Origins of pyrites in the ∼2.5 Ga Mt. McRae Shale, the Hamersley District, Western Australia

Geochimica et Cosmochimica Acta, Vol. 62, No. 19/20, pp. 3205–3220, 1999 Copyright © 1998 Elsevier Science Ltd Printed in the USA. All rights reserved 0016-7037/99 $19.00 1 .00

Pergamon

PII S0016-7037(98)00229-4

Origins of pyrites in the ;2.5 Ga Mt. McRae Shale, the Hamersley District, Western Australia TAKESHI KAKEGAWA,1,2,* HAJIME KAWAI,1 and HIROSHI OHMOTO2 2

1 Tohoku University, Graduate School of Science, Sendai 980-77, Japan The Pennsylvania State University, Department of Geosciences, University Park, Pennsylvania 16802, USA

(Received October 15, 1997; accepted in revised form June 23, 1998)

Abstract—The Mt. McRae Shale is the footwall rock unit of the Brockman Iron Formation (;2.5 Ga in age) in the Hamersley Basin, Australia. It is characterized by a high concentration of organic carbon (;2– 8 wt%), an abundance of disseminated pyrite (;1 to ;10 wt% S in the bulk rocks), and abundant pyrite nodules (;1–10 cm radius). We have examined microscale (;200 mm to 1 cm) variations in sulfur isotopic compositions of pyrite and S and C contents in six rock samples from a ;27 m drill core section of the base of the Mt. McRae Shale at the Whaleback Mine. The microanalyses of sulfur isotope compositions were performed in situ on single or aggregates of pyrite crystals (eighty-four analyses) using a Nd-YAG laser microprobe. The carbon contents of thirty powdered samples, drilled from different parts of the six rock samples, are similar (;2 to ;8 wt% C) with the exception of one sample (0.5 wt%). However, the occurrence, morphology, abundance, and d34S value of pyrite reveal distinct differences between the two samples from the upper part and the four samples from the lower part of the drill core section. Pyrite crystals in the upper part occur mostly as disseminated fine grains (;5 mm). The pyrite S contents are uniform within each sample, but the d34S values vary from 26.3 to 17.1‰. These data suggest that pyrite crystals in the upper section were formed by bacterial reduction of seawater sulfate. Pyrite crystals of the lower section occur in the form of veinlets, nodules, or laminae of coarse grains (;200 mm): the pyrite contents are highly variable within each specimen (0 to .10 wt%), and the d34S values vary from 12.2 to 111.8‰. The formation process of pyrites in the lower section appears to have been complicated: pyrite laminae were first formed by bacterial sulfate reduction during early diagenesis, and then some of the early pyrites were dissolved and reprecipitated to form pyrite nodules by later diagenetic or hydrothermal solutions in a closed-system. The sulfur isotope data obtained in this study can be best explained by a model postulating that the seawater about 2.5 Ga ago in the Hamersley Basin had the d34S value of 110 to 115‰ and that the kinetic isotope effects accompanying bacterial sulfate reduction were 8 –13‰ and 16 –21‰, respectively, during the deposition of the lower and upper sections of the McRae Shale. The variable d34S values of microscale area and the magnitudes of the kinetic isotope effects suggest that: (1) the sulfate concentration of the 2.5 Ga seawater was already more than one-third of the present seawater value, and (2) the activity of sulfate-reducing bacteria in the 2.5 Ga ocean was generally higher than that in the modern ocean. Suggestion 1 further implies that, by 2.5 Ga, (3) the atmosphere became oxic, and (4) the chemical and isotopic characteristics of sulfur in the earth’s near surface reservoirs (oceans and sediments) were controlled by the Phanerozoic-style biogeochemical cycles, rather than by the mantle-buffer (i.e., magmatic) mechanism. Suggestion 2 may imply that (5) the Archean oceans were generally warmer compared to the modern oceans, and (6) they produced a higher abundance of organic matter that were easily metabolized by sulfate-reducing bacteria. Copyright © 1998 Elsevier Science Ltd Previous analyses of sulfur isotope compositions of pyrite in sedimentary rocks have been made on bulk rock sulfides. A measured d34S value typically represented the average d34S value of more than 100 pyrite grains of ,1 mm in size that occurred in a hand specimen sized rock. This is because more than 10 micromoles of sulfur were required in a conventional sulfur isotope analysis. Using newly developed laser microprobe techniques that allow d34S analysis of small (,100 mm) individual grains of pyrite, Ohmoto et al. (1993) were able to detect as much as 10‰ variations in the d34S values among the small pyrite crystals occurring in ;2 cm3 samples of ;3.4 Ga old shales and cherts from the Barberton Greenstone Belt, South Africa. On the basis of these data, they suggested that these pyrites were formed by sulfate-reducing bacteria and that the ;3.4 Ga ocean already contained more than 10% of the present concentration of sulfate with a d34S value of approximately 12‰.

1. INTRODUCTION

Pyrite formation in modern marine sediments is initiated by the action of sulfate-reducing bacteria (e.g., Desulfovibrio desulfuricans; cf., Odom and Singleton, 1993). However, whether or not sulfate was abundant and sulfate-reducing bacteria were active in Archean oceans is still debated. Most previous researchers have argued that pyrites in Archean sedimentary rocks were formed by inorganic processes because their sulfur isotope compositions are less fractionated than their modern counterparts (e.g., Cameron, 1982; Hattori and Cameron, 1986; Canfield and Teske, 1996). In particular, Veizer et al. (1989) suggested that mantle flux (i.e., magmatic activity) controlled the chemical and isotopic characteristics of ocean water and marine sediments during the Archean. *Author to whom correspondence should be addressed (kakegawa @mail.cc.tohoku.ac.jp). 3205

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Fig. 1. Geological map of the Hamersley Basin (after Barley et al., 1992).

The Mt. McRae Shale in the Hamersley Basin, Australia is about 2.5 Ga in age, and it contains abundant organic carbon (up to 10 wt%), indicating that biological activity was high in the ocean during deposition of the shale. Compared to other Archean shales, the Mt. McRae Shale is less silicified and less metamorphosed. This shale has also been studied from a variety of perspectives by a research group at Tohoku University, including the inorganic and organic chemistry (Naraoka et al., in prep.), Rb-Sr and Nd-Sm isotopes (Uyeda, 1993), and oxygen isotopes and fluid inclusions (Haruna and Ohmoto, 1992). Therefore, the Mt. McRae Shale is well suited for examination if the pyrites, occurring in various morphologies, were formed by sulfate-reducing bacteria. If so, this shale is also suitable for development of constraints on the content and d34S value of sulfate in seawater about 2.5 Ga ago. In the present study, in situ analyses of sulfur isotopes using a laser microprobe were performed on different types of pyrite from the Mt. McRae Shale. Sulfur and carbon contents were also determined for microsized samples separated using a dental drill in order to examine the variation of these elements on a millimeter scale. 2. GEOLOGY

The Hamersley Basin is dominated by the Mt. Bruce Supergroup (or the Mt. Bruce Megasequence Set defined by Blake and Barley, 1992) which is traditionally divided into three lithostratigraphic groups: the Fortescue Group, the Hamersley Group, and the Turee Creek Group (Trendall, 1979; see Fig. 1). The Mt. Bruce Megasequence Set was deposited in the Hamersley Basin from ;2.77 to 2.44 Ga (Arndt et al., 1991; Trendall et al., 1990; Pidgeon and Horwitz, 1991; Barley et al., 1997). The tectonic evolution of the Mt. Bruce Megasequence Set has been discussed by previous investigators (e.g., Blake and Barley, 1992; Barley et al., 1992). One of the important features of the Hamersley Group is the occurrence of large volumes of banded iron formation (BIF): the Marra Mamba, the Brockman, and the Boolgeeda Iron

Fig. 2. Generalized stratigraphic succession of the Hamersley Group.

