. However,using salinity and temperature data gathered from fluid in-clusions in the veins of the underlying feeder zone (Gemmelland Large, 1992; Khin Zaw et al., 1996), Solomon and KhinZaw (1997) showed that the early ore fluids were likely tohave become denser than seawater shortly after emergingfrom the underlying feeder veins, thereby ponding in thebasin and establishing a brine pool within which sulfides weredeposited.

In this paper we describe some of the textures seen in themassive ore and how they, and other properties of the ore,throw light on the processes that may have operated in thebasin after the ore fluids emerged from the vent.

The Geology of the Hellyer AreaThe Hellyer massive sulfide deposit is a stratiform lens

lying within the Que-Hellyer Volcanics, a unit within theMount Read Volcanics that consists of basaltic lavas, breccias,and volcaniclastics with minor dacitic units and rare sedimen-tary lenses, aged about 500 to 510 Ma (Corbett, 1992; McPhieand Allen, 1992; Black et al., 1997). The Que-Hellyer Vol-canics has a maximum thickness of 1 km near Hellyer and isoverlain by medium to dark gray, marine, carbonaceous mud-stones and siltstones of the Que River Shale, which containslate Middle Cambrian trilobites and dendroidea (Jago, 1979;Corbett, 1992). Waters and Wallace (1992) described theQue-Hellyer Volcanics in the Que River and Hellyer mineareas as follows, from the base upward:

1. Basalt, consisting dominantly of massive and pillowedlavas and hyaloclastite rocks.

Textures of the Hellyer Volcanic-Hosted Massive Sulfide Deposit, Tasmania—the Aging of a Sulfide Sediment on the Sea Floor

MICHAEL SOLOMON†

Centre for Ore Deposit Research, University of Tasmania, G. P. O. Box 252-79, Hobart, Tasmania, Australia 7001

AND ORLANDO C. GASPAR

Rua Marechal Saldanha, 935-2° Dt, 4150-659 Porto, Portugal

AbstractThe relatively undeformed, polymetallic, massive sulfide ore of the Hellyer deposit retains many textures de-

veloped during the early crystallization history. Common textures in the least modified ore away from the mainfeeder zone include pyrite framboids, concentric- and radial-textured spheres, dendritic and fibrous forms,breccias, mineral banding, veins, and reniform and crustification textures. Many are typical of early crystalliza-tion within a medium at high degrees of supersaturation, some result from open-space filling, and others indi-cate disturbances of the sulfide mass possibly due to tectonic activity and volume adjustment; the bandingcould be due to current winnowing. Most of these textures are found in modern midocean ridge and Kurokomassive sulfide deposits, but at Hellyer the framboids are much more common. Hexagonal outlines in oresphalerite indicate derivation of much of the sphalerite by inversion of wurtzite. The high content of chal-copyrite in the wurtzite-derived sphalerite appears in most cases to have developed by exsolution rather thanreplacement or coprecipitation (although there is local evidence of both processes having operated), possiblyresulting from early asymmetrical crystal growth. Comparison of Hellyer ore and typical Kuroko ores revealsother differences than framboid abundance (e.g., metal contents and ratios, mineralogy, and sulfur isotopecompositions), the most significant of which are the relatively high salinities of the Hellyer ore fluids. These,together with evidence of a coeval basin, imply sulfide deposition in a brine pool, as previously proposed. Thedevelopment of a quenched sulfide mud at the bottom of a basin filled with spent ore fluid, with little or no in-volvement of seawater, explains the different mineral content (e.g., lack of Fe oxides, scarcity of barite in themassive sulfide, and the pyrite-arsenopyrite assemblage), the high metal content, lack of chimney structures,and differences in sulfur isotope compositions. The sulfide mud, dominated by Fe-S phases, would be an idealgrowth medium for framboidal pyrite.

The weight of experimental evidence concerning the development of framboidal pyrite points to involve-ment of metastable precursor Fe-S phases, such as mackinawite and greigite, in a weakly oxidized environment.Oxygen may have been added to the reduced, acid brine pool as a result of inward diffusion of oxygen fromoverlying seawater, which is shown to have been oxic by the presence of sessile benthic dendroids in the shalesoverlying the ore deposit. However, the level of oxidation was insufficient to form marcasite.

Economic GeologyVol. 96, 2001, pp. 1513–1534

†Corresponding author: e-mail, [email protected]

2. Dominantly andesitic and dacitic lavas, breccias, and epi-clastic rocks.

3. The mixed sequence (which hosts the massive sulfide de-posit, Fig. 1), with a maximum thickness of 150 m, compris-ing a highly variable assemblage of volcanic breccias andsandstones, dacitic volcanic rocks, and minor shales. Thepolymict volcanic breccias of the mixed sequence, thought tobe derived from debris flows, locally contain fragments ofmassive sulfide and barite, which, however, are not necessar-ily derived from Hellyer ore. On a regional scale, the mixedsequence thickens toward the Hellyer mine, indicating that abroad basin centered on the deposit.

4. The Hellyer basalt (Fig. 1), up to 220 m thick, consistingof massive and pillowed basaltic lavas, hyaloclastite breccias,peperites, and minor sedimentary lenses.

The Hellyer basalts belong to a different chemical suitefrom those of units (1) and (2), but all are subduction-related,

calc-alkaline rocks of high K and/or shoshonitic character(Whitford et al., 1989; Crawford et al., 1992). They probablydeveloped in a postcollisional, marine volcanic arc (Berry andCrawford, 1988) or a back-arc basin (Solomon and Groves,2000).

The Hellyer orebody lies on the crest of a major north-northwest–trending, faulted anticline that developed duringthe Tasmania-wide, mid-Devonian Tabberrabberan orogeny(Williams et al., 1989). An associated axial surface cleavagewas developed in the hydrothermally altered footwall rocks,and metamorphism in this part of the Mount Read Volcanicsbelt was of prehnite-pumpellyite facies (Offler and Whitfordet al., 1992). Although the volcanic succession would almostcertainly have undergone alteration as a result of convectionof Cambrian seawater, it has not yet been possible to separatethis effect from that related to Devonian regional metamor-phism. Nor has any of the folding and faulting that occurredelsewhere in western Tasmania toward the close of the Cam-brian (the Tyennann orogeny; Corbett and Turner, 1989)been recognized in the Hellyer area.

The Geology of the Hellyer OrebodyThe mined orebody consisted of a massive sulfide lens

overlain in part by a barite-sulfide cap and localized lenses ofsilica-sulfide rock. Fragments of barite in the overlying vol-caniclastic rocks of the mixed sequence (or hanging-wall vol-caniclastics in mine terminology) indicate exposure of the sul-fide-barite orebody on the sea floor. The orebody, includingthe barite and silica caps, originally contained 16.2 millionmetric tons (Mt) and consisted of, by weight, 53.1 percentpyrite, 21.8 percent sphalerite, 8.7 percent galena, 1.4 per-cent arsenopyrite, 0.9 percent chalcopyrite, 4.7 percentbarite, and 4 percent quartz (McArthur, 1996). Previous pa-pers and theses have displayed present-day cross sections andplans of the Hellyer orebody (McArthur and Dronseika,1990; Sharpe, 1991; Gemmell and Large, 1992; McArthur,1996). Figure 1 illustrates the stratiform, lens-like nature ofthe deposit, the local stratigraphic succession, and a majorfault, the Jack fault, which cuts the orebody almost into twoseparate parts. The orientation of the fault is shown in plan inFigure 2 (from Gemmell and Large, 1992), and the two partsof the orebody can be seen in longitudinal section in Figure3. The effects of Devonian folding and faulting were un-wound by Downs (1993), on the basis of an originally hori-zontal surface at the top of the mixed sequence, to reveal anorebody lying in a fault-related basin, presumably of Cam-brian age (fig. 2 of Solomon and Khin Zaw, 1997). Downs(1993) identified a basin that initially resulted from east-westextension, with related north-south faults, and subsequentlyone under the influence of northwest-southeast extensionthat involved reactivation of early faults and development ofnew northeast-southwest faults. The resulting fault complexmay be seen in figure 1 of Solomon and Khin Zaw (1997).

The footwall feeder zone described by Gemmell and Large(1992) extends along this fault complex for almost the entirelength of the orebody and includes a silicified core that bulgesupward into the massive sulfide between about 10650N and10850N (Downs, 1993). This bulge or spine may have beenpresent for the entire period of, or have grown during, themineralization phase. By the close of mineralization the top of

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25 m

pillow basalts

massivesulfide

stringerzone stringer

envelopezone

barite

silica

mixedsequence

footwallvolcani-clastics

JackFault

Tasmania

Hellyer

FIG. 1. East-west cross section 10800N (see Fig. 2 for location) of theHellyer massive sulfide deposit, illustrating the stratigraphy of the mine ge-ology, the form of occurrence of the massive sulfide cut in two by the De-vonian Jack fault, and the barite and silica caps, and the distribution of alter-ation assemblages in the volcanic rocks of the footwall. The mixed sequencewas known in the mine as the hanging-wall volcaniclastics and reappears tothe east of the Jack fault in sections 10785 N and 10830N. The volcanic rocksbelow the mixed sequence were known in the mine as the footwall volcani-clastics. The stringer zone of Gemmell and Large (1992), consisting largelyof sericite and quartz; The stringer zone, consisting largely of chlorite withquartz, sericite, and carbonate, and having a silicified core, and the stringerenvelope zone, consisting largely of sericite and quartz, are from Gemmelland Large (1992).

the orebody must have bulged upward slightly because themixed sequence is absent over the center of the orebody andthickens radially outward (McArthur and Dronseika, 1990);the average slope on the upper surface was about 2.5° (fromDowns, 1993; McArthur, 1996).

The silicified core in the footwall contains a sequence ofveins thought to be more or less of the same age as the ore-body above, namely, stage 2A (earliest, mostly quartz, pyrite,

and carbonate), stage 2B (mostly base metal sulfides, particu-larly pyrite), and stage 2C (mostly barite). Because the stage2B and 2C veins have dominantly northeast strike and steepdip, Downs (1993) suggested that they developed during thesecond deformation phase. The orientation of the stage 2Aveins is not known. The stage 2B veins are parallel to basaltdikes, up to 2 m wide, and of composition similar to theHellyer basalt, which cuts the footwall rocks and the massivesulfide orebody (Downs, 1993). Further evidence of tectonicdisturbance during mineralization is seen in two wedges ofclastic material that extended into the upper parts of the mas-sive sulfide body from the basin margin (McArthur, 1996;Solomon and Khin Zaw, 1997).

Above the silicified footwall zone along almost the entirelength of the orebody the massive sulfide has Zn and Pb con-tents below that of the remainder of the orebody, as well as aslightly higher Cu content and a higher pyrite content(McArthur, 1996). This suprafeeder zone is also depleted inAg, the boundary between the depleted zone and the re-mainder being arbitrarily set at 100 ppm Ag by McArthur andDronseika (1990). These authors described the low Ag ore asthe depleted footwall zone, and the remainder (ca. 60% of thetotal orebody) the hanging-wall enriched zone, whileSolomon and Khin Zaw (1997) used the terms the Cu coreand the Zn-Pb ore, respectively, to avoid genetic connotations(Fig. 4). Average grades for these ore types are given in Table1. Smaller zones of Ag depletion occur locally at the base ofthe orebody away from the silicified zone, testifying to up-ward passage of fluids through the orebody from the alteredfootwall (Fig. 4). The ore of the Cu core tends to be more re-crystallized than the remainder of the orebody, particularlyabove the main feeder veins, where it has been brecciatedand intensively pyritized (McArthur, 1996).

In the Zn-Pb ore, the subject of the present study, the dis-tribution of Zn, Pb, and Ag, and the Zn/Pb ratio, as measuredin 2-m drill core intervals, vary irregularly both vertically andhorizontally (from McArthur, 1996). Some of the highest Znand Pb grades (2 m @ 33.0 wt % Zn, and 2 m @ 23.0 wt % Pb)are found in the upper part of the western shelf in present-day section 10730N (Fig. 4). Between sections 10550N and10950N, the top of the massive sulfide is locally enriched inCu (in chalcopyrite and minor tetrahedrite), with gradesabove 0.5 wt percent (McArthur, 1996; Solomon and KhinZaw, 1997; Fig. 4). In some sections (e.g., 10450N, Fig. 5) thisCu-enriched zone extends several meters up into the baritelayer. Gold tends to be enriched near the top of the massivesulfide.

The barite and silica caps were mined along with the mas-sive sulfide ore, but between about 9850N and10300N, a sep-arate lens of uneconomic, sulfide-poor barite lies at the orehorizon and slightly offset from the line of the massive sul-fide-barite orebody (Figs. 2, 3). It overlies veined and alteredvolcaniclastic rock like that seen in the stringer envelope zoneof Gemmell and Large (1992), indicating that it did not growby zone refining of a sulfide body beneath. In the mineralparagenesis of the footwall veins established by Gemmell andLarge (1992), barite first makes its appearance in the cores ofsome stage 2B veins and is the dominant mineral in the suc-ceeding stage 2C veins. Similar, steep to vertical barite veinswere found cutting the massive sulfide ore above the veined,

TEXTURES: HELLYER VHMS DEPOSIT, TASMANIA 1515

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FIG. 2. Plan view of the Hellyer massive sulfide orebody, which includesthe barite and silica caps, projected to the surface, from Gemmell and Large(1992). The southern end is closest to the surface. The separate barite lenshas not been mined. The small barite lens actually extends south to 9850N(see Fig. 3). Mine grid in meters.

plungeof

orebody

outline ofHellyerorebody

barite lens

silicified core of the footwall alteration zone (Downs, 1993),and some link up with the overlying barite layer (J. B. Gem-mell, University of Tasmania, pers. commun., 1999). Thebarites of these veins have sulfur isotope compositions similarto barites in the barite cap and in the footwall veins (Sharpe,1991; Gemmell and Large, 1992, 1993). The barite veins andfragments of barite in the silica cap (Sharpe, 1991) confirmthe upward younging of the deposit proposed by Solomonand Khin Zaw (1997).

The sulfur of the barites is strongly enriched in 34S relativeto coeval seawater sulfate, indicating to Solomon et al. (1988)and Gemmell and Large (1992) that the sulfate was derivedby partial reduction of seawater sulfate during the late stagesof subsea-floor circulation of seawater. For reduction to haveoccurred the temperatures involved were probably aboveabout 200°C (Ohmoto, 1996). This circulation must havepenetrated deep into the underlying Precambrian basementbecause the Sr in the barite is more radiogenic than eitherambient seawater or the host volcanic rocks (Whitford et al.,1992). The possibility raised by Gemmell and Large (1992)that the 34S-enriched barite sulfur was derived from ambientanoxic seawater containing partially reduced sulfate is dis-counted by the evidence given in the next section.


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FIG. 3. Longitudinal section of the Hellyer orebody, which includes the massive sulfide and barite and silica caps, show-ing the separate barite lens in the south and the manner in which the orebody is split into two subparallel units by the Jackfault. The vertical lines in the mine grid are the northings in meters, and the horizontal lines refer to height above sea levelin meters.

FIG. 4. Section 10730N across the Hellyer orebody, showing the majorfaults, the depleted footwall zone (or Cu core) in which Ag <100 ppm, thedistribution of Cu enrichment (>0.5 wt % Cu) near the top of the massivesulfide, and the localized silica cap. There is no barite layer in this section.SEZ = stringer envelope zone and STZ = stringer zone within the alteredfootwall rocks (see Gemmell and Large (1992) and Gemmell and Fulton(2001) for details). Eastings refer to the mine grid, and the vertical refers toheight above sea level in meters.

TABLE 1. Average Metal and Barium Contents of the Hellyer, Kosaka, and Hanaoka Mines (from Shimazaki, 1983; McArthur, 1996)

Ba (wt %) Zn (wt %) Pb (wt %) Cu (wt %)

Hellyer, barite 30.3 5.7 3.5 0.27Hellyer, Zn-Pb 1.3 15.5 8.2 0.34Hellyer, Cu core 1.1 11.7 5.6 0.48Kuroko 13.6 22.5 16.0 2.2Oko 0.24 2.0 0.2 7.4

barite

massive sulfide & barite

The oxidation potential of ambient seawater

The hanging-wall basalts at Hellyer are overlain by gray,carbonaceous marine shales, the Que River Shales. Theymostly contain <<1 wt percent C and a preliminary chemicalstudy by Gee (1970) of samples collected several kilometerssouth of Hellyer indicated that they have C/S values of about0.5, like those of normal Cambrian marine shales (Raiswelland Berner, 1986), with no excess sulfur (i.e., a positive inter-cept on the S axis at zero C values). Another study of similarscale by Sinclair (1994), using samples up to 2 km from theHellyer deposit, indicated an excess of about 0.5 wt percentS, pointing to the possibility that some aqueous sulfideformed in the water column (from Leventhal, 1983). This pat-tern is found in the central Black Sea sediments, but there theexcess S content is greater (1–2% S) and both C and S con-tents are much higher (Leventhal, 1983). The S contents ofthe Que River Shales near the mine are not significantly dif-ferent from elsewhere, but it seems probable that some ex-cess S was derived from hydrothermal activity that persistedduring deposition of the pillow lavas and basal shales overly-ing the Hellyer orebody. Jack (1989) found that above theorebody these rocks have been altered to fuchsite-sericite-chlorite-barite-pyrite assemblages, and immediately beneaththe Que River Shales the basalts for several hundred metersaway from the Hellyer orebody are enriched in pyrite.

The chemical data may not be conclusive but the presenceof agnostid trilobites and dendroids in the shales about 3 kmwest of the mine (Quilty, 1971; Jago, 1979) indicates that theentire water column was oxic, because the dendroid grapto-lites were hemichordates mostly of sessile benthos type, i.e.,

they were attached to the sea floor (Clarkson, 1986; Berry,1987), while the trilobites lived in the water column. Bothgroups required significant oxygen to sustain life. This con-clusion does not relate directly to the time of ore formationbut the period of basalt volcanism intervening between oreand shale deposition may well have been brief, because hy-drothermal alteration continued in the pillow lavas above theorebody until volcanism ceased (Jack, 1989).

Previous work and the scope of the present study

The barite and silica layers have been studied by Sharpe(1991) and are not dealt with in any detail here. McArthur(1996) provided general information based on a ten-year as-sociation with the orebody, as well as descriptions of the min-erals in the massive sulfide, and a detailed pictorial display ofthe textures, based on a three-dimensional sampling grid ex-tending from the 11030N to 10550N (Fig. 2). Additionally,Ramsden et al. (1990) described in some detail the mineralsand chemical composition of three early drill hole intersec-tions of the massive ore. For the present study the Zn-Pborebody was investigated by sampling of drill core, with sam-ple locations in sections 10430N to 11030N, and on 10300N,but with particular focus on 10450N (Fig. 5) and 10730N(Fig. 6). The locations of samples referred to in the text aregiven in Table 2 and Figures 5 and 6. Microscopic studies in-

volved a little over one hundred 5- to 10-cm lengths of pol-ished core mostly cut longitudinally, and reexamination ofabout 80 of the polished sections of Ramsden et al. (1990).


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FIG. 5. Cross section 10450N, showing locations of samples, mostly fromdiamond drill core. The numbers are the sample numbers mentioned in thetext (see Table 2 for locations). The “hanging-wall volcaniclastics” is the mineterm for the regional “mixed sequence.”

FIG. 6. Cross section 10730N, showing locations of samples, mostly fromdiamond drill core; see Figure 4 for geology. The numbers are the samplenumbers mentioned in the text (see Table 2 for locations).

Sample numbers are those of the Department of Geology,University of Tasmania (112...) and of the CSIRO, NorthRyde, Sydney, which used Hellyer mine sample numbers(315...).

The Textures of the Zn-Pb Massive Sulfide OreAlthough a steeply inclined foliation related to Devonian

folding is present in the phyllosilicate-rich parts of the foot-wall, it is only developed in parts of the Zn-Pb ore, in partic-ular where there is a preexisting mineral banding. The ore isalso locally sheared and brecciated near faults, and these fea-tures have been described by McArthur (1996). Otherwise,the Zn-Pb ore is typically massive, fine grained, and highlyvariable in texture, with proportions of minerals varyingrapidly and apparently erratically (McArthur, 1996). Grain di-ameter is mostly from submicron to centimeter scale, in manyplaces with sharp changes over millimeters (Fig. 7A); the oretends to be slightly more recrystallized toward the base.Prominent local features at the mesoscale not related to tec-tonic deformation include breccia or fragmental textures,banding, and veining, and these are described below along


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FIG. 7. A. Typical appearance of the massive Zn-Pb ore at Hellyer—pyritic, highly variable in texture, and very finegrained; sample 112529; scale bar = 1.0 cm. B. Typical microscopic appearance of the relatively undeformed Zn-Pb ore atlow magnfication; sample 112568; scale bar = 300 µm. C. Enlargement of some of the specks of B. Early pyrite kernels havebeen overgrown by later pyrite, with an outer fringe of fibrous, radial pyrite; the remainder is sphalerite and galena; sample112568, etched with concentrated HNO3; scale bar = 15 µm. D. A sample dominated by circular features that consist of cen-ters of early radial, fine-grained pyrite that has been variably recrystallized (?) and/or replaced during later pyritization; sam-ple 112539, etched with HNO3; scale bar = 250 µm. E. Growth of galena at dispersed centers, in several cases the galena al-ternating with pyrite, within a pyrite matrix that may be later than the galena; sample 315866; scale bar = 300 µm. F. Typicalpyrite framboids, together with less common pyrite euhedra in a matrix largely of quartz; sample 112884; scale bar = 30 µm.G. Assemblage of pyrite framboids and pyrite euhedra overgrown by pyrite, in a matrix largely of sphalerite, showing con-siderable variation in grain size of the framboid crystallites; sample 112552; etched with HNO3; scale bar = 20 µm. H. Earlyrecrystallized (?) pyrite framboids enclosed in a mosaic of later pyrite; dark areas are sphalerite; sample 112873, etched withHNO3; scale bar = 20 µm.

TABLE 2. Locations of Hellyer Ore Samples Referred to in the Text

Sample no. Northing (mine grid) Drill hole (m)

112375 10730 HL348/42.2112527 11030 HL040/464.1112529 11030 HL127/107112539 11030 HL141/141.5112552 10730 HL341/100.4112568 10730 HL348/17.75112577 10730 HL348/40112579 10730 HL351/16.8112837 10390 HL3/205.2112856 10490 HL493/196.2112873 10350 HL522/172.5112881 10490 HL651/86.5112884 10490 HL652/60.2112895 10490 HL653/88.65112908 10490 HL655/100.1315850 10450 HL38A/162.5315852 10450 HL38A/164.5315853 10450 HL38A/165.5315858 10450 HL38A/170.5315866 10450 HL38A/178.5

A

C

B

D

with microscale features. Despite intensive search by our-selves and previous workers, no relict chimney structureshave been found.

Dispersed mineral nuclei

Finely dispersed pyrite nuclei within a sulfide or silicatematrix, as seen in Figure 7B, are fairly common in the Zn-Pbore. Etching shows that the nuclei are commonly framboidal(see below) but may also be near-circular, apparently uni-form, crystals spaced some 10 to 20 µm apart (Fig. 7C). Mostshow a thin overgrowth of pyrite and/or a fringe of fibrous, ra-dial pyrite. Sections at right angles indicate that the centersare crudely spherical. Growth at these numerous, dispersedcenters suggests rapid, uniform diffusion through the hostmedium. In many samples, relict outlines of such centers ofcrystallization can be seen within large areas of later, morecoarsely crystalline pyrite. In sample 112539 (Fig. 7D) the cen-ters, also clearly more or less spherical, consist of very fine ra-dial pyrite, partly replaced by later, or recrystallized, pyrite.Sharpe (1991) illustrated similar dispersed pyrite nuclei in the

silica-pyrite cap within a matrix of microcrystalline (1- to 10-µm-diam) quartz. Evidence of the early growth of sphalerite is seenin the dispersed nuclei of some samples (e.g., 112579), and in315853 they have a fringe of fibrous, radial pyrite. Althoughgalena seldom shows such texture, the circular galena bodies ofFigure 7E may represent this style of early crystal growth.

Framboids

In many cases the dispersed centers mentioned above,prove, on etching, to be recrystallized framboids. These areabundant in both Zn-Pb and Cu core types and also the bariteand siliceous caps; in a random inspection of 60 samples, in-cluding veins and deformed and strongly recrystallized sam-ples, half contained framboids. They are exclusively of pyriteand hosted by nonframboidal pyrite, sphalerite, chlorite, andquartz, and in and just below the barite layer, by barite; theyare also a feature of some pyritic veins (see below). As withframboids in sedimentary rocks, they are commonly associ-ated with apparently coeval pyrite euhedra (Fig. 7F, G). Theirdiameter varies within samples but is generally <40 µm,


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FIG. 7. (Cont.)

E

G

F

H

though there are a few examples up to 100 µm (e.g., sample112552). In some samples, like 112568 and 112539, the fram-boids are locally numerous, occupying from 50 to 90 percentof the field of view. The larger diameters are greater thanthose recorded in most sediments or sedimentary rocks (e.g.,Sawlowicz, 1993), possibly due to recrystallization. The inter-nal crystallites are commonly cubes, with diameters mostlybetween about 0.3 and 1 µm but up to 3 µm. In Figure 7H,relict framboids are visible within a complex mosaic of pyritecrystals and interstitial galena. Similar pyrite growing aroundand gradually obliterating framboids is seen in Figure 8A. Atleast some of the later pyrite must be derived by reorganiza-tion of earlier pyrite. In many samples, apparently texturelesspyrite shows faint relics of framboidal texture after etching.As discussed below, the framboids probably indicate the exis-tence of precursor Fe monosulfides during mineralization.

Bladed, dendritic, reticulate, and fibrous texturesPyrite, sphalerite, and galena commonly occur as clusters of

blades, at various scales up to about 500 µm in width and >1

mm long (Fig. 8B, C). A common occurrence is of galenablades partially replaced and/or rimmed by pyrite (Fig. 8B).They may show a crudely radial pattern or form a network(reticulate or dendritic textures), and the blades are com-monly partially replaced and/or are rimmed by another min-eral. All the observed minerals may be replacements of ear-lier phases such as pyrrhotite or barite, and McArthur (1996)interpreted the bladed forms as pseudomorphs of sulfates.However, no incomplete sulfide-barite replacements werefound, even in the barite cap (Sharpe, 1991). Davis et al.(1987) described hexagonal pyrrhotite laths partially replacedby pyrite in a black smoker sulfide mound, and the Hellyerpyrite laths may well have a similar history, although McArthur(1996) only recorded rare occurrences of pyrrhotite, as blebsin sphalerite. Fibrous arrays, commonly radial or part radial,were observed in pyrite and sphalerite (Fig. 8D).