Formations (Fig. 2). This group also contains thick formations comprised mostly of carbonates (the Wittenoom Formation; Simonson et al., 1993), shales (the Mt. McRae Shale and the Sylvia Formation), and volcanic rocks (the Woongarra Volcanics). These BIFs and other formations are widespread throughout the Hamersley Basin (Fig. 1). The Mt. McRae Shale is directly overlain by the Brockman Iron Formation and predominantly comprised of carbonaceous and pyritic shales with interbedded minor cherts. The average thickness of this shale is 50 m (Morris, 1993). The age of the shale can only be indirectly estimated. A U-Pb zircon age of 2470 6 4 Ma has been reported for the S13 macroband of the Dales Gorge Member of the Brockman Iron Formation based on SHRIMP analyses on separated zircon crystals (Trendall et al., 1990). A U-Pb zircon age of 2603 6 7 Ma has been obtained for a crystal-rich tuff in the upper part of the Wittenoom Formation (Hassler, 1993). Therefore, the depositional age of the Mt. McRae Shale (and the Mt. Sylvia Formation) is

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493-6), fine grains (;10 mm) of subhedral and euhedral pyrite are the dominant crystal forms (Fig. 4a,b). These fine grains of pyrite, forming laminae, are concordant with the sedimentary bedding. Some of the pyrite in the upper part shows concordant microveinlets or lenses composed of several crystals of subhedral coarse-grained pyrite (;100 mm; cf. Fig. 4b). Coarse grains (.100 mm) of pyrite are the dominant morphology in the lower section (Fig. 4d) and locally form concordant laminae. In the lower part of the drill core, pyrite nodules (;1–10 cm in diameter) appear in great quantities (Fig. 4c). A small amount of disseminated anhedral pyrite is also observed in the lower section (Fig. 4e). Pyrite veinlets crosscutting sedimentary bedding developed in the lowermost sample from the Mt. Sylvia Formation (Fig. 4f). Fine grains of quartz and clay minerals (illite) associated with abundant organic matter are observed throughout the drill core. Chlorite and sericite are also abundant in the lower part of the Mt. McRae Shale. Carbonate minerals were not observed under the microscope, with the exception of the sample (49328) from the Mt. Sylvia Formation. The absence of biotite (typical low grade metamorphic mineral in sedimentary rocks) and the abundance of clay minerals indicate zeolite facies to lower greenschist metamorphic grade for these samples. Fig. 3. Stratigraphy of a part of drill core No. 493. The core section encompasses the lowermost Mt. McRae Shale and the uppermost part of the Mt. Sylvia Formation. The top of the drill core is set as 0 m. A chert layer occurs at about 6 m below the surface. Below this chert, pyrite nodules are more abundant. Bruno’s Band, comprised of massive hematite, occurs at the top of the Mt. Sylvia Formation. For analyses, five samples from the Mt. McRae Shale and one sample from the Mt. Sylvia Formation were selected. py 5 pyrite.

bracketed between 2603 Ma and 2470 Ma. In the Mt. Newman area, the thickness of the Mt. McRae Shale is approximately 60 m. The metamorphic grade of the Hamersley Group is reported to be subgreenschist facies (Trendall, 1979). 3. SAMPLES

4. ANALYSES

Samples used in the present study contain various types of pyrite, some of which show a range of morphologies in the same thin section (e.g., Fig. 4b). The primary interest of the present study is to determine the d34S values of each type of pyrite and to discuss their origins. It is difficult, however, to physically separate one type of pyrite from the others. In addition, the conventional Cu2O combustion method to prepare SO2 gas for sulfur isotope analysis (Robinson and Kusakabe, 1975) requires a concentrate of approximately 3 mg of pyrite. This is difficult to achieve if the pyrite is disseminated in very fine grains. Because the laser microprobe method can accommodate much smaller amounts of pyrite, we elected to use the laser microprobe system at Tohoku University to determine the d34S values for each type of pyrite.

3.1. Stratigraphy of the Samples

A 35 m drill core from the Mt. Whaleback Mine of the Mt. Newman area (Fig. 3) was examined in the present study. The Mt. Whaleback mine hosts one of the largest iron ore deposits in Australia with an approximate ore reserve of 1.5 billion tons of high grade hematitic iron ores. The drill core used in the present study penetrated strata ranging from the lower Mt. McRae Shale to the upper Mt. Sylvia Formation. Pyrite nodules occur in the middle 20 m section between the upper chert/shale zone and Bruno’s Band (a hematitic iron bed). The hand specimens were selected from the upper (two samples) and lower parts (three samples) of the Mt. McRae Shale (Fig. 3). The core material of the Mt. Sylvia Formation ranges from siltstone to shale. Bruno’s Band occurs at the top part of the Mt. Sylvia Formation. A sample (493-28) of the Mt. Sylvia Formation was also examined. 3.2. Petrography

Various types of pyrite are observed in the six samples. In the upper part of the drill core section (samples 493-1 and

4.1. Sulfur Isotope Analyses

The Nd-YAG laser system used at Tohoku University (Fig. 5) consists of: (a) a Nd-YAG laser microprobe unit, (b) a sample chamber, (c) an optical system (binocular microscope and TV monitor), and (d) an oxygen flask. The selected samples were cut into rock billets with a volume of ;2 cm3 and then placed in the sample chamber. The laser combustion was carried out using a continuous-wave (C.W.) Nd-YAG laser in an O2 atmosphere (8 torr of PO2). The diameter of the laser beam can be changed from the multiple mode (;150 mm in laser diameter) to the TEM00 mode (;30 mm in laser diameter) using an adjustable aperture. Only the TEM00 mode was adopted for laser microprobe analyses because amounts of pyrite in the examined samples are smaller than the laser diameter in the multiple mode. The Nd-YAG laser was fired on a sample for 5 s. The typical output power used in this study was ;5 watts. The laser beam typically created a pit ;100 mm in diameter and depth on the surface of the pyrite sample. When the pyrite grain sizes were