This group of textures is commonly found in isolated blebswithin a pyrite-galena-sphalerite matrix or in zones liningsphalerite lenses of centimeter scale, as in sample 112568(Fig. 8B), or within veins.


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FIG. 8. A. Framboids enclosed in pyrite matrix; dark areas are holes in the section; sample 112552; etched with HNO3;scale bar = 20 µm. B. Dendritic or reticulate texture of pyrite laths in sphalerite (s); some galena is intimately mixed withpyrite in the laths; sample 112568; scale bar = 350 µm. C. Cluster of pyrite laths of varying size in a largely sphalerite matrix;dark areas are holes in the section; sample 112881; scale bar = 400 µm. D. Fibrous texture with pyrite-galena laths in a ma-trix of sphalerite; the galena may have replaced pyrite; sample 315852; scale bar = 20 µm. E. Reniform cluster with concen-tric bands of sphalerite (dark) within pyrite, possibly growing in open space on the material at the base of the photograph;cavity at upper left has been filled with galena; sample 112375; scale bar = 2.0 mm. F. Typical highly irregular veinlet, dom-inantly of pyrite, in a fine-grained matrix of sphalerite; sample 112552; scale bar = 2.0 mm. G. Massive pyrite vein cut at rightangles by carbonate and carbonate-quartz ladder veins (dark in photograph). Less regular, thinner veins are largely pyritewith complex textures. The matrix shows fine bands that are dominantly of pyrite, or sphalerite, or galena and are probablycomponents of earlier veins; sample 112856; scale bar = 1.0 cm. H. Banded pyritic vein in the core of a much wider veincomposed of subparallel, pale bands largely of pyrite or arsenopyrite (a) and dark bands of sphalerite with galena; sample112884; scale bar = 1 cm.

A

C

B

D

Open-space fillings

The microbotryoidal or reniform texture of Figure 8E oc-cupies about 1 cm2 lying within a typically inhomogeneouspyritic matrix. The domelike crystals of pyrite and sphaleriteindicate rapid growth and uniform supply of crystal-buildingcomponents. They appear to be growing unidirectionally intoopen space. Sphalerite was probably precipitated intermit-tently, forming discontinuous bands in pyrite, whereas galenais interstitial and apparently late; similar material is present insample 112908. In other samples there are lensoid or irregu-lar patches up to 2 to 3 cm across, filled by crudely concentricbands of pyrite and low Fe sphalerite (e.g., samples 112568,112577), and these seem to have grown in cavities within thepyritic massive sulfide. Fragments of such finely banded ma-terial occur erratically (e.g., sample 315858) and as compo-nents of breccias, and some may have originally been part ofthin sulfide crusts formed in open space. Fragmentedcrusts(?) of this type occur in the Kamakita mine in northernHonshu (Lee et al., 1974).

Veins

Pyritic veins and veinlets having random orientation occurthroughout the Zn-Pb ore, varying in thickness from severalcentimeters to <0.1 mm; they meander, split, thin over shortdistances, and may cut other veins (Fig. 8F, G). Larger veinscommonly have branching offshoots (sample 112881). Theveins have well-defined margins and generally show complexinternal structure and crude layering parallel to the walls, in-dicating growth inward from the walls (Figs. 8H, 9A, B). Theinternal bands generally show reniform inner margins andvery fine grain size. The main vein in Figure 8G is unusual inshowing no internal structure, even after etching; the quartz-carbonate veinlets oriented perpendicular to the vein walls inthis case suggest internal shrinkage after crystallization or lo-calized extension. Circular structures showing very fine radialfibrous growth are common in layers within other veins andappear to have grown within a medium (samples 112895,112884). Framboids similar to those already described, butwithout associated pyrite euhedra, are also common in the


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FIG. 8. (Cont.)

E

G

F

H

veins, occurring in irregular patches and bands parallel tovein walls (Fig. 9B, C).

There are also abundant examples within the veins of finelybanded encrustations on the walls of open spaces, with indi-vidual bands continuous over distances up to about 5 cm (Fig.9D, F, and G). In these examples sphalerite and galena makeup fine bands within the dominant pyrite, and the latest, in-terstitial material is sphalerite or quartz. In other samples,galena is a common interstitial phase. The encrustations com-monly fill space between very fine grained concentric and/orradial growths at dispersed centers or framboids (Fig.9E). Inparts of the vein of Figure 8H there has been fragmentationof earlier banded growth and recementation by later finelybanded encrustation (Fig. 9F, G), indicating vein reworking.

The relationship of the veins described above to those ofthe footwall is not clear. Stage 2B veins coalesce and mergeupward into the pyritic zone at the base of the depleted foot-wall zone above the siliceous core, according to Gemmell and

Large (1992), and there does not appear to be any physicalconnection to the veins in the Zn-Pb ore.

Breccias

The breccias consist of fragments of pyritic ore of variableshape and size, up to 10 to 20 cm in diameter, in a pyritic ma-trix. They may extend for tens of centimeters or thin to non-fragmented material over centimeters. Some fragments haverounded and irregular, even wispy, shapes, indicating thatthey were unlithified at the time of brecciation. Pseudobrec-cias are created where complex vein sets intersect. In Figure9H it appears that microfragmentation followed pyrite over-growth of a fractured, finely banded pyrite aggregate or crust.

Mineral banding

Mineral banding is most common near the top of the Zn-Pbore and is typically of sphalerite-galena-pyrite bands up to 1cm thick, alternating with pyrite-rich bands (McArthur, 1996;


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FIG. 9. A. Total width of vein in Figure 8H, showing banding within the vein; dominantly pyrite of varying grain and tex-ture; scale bar overlies sphalerite matrix; sample 112884; etched with HNO3; scale bar = 1.5 mm. B. Composite vein show-ing bands dominated by pyrite framboids; galena infilling (g); sample 112837, etched with HNO3; scale bar = 100 µm. C.Area of pyrite framboids within the vein of Figure 8H; etched with HNO3; sample 112884; scale bar = 200 µm. D. Open-space filling within the vein of Figure 8H. Within the infilled lens the dark material is sphalerite (s), the remainder is largelypyrite. The upper part of the picture shows fragments of fine-grained pyritic material cemented by galena (g); sample112884; scale bar = 200 µm. E. In vein of Figure 8H, framboids are overgrown by galena (etched, g) and lined by finelybanded sphalerite and pyrite; sample 112884, etched with HNO3; scale bar = 20 mm. F. In vein of Figure 8H, internally brec-ciated pyritic vein material, cemented by banded pyrite, galena infilling of the cavities (g); black areas are holes in the sec-tion; sample 112884, etched with HNO3; scale bar = 80 µm. G. In vein of Figure 8H, brecciated radial and concentricallybanded pyrite vein filling, cemented by fine-grained pyrite; dark gray is very fine grained radial pyrite, with sphalerite (gray)and galena (g) infilling; sample 112884, etched with HNO3; scale bar = 150 µm. H. Fragment of banded pyrite-sphaleritematerial overgrown by pyrite within a highly variably textured, partly fragmental, silica-pyrite-sphalerite matrix; sample112568; etched with HNO3; scale bar = 150 µm.

A

C

B

D

Fig. 10A). The sphalerite-galena bands are commonlysheared and internally annealed, while the pyritic bands areboudinaged, with prominent fine fractures (commonly filledwith galena or other sulfides) perpendicular to the banding.However, relics of circular textures are still visible, and thepyrite is locally very fine grained. At the boudin nodes (thethinnest parts of the pyrite bands) there is an increased con-tent of quartz. These features suggest that the pyritic bandswere present prior to development of the foliation, as alsoconcluded by McArthur (1996) who referred to obviously tec-tonically remobilized zones of galena and/or tetrahedrite thatcut across earlier banding of the type shown in Figure 10A.Similar banding in the Rosebery deposit, lying about 25 km tothe south-southwest of Hellyer, is clearly earlier than the De-vonian cleavage (Solomon et al., 1987; Solomon and Groves,2000).

Comparison to other massive sulfide deposits

All the textures discussed above are found in modern sea-floor massive sulfide deposits. Bladed, reticulate, dendritic,and fibrous textures are particularly common (e.g., Paradis etal., 1988; Marchig, 1991) and were probably formed as a re-sult of quenching by mixing with cold seawater. Framboids ofpyrite and marcasite have been reported from several modernsea-floor black smoker-type deposits, for example, Hanning-ton and Scott (1988) and Peter and Scott (1988) referred topyrite framboids in amorphous silica and lithified sediment,

respectively, whereas Halbach et al. (1993) pictured themwithin sphalerite. However, in all these modern deposits andalso in most Kuroko ores (an exception is the Daikoku depositof the Ainai mine, Watanabe, 1974), framboidal texture doesnot seem as common as it is in the Hellyer massive sulfide.The only group of deposits in which framboids seem to beanywhere near as common as at Hellyer are the massive sul-fide ores of the Iberian pyrite belt (e.g., fig. 14 in Almodóvaret al. (1998) in the Aznalcóllar and Los Frailes deposits, andobservations by OG).

Banding of the type shown in Figure 10A, common nearthe top of the massive sulfide, appears to be different frombanding described in Kuroko orebodies, which mostly showgrading, load casts, and other evidence of a sedimentary ori-gin or coarse grain size (e.g., Ito et al., 1974; Watanabe, 1974).Sulfide veins occur in most massive sulfide ores but in partsof the Hellyer orebody (e.g., the southern end) are unusuallycommon and complex.

The Formation of the Textures in the Zn-Pb OreThe textures of microcrystals growing from numerous cen-

ters (Fig. 7B, C and D) are probably examples of early crystalgrowth from fine-grained precursor material. Their near-spherical shapes indicate rapid, radial growth in a uniformmedium providing little or no hindrance to diffusive supplyof components; there is no evidence that they formed bycorrosion.


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FIG. 9. (Cont.)

E

G

F

H

Bladed, skeletal, dendritic, and reticulate textures, such asthose involving pyrite, sphalerite, and galena, originated byearly preferential growth due to the presence of impurities oncrystal faces, with subsequent growth continuing at the pointof maximum curvature so as to produce elongate forms(Mullin, 1993, p. 292). The crystals become more equant asthe rate of growth slows and tend to disappear from view ifthe infill is of the same composition. Figure 8B and C are ex-amples in which the dendritic phase ceased to grow, and theinfills have different composition. The isolated blebs of den-dritic and bladed textures, and of sphalerite, may well repre-sent the filling of small fluid conduits. The bladed, skeletal,and fibrous textures are typical of early, rapid crystal growthaccording to Lebedev (1967) and Mullin (1993, p. 21). Theymay develop in open space or within a medium and probablyalso require high degrees of supersaturation (Rimstidt, 1997).Kojima and Ohmoto (1991) observed fibrous textures inwurtzite and sphalerite grown at the cooler end of a vesselhaving a temperature gradient sufficient to induce supersatu-ration. They also produced spherical shells and globular tex-tures 250 to 300 µm in diameter, and there are similar fea-tures in Hellyer pyrite. Given the temperatures of the orefluids at Hellyer (up to 325°C; Khin Zaw et al., 1996), and the

evidence given above that the orebody formed on the seafloor, the most likely cause of the supersaturation in all ofthese cases is quenching.

Breccias may represent wasting of sulfide material from thecentral spine into the basin (McArthur, 1996), possibly trig-gered by tectonic disturbance. The changing stress field dur-ing mineralization (Downs, 1993) may have disturbed theearly basin, causing erosion, and also possibly fracturing, ofthe accumulated sulfide. Movements along the basin-relatedfaults may also have been responsible for the two incursionsof volcaniclastic debris from the basin walls described byMcArthur (1996) and Solomon and Khin Zaw (1997). Thereare no well-worked, obviously sedimentary textures in frag-mental material such as occur, for example, in the Matsumineand Ainai deposits in the Hokuroku basin (Ito et al., 1974; andIshikawa and Yanagisawa, 1974, respectively).

Some of the veins may have been initiated by tectonic frac-turing of the sulfide, others may have resulted from shrinkageand internal settling of the sulfide mass. However, in thesouthern parts of the orebody the veining is too intense to besolely due to shrinkage. The veins presumably filled by pre-cipitation from ore fluids rising from the altered footwall. Thezones of Ag depletion in section 10730N, for example (Fig. 4),


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FIG. 10. A. Typical coarsely banded, sheared ore; pale bands are largely pyrite, dark areas are largely sphalerite and galenawith minor pyrite; sample 112527; scale bar = 1 cm. B. Typical hexagonal sphalerite derived from wurtzite, containing chal-copyrite bands parallel to crystal outlines and also oriented pyrite crystals. The crystal is rimmed by pyrite with minor galenaand the matrix is of silicate with pyrite; sample 315858; scale bar = 200 µm. C. Sphalerite with bands of chalcopyrite, out-lining a probable earlier wurtzite crystal and enclosing oriented pyrite crystals in a matrix of sphalerite; sample 315852; scalebar = 80 µm. D. Local crust of sphalerite peppered with very fine, irregular, pale gray chalcopyrite in crudely defined bands,along with erratically distributed, larger, white crystals of pyrite. The sphalerite crust is rimmed by very fine, radial pyrite;sample 315853; scale bar is 500 µm.

A

C

B

D

are evidence that fluids penetrated the Zn-Pb ore, and thebrecciation of vein material (Fig. 9F, G) indicates continueddisturbance of the sulfide body.

Mineral banding can develop by differential gravity settlingof sulfide particles and would be effective in separatinggalena from sphalerite and pyrite. However, it is unlikely thatpyrite and sphalerite could be separated unless the particlesof each phase are of different radius. The latter appears tohave been the controlling factor in the development of agalena-rich layer in a chemically precipitated, mixed PbS-ZnSgel (?) or mush prepared by Lebedev (1967, p. 227). Thepaucity of pyrite and the concentration of galena and spha-lerite in the western shelf (section 10730N, Fig. 4) may be theresult of currents sweeping relatively fine particles of Zn andPb sulfides onto the ledge. However, the ore does not showthe fine laminae, slump structures, and synsedimentary faultsthat characterize the upper part of the Filon Norte orebodyat Tharsis in Spain (Tornos et al., 1998). Another possiblemechanism for forming the banding is differential pressuresolution during aging, as proposed for parts of the LayeredSeries of the Skaergaard intrusion by Boudreau and McBir-ney (1997), and McArthur (1996) proposed an origin by lat-eral injection of fluids beneath a sulfide or barite crust.

The significance of framboids

Wilkin and Barnes (1997) concluded from an extended re-view of the origin of framboidal pyrite, and their own hy-drothermal experiments, that its development in acid, re-duced fluids involves the nucleation and growth of Femonosulfide (e.g., mackinawite) that then undergoes sulfida-tion or Fe loss to greigite, the crystallites of which aggregateto form the typical texture; the greigite is then replaced bypyrite. This process, involving the formation of successivemetastable phases, is an example of Ostwald’s step rules (Ost-wald, 1897). Framboidal greigite is unlikely to form at tem-peratures above 200°C, because direct nucleation of Fe disul-fides then proceeds faster than precursor formation (Wilkinand Barnes, 1996, 1997). Once pyrite or marcasite hasformed, it can act as a nucleus for further growth of Fe disul-fide direct from solution. The development of framboids re-quires supersaturation of Fe monosulfides, which have solu-bilities more than ten orders of magnitude greater than pyriteat 25°C (Benning et al., 2000), but this may well have beenachieved during quenching of the venting ore fluids atHellyer. Schoonen and Barnes (1991a, b) reported amor-phous Fe monosulfide, cubic FeS, and mackinawite as theearliest precursor phases below 100°C, and Sweeney and Ka-plan (1973) found evidence of hexagonal pyrrhotite precedingframboid formation in protracted experiments at 60° and85°C. However, this step seems not to have been duplicatedby other workers, Lennie et al. (1995) concluding thatpyrrhotite formation below 100°C is extremely slow. Schoo-nen and Barnes (1991c) found that hexagonal pyrrhotite wasthe dominant initial product of quenching experiments above180°C, mackinawite the major phase below that temperature,and mixtures of monosulfides below 100°C (see above).Hexagonal pyrrhotite partially converted to pyrite or marca-site is commonly observed in modern sea-floor sulfide de-posits (e.g., Davis et al., 1987; Marchig et al., 1988; Peter andScott, 1988), and as already noted, it may have been present

at Hellyer, because McArthur (1996) found pyrrhotite blebsin sphalerite.

Schoonen and Barnes (1991a, b) and Benning et al. (2000),after reviewing earlier work and performing additional exper-iments, concluded that pyrite formation at low temperatures(<100°C) from acid solutions proceeded only via monosulfideprecursors and greigite under weakly reducing conditions,i.e., in the presence of an oxidant such as molecular oxygen.Thus framboidal pyrite will only form at low temperatureunder similar conditions (Wilkin and Barnes 1997). Althoughthere is considerable support for Fe loss as the main processinvolved in forming pyrite from monosulfides below 100°C(e.g., Furukawa and Barnes, 1995), the experiments con-firmed earlier conclusions that mackinawite remains the sta-ble phase in anoxic solutions, conversion only proceeding inthe presence of zero-valent sulfur, i.e., the process is one ofsulfidation (Benning et al., 2000). However, Butler andRickard (2000) recently produced framboids in experimentsat 60° to 140°C that did not involve any sulfur species otherthan H2S, and without producing greigite. Framboids and eu-hedral crystals were produced from freeze-dried FeS in theabsence of oxygen, and the authors suggested that the reac-tion proceeded via dissolution of FeS and reaction of aqueousFeS with H2S. Their experimental solutions were maintainedat a pH of 6, but they suggested that more acid conditionsshould accelerate the reaction. Although the weight of earlierexperimental evidence clearly favors weakly reducing condi-tions and the formation of ephemeral greigite, the conflictingdata of Butler and Rickard (2000) casts some doubt on theprocesses of pyrite nucleation and framboid formation, af-fecting the ensuing discussion in this paper. The additionalsulfur required for either pathway to be followed at Hellyercould come from the ore fluid provided that ΣS > Σmetals.

Framboids typically have internal crystals of about the samesize and form, indicating rapid nucleation of microcrystals fol-lowed by a short period of growth (Wilkin and Barnes, 1997).At Hellyer this appears to have taken place at centers dis-persed in a sulfide medium, the centers probably being sup-plied by uniform diffusion from the surrounding mediumuntil the supply was exhausted. The entire disequilibriummackinawite-pyrite conversion and the growth of framboidsin any one place could take no more than a few days, andframboid growth appears to be favored when the mineralconversions are rapid (Wilkin and Barnes, 1996, 1997). Theoccurrence of euhedra apparently coeval with framboids (Fig.7F, G) is common in sediments and is seen in experiments(Wilkin and Barnes, 1996). In the Hellyer ores framboidsclearly made ideal precursors for the subsequent growth ofpyrite and much of the ore may have evolved in this way.

Textures in SphaleriteA significant proportion of the sphalerite in the massive sul-

fide displays evidence of derivation from wurtzite. Hexagonaloutlines to crystals, chalcopyrite bands of hexagonal shape,hexagonal crusts of pyrite around sphalerite, and partialhexagonal outlines marked by pyrite or galena are seenthroughout the orebody, including the veins, and all point toderivation from wurtzite (Fig. 10B, C), as suggested by Rams-den et al. (1990) and McArthur (1996). They also occur in thebarite cap (Sharpe, 1991; Solomon and Groves, 2000, fig. SC


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8.6C). Etching reveals that the sphalerite in many cases hascrystal outlines that cut across the precursor crystal bound-aries, presumably due to recrystallization subsequent to de-position. Most complete hexagonal forms have diameters upto about 100 µm, but Ramsden et al. (1990) displayed a near-perfect hexagon about 350 µm across in which the Fe contentof the sphalerite bands had near-constant composition acrossthe crystal; the presumably inclined chalcopyrite zones ren-dered the intervening sphalerite almost opaque. Surroundingsphalerite has a high Fe content in some cases, a low Fe con-tent in others.

Linear integrations of photographs of several examplesshowed that the volume of chalcopyrite in the wurtzite-de-rived sphalerite is much greater than can be explained by itsequilibrium solubility, which is no more than 2.4 mole per-cent CuS below 500°C (Kojima and Sugaki, 1985). Values ob-tained from the Hellyer samples by linear integration rangedupward from about 6 mole percent CuS. Ramsden et al.(1990) and McArthur (1996) assumed that the chalcopyritehad replaced sphalerite (the “chalcopyrite disease” of Barton,1978). It is clearly possible that replacement may have oc-curred (see below), but the regular orientation parallel tocrystal outlines of most of the lines or strings of chalcopyrite,and the fact that the the lines or strings are confined to whatappear to be primary crystals, hints strongly at exsolution fol-lowing cooling and/or inversion to sphalerite. While it is notpossible to exclude with certainty either coprecipitation (e.g.,Kojima et al., 1995) or replacement of chalcopyrite it is sug-gested that the early, metastable wurtzite crystals, perhapsgrowing asymmetrically, have been able to include muchhigher CuS and FeS contents than indicated by equilibriumexperimentation.

In some sphalerite that appears to have grown in open space,forming crusts with pyrite and other phases, chalcopyrite occursas dispersed fine blebs (Fig. 10D) and has a very similar ap-pearance to figure 2C of Kojima (1990), which displays a tex-ture produced by coprecipitation of sphalerite and chalcopy-rite brought about by cooling of supersaturated solutions.

There is clear evidence that some of the chalcopyrite of theZn-Pb ore locally replaced earlier phases. In samples fromthe Cu-rich zone at the top of the massive sulfide in section10450N, chalcopyrite occurs in fractures and some appears tohave locally replaced sphalerite and pyrite (e.g., sample315850). Chalcopyrite microveinlets cut across part of thevein in Figure 8H (though not visible in the photograph), andchalcopyrite has locally and irregularly replaced pyrite withinthe vein. This late chalcopyrite may belong to the phase of Cuenrichment of the depleted footwall zone. In sample 112674,from section 10450N (Fig. 5), elongate chalcopyrite blebs oc-cupy fine fractures crosscutting growth zones in nonhexago-nal, almost colorless sphalerite, or form radial arrays perpen-dicular to, and cutting, the primary zoning. However, ingeneral such evidence of replacement by chalcopyrite is notcommon in the Zn-Pb ore. Where annealing has occurred insphalerite-rich ore, due to deformation and metamorphism,the chalcopyrite is dispersed regularly through the sphalerite,with the boundaries of these phases and associated galenashowing equilibrium interfacial angles.

Ramsden et al. (1990) systematically analyzed the Fe con-tent of sphalerite every 2 m along three drill holes cutting the

Zn-Pb ore, and McArthur (1996) analyzed samples taken ona regular grid from three sections. Sphalerite in the Zn-Pb oreranges from 1.7 to a maximum 10.5 wt percent FeS, the high-est values tending to lie within the western and lower parts ofthe ore, and the lowest values mostly close to the barite layer.Although Ramsden et al. (1990) suggested there was a hori-zon within the massive sulfide of Fe-rich, wurtzite-derivedsphalerite, subsequent work by McArthur (1996) and our-selves indicates that they occur in limited and haphazardlyarranged zones, mostly in the Zn-Pb ore. Judging by color,sphalerite filling or lining centimeter-scale cavities generallyhas a relatively low Fe content (see earlier comments onopen-space filling and dendritic textures).

Although the range of FeS contents in the sphalerites fromthe Cu core is similar (1.4–6.3 and 1–10.8 wt %; McArthur,1996), the average is lower (4.9 in the Zn-Pb ore and 3.1 inthe Cu core), contrasting with the variation seen in Hokurokuand other volcanic-hosted massive sulfide ores (e.g., Urabe,1974). Sphalerite in the barite and silica caps has an FeS con-tent similar to the Cu core, namely, 0.9 to 5.6 wt percent, av-erage 3.6 wt percent (Sharpe, 1991). The stage 2B veins inthe footwall have yielded the highest fluid inclusion homoge-nization temperatures (Khin Zaw et al., 1996) and containpyrite and fluid-nucleated chalcopyrite. Solomon and KhinZaw (1997) surmised that the leaching of Zn and Pb from theCu core, and deposition of the stratiform Cu-rich zone at thetop of the Zn-Pb ore, occurred during this stage. Depletion ofFeS in sphalerite may have taken place in response to a fall inaS2

or it might be related to the later part of the stage 2B veinand/or the stage 2C of Gemmell and Large (1992), when in-coming sulfate in the traversing fluid may have raised theoverall oxidation potential, and temperature was declining.

The Hellyer Textures in Relation to Ore Formation on the Sea Floor

The presence of barite fragments in the hanging-wall vol-caniclastics or mine sequence (Sharpe, 1991) demonstratesexposure of the Hellyer orebody on the sea floor. Except pos-sibly at the base of the Cu core (depleted footwall zone),where intense pyritization has occurred, there is little evi-dence of replacement of the footwall volcanic rocks, so it ishighly probable that the orebody formed largely on the seafloor. To date, two different hypotheses have been proposedfor sulfide deposition at Hellyer, namely, the Kuroko-type ormound model and the brine pool model.

No previous authors have drawn comparison to the RedSea deeps, not because the textures are dissimilar, but be-cause the tectonic environment, the metal-bearing basinalsediments (clays, carbonates, Fe and Mn oxides, sulfates, andphylloslicates), and the metal contents and ratios (Bischoff,1969; Pottorf and Barnes, 1983) are quite different fromthose in volcanic-hosted massive sulfide deposits. The At-lantis II Deep has, however, been compared to the environ-ments in which sediment-hosted Pb-Zn ores were formed(e.g., Large et al., 2000).