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smaller than 100 mm, the laser beam also combusted the surrounding rock matrix. In addition to SO2, variable amounts of H2O, CO2, and various hydrocarbon gases are usually generated during laser microprobe analyses of sedimentary rocks. The SO2 was purified before introduction to the mass spectrometer through the following procedures. After laser combustion, the sample chamber was opened to cryogenic traps 1 and 2 (Fig. 5). The H2O was condensed in traps 1 and 2 cooled by an acetone slush (T ' 280°C). A platinum furnace, located between traps 1 and 2, was heated to 850°C to convert hydrocarbons to CO2 and H2O. Another cryogenic trap (trap 3), located next to the platinum furnace, condensed SO2 and CO2 gases using liquid nitrogen (T 5 2195°C). After freezing, the non-condensable gases (mostly O2) were pumped away. Next, trap 3 was isolated from trap 2 and the vacuum pump system, and CO2 was pumped out as trap 3 was cooled by a liquid nitrogen/pentane slush (T ' 2120°C). The purified SO2 gas was transferred from the cryogenic pentane trap of the purification line to the cold finger of the micro-inlet line of the mass spectrometer by cooling the cold finger with liquid nitrogen (Fig. 5). The cold finger was then isolated and heated to 40°C. The micro-inlet line was heated constantly at 40°C to avoid absorption of SO2 on the metal surface. A Finnigan MAT252 mass spectrometer analyzed the abundance ratio of masses 66 and 64. These ratios were converted to the d34S value, which are reported with respect to the CDT (Can˜on Diablo troilite) standard. The reproducibility of the above analyses was estimated to be 60.2‰ from replicate analyses of several pyrite standards. In general, d34S values by the laser microprobe are lower than those by the conventional method by 1.0 6 0.2‰, probably due to the oxygen isotope fractionation between O2 and SO2 during the combustion of pyrite (Kakegawa, 1993). This 1‰ fractionation factor appears to be constant for all types of pyrite crystals. The d34S values shown in the succeeding figures are those corrected by adding 1.0‰ to the measured values. 4.2. Analyses of Sulfur and Carbon Contents

Sample powders for the analyses of S and C contents were prepared from the same rock chips used for laser microprobe analyses. A dental drill was used in order to investigate variations in S and C contents on a millimeter scale. Typically, a dental drill creates a hole with an area of ;2 mm2. The laser spots were immediately adjacent to the drilled holes. Analyses of the S and C contents were performed on 10 mg of powdered sample using a Carlo Erba Elemental Analyzer. In addition to the microanalyses of S and C contents performed in the present study, bulk analyses of S and C contents of the same drill core samples have been carried out by Naraoka et al. (in prep.) as a part of a separate study. 5. RESULTS

Fig. 4. Photomicrographs showing the occurrences of pyrite in the Mt. McRae Shale. (a) pyrite laminae comprised of fine-grained pyrite (;10 mm in diameter). (b) pyrite laminae with coarse-grained pyrites. Both types of pyrites occur concordantly with the sedimentary bedding. The coarse-grained pyrites were formed as a result of recrystallization of the pyrite laminae (see text). (c) part of a pyrite nodule. (d) coarsegrained disseminated pyrite (;100 mm) in the lower section. (e) disseminated anhedral pyrite. (f) microveinlet of pyrite crosscutting sedimentary laminae.

5.1. Sulfur Isotope Compositions 5.1.1. Isotopic similarities among different types of pyrite

Laser microprobe analyses were performed on eighty-four spots on six rock billets. Each spot is an area of ;100 mm radius that represents either a part of a larger individual pyrite crystal or an area of shale containing fine-grained disseminated

pyrite crystals. Typically, the analyses were performed on several different spots of single sedimentary layers of 1–10 mm in thickness and of microveinlets and nodules to examine the variation in microscale spatial variations. The analyses were also performed on several sedimentary layers in each rock billet to examine the temporal variations. Isotopic analyses were

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Fig. 4. (Continued)

done on three to twenty spots of each rock billet. The analytical results are presented in Table 1. The total range of d34S values (eighty-four analyses) is from 26.3 to 111.8‰. Lambert and Donnelly (1991) described two populations of disseminated pyrite that occur in sedimentary rocks of the Fortescue and Hamersley groups: one has d34S values close to 0‰, and the other has more variable d34S (25 to 114‰). They suggested that the pyrite with variable isotopic compositions was formed as a result of sulfate reduction, although they were not certain whether the sulfate reduction was by a microbial or a thermochemical process. The pyrite of the Mt. McRae Shale from the present study may correspond to Lambert and Donelly’s second population, because the d34S values are variable and range from 26 to 112‰. All the data from the present study are compared in Fig. 6 in terms of pyrite type and stratigraphic position. Rock chips used for analysis often contain different types of pyrite: dissemi-

nated coarse grains and laminae in 493-1 (Fig. 4b), laminae and nodules in 493-11, and laminae and disseminated grains in 493-18. Regardless of the types of pyrite, isotopic compositions are similar for the given sample (see Fig. 6). In particular, sample 493-11 contains both pyrite nodules and pyrite laminae. The d34S values of the laminae (14.5 to 110.0‰) are similar to those of the nodule (13.7 to 17.6‰), suggesting that they are analogous in mode and/or timing of formation. 5.1.2. Lateral variations in microsedimentary layers

To investigate microscale variations in d34S values among the pyrite crystals formed at the same time, laser microprobe analyses were carried out on several spots along a ;5 mm distance on each single sedimentary layer of ;1 mm in thickness (Table 1). Pyrites in the upper and lower portions of the section show variable d34S values. In general, a greater than

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Fig. 5. The laser microprobe system at Tohoku University.

5‰ variation is detected among different spots in each sedimentary layer (Fig. 7). For example, a d34S variation of 7‰ (from 26.3 to 10.7‰) is recognized among the five spots in layer 5 of sample 493-1 in the upper section, and a d34S variation of 6.1‰ (from 15.7 to 111.8‰) among the ten spots in layer 5 of sample 493-18 in the lower section (Table 1). 5.1.3. Stratigraphic variations

The upper part of the drill core has distinctive isotopic compositions compared to the lower part. Laminae and coarse grains of pyrite in the two samples from the upper part of the section (Group 1 in Fig. 6) show a similar range of d34S values, from 26.3 to 17.1‰. The four samples from the lower section (Group 2 in Fig. 6) have positive d34S values ranging from 12.2 to 111.8‰ (Group 2 in Fig. 6). 5.2. Carbon and Sulfur Contents