The Kuroko mound model

McArthur (1989), McArthur and Dronseika (1990), Large(1990), and McArthur (1996) proposed that the Hellyer mas-sive sulfide grew from fluids venting onto the sea floor during


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footwall alteration and veining, initially forming a sulfate-sul-fide-silica crust or mound beneath which subsequent sulfidemineralization occurred by addition and replacement. Ac-cording to McArthur (1996), zone refining of early massivesulfide resulted in formation of the Cu core and Pb and Znenrichment of the overlying and surrounding massive sulfide(the hanging-wall–enriched zone or Zn-Pb ore). In this sce-nario, the early products of mixing with seawater underwentsubstantial and complex modification during the life of thesystem. These models are based on hypotheses for formationof the ores of the Hokuroku basin and of modern sea-floorsulfide mounds. Ohmoto (1996) combined the propositions ofGoldfarb et al. (1983), Eldridge et al. (1983), and Lydon(1988) to suggest that two major processes operated duringKuroko ore formation: first, rapid mixing of ore fluids inchimneys or at the exterior of sulfide mounds with local sea-water, followed by incorporation of quenched material intothe mound; and second, reaction of the early products withlater, hotter, rising ore fluids to produce the vertical zoningtypical of Kuroko ores. Neither Large (1992) nor McArthur(1996) included chimney formation in their models (in theformer case because chimneys have not been found in Aus-tralian deposits), nor did they deal with the problem of dis-

posing of the large volumes of buoyant ore fluid that musthave passed through the mound and into the ocean.

The textures in the Hellyer deposit appear to be compati-ble with the Kuroko mound hypothesis, because they are sim-ilar to those in black smoker- and Kuroko-type deposits ex-cept for the abundance of framboids, the banding, and theintensity of veining. The sulfide banding could have resultedfrom plume fallout, and the veining might simply indicate rel-atively intense fault movements during mineralization. How-ever, there does not seem to be a ready explanation for theabundance of framboids. Comparison of the geology of theKuroko and Hellyer deposits reveals a number of other dif-ferences (Table 3; see also Solomon and Groves, 2000) thatsuggest a different interpretation to the Kuroko moundmodel for ore formation on the sea floor. A striking differenceis the high Zn/Cu and Pb/Cu value of the Hellyer depositcompared to Kuroko ores. This might be explained by a phaseof high-temperature zone refining that was not as protractedat Hellyer as in Kuroko ores, resulting in lower Cu content,and the early Zn-Pb phase may have been prolonged. The lessintense zone refining could also account for the relatively finegrain of Hellyer ore, and the relatively minor loss of Pb andZn from the Cu core. However, if the Zn-Pb ore is the early


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TABLE 3. Comparisons between Hellyer and Kuroko Ores

Hellyer1 Kuroko Source for kuroko

Size, total tonnage 16.2 Mt Only one greater (Matsumine @ 30 Mt), Tanimura et al. (1983)and one similar (Motoyama @ 15 Mt), the rest much smaller

Metal content, Zn/Cu 2.2 Mt Zn, 1.1 Mt Pb, Matsumine 1.1 Mt Zn, 0.3 Mt Pb, 0.7 Mt Cu, Tanimura et al. (1983)0.06 Mt Cu, Zn/Cu = 37; Zn/Cu = 1.5; Motoyama 0.7 Mt Zn, 0.1 Mt Pb,

0.3 Mt Cu, Zn/Cu = 2

Ore types Barite-sulfide ore Kuroko (black ore) has less barite, more sulfide Eldridge et al. (1983)Zn-Pb ore No equivalentCu core Oko (yellow ore) has higher Cu, lower Zn and Pb Shimazaki (1974)

Grain size Variable, submicron upward More coarsely grained overall Eldridge et al. (1983)

Relative ages of ore types Cu core replaces Zn-Pb ore but Oko (yellow ore) replaces kuroko, Eldridge et al. (1983)barite-sulfide later which forms first

Mineral content/fO2Arsenopyrite-pyrite/low fO2

Pyrite (As in tetrahedrite-tennantite)/ Shimazaki (1974)relatively high fO2

Fe oxides Not present Hematite in tetsusekiei Kalogeropoulis and Scott (1983)

Chimney evidence None Probable fragments Eldridge et al. (1983): observations by MS (1981)

FeS content of sphalerite Lower in Cu core than in Higher in oko than kuroko but overall Hellyer: Urabe (1974), Zn-Pb ore much lower than Ohmoto et al. (1983)

δ34S in barite >> coeval seawater sulfate Slightly greater than coeval seawater sulfate Ohmoto et al. (1983)

δ34S trend Increasing upward Decreasing upward in Shakanai ores Sasaki (1974)

Textures Framboids, banding, and Framboids, banding, and veinlets generally For example, Sato (1974), veinlets common not so common Watanabe (1974),

Eldridge et al. (1983)

Fluid inclusion salinities Up to 15 wt % ≤5.5 wt % Pisutha-Arnond and Ohmoto (1983)

1 Data for Hellyer from Sharpe (1991), Gemmell and Large (1992, 1993), McArthur (1996), Solomon and Khin Zaw (1997)

material then it would be expected to contain barite and havesimilarities to kuroko (black ore). Barium was present in theore fluids because Ba is enriched in the sericitized parts of thefootwall (Gemmell and Fulton, 2001). Given the oxic natureof ambient seawater, it would seem very difficult to avoid pre-cipitation of barite from seawater sulfate (with sulfur isotopecompositions close to that of ambient seawater sulfate) as thebuoyant ore fluids passed from the mound, and also the for-mation of sulfate chimneys. The barite-sulfur appears to bederived from ambient seawater sulfate but only after deepcirculation and partial inorganic reduction, as discussed above(“Geology of the Hellyer Orebody”). There is no clear evi-dence of barite replacement and the supposed sulfate re-placement textures can also be interpreted as products ofearly, rapid, crystal growth with equal validity. The barite-sul-fide cap is younger than the massive sulfide so it cannot haveacted as a seal.

The lack of Fe oxides in the Hellyer deposit, and particu-larly in the silica lenses, also indicates little or no contact withambient seawater, as does the pyrite-arsenopyrite assemblagein the massive sulfide, which indicates retention of the low ox-idation potential of the ore fluids even after reaching the seafloor. Further support for negligible oxidation is seen in thesulfur isotope values of sulfides, which show little or no up-ward variation at Hellyer (Gemmell and Large, 1992, 1993).In the Shakanai 1 ores the δ34S values decrease upward byseveral per mil, probably due to an increase in the SO4

2–/H2Sratio of the ore fluids (Sasaki, 1974) as they rose and mixedwith seawater.

The significance of the saline ore fluids at Hellyer

The most important difference in Table 3 relates to thesalinity of the ore fluids. Khin Zaw et al. (1996) determinedthat fluid inclusions in quartz in the stage 2 footwall veinshave salinities ranging to much higher values than found inthe Kuroko ores, which are close to those of seawater(Pisutha-Arnond and Ohmoto, 1983). As shown by Solomonand Khin Zaw (1997), using temperature-density data forNaCl solutions, if the saline fluids of the stage 2A veins werevented into seawater they would have reversed buoyancy andponded in a suitable trap, such as the basin revealed byDowns (1993). Some stage 2B veins penetrated the massivesulfide (Gemmell and Large, 1992) so at least several metershad accumulated during stage 2A and earliest stage 2B, im-plying substantial volumes of fluid in the basin by that time.During stage 2B the saline fluids were joined by, but appar-ently did not mix with, fluids of low salinity but these are un-likely to have dramatically disturbed the brine pool. Stage 2Cfluids were similar to those of stage 2A. Fluid inclusions frommassive sulfide stockworks proved powerful tools in predict-ing black smoker behavior before the modern sea-floor dis-coveries (Solomon and Walshe, 1979), and there seems noreason why they should not be similarly important at Hellyer.The brine pool hypothesis could be vitiated if it was foundthat the inclusions did not represent the ore fluids or if thetiming with respect to mineralization had been misread.However, as yet no conflicting data or interpretations havebeen reported, and the inclusion characteristics provide avery strong argument for the depositional environment atHellyer. As discussed below, the model can account for, or is

at least compatible with, most of the differences listed inTable 3, including the slight differences in textures comparedto other massive sulfide deposits.

Sulfide deposition in the brine pool

High-temperature midocean ridge fluids emerging fromanhydrite chimneys are clear until mixing with seawater com-mences at a height of a few centimeters, at which point theplume blackens with contained sulfides. Feely et al. (1987,1994) and Mottl and McConachy (1990) sampled blacksmoker plumes at midocean ridges, recording pyrrhotite,pyrite, sphalerite, and chalcopyrite as the major sulfide parti-cles, along with an unidentified Fe-Si phase, amorphous sil-ica, Fe oxyhydroxides, and anhydrite. No amorphous sulfideswere found but may have been present at the time of trap-ping, because recrystallization of Fe sulfides could have pre-ceded the subsequent, onshore, X-ray analysis; crystallizationfrom amorphous Fe-S material prepared chemically at 65°Cin the laboratory by Schoonen and Barnes (1991b) com-menced within hours of production. Black smoker particlesranged in size up to >100 µm, and the larger sulfide particlessettled close to the vents (Feely et al., 1987). However, thebulk of the solid material, 99 to 99.9 percent of the total mass,and with diameters of <0.4 to 10 µm, tended to be carried upinto the plumes and distributed laterally (Nelsen et al., 1986;German and Sparks, 1993; Feely et al., 1994).

Ore fluids venting into the Hellyer brine pool would haveundergone similar supersaturation due to quenching, pro-vided the pool temperature was much lower than that of thevent fluids. Using the expressions of McDougall (1984) for abrine pool fed from below, and knowing the range of temper-atures and salinities of the ore fluids from the fluid inclusiondata, the temperatures and salinities of the pool at steadystate can be calculated for various basin areas and mass fluxes.In Figure 11 it is assumed that the area of the pool surface is6.105 m2 (about twice the area of the orebody) and the salini-ties either 11 or 8 wt percent. The volume fluxes used rangefrom 1 l/s–1 (a typical black smoker vent flux at midoceanridges, according to McDuff, 1995) to 100 l/s–1, a figure de-rived from the estimate of 150 ± 60 kg/s–1 by Converse et al.(1984) for the vent flux from1 km of midocean ridge at 21° N.The Hellyer feeder line is about half that length. Though theassumption of steady state and volume flux allow considerablevariation, it seems unlikely from the calculations that brinepool temperatures would have exceeded more than about70°C. If the pool area was larger, then the temperature wouldhave been lower. Rapid cooling of ore fluids at or near satu-ration of Fe, Zn, Pb, and Cu sulfides from 250° to 325°C tobrine pool temperature would result in high degrees of su-persaturation (e.g., Seward and Barnes, 1997).

The mineral content of the veins in the footwall feederzone indicates that subsurface saturation occurred duringstage 2A of pyrite, during stage 2B of pyrite, sphalerite,galena, and chalcopyrite, and in stage 2C of pyrite and barite(Gemmell and Large, 1992). The erratic Cu analyses recordedduring PIXE analyses by Khin Zaw et al. (1996) of fluid in-clusions from stage 2B veins may reflect accidental trappingof effectively invisible, nanometer-sized, chalcopyrite parti-cles. Precipitation in the veins may have been caused initiallyby conductive cooling against the wall rocks and/or fluid-rock


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reactions that increased pH but perhaps more likely as a re-sult of entrainment of more dilute and cooler seawater.

As the ore fluids entered the Hellyer pool and mixed tur-bulently with the cooler, spent ore fluid, disequilibriumquenching of mineral sulfides would have been close to in-stantaneous. The higher the degree of supersaturation (andthe solubility) the faster the rate of nucleation and the greaterthe number of nuclei, and hence the finer the grain size(Morse and Casey, 1988; Mullin, 1993, p. 175). The metal sul-fur particles would have been swept by the currents towardthe bottom of the pool, where they would have settled to thepool floor under gravity, as shown schematically in Figure 12.The current pattern is taken from the experiments of Mc-Dougall (1984), and the flow and quenching pattern is prob-ably similar to that proposed by Pottorf and Barnes (1983) for

the Atlantis II Deep. Whether or not the particles underwentflocculation (= aggregation or agglomeration) as they settledwould have depended partly on the number of particles perunit volume (Mullin, 1993, p. 286). Given the likely low metalcontent of the ore fluids, and the rate of mixing with poolwater in the turbulent plume, the particles are likely to havebeen widely dispersed, limiting flocculation in the early stagesof settling but allowing it as they became concentrated towardthe basin floor. Judging by the black smoker plumes, the par-ticles reaching the basin floor are likely to have had a widerange in diameter, from submicron upward. Micron-sizedparticles of pyrite would settle extremely slowly; Stokes’s Law;(Stokes, 1851), indicating that a 1-µm particle in pool waterwould descend at the rate of about 2 µm/s–1 (17 cm/d–1). It isnot clear how much of the crystal growth and compositionalchange (for example, framboid development) occurs in thebrine pool during quenching and transport, as distinct fromwithin the settled sulfide mud, discussed below. For the pur-poses of discussion it is assumed that most occurs within thesettled mud. If the sulfide particles did not grow beyondabout a 1 µm diameter, on settling they might have aggre-gated to form a network structure, thereby trapping includedfluid, and the resulting material is likely to have had the con-sistency of a dense mud, possibly with gel-like properties.Judging by the rates of crystallization in chemically prepared,very fine grained precipitates, however, this colloidal phase, ifit occurred, seems likely to have been short, perhaps hours todays (Lebedev, 1967; Schoonen and Barnes, 1991b).

The aging of the sulfide mud

In the settled sulfide mud, internal adjustments would haveinvolved the growth of larger crystals at the expense ofsmaller (the process of Ostwald ripening, defined byLiesegang, 1911). For diffusion-controlled crystal growth, therate is proportional to the diffusion coefficients of the com-ponents and the degree of supersaturation and is inverselyproportional to the grain radius (Mullin, 1993, p. 288). Ifgrowth is controlled by surface reactions, then the kinetics area function of the degree of supersaturation (Mullin, 1993).


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FIG. 11. Variation of brine pool temperature and salinity for various ventfluxes (in liters per second) and fluid parameters, taking an area of the poolsurface as 6.105 m2 (see text) and assuming a steady state; calculated fromequations in McDougall (1984). Figures shown in the lines for fluids with 11wt percent salinity are calculated salinities of the pool. The pool tempera-tures give some indication of the degree of quenching of the incoming orefluids during mineralization.

FIG. 12. An east-west sketch section, showing the proposed brine pool, the upstanding veined siliceous feeder pipe, thelikely current flow within the pool, and the important molecules undergoing diffusion across the brine-seawater interface.The seawater is oxic because it supports a hemichordate sessile benthic fauna during deposition of the Que River Shale, whilethe brine pool, filled with spent ore fluid, is close to anoxic (as indicated by the mineral content of the massive sulfide). Over-flow from the pool is likely to have been to north and/or south along the main fault lines. Drawn from Hellyer geology andthe experiments of McDougall (1984).

Rapid and uniform diffusion through the settled mud atHellyer would have allowed growth of near-spherical struc-tures (e.g., Fig. 7C-E), while slower rates may have led togrowth of euhedral crystals, simulating the industrial methodsof growing high-quality crystals (e.g., Lefaucheux and Robert,1994). Nonuniform growth within the mud could have givenrise to the dendritic and fibrous textures of Figure 8B-D.These textures are similar to those of modern black smoker-type deposits but in these cases the crystals grow in a frame-work of sphalerite and silica or anhydrite and sphalerite (theearly “scaffolding” of Paradis et al., 1988), rather than amedium dominated by sulfide minerals. Crystal growth mayalso have been influenced by ore fluids traversing the sulfidemass from below, as shown by the hydrothermally alteredfootwall, which is as laterally extensive as the orebody, also thestringer veins in some of the immediate footwall, and the localdepletions in Ag, as shown, for example, in the western shelfof section 10730N (Fig. 4). The Zn-Pb ore has clearly not un-dergone the hydrothermal flushing and more intense recrys-tallization that is common in many massive sulfide oresformed from buoyant fluids; however, some fraction of therising ore fluids clearly permeated the wall rocks beneath themassive sulfide body during mineralization. These fluids mayhave been responsible for the apparent conduit and open-space fillings (Fig. 8E). The Zn-Pb ore contains only a fewpercent silica as fine-grained quartz, yet the ore fluids in thefootwall veins were probably saturated in silica becausequartz is abundant in stage 2A veins, it occurs in stage 2Bveins, and the central part of the alteration pipe was silicifiedlargely before vein formation (Gemmell and Large, 1992).Quenching of such fluids on venting would have resulted inhigh degrees of silica supersaturation but the rate of nucle-ation is likely to have been slow due to the rapid fall in tem-perature on exhalation and the scarcity of suitable particles onwhich to nucleate (Rimstidt and Barnes, 1980; Ohmoto, 1996),resulting in most silica deposition occurring only in the dyingstages of orebody formation. The silica remains very finegrained, particles commonly having diameters of 1 to 2 µm(Sharpe, 1991), and may have been colloidal on deposition.

There is unlikely to have been entrainment of seawateracross the surface of the brine pool at steady state (Mc-Dougall, 1984), but seawater sulfate may have been incorpo-rated as a result of mixing with ore fluid during the early fill-ing of the basin and also subsequently due to ore fluidentering the pool and overshooting the brine-seawater inter-face (see Solomon and Khin Zaw, 1997). As previously men-tioned, there is a possibility that some of the bladed texturescould be replacements of sulfate minerals. However, the lowcontent of barite in the Zn-Pb ore (1.3 wt % Ba, McArthur(1996), including some later veins and barite just below thebarite cap) points to a low content of sulfate in the pool fluidduring massive sulfide mineralization, because Ba is enrichedin the altered footwall rocks (Gemmell and Fulton, 2001). Ifsubstantial sulfate had been present, barite would have beendeposited on mixing with the pool fluid (see earlier discussionof the Kuroko mound model). The exclusion of seawater sul-fate is also indicated by the apparent lack of chimney frag-ments in the orebody.

The pyrite-arsenopyrite assemblage, which is present through-out the Zn-Pb ore, suggests that at no time during massive

sulfide mineralization was the ore fluid oxidized and that sub-sequent events did not allow oxidation of the sulfide body.This is also supported by the scarcity of Fe oxide (0.03 wt %magnetite and trace hematite; McArthur, 1996), and the lackof Fe oxides in the silica cap. It is also likely that the brinepool was acid, because feldspar is unstable and sericite stablein the altered footwall and within the massive sulfide. Atsalinities near 2 m NaCl equiv, the pH of the footwall fluidsat 300°C was probably <5 (see Henley, 1984). Cooling of thefluid in the pool would cause dissociation of HCl and furtherincrease the aH+ (see discussion in Solomon and Khin Zaw,1997). The likely pH in the Hellyer pool would have favoredprecipitation of marcasite rather than pyrite (Murowchik andBarnes, 1986), but marcasite also requires the presence ofweakly oxidized sulfur species, such as polysulfides, to con-centrations estimated to be at least 10–5 to 10–6 m (Benning etal., 2000). Thus the brine pool is assumed to have had con-centrations of oxidized sulfur species below that level. Thissituation contrasts with the ubiquity of marcasite in modernmidocean ridge sulfide deposits in which there has been sub-stantial interaction with local seawater (e.g., Paradis et al.,1988; Peter and Scott, 1988).

Arsenopyrite precipitation ceased during formation of thebarite cap (Sharpe, 1991), suggesting an increase in the oxi-dation potential of the pool fluid. There is a slight increase inthe average δ34S value of the sulfides in the barite (8.2‰)compared to that in the massive sulfide (7.2‰), from analy-ses by Gemmell and Large (1992, 1993), Sharpe (1991), andone of us (MS). This could represent an approach to sulfate-sulfide isotope equilibrium but is still far removed from likelyequilibrium values, given the average 40 per mil for thebarite-sulfur (Sharpe, 1991; Gemmell and Large, 1992). Thusat the barite stage there may have been an increase in fO2

tolevels above pyrite-arsenopyrite equilibrium but not to thelevels of Fe oxide stability; there is no indication that ar-senopyrite was dissolved at this time, and arsenopyrite depo-sition was resumed in the silica cap (Sharpe, 1991).

Whatever its physical state, the settled sulfide mass mustinitially have included at least the minimum volume of brineconsistent with cubic close-packed stacking (33%) and verylikely contained considerably more. As the sulfide mass crys-tallized and pore space was filled, contained water would havebeen largely expelled, possibly leading to overall shrinkageand consequent cracking. Deposition in cracks from footwall-sourced fluids may account for some of the irregular veinsand veinlets. However, many veins may be the result of re-peated movements in the basin-related faults, continuousbands lining the vein interiors indicating intermittent satura-tion of pyrite, sphalerite, and galena (Fig. 9F, G). It is alsopossible that some veins grew by diffusion toward incipientcracks developing early in the mud history.

Framboids and oxidation of the brine pool

It appears from the weight of experimental evidence that atlow temperature framboidal pyrite is derived by stepwise sul-fidation of Fe monosulfides via greigite, in weakly oxidizedsolutions (see above). If the pool is initially anoxic, then asource of oxidized sulfur species must be sought, but at a levelless than 10–5 to 10–6 m polysufide to account for the lack ofmarcasite (see above). A possible source is the precipitation of


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arsenopyrite, as this proceeds by release of O2 during the fol-lowing reaction (from Heinrich and Eadington, 1986):4H3AsO3 + 4H2S + 4FeCl+ → 4FeAsS + 3O2 + 4Cl– + 8H+ +6H2O. However, with only about 1.4 wt percent arsenopyritein the orebody (McArthur, 1996), this contribution would beminor. A more significant source might result from the diffu-sion of O2 across the brine-seawater interface, fossil remnantsindicating that ambient seawater was oxic. If 500 Ma seawa-ter had a composition similar to that of the present day, thebrine pool-seawater interface would inevitably have alloweddiffusion, particularly of O2 downward and H2 and H2S up-ward (Fig. 12), because of the contrasting compositions ofseawater and Hellyer ore fluid. A semiquantitative estimate ofthe likely significance of gaseous diffusion of O2 can be gainedfrom the expression (from Cussler, 1984): mass flux =D/l(c1–c2), where D = diffusion coefficient, l = thickness offilm, and c1 and c2 are the concentrations of O2 in seawaterand brine, respectively. The fluids on either side of the sea-water-brine interface reasonably could be considered wellmixed, the seawater due to convection driven by loss of heatfrom the pool, and the brine by circulation due to plume ac-tivity (Fig. 12). Hence, if l = 1 cm, c1 and c2 are 5.10–3 g/l–1 andzero, and D = 2.1 × 10–5 cm–2 /s–1 at 25°C, then for the as-sumed area of the pool surface, the flux across the interface isapproximately 500 g/s–1. This flux would be incorporated intothe incoming supply from the vent, allowing some oxygen-re-quiring reactions to proceed, as discussed above. If Butlerand Rickard (2000) are correct, then the diffusive contribu-tion of O2 is not required (but would still be made).

Review

The brine pool hypothesis can account for the large sizeand metal content of the Hellyer ore compared to mostHokuroku deposits (Table 3) without having to call on in-creased thermal energy. Basins combined with negativelybuoyant fluids are much more efficient trapping systems thanthose related to positively buoyant fluids like black smokersystems; other factors being equal, the larger the basin andthe deeper the pool, the larger the deposit. Differences in theintensity and/or duration of zone refining compared toKuroko deposits can account for the relatively low Cu con-tent, as discussed for the Kuroko model. There would seemto be no reason why in other brine pool-derived ores the Cucontent could not be much higher due to mineral replace-ment in the lower parts of the massive sulfide. For Hellyer,the lack of barite with seawater sulfate signatures, the pyrite-arsenopyrite assemblage, the lack of Fe oxides and chimneyfragments, the absence of variation in δ34S values of the sul-fides, and the distribution of the aFeS in sphalerite all fit a hy-pothesis that excludes seawater during ore formation. Finally,the brine pool follows directly from the difference in salinitiesbetween Kuroko and Hellyer ore fluids, assuming the foot-wall fluids represent the ore fluids.

The sedimented sulfide would have provided an idealmedium for framboid growth, and framboids within veinsmay have grown within layers of Fe monosulfides quenchedonto vein walls, accounting for the frequency of occurrencecompared to Kuroko and black smoker-type deposits. Sedi-mented sulfides are relatively uncommon in modern blacksmoker systems. The conditions within the Hellyer sulfide

mud would not have been dissimilar to those of laboratory ex-periments studying crystallization from quenched sulfides orsulfides produced by fluid mixing. Of the other unusual tex-tural features noted in Hellyer ore, the banding common nearthe top of the orebody might reflect differential particle set-tling in the pool, and some of the relatively common veins andveinlets might have been initiated by shrinkage and settling ofthe aging sulfide mud.

Conclusions Many textures typical of the early stages of crystallization

are still preserved in Hellyer ore as a result of relatively minorsubsequent deformation and metamorphism. Early crystalgrowth appears to have been dominated by nucleation andgrowth of Fe monosufides and sulfides at centers more or lessevenly dispersed in a particulate sulfide medium. Many ofthese centers are framboids, others show fine radial and con-centric patterns. Their presence indicates uniform three-di-mensional growth, probably reflecting high rates of volumediffusion through the host medium. Locally, the nuclei havemerged to form large masses of pyrite with minor sphaleriteand galena. Sphalerite and galena also grew at dispersed cen-ters, though less commonly.

Bladed, fibrous, and dendritic textures in pyrite, and lesscommonly in sphalerite and galena, were also generated inthe earlier growth stages under conditions of high supersatu-ration and preserved due to later infill having a different com-position. Many examples appear to have formed in small-scaleconduits. Other centimeter-scale open spaces were filled bymicrobotryoidal, reniform pyrite, sphalerite, and galena.There is a possibility that all or some of the bladed and fibroustextures are due to replacement of sulfates but there is no ev-idence of partial replacement, and the barite content of theore is very low, despite the presence of Ba in the ore fluids.Relics of pyrrhotite suggest this may have been an early prod-uct of quenching. Some of the discontinuous, highly irregularveins and veinlets that are common in the Zn-Pb ore mayhave originated as a result of cracking during volume shrink-age and/or tectonic disturbance, and ore fluids rising from thefootwall through the cracks probably contributed to vein fill-ing, which was largely by open-space filling. Near the top ofthe orebody centimeter-scale pyritic bands alternate withsphalerite-galena-rich bands. The banding is precleavage andit may have resulted from winnowing and sedimentation.