To find the spatial and temporal variations in S and C contents of the shale samples, several different fine-scale (1–10 mm thick) sedimentary layers were analyzed for each rock sample. The results of the analyses are summarized in Table 2 and Fig. 7. The thirty microlayers in six samples show similar ranges of carbon content, 1.6 – 8.1 wt% C. An exception is layer 1 of sample 493-14 (0.5 wt% C) which represents a pyrite-rich layer. These data indicate that the carbon contents of the shales are generally homogeneous on a scale of millimeters (Fig. 7). Naraoka et al. (in prep.) confirmed that no carbonate minerals existed in the samples studied, except for minor siderite in one sample from the Mt. Sylvia Formation (493-28). Therefore, the measured carbon contents for the Mt. McRae Shale represent the contents of organic carbon. The high contents of organic carbon suggest that the Mt. McRae Shale deposited in an environment with high biological productivity and preservation. The sulfur contents of the thirty microlayers in the six shale samples are more variable (0.0 –24.2 wt%) than the carbon contents (Table 2). Two samples (493-6 and -11) show relatively constant sulfur contents within each rock chip, ;2% in sample 493-11 and ;20% in sample 493-6 (Fig. 7). A remark-

able feature is the sulfur depletion (almost to zero) found in some parts of samples 493-14, -18, and -28. In sample 493-28, pyrite occurs only as microveinlets (see Fig. 4f). The sulfur contents of the matrix are below detection. Sulfur depletion was also found around a pyrite nodule as described below. 5.3. Sulfur and Carbon Contents of Shale around a Pyrite Nodule

In order to understand the formation mechanism of pyrite nodules, we analyzed the sulfur and carbon contents of shales at thirteen spots in a ;7 cm thick zone adjacent to a pyrite nodule in sample 493-11(Table 3 and Fig. 8). In this sample, pyrite occurs as a nodule of approximately 3 cm in diameter (zone C in Fig. 8), which is bordered by a ;4 cm zone containing very small amounts of disseminated fine-grains of pyrite (zone B) and a ;3 cm zone of pyrite laminae (zone A). The carbon contents are found to be constant at 4.5–5.5 wt% within zone A but slightly variable within zone B from 4.7 to 8.1 wt%. Sulfur contents are more variable (0.2–9.2 wt%): 3.5–9.2 wt% in zone A (Fig. 8) but essentially zero in zone B. The sulfur content increases only slightly (from 0.2 to 0.5 wt%) immediately adjacent to the pyrite nodule. 6. DISCUSSION 6.1. Early Diagenetic Pyrite

The environmental systems for the formation of pyrite by bacterial sulfate reduction may be divided in two based on the relative rates of supply and consumption of sulfate in the system: (1) a closed-bottom system with respect to sulfate where the rate of reduction is equal to or faster than the rate of sulfate supply such that all the sulfate that enters the system is converted to pyrite in the system, and (2) an open-bottom system with respect to sulfate where the rate of sulfate supply is greater than the rate of sulfate reduction such that only a small portion of sulfate that enters the system at a time is converted to pyrite (Schwarcz and Burnie, 1973; Ohmoto, 1992). Note that, isotopically, natural systems are almost always open with respect to H2S (see Ohmoto, 1992).

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Table 1. d34S values of pyrite crystals in 6 core samples of the Mt. McRae Shale determined by the laser microprobe method. Chip no.

Layer no.

Location of layer (mm)

493-1 493-1 493-1 493-1 493-1 493-1 493-1 493-1 493-1 493-1 493-1 493-6 493-6 493-6 493-6 493-6 493-6 493-6 493-6 493-6 493-6 493-6 493-6 493-6 493-6 493-6 493-6 493-6 493-6 493-6 493-11 493-11 493-11 493-11 493-11 493-11 493-11 493-11 493-11 493-11 493-11 493-11 493-11 493-11 493-11 493-11 493-11 493-11 493-11 493-11 493-11 493-11 493-11 493-11 493-14 493-14 493-14 493-14 493-14 493-14 493-14 493-14 493-14 493-14 493-14 493-14 493-14

1 1 1 2 2 4 5 5 5 5 5 1 1 2 2 2 4 5 5 5 5 5 6 n.d. n.d. n.d. n.d. n.d. n.d. n.d. 3 4 7 8 9 10 11 13 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. 1 1 1 1 1 1 1 1 2 2 2 5 6 7 8 8 8 8

4.0 4.0 4.0 8.0 8.0 12.5 14.0 14.0 14.0 14.0 14.0 4.0 4.0 6.5 6.5 6.5 8.5 12.0 12.0 12.0 12.0 12.0 15.5 n.d. n.d. n.d. n.d. n.d. n.d. n.d. 6.0 8.5 14.0 14.5 15.0 16.5 18.5 19.5 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. 3.0 3.0 3.0 4.5 4.5 4.5 14.0 15.0 15.5 16.5 16.5 16.5 16.5

Spot no. (in each layer)

d34S(CDT) (‰)

Dominant Py occurrence

1 2 3 1 2 1 1 2 3 4 5 1 2 1 2 3 1 1 2 3 4 5 1 n.d. n.d. n.d. n.d. n.d. n.d. n.d. 1 1 1 1 1 1 1 1 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. 1 2 3 4 5 1 2 3 1 2 3 1 1 1 1 2 3 4

20.3 24.9 10.2 23.7 12.5 25.2 20.3 26.3 21.9 10.7 20.4 21.9 20.6 17.1 20.6 11.2 20.4 24.5 23.8 24.2 24.6 20.6 21.1 10.1 11.0 20.3 21.8 21.0 21.7 21.5 16.0 17.3 19.1 15.9 16.0 15.2 16.1 16.3 14.6 16.8 18.5 19.4 14.5 19.1 19.3 19.4 18.5 110.0 17.8 17.6 16.5 16.1 16.4 13.7 16.4 15.6 14.5 19.0 17.4 15.5 12.3 12.6 16.5 14.7 12.7 15.3 12.2

Coarse grains Coarse grains Coarse grains Coarse grains Coarse grains Laminae Coarse grains Coarse grains Coarse grains Coarse grains Coarse grains Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Nodule Nodule Nodule Nodule Nodule Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae

Table 1. (Continued) Chip no.

Layer no.

Location of layer (mm)

Spot no. (in each layer)

d34S(CDT) (‰)

Dominant Py occurrence

493-14 493-18 493-18 493-18 493-18 493-18 493-18 493-18 493-18 493-18 493-18 493-18 493-28 493-28 493-28 493-28 493-28

8 2 5 5 5 5 5 5 5 5 5 5 1 1 1 1 1

16.5 8.0 11.0 11.0 11.0 11.0 11.0 11.0 11.0 11.0 11.0 11.0 9.5 6.0 8.0 7.5 7.0

5 1 1 2 3 4 5 6 7 8 9 10 1 2 3 4 5

14.7 110.5 18.5 111.8 17.5 18.3 15.7 18.5 110.3 110.8 110.5 19.7 18.7 17.3 111.7 110.4 110.2

Laminae Disseminated py Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Laminae Veinlet Veinlet Veinlet Veinlet Veinlet

Different spots on several sedimentary layers were analyzed for each sample. The locations of the sedimentary layer are indicated by the distance (mm scale) from an edge of the rock chip. n.d. 5 not determined. d34S 5 d34SCDT (‰).

The open-bottom system is found in: (1) the upper part of a sediment column ('bioturbation zone) in modern normal marine sediments and (2) in the anoxic water column of a euxinic basin, such as the Black Sea and the Baltic Sea (e.g., Sweeney and Kaplan, 1980; Boesen and Postma, 1988; Lyons, 1997). The closed-bottom system is represented by interstitial water in a saturated sediment column where the diffusive supply of seawater sulfate is limited. Sulfate reduction in a closed-bottom system is analogous to a Rayleigh distillation process, and the isotopic change during sulfate reduction can be evaluated from the following equations: F(a- 10) 5

a5

d34S(H2S) 1 1000 a(d34S(i) 1 1000)

(34S/32S)H2S 1000 ' (34S/32S)SO4 DSO22 2 H2S11000 4 22

(1) (2)

where F is the fraction of sulfate remaining in the system, d34S (H2S) is the d34S value of sulfide produced at time t, d34S(i) is the initial isotopic composition of the sulfate (i.e., seawater sulfate), and a is the kinetic isotopic fractionation factor. The d34S values of pyrite should be the same compositions as those of HS2 (H2S), because isotope fractionation during the transformation of HS2 (H2S) to FeS2 in sedimentary systems is negligible (Wilkin and Barnes, 1996). Because of the Rayleigh fractionation, individual crystals of pyrite formed in a closedbottom system have more variable d34S values compared to those formed in an open-bottom system. Three possibilities may be considered for the origin of early pyrite (i.e., disseminated and laminae pyrite) in the Mt. McRae Shale: (1) direct precipitation from a H2S (or HS2) rich water column (i.e., syngenetic pyrite), in which the H2S (or HS2) could have been produced by hydrothermal or biological processes, (2) bacterial reduction of seawater sulfate during early

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Fig. 6. Comparisons of d34S values according to the stratigraphic positions and types of pyrites.

stages of sediments, and (3) hydrothermal (thermochemical) reduction of seawater sulfate in sediments. An example of syngenetic pyrite formation is found in the anoxic water body of the Black Sea, where the H2S is produced by bacterial sulfate reduction in the water body (cf. Lyons, 1997; Ohmoto and Goldhaber, 1997). Syngenetically formed pyrite crystals most likely have uniform d34S values in microscale area (Ohmoto et al., 1991; Lyons, 1997), unless sulfur isotope compositions were modified significantly by a diagenetic process. As shown in Fig. 7, d34S values vary by 4.4 – 11.7‰ among the different spots in the same microlayers, suggesting that sulfate reduction in pore volumes of sediments was responsible for pyrite formation rather than syngenetic process. High temperature hydrothermal processes, including thermochemical (abiological) sulfate reduction at the deposition sites, typically result in uniform isotopic composition for sulfides of the same generation (Ohmoto and Rye, 1979; Ohmoto and Goldhaber, 1997). The large d34S variations observed among pyrites in the same stratigraphic layers and those among different stratigraphic units are difficult to explain by thermochemical reduction at the deposition sites, but can be easily explained by the model of bacterial sulfate reduction in the sediments. Abundant carbonate minerals may also be expected as products of oxidation of organic matter, if the thermochemical sulfate reduction of the organic matter in sediments was responsible for sulfide formation (Machel et al., 1995). However, carbonate minerals are not common in the Mt. McRae Shale.

Sulfur and organic carbon contents are generally positively correlated in Phanerozoic normal marine sediments (i.e., sediments under an oxic seawater) with a regression equation of S (wt%) 5 0.37 3 C (wt%) (e.g., Berner, 1984; Ohmoto et al., 1991). Sulfate-reducing bacteria in combination with organic carbon limited conditions are responsible for such a correlation of S and C contents in marine sedimentary rocks (Raiswell and Berner, 1985). The regression line for sediments formed in a euxinic basin may have a measurable S content at 0 wt% C because they may contain appreciable amounts of pyrite formed in the water column (Leventhal, 1982; Raiswell and Berner, 1985; Ohmoto et al., 1991; Lyons and Berner, 1993; Lyons, 1997). Figure 9a shows the relationship between C and S contents of micro-drilled samples in two shale samples from the upper part of the drill core. Some variations are recognized in S and C contents among the microdrilled powders. The S-C contents of sample 493-1 fall on the regression line for modern normal marine sediments. However, the S-C contents of sample 493-6 plot far above the regression line because only pyrite-rich layers (pyrite laminae) were microdrilled for this sample. Hence, the high S contents of sample 493-6 are probably due to a sampling bias rather than indicating late pyrite formation. In this sample, coarse-grained pyrite crystals, probably recrystallized from fine-grained crystals, are common. Remobilization and recrystallization of fine-grained pyrite could have been a reason for variable sulfur contents. The relationship between S and C contents in shales from a lower part of the drill core is illustrated in Fig. 9b. The S contents of the lower section are

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Fig. 7. Variations in the S and C contents and d34S in each rock chip sample. Different spots on several sedimentary layers were analyzed for each sample. The locations of the sedimentary layers are indicated by the distance (mm scale) from an edge of the rock chip.

more variable (0.2–24.2%) compared to the upper two samples. The probable reason for the variation will be discussed in the following section. Analyses of C and S contents of bulk rocks on a hand specimen scale have been performed on the same drill core samples (Naraoka et al., unpubl. data). Their data are similar to those for some Phanerozoic shales (Fig. 9c). Abnormally high sulfur contents were not detected in the bulk scale, suggesting that hydrothermal input of sulfur into the Mt. McRae Shale was not important. 6.2. Remobilization of Pyrite

Pyrite nodules and pyrite-bearing carbonate concretions in the Phanerozoic shales have been studied by many investigators (Maynard, 1980; Berner, 1980; Coleman and Raiswell, 1981; Carrigan and Cameron, 1991). Previous investigators have stressed the importance of the compaction of shales for the formation of pyrite nodules. In the case of the Mt. McRae Shale, the host sediments are differentially compacted around the pyrite nodules, indicating that the pyrite nodules formed before the compaction was completed. The sulfur contents of shales are extremely low (,0.2 wt%) around the pyrite nodule (see zone B of Fig. 8). The deficiency of sulfur in zone B is interpreted as a result of remobilization of pyrite during compaction of sediments. The dissolved pyrite

may have reprecipitated as pyrite nodules. This is suggested by the similar d34S values between pyrite laminae and pyrite nodules occurring in the same hand specimen (see Fig. 6). In each rock chip of the lower group (Fig. 9b), the organic carbon contents are fairly constant (2– 4 wt%) on a millimeter scale, but the S contents are extremely variable. This variation may have been caused by dissolution and reprecipitation of pyrite. In sample 493-28, the sediment matrix contains essentially no sulfur (Fig. 7f). Most of the S in this sample is localized in microveinlets or aggregates of coarse-grained pyrite that were most likely formed during compaction of the shale, coeval with the pyrite nodule formation. Their similar d34S values support this model (Fig. 6). 6.3. Constraints on the Formation Mechanisms of Pyrite Nodules 6.3.1. Formation depth and timing