Pyrite framboids up to 100 µm in diameter are typical ofthe Zn-Pb, barite, and silica ore types and indicate the pres-ence of metastable mackinawite and greigite as stepwise pre-cursors to framboid formation (Wilkin and Barnes, 1997,though see contrary arguments by Butler and Rickard, 2000).The framboids appear to have grown in situ within the fine-grained sulfide body and also in some veins. They are muchmore common than in modern black smoker and mostKuroko deposits.

Relict hexagonal crystal outlines indicate that a substantialpart of the sphalerite in the ore was derived from wurtzite,the inversion being another example of Ostwald’s step rules.Much of the abundant chalcopyrite occurring in growth zonesin the hexagons may have developed by exsolution of Cu andFe from early asymmetrical crystals. In the high Cu zone atthe top of the massive sulfide, and locally internally, there is


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evidence of replacement by chalcopyrite and of coprecipita-tion of sphalerite and chalcopyrite.

All the textures seen in the Hellyer deposit are present inmodern ocean-floor, black smoker-type, sulfide deposits andin Kuroko ores, particularly the bladed, dendritic, and fibroustypes, and probably resulted from high degrees of supersatu-ration induced by quenching of the ore fluids reaching the seafloor. However, framboids appear to be much more commonin the Hellyer ore, whereas open-space textures are less so. Inaddition, the abundant intricate vein and veinlet systems atHellyer appear to be more common than in modern andKuroko ores, as is the banding seen near the top of the mas-sive sulfide.

Although Large (1992) and McArthur (1996) consideredthe Hellyer deposit to have formed in much the same way asKuroko ores, the pyrite-arsenopyrite assemblage, scarcity ofbarite in the massive sulfide, absence of Fe oxides, lack ofchimney structures, and the sulfur isotope compositions indi-cate minor to negligible involvement of ambient seawaterduring massive sulfide formation on the sea floor. The pyrite-arsenopyrite assemblage indicates reduced ore fluids for theZn-Pb ore and maintenance of reduced conditions duringmineralization. More importantly, the highly saline inclusionsin the footwall veins, and the existence of a coeval basin,strongly indicate a derivation by quenching of fluids in a brinepool (Solomon and Khin Zaw, 1997). Sedimentation ofquenched sulfides in the pool would have formed a sulfidemud containing submicron and larger particles in which Ost-wald ripening would have proceeded, forming most of thetextures described. The brine pool would have provided idealconditions in which to grow framboidal pyrite and might haveallowed differential particle settling to form sulfide banding.Away from the main feeder pipe the upward flow of ore flu-ids from beneath the sulfide mud was probably limited com-pared to such flow in Kuroko deposits but was probably im-portant in precipitating Fe sulfides, sphalerite, and galenawithin minor conduits and veins and veinlets. Although manyof the Hellyer textures could be, and have been, described ascolloform, there is little evidence that either the brine pool orthe mud behaved as colloids for any length of time.

The weight of experimental and observational evidencepoints to weak oxidation of the reduced fluids in the brinepool in order to develop framboidal pyrite at likely pool tem-peratures. The presence in the shales overlying the ore de-posit of hemichordate dendroids having sessile benthic life-style shows that the entire water column was oxic, anddiffusion of O2 across the brine-seawater interface could haveallowed adequate oxidation. The absence of marcasite in whatwas probably a highly acid pool indicates that the oxidationpotential was too low to support marcasite.

AcknowledgmentsWe are grateful to Peter Williams of the University of West-

ern Sydney, Maria Ondina of the Centro de Crystalografia eMineralogia, Lisbon, Clive Burrett of the School of Earth Sci-ences, University of Tasmania, and Ross Large, Bruce Gem-mell, and Christian Schardt of the Centre for Ore DepositResearch, University of Tasmania, for many useful sugges-tions. Two anonymous reviewers also contributed to the formof the final manuscript. Gary McArthur and Chris Davies

(previous senior mine geologists at Hellyer) have been mosthelpful with data and comments. June Pongratz assisted withpreparation of the figures. We also thank David French ofCSIRO, Division of Exploration Geoscience, for arrangingthe loan of polished sections made for the study by Ramsdenet al. (1990). The research was supported by the AustralianResearch Council’s Research Centre’s program.May 3, 2000; May 18, 2001

REFERENCESAlmodóvar, G.R., Sáez, R., Pons, J.M., Maestre, A., Toscano, M., and Pas-

cual, E., 1998, Geology and genesis of the Aznalcóllar massive sulphide de-posits, Iberian pyrite belt, Spain: Mineralium Deposita, v. 33, p. 111–136.

Barton, P.B., Jr., 1978, Some ore textures involving sphalerite from the Furo-tobe mine, Akita Prefecture, Japan: Mining Geology, v. 28, p. 64–72.

Benning, L.G., Wilkin, R.T., and Barnes, H.L., 2000, Reaction pathways inthe Fe-S system below 100°C: Chemical Geology, v. 167, p. 25–51.

Berry, R.F., and Crawford, A.J., 1988, The tectonic significance of Cambrianallochthonous mafic-ultramafic complexes in Tasmania: Australian Journalof Earth Sciences, v. 35, p. 523–533.

Berry, W.B.N., 1987, Phylum Hemichordata (including Graptolithina), inBoardman, R.S., Cheetham, A.H., and Rowell, A.J., eds., Fossil inverte-brates: Oxford, Blackwell, p. 612–635,

Bischoff, J.L., 1969, Red Sea geothermal brine deposits: Their mineralogy,chemistry, and genesis, in Degens, E.T., and Ross, R.A., eds., Hot brinesand recent heavy metal deposits in the Red Sea: New York, Springer-Ver-lag, p. 368–401.

Black, L.P., Seymour, D.B., Corbett, K.D., Cox, S.E., Streit, J.E., Bottrill,R.S., Calver, C.R., Everard, J.L., Green, G.R., McLenaghan, M.P., Pem-berton, J., Taheri, J., and Turner, N. J., 1997, Dating Tasmania’s oldest ge-ological events: Tasmanian Geological Survey, Record 1997/15.

Boudreau, A.E., and McBirney, A.R., 1997, The Skaergaard layered series.Part III. Non-dynamic layering: Journal of Petrology, v. 38, p. 1003–1020.

Butler, I.B., and Rickard, D., 2000, Framboidal pyrite formation via the oxi-dation of iron (II) monosulfide by hydrogen sulfide: Geochimica et Cos-mochimica Acta, v. 64, p. 2665–2672.

Clarkson, E.N.K., 1986, Invertebrate palaeontology and evolution, 2nd ed.:London, Allen and Unwin, 382 p.

Converse, D.R., Holland, H.D., and Edmond, J.M., 1984, Flow rates in theaxial hot springs of the East Pacific Rise (21°N): Implications for the heatbudget and the formation of massive sulfide deposits: Earth and PlanetaryScience Letters, v. 69, p. 159–175.

Corbett, K.D., 1992, Stratigraphic-volcanic setting of massive sulfide de-posits in the Cambrian Mount Read Volcanics, Tasmania: ECONOMIC GE-OLOGY, v. 87, p. 564–586.

Corbett, K.D., and Turner, N.J., 1989, Early Palaeozoic deformation and tec-tonics: Geological Society of Australia Special Publication 15, p. 154–181.

Crawford, A.J., Corbett, K.D., and Everard, J.L., 1992, Geochemistry of theCambrian volcanic-hosted massive sulfide-rich Mount Read Volcanics, Tas-mania, and some tectonic implications: ECONOMIC GEOLOGY, v. 87, p.597–619.

Cussler, E.L., 1984, Diffusion mass transfer in fluid systems: Cambridge,University of Cambridge Press, 525 p.

Davis, E.E., Goodfellow, W.D., Bornhold, B.D., Adshead, L., Blaise, B.,Villinger, H., and Cheminant, G.M., 1987, Massive sulfides in a sedimentedrift valley, northern Juan de Fuca Ridge: Earth and Planetary Sciences, v.82, p. 49–60.

Downs, R.C., 1993, Syn-depositional fault controls on the Hellyer volcanic-hosted massive sulphide deposit: Unpublished M.Sc. thesis, Hobart, Uni-versity of Tasmania, 63 p.

Eldridge, C.S., Barton, P.B., Jr., and Ohmoto, H., 1983, Mineral textures andtheir bearing on formation of the Kuroko orebodies: ECONOMIC GEOLOGYMONOGRAPH 5, p. 241–281.

Feely, R.A., Lewison, M., Massoth, G.J., Robart-Baldo, G., Lavelle, J.W.,Byrne, R.H., von Damm, K.L., and Curl, H.C., 1987, Composition and dis-solution of black smoker particulates from active vents on the Juan de FucaRidge: Journal of Geophysical Research, v. 92, p. 11,347–11,363.

Feely, R.A., Massoth, G.J., Trefry, J.H., Baker, E.T., Paulson, A.J., and Lebon,G.T., 1994, Composition and sedimentation of hydrothermal plumes parti-cles from North Cleft Segment, Juan de Fuca Ridge: Journal of Geophysi-cal Research, v. 99, p. 4985–5006.


0361-0128/98/000/000-00 $6.00 1532

Furukawa, Y., and Barnes, H.L., 1995, Reactions forming pyrite from pre-cipitated amorphous ferrous sulfide, in Vairavamurthy, M.A., and Schoo-nen, M.A.A., eds., Geochemical transformations of sedimentary sulfur:Washington, DC, American Chemical Society, p. 194–205.

Gee, C.E., 1970, The geochemistry of some black shales in relation to the ori-gin of certain stratiform orebodies in Tasmania: Unpublished Ph.D. thesis,Hobart, University of Tasmania, 153 p.

Gemmell, J.B., and Fulton, R., 2001, Geology, genesis, and exploration ap-plications of the footwall and hanging-wall alteration associated with theHellyer volcanic-hosted massive sulfide deposit, Tasmania, Australia: ECO-NOMIC GEOLOGY, v. 96, p. 1003–1035.

Gemmell, J.B., and Large, R.R., 1992. Stringer system and alteration zonesunderlying the Hellyer volcanogenic massive sulfide deposit, Tasmania,Australia: ECONOMIC GEOLOGY, v. 87, p. 620–649.

——1993, Evolution of a VHMS hydrothermal system, Hellyer deposit, Tas-mania, Australia: Sulfur isotope evidence: Resource Geology Special Issue17, p. 108–119.

German, C.R., and Sparks, R.S.J., 1993, Particle recycling in the TAG hydro-thermal plume: Earth and Planetary Science Letters, v. 116, p. 129–134.

Goldfarb, M.S., Converse, D.R., Holland, H.D., and Edmond, J.M., 1983,The genesis of hot spring deposits on the East Pacific Rise, 21°N: ECO-NOMIC GEOLOGY MONOGRAPH 5, p. 184–197.

Halbach, P., Pracejus, B., and Märten, A., 1993, Geology and mineralogy ofmassive sulfide ores from the central Okinawa trough, Japan: ECONOMICGEOLOGY, v. 88, p. 2206–2221.

Hannington, M.D., and Scott, S.D., 1988, Mineralogy and geochemistry of ahydrothermal silica-sulfide-sulfate spire in the caldera of Axial Seamount,Juan de Fuca Ridge: Canadian Mineralogist, v. 26, p. 603–625.

Heinrich, C.A., and Eadington, P.J., 1986, Thermodynamic predictions of thehydrothermal chemistry of arsenic, and their significance for the parage-netic sequence of some cassiterite-arsenopyrite base metal sulfide deposits:ECONOMIC GEOLOGY, v. 81, p. 511–529.

Henley, R.W., 1984, pH calculations for hydrothermal fluids: Reviews in Eco-nomic Geology, v. 1, p. 83–98.

Ishikawa, Y., and Yanagisawa, Y., 1974. Geology of the Ainai mine, with spe-cial reference to syngenetic origin of the Daikoku deposits: Mining Geol-ogy Special Issue 6, p. 79–87.

Ito, T., Takahashi, T., and Omori, Y., 1974, Submarine volcanic-sedimentaryfeatures in the Matsumine Kuroko deposits, Hanaoka mine, Japan: MiningGeology Special Issue 6, p. 115–130.

Jack, D.J., 1989, Hellyer host rock alteration: Unpublished M.Sc. thesis, Ho-bart, University of Tasmania, 182 p.

Jago, J.B., 1979, Tasmanian Cambrian biostratigraphy—a preliminary report:Australian Journal of Earth Sciences, v. 26, p. 223–230.

Kalogeropoulis, S.I., and Scott, S.D., 1983, Mineralogy and geochemistry oftuffaceous exhalites (tetsusekiei) of the Fukazawa mine, Hokuroku district,Japan: ECONOMIC GEOLOGY MONOGRAPH 5, p. 412–432.

Khin Zaw, Gemmell, J.B., Large, R.R., Mernagh, T.P., and Ryan, C.G., 1996,Evolution and source of ore fluids in the stringer system, Hellyer VHMSdeposit, Tasmania, Australia: Evidence from fluid inclusion microther-mometry and geochemistry: Ore Geology Reviews, v. 10, p. 251–278.

Kojima, S., 1990, A coprecipitation experiment on intimate association ofsphalerite and chalcopyrite and its bearings on the genesis of Kuroko ores:Mining Geology, v. 40, p. 147–158.

Kojima, S., and Ohmoto, H.L., 1991, Hydrothermal synthesis of wurtzite andsphalerite at T = 350–250°C: Mining Geology, v. 41, p. 313–327.