The differentially compacted shaley matrix and the shapes of pyrite nodules support the hypothesis that pyrite nodules formed before the complete compaction of sediments. That is, the growth of pyrite nodules took place in soft and wet sediments. Muds contain as much as 90% water when deposited (Potter et al., 1980) and continue to be compacted during burial due to increased sediment loading. The decrease in the porosity of

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Table 2. Sulfur and carbon contents of 6 core samples of the Mt. McRae Shale

Chip no. 493-1 493-1 493-6 493-6 493-6 493-6 493-11 493-11 393-11 493-11 493-11 493-11 493-11 494-11 493-11 493-11 493-11 493-11 493-11 493-14 493-14 493-14 493-14 493-18 493-18 493-18 493-28 493-28 493-28 493-28 493-28

Layer no.

Location of layer (mm)

Sulfur (wt%)

Carbon (wt%)

1 3 1 5 3 6 1 2 6 9 12 14 n.d. n.d. n.d. n.d. n.d. n.d. n.d. 1 3 4 6 1 3 6 1 2 3 4 2

4.0 11.5 4.0 12.0 7.5 15.5 3.0 6.0 11.0 15.0 19.0 24.0 n.d. n.d. n.d. n.d. n.d. n.d. n.d. 3.0 7.0 12.0 15.0 3.0 9.0 12.0 3.0 6.0 6.5 8.0 6.5

1.9 2.8 22.5 18.6 18.2 19.7 8.9 5.1 3.5 9.2 4.4 4.4 0.9 0.2 0.2 0.2 0.2 0.2 0.5 24.2 0.2 3.0 2.1 0.2 17.5 0.6 0.0 0.0 0.0 0.0 0.0

4.5 4.4 5.0 5.3 5.0 5.4 4.5 4.7 5.5 4.9 5.0 5.5 4.7 5.4 5.9 6.5 8.1 7.3 6.1 0.5 3.3 4.6 4.1 2.7 1.6 3.4 2.0 2.4 2.6 3.4 2.3

n.d. 5 not determined

muds continues to a few hundred meters of burial with 75% of the water being expelled during burial to 100 m (Hanor, 1979). Most previous investigators have concluded that carbonate nodules (concretions) in modern shales form at a shallow depth of burial, possibly just below the seawater-sediment interface (e.g., Raiswell, 1976; Curtis et al., 1986; Mozley and Burns, 1993). There exist various differences in mineralogy and size

Table 3. Sulfur and carbon contents of shales around a pyrite nodule in sample 493-11 Analyses

Distance (mm)

Sulfur (wt%)

Carbon (wt%)

1 2 3 4 5 6 7 8 9 10 11 12 13

6.8 6.5 6.0 5.6 5.2 4.7 4.2 3.6 3.0 2.5 1.7 1.0 0.0

8.9 5.1 3.5 9.2 4.4 4.4 0.9 0.2 0.2 0.2 0.2 0.2 0.5

4.5 4.7 5.5 4.9 5.0 5.5 4.7 5.4 5.9 6.5 8.1 7.3 6.1

Fig. 8. Variations in S and C contents of shales around a pyrite nodule in sample AUS493-11.

between carbonate nodules and pyrite nodules. Carbonate nodules are typically composed of carbonate minerals with an abundance of incorporated clay and pyrite, and their sizes often exceed 50 cm in diameter (Hennessy and Knauth, 1985; Mozley and Burns, 1993). By contrast, pyrite nodules are typically monomineralic and smaller than 10 cm in diameter. However, the deformation patterns of sedimentary laminae around carbonate concretions are similar to those around pyrite nodules, suggesting that the pyrite nodules and carbonate concretions most likely formed during similar stages of diagenesis, i.e., during shallow burial depth (,100 m from the sediment/water interface). 6.3.2. Dissolution and reprecipitation of early sulfides

Hydrogen sulfide generated by sulfate-reducing bacteria is commonly fixed by Fe in modern marine sediments unless Fe is limited. Initially, mackinawite (FeS0.9) is formed (Berner, 1967). This phase reacts with elemental sulfur and/or other intermediates, such as polysulfides, to form greigite (Fe3S4; Skinner et al., 1964; Goldhaber and Kaplan, 1974; Schoonen and Barnes, 1991a; Wilkin and Barnes, 1996). The mackinawite and greigite are highly unstable, and the transition to FeS2 may take place within a short period of time (Jørgensen, 1977). The solubility products of mackinawite and greigite are higher than that of pyrite (Berner, 1967; Goldhaber and Kaplan, 1974; Schoonen and Barnes, 1991a,b), suggesting that greigite and mackinawite could be the major sources of sulfur and iron for the pyrite nodules. Remobilization of early sulfides may have

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Fig. 9. Plots of S contents vs. C contents of the Mt. McRae Shale. (a) microarea analyses of Group 1 samples (493-1 and -6); (b) microarea analyses of Group 2 samples (493-11, -14, -18, and - 28); (c) bulk-rock analyses of drill core No. 493 (Naraoka et al., unpubl. data).

been caused either by diagenetic or hydrothermal solutions. However, it is not clear which process (diagenetic or hydrothermal) was responsible for the remobilization of early sulfides, or whether sulfate-reducing bacteria were directly in-

volved in the formation of the pyrite nodule. Hydrothermal activity during and/or shortly after the formation of pyrite nodules has been suggested by Haruna and Ohmoto (1992). They have recognized that some pyrite nodules in the McRae

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is the pyrite content (wt%) in shale before the nodule formation. Equation (3) can be simplified as (R1/R2)35D2/D1 * X/100

(4)

The typical radius dimensions of nodules in the Mt. McRae Shale are 1–2 cm. The shales away from the nodule typically contain 4 wt% sulfur or 8 wt% pyrite with D1 being 3 g/cm3. It follows that a pyrite nodule with a radius of 2 cm requires reconcentration of disseminated pyrites from a sphere of only a 5 cm radius. The agreement between the results of calculation and observation (Fig. 8) supports a suggestion that a typical pyrite nodule in the McRae Shale was formed by migration of Fe and S from a region within a few centimeters of the nodule. 6.4. Seawater Sulfate in the Hamersley Basin, ;2.5 Ga Ago

Fig. 10. Relationship between the radius of a pyrite nodule (R1) and the radius of a zone of migration of Fe and S (R2). See text for explanation.

Shale are rimed by quartz crystals that contain fluid inclusions with homogenization temperatures of 120 –220°C. The importance of decomposition of organic matter has been emphasized for the formation of carbonate concretions by previous investigators (e.g., Mozley and Burns, 1993; Berner, 1984). Decomposition of organic matter, which generates CO2 (HCO2 3 ) to form carbonate minerals, may be caused either by microbial or thermal processes during early diagenesis. Various organic acids, typically humic, fulvic and acetic acids are also generated from organic matter at or near the seawater-sediment interface (Giordano, 1993; Nissenbaum and Swaine, 1976). Organic acids contribute significantly to the dissolution of silicates and oxides (e.g., Bruno et al., 1992). Although experimental data are not available, organic acids may also help to dissolve pyrite (or iron monosulfide) and transport iron during the early stage of diagenesis. 6.3.3. Migration distance of iron and sulfur

The following mass balance calculations are made in order to examine the relationship among (1) the size of a pyrite nodule, (2) the amount of disseminated pyrite dissolved from the host shale, and (3) the migration distance of pyrite from the host shale to the site of nodule formation (Fig. 10). It is hypothesized that the earlier disseminated pyrite and/or iron monosulfides were completely dissolved in the vicinity of a nodule. The mass balance of pyrite before and after the nodule formation in a spherical volume of shale (radius of R1) can be expressed as 4/3 p R31 * D1 * X/100 5 4/3 p R32 * D2 * 100/100,

(3)

where D1 is the density of shale before the nodule formation, D2 is the density of a pyrite nodule, R1 is the radius of sulfur-depleted shale, R2 is the radius of a pyrite nodule, and X

Because benthic organisms were absent in Archean sediments (Schopf and Walter, 1983), most diagenetic sulfides in Archean sedimentary rocks may be assumed to have formed in closed-bottom systems with respect to sulfate where sulfate was supplied to a packet of sediments by diffusion, rather than by advection, and the sulfate was all consumed within the packet of sediments (Ohmoto and Felder, 1987; Ohmoto et al., 1993). The thickness of a closed-bottom system is typically 200 cm in modern sediments (Goldhaber and Kaplan, 1974; Ohmoto and Goldhaber, 1997) but could have been less than ;10 cm in Archean sediments if the rates of bacterial sulfate reduction in Archean oceans were more than twenty times faster than those in the modern ocean (Ohmoto et al., 1991; Ohmoto, 1992). Higher reduction rates could have been caused if the Archean oceans were warmer and/or richer in organic matter that was more easily metabolized by sulfate-reducing bacteria. In an ideal closed-bottom system where one pore volume of seawater sulfate is successively converted to a certain number of grains of pyrite crystals of an equal size (e.g., 100 grains of 10 mm size), when all pyrite crystals in the systems were individually analyzed, the frequency diagram for the d34S values of individual pyrite crystals would exhibit the following characteristics (cf., Ohmoto et al., 1993; Ohmoto and Goldhaber, 1997): (1) the average of the measured d34Spyrite values coincides with the d34Sseawater value, (2) the difference between the d34Sseawater and the minimum d34Spyrite coincides with the kinetic isotope effect (D value) accompanying bacterial sulfate reduction, (3) the most frequently occurring d34Spyrite value is very close to the minimum d34Spyrite value, and (4) the maximum d34Spyrite value exceeds the d34Sseawater value, but the number of pyrite crystals with the d34Spyrite .d34Sseawater relationship is much less than those with the d34Spyrite , d34Sseawater relationship. However, when only some pyrite crystals in a closedbottom system were analyzed, as is the case in the present study, (1) the average of the measured d34Spyrite values tends to be less than the d34Sseawater value, (2) the difference between the d34S seawater and the minimum d34S pyrite tends to be smaller than, but close to, the true kinetic isotope effect (D value) accompanying bacterial sulfate reduction, and (3) the maximum d34S pyrite value tends to be less than the d34Sseawater value.

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Fig. 11. d34S frequency diagrams for pyrite crystals of Group 2 samples. The d34S variations and frequency patterns for micro areas of samples AUS 493-11(l) (pyrite laminae dominated sample) and AUS 493-11(n) (a nodule in the sample) are shown in (a) and (b), respectively. The d34S variations and frequency patterns of other samples (AUS493-14, -18, and -28) are shown in (c) to (e), respectively. Note that sample AUS 493-28 was collected from the Mt. Sylvia Formation, and their d34S values were determined only on a microveinlet of pyrite. ave. 5 average; n 5 number of analyses; No. 5 number.

The d34S frequency patterns for each sample analyzed in this study are illustrated in Figs. 11 and 12. The average d34S values of Group 2 samples range from 15.0‰ (sample 493-14) to 19.7‰ (sample 493-28). Two samples (sample 493-18 and -28) have average d34S values of around 19.5‰ (Fig. 11). The maximum d34Spyrite value is similar for four Group 2 samples at between 110 to 112‰. Strauss and Beukes (1996) reported a d34S value of 115‰ for trace sulfate minerals in carbonate rocks of the 2.5 Ga Malmani Subgroup in South Africa. Taking all these data into consideration, we may suggest that the d34S value of the 2.5 Ga seawater was higher than 110‰ and possibly as high as 115‰. The minimum d34Spyrite value for Group 2 samples is 12‰ (see Fig. 11). This value, together with the above seawater value, suggests that the kinetic isotope effect (D) was greater than 8‰ and possibly as much as 13‰ during the formation of Group 2 pyrite. The minimum d34S value among Group 1 pyrite crystals in Fig. 12 is 26‰. This value, together with the d34Sseawater value of 110 to 115‰, suggests the kinetic isotope effect (D) of 16 –21‰ for the bacterial sulfate reduction during the deposition of Group 1 pyrite (Fig. 12). The Rayleigh distillation curve for pyrite crystals formed in a simple closed-bottom system is

illustrated in Fig. 12c where the d34S of seawater sulfate was assumed to be 110‰, the D value was 16‰, and a single batch of seawater sulfate was supplied into the system (Eqns. 1 and 2). In this model, most pyrite would have d34S values between 27‰ and 11‰, agreeing with the d34S variations found in both samples (Fig. 12c). The estimated values for the kinetic isotope effect accompanying pyrite formation in the Mt. McRae Shale (21‰) are much smaller than the Phanerozoic D values of about 20 –70‰ (Schwarcz and Burnie, 1973; Ohmoto, 1992; Canfield and Teske, 1996). Two different models have been proposed to explain the small kinetic isotope effects in the Archean sediments. The first model attributes the small kinetic isotope effects to lower concentrations of seawater sulfate, i.e., ,1/10 of the present seawater level of 28 mM (Canfield and Teske, 1996; Habicht and Canfield, 1996). This model is mostly based on the results of bacteria culture experiments by Harrison and Thode (1958) in which the D value became smaller than 5‰ when the concentration of sulfate in water was decreased to less than about 1 mM. However, a serious problem with this model is its difficulty in explaining the occurrences of pyrite-rich shales, such as the Mt. McRae Shale, of Archean age. Sedi-