Kojima, S., and Sugaki, A., 1985, Phase relations in the Cu-Fe-Zn-S systembetween 500° and 300°C under hydrothermal conditions: ECONOMIC GE-OLOGY, v. 80, p. 158–171.

Kojima, S., Nagase, T., and Inoue, T., 1995, A coprecipitation experiment onthe chalcopyrite disease texture involving Fe-bearing sphalerite: Journal ofMineralogy, Petrology and Economic Geology, v. 90, p. 261–267.

Large, R.R., 1992, Australian volcanic-hosted massive sulfide deposits: Fea-tures, styles, and genetic models: ECONOMIC GEOLOGY, v. 87, p. 471–510.

Large, R.R., Bull, S.W., and McGoldrick, P.J., 2000, Lithogeochemical halosand geochemical vectors to stratiform sediment hosted Zn-Pb-Ag deposits,Part 2. HYC deposit, McArthur River, Northern Territory: Journal of Geo-chemical Exploration, v. 68, p. 105–126.

Lebedev, L.M., 1967, Metacolloids in endogenic deposits: New York,Plenum Press, 298 p.

Lee, M.S., Miyajima, T., and Mizumoto, H., 1974, Geology of the Kamikitamine, Aomori Prefecture, with special reference to genesis of fragmentalores: Mining Geology Special Issue 6, p. 53–66.

Lefaucheux, F., and Robert, M.C., 1994, Crystal growth in gels, in Hurle,D.T.J., ed., Handbook of crystal growth: Amsterdam, Elsevier, v. 2, p.1271–1303.

Lennie, A.R., England, K.E.R., and Vaughan, D.J., 1995, Transformation ofsynthetic mackinawite to hexagonal pyrrhotite: A kinetic study: AmericanMineralogist, v. 80, p. 960–967.

Leventhal, J.S., 1983, An interpretation of carbon and sulfur relationships inBlack Sea sediments as indicators of environments of deposition: Geochim-ica et Cosmochimica Acta, v. 47, p. 133–137.

Liesegang, R.E., 1911, Über die Reifung von Silberhaloidemulsionen:Zeitschrift fur Physikalische Chemie, v. 75, p. 374–377.

Lydon, J.W., 1988, Ore deposit models 14, volcanogenic massive sulfide de-posits, Part 2: Genetic models: Geoscience Canada, v. 15, p. 43–65.

Marchig, V., 1991, Hydrothermale Aktivität am Ostpazifischen Rücken: Ge-ologisches Jahrbuch Reihe D, Heft 93, p. 125–197.

Marchig, V., Gundlach, H., Holler, G., and Wilke, M., 1988, New discoveriesof massive sulfides on the East Pacific Rise: Marine Geology, v. 84, p.179–190.

McArthur, G.J., 1989, Hellyer: Geological Society of Australia Special Publi-cation 15, p. 144–148.

——1996, Textural evolution of the Hellyer massive sulphide deposit: Un-published Ph.D. thesis, Hobart, University of Tasmania, 272 p.

McArthur, G.J., and Dronseika, E.V., 1990, Que River and Hellyer zinc-lead-silver deposits: Australasian Institute of Mining and Metallurgy Monograph14, p. 1229–1239.

McDougall, T.J., 1984, Fluid dynamic implications for massive sulphide de-posits of hot saline fluid flowing into a submarine depression from below:Deep-Sea Research, v. 31, p. 145–170.

McDuff, R.E., 1995, Physical dynamics of deep-sea hydrothermal plumes:Geophysical Monograph 91, p. 357–368.

McPhie, J., and Allen, R.L., 1992, Facies architecture of mineralized subma-rine volcanic sequences: Cambrian Mount Read Volcanics, western Tasma-nia: ECONOMIC GEOLOGY, v. 87, p. 587–596.

Morse, J.W., and Casey, W.H., 1988, Ostwald processes and mineral parage-nesis in sediments: American Journal of Science, v. 288, p. 537–560.

Mottl, M.J., and McConachy, T.F., 1990, Chemical processes in buoyantplumes on the East Pacific Rise near 21°N: Geochimica et CosmochimicaActa, v. 54, p. 1911–1927.

Mullin, J.W., 1993, Crystallisation, 3rd ed.: Oxford, England, Butterworthand Heinemann, 527 p.

Murowchik, J.B., and Barnes, H.L., 1986, Marcasite precipitation from hydro-thermal solutions: Geochimica et Cosmochimica Acta, v. 50, p. 2615–2629.

Nelsen, T.A., Klinkhammer, G., and Trefry, J., 1986, Real-time observationand tracking of dispersed hydrothermal plumes using nephelometry: Ex-amples from the Mid-Atlantic Ridge: Earth and Planetary Science Letters,v. 81, p. 245–252.

Offler, R., and Whitford, D.J., 1992, Wall-rock alteration and metamorphismof a volcanogenic massive sulfide deposit at Que River, Tasmania: Petrologyand mineralogy: ECONOMIC GEOLOGY, v. 87, p. 686–705.

Ohmoto, H., 1996, Formation of volcanogenic massive sulfide deposits: theKuroko perspective: Ore Geology Reviews, v. 10, p. 135–177.

Ohmoto, H., Mizukami, M., Drummond, S.E., Eldridge, C.S., Pisutha-Arnond, V., and Lenagh, T.C., 1983, Chemical processes of Kuroko forma-tion: ECONOMIC GEOLOGY MONOGRAPH 5, p. 570–604.

Ostwald, W.Z., 1897, Studien über die Bildung und Umwaldung fester Kor-per. I. Abhandlung: Übersattigung and Überkaltung: Zeitschrift furPhysikalische Chemie, v. 22, p. 289–330.

Paradis, S., Jonasson, I.R., Le Cheminant, G.M., and Watkinson, D.H., 1988,Two zinc-rich chimneys from the southern Juan de Fuca Ridge: CanadianMineralogist, v. 26, p. 637–654.

Peter, J.M., and Scott, S.D., 1988, Mineralogy, composition, and fluid-inclu-sion microthermometry of seafloor hydrothermal deposits in the southerntrough of the Guayamas basin, Gulf of California: Canadian Mineralogist,v. 26, p. 567–587.

Pisutha-Arnond, V., and Ohmoto, H., 1983, Thermal history, and chemicaland isotopic compositions of the ore-forming fluids responsible for theKuroko massive sulfide deposits in the Hokuroku district of Japan: ECO-NOMIC GEOLOGY MONOGRAPH 5, p. 523–558.

Pottorf, R.J., and Barnes, H.L., 1983, Mineralogy, geochemistry, and ore gen-esis of hydrothermal sediments from the Atlantis II Deep, Red Sea: ECO-NOMIC GEOLOGY MONOGRAPH 5, p. 198–223.

Quilty, P.G., 1971, Cambrian and Ordovician hydroids and dendroids of Tas-mania: Australian Journal of Earth Sciences, v. 17, p. 171–189.


0361-0128/98/000/000-00 $6.00 1533

Raiswell, R., and Berner, R.A., 1986. Pyrite and organic matter in Phanero-zoic normal marine shales: Geochimica et Cosmochimica Acta, v. 50, p.1967–1976.

Ramsden, A.R., Kinealy, K.M., Creelman, R.A., and French, D.H., 1990,Precious and base metal mineralogy of the Hellyer volcanogenic massivesulphide deposit, N. W. Tasmania—a study by electron microprobe, inGray, P.M.S., Castle, J.F., Vaughan, D.J., and Warner, N.A., eds., Sulphidedeposits—their origin and processing: London, Institute of Mining andMetallurgy, p. 49–71,

Rimstidt, J.D., 1997, Gangue mineral transport and deposition, in Barnes,H.L., ed., Geochemistry of hydrothermal ore deposits, 3rd ed.: New York,Wiley, p. 487–515.

Rimstidt, J.D., and Barnes, H.L., 1980, The kinetics of silica-water reactions:Geochimica et Cosmochimica Acta, v. 44, p. 1683–1699.

Sasaki, A., 1974, Isotopic data of Kuroko deposits: Mining Geology SpecialIssue 6, p. 389–397.

Sato, J., 1974, Ores and minerals from the Shakanai mine, Akita Prefecture,Japan: Mining Geology Special Issue 6, p. 323–335.

Sawlowicz, Z., 1993, Pyrite framboids and their development: A new con-ceptual mechanism: Geologische Rundschau, v. 82, p. 148–156.

Schoonen, M.A.A., and Barnes, H.L., 1991a, Reactions forming pyrite andmarcasite from solution: I. Nucleation of FeS2 below 100°C: Geochimica etCosmochimica Acta, v. 55, p. 1495–1504.

——1991b, Reactions forming pyrite and marcasite from solution: II. ViaFeS precursors below 100°C: Geochimica et Cosmochimica Acta, v. 55, p.1505–1514.

——1991c, Mechanisms of pyrite and marcasite formation from solution: III.Hydrothermal processes: Geochimica et Cosmochimica Acta, v. 55, p.3491–3504.

Seward, T.M., and Barnes, H.L., 1997, Metal transport by hydrothermal orefluids, in Barnes, H.L., ed., Geochemistry of hydrothermal ore deposits, 3rd

ed.: New York, Wiley, p. 435–486,Sharpe, R., 1991, The Hellyer baritic and siliceous caps: Unpublished B.Sc.

Honors thesis, Hobart, University of Tasmania, 114 p.Shimazaki, Y., 1974, Ore minerals of the Kuroko-type deposits: Mining Ge-

ology Special Issue 6, p. 311–322.Sinclair, B.J., 1994. Geology and geochemistry of the Que River Shale, west-

ern Tasmania: Unpublished B.Sc. Honors thesis, Hobart, University of Tas-mania, 106 p.

Solomon, M., and Groves, D.I., 2000, The geology and origin of Australia’smineral deposits, 2nd ed.: Hobart and Perth, Australia, Centre for Ore De-posit Research and Centre for Global Metallogeny, 1002 p.

Solomon, M., and Khin Zaw, 1997, Formation on the sea floor of the Hellyervolcanogenic massive sulfide deposit: ECONOMIC GEOLOGY, v. 92, p.686–695.

Solomon, M., and Walshe, J.L., 1979, The formation of massive sulfide de-posits on the sea floor: ECONOMIC GEOLOGY, v. 74, p. 797–813.

Solomon, M., Vokes, F., and Walshe, J.L., 1987, Chemical remobilization ofvolcanic-hosted massive sulphide deposits at Rosebery and Mt. Lyell, Tas-mania: Ore Geology Reviews, v. 2, p. 173–190.

Solomon, M., Eastoe, C.J., Walshe, J.L., and Green, G.R., 1988, Mineral de-posits and sulfur isotope abundances in the Mount Read Volcanics betweenQue River and Mount Darwin: ECONOMIC GEOLOGY, v. 83, p. 1307–1328.

Stokes, G.G., 1851, On the effect of the internal friction on the motion ofpendulums: Transactions of the Cambridge Philosophical Society, v. 9, p.8–106.

Sweeney, R.E., and Kaplan, I.R., 1973, Pyrite framboid formation: Labora-tory synthesis and marine sediments: ECONOMIC GEOLOGY, v. 68, p.618–634.

Tanimura, S., Date, J., Takahashi, T., and Ohmoto, H., 1983, Geologic settingof the Kuroko deposits, Japan. Part II. Stratigraphy and structure of theHokuroku district: ECONOMIC GEOLOGY MONOGRAPH 5, p. 24–38.

Tornos, F., Clavijo, E.G., and Spiro, B.F., 1998, The Filon Norte orebody(Tharsis, Iberian pyrite belt): A proximal low-temperature shale-hostedmassive sulphide in a thin-skinned tectonic belt: Mineralium Deposita, v.33, p. 150–169.

Urabe, T., 1974, Iron content of sphalerite coexisting with pyrite from someKuroko deposits: Mining Geology Special Issue 6, p. 377–384.

Watanabe, M., 1974, On the textures of ores from the Daikoku ore deposit,Ainai mine, Akita Prefecture, northeast Japan, and their implications in theore genesis: Mining Geology Special Issue 6, p. 337–348.

Waters, J.C., and Wallace, D.B., 1992, Volcanology and sedimentology of thehost succession to the Hellyer and Que River volcanic-hosted massivesulfide deposits, northwestern Tasmania: ECONOMIC GEOLOGY, v. 87, p.650–666.

Whitford, D.J., McPherson, W.P.A., and Wallace, D.B., 1989, Geochemistryof the host rocks of the volcanogenic massive sulfide deposit at Que River,Tasmania: ECONOMIC GEOLOGY, v. 84, p. 1–21.

Whitford, D.J., Korsch, M.J., and Solomon, M., 1992, Strontium isotopestudies of barites: Implications for the origin of base metal mineralizationin Tasmania: ECONOMIC GEOLOGY, v. 87, p. 953–959.

Wilkin, R.T., and Barnes, H.L., 1996, Pyrite formation by reactions of ironmonosulfides with dissolved inorganic and organic sulfur species:Geochimica et Cosmochimica Acta, v. 60, p. 4167–4179.

——1997, Formation processes of framboidal pyrite: Geochimica et Cos-mochimica Acta, v. 61, p. 323–339.

Williams, E., McLenaghan, M.P., and Collins, P.L.F., 1989, Mid-Palaeozoicdeformation, granitoids and ore deposits: Geological Society of AustraliaSpecial Publication 15, p. 238–292.


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