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T. Kakegawa, H. Kawai, and H. Ohmoto

solutions containing 10 –100 mM sulfate, and (2) the data of modern marine sediments where the D values are negatively correlated to the rate of sulfate reduction (e.g., Goldhaber and Kaplan, 1975). Ohmoto (1992) has also shown that sulfate minerals in marine sediments and in massive sulfide ore deposits of Archean age are much more common than previously believed and cited them as evidence for sulfate-rich oceans. The kinetic isotope effects observed in bacteria culture experiments by Habicht and Canfield (1996) were mostly larger than 20‰ even at high rates of sulfate reduction. Based on these data, they have questioned the model by Ohmoto et al. (1993). However, a close examination of the experimental data by Habicht and Canfield (1996) suggests that the rates of sulfate reduction in their experimental systems were not as high as those of Harrison and Thode (1958), where D values less than 20‰ were observed. Therefore, we interpret that the small kinetic isotope effects accompanying the pyrite formation in the Mt. McRae Shale occurred because the 2.5 Ga ocean already contained as much sulfate as modern oceans and sustained much higher activity levels of sulfate-reducing bacteria than do modern oceans. The higher activity of sulfate-reducing bacteria could have caused because the Archean oceans were warmer, and the organic matter was more easily metabolized by sulfate-reducing bacteria (Ohmoto and Felder, 1987; Kasting, 1993). A chert layer, ;1 mm thick, separates the older Group 2 pyrite from the younger Group 1 pyrite (Fig. 3). There was a distinct increase in the kinetic isotope effect accompanying bacterial sulfate reduction from 8 to 13‰ during the deposition of Group 2 pyrite to 16 –21‰ during the deposition of Group 1 pyrite. Such an increase in the kinetic isotope effect suggests that some changes occurred in the environment of the Hamersley Basin following the deposition of the chert layer, decreasing the rate of bacterial sulfate reduction. For example, a decrease in the rate of clastic sedimentation (Zaback et al., 1993) or in the seawater temperature could have increased the D value. Fig. 12. d34S frequency diagrams for pyrite crystals of Group 1 samples. The d34S variations and frequency patterns for micro areas of samples AUS 493-1 and -6 are shown in (a) and (b), respectively. In order to compare the observed d34S variations with those expected from the Rayleigh fractionation process, a theoretical frequency curve (c) was calculated using Eqns. 1 and 2 in the text. In these calculations, it was assumed that the d34S value of the seawater sulfate was 110‰ and kinetic isotope effect (D 1) was 16‰. ave. 5 average; n 5 number of analyses; No. 5 number.

ments formed in sulfate-poor environments, such as those in fresh water environments, generally contain much less than 0.5 wt% pyrite sulfur (e.g., Raisewell and Berner, 1985). The second model to explain the small kinetic isotope effects in Archean sediments postulates that the Archean seawater was sulfate-rich, containing $ 10 mM (i.e., $ one-third of the present seawater level), and that the activity of sulfate-reducing bacteria was much higher than that in the modern ocean (Ohmoto et al., 1993). This model is mainly based on: (1) the results from bacteria culture experiments by Harrison and Thode (1958), Kaplan and Rittenberg (1964), and Chambers et al. (1975) which showed increases in the D values from about 5‰ to about 45‰ with decreasing rates of sulfate reduction in

7. SUMMARY AND CONCLUSIONS

Analyses of sulfur isotopes by the laser microprobe were performed on pyrite grains in five core samples from the Mt. McRae Shale and one sample from the Mt. Sylvia Formation (;2.5 Ga), Hamersley Group. Sulfur and carbon contents were also determined on micro-drilled powders from these samples. Based on the occurrence, morphology, abundance, and d34S values of pyrite, the six samples were divided into two groups. Group 1 is shales from the upper section that contains finegrained disseminated pyrite (,10 mm), locally showing laminated texture. This pyrite appears to have been formed by bacterial reduction of seawater sulfate during early diagenesis (Fig. 13). Pyrite in Group 2 (lower section) consists of nodules and laminae and appears to have formed through more complicated processes. Pyrite laminae were probably the primary phases. They may have been remobilized to form pyrite nodules, pyrite veins, and disseminated pyrite. This remobilization of pyrite occurred during the compaction of the Mt. McRae Shale either by diagenetic or hydrothermal solutions (Fig. 13). Bacterial reduction of multiple supplies of seawater sulfate may have been important in the formation of pyrite laminae in Group 2.

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Cook-Gibson for their technical assistance. This manuscript benefited greatly from comments and suggestions by T. Lyons and an anonymous reviewer. This study was supported by the National Science Foundation under Grant Nos. EAR-9003554 and EAR-9706279 to H. Ohmoto and by the Japanese Ministry of Science, Culture and Education under Grant No. 03102002 to H. Ohmoto. REFERENCES

Fig. 13. Model for the origin of pyrites in the Mt. McRae Shale. The upper and lower parts of the drill core sections are classified into Group 1 and Group 2, respectively. Most pyrite crystals in Groups 1 and 2 were probably formed by bacterial sulfate reduction during early diagenesis. Pyrite nodules were most likely formed during diagenetic stages by remobilization of disseminated pyrite by diagenetic or hydrothermal solutions.

Variations in d34S values of pyrite in the Mt. McRae Shale, such as those within single micro layers and those among different stratigraphic units, indicate that most of the pyrite crystals were formed by bacterial reduction of seawater sulfate during the early diagenesis of sediments, and that the kinetic isotope effects accompanying bacterial sulfate reduction varied both in time and space. Pyrite nodules in the Mt. McRae Shale were formed during diagenesis of the host sediments through remobilization of disseminated pyrite by diagenetic or hydrothermal solutions. Analyses of d34S values of the disseminated and laminae pyrites suggest that the 2.5 Ga seawater in the Hamersley Basin was already rich in sulfate (greater than one-third of the present seawater level) and that the d34Sseawater value had already increased to .110‰ (possibly around 115‰). Some of the important implications of these suggestions are that, by 2.5 Ga ago, the atmosphere had already become O2 rich, and the global cycles of sulfur had been dictated by the Phanerozoic-style biogeochemical processes, rather than by the mantle-buffer processes of Veizer et al. (1989). The activity of the sulfate-reducing bacteria in the Archean ocean was probably higher than that in the modern ocean, resulting in a small kinetic isotope effect (D) during sulfate reduction. The high activity of sulfate-reducing bacteria was probably caused by generally warm Archean weather and the highly metabolizable nature of organic matter for sulfate-reducing bacteria. Acknowledgments—The authors wish to thank M. Arthur, D. H. Eggler, D. P. Gold, A. W. Rose, and K. Osseo-Asare for their comments on earlier manuscripts. Thanks are extended to K. Hayashi, H. Naraoka, Y. Watanabe, and S. R. Poulson for discussions and support in various phases of this research. Special thanks go to D. P. Walizer and L. I.

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