Carbon and sulfur isotopic compositions of organic carbon and pyrite in sediments from the Transvaal Supergroup, South Africa

Pre/umbrinn Resenrth ELSEVIER

Precambrian Research 79 (1996) 57-71

Carbon and sulfur isotopic compositions of organic carbon and pyrite in sediments from the Transvaal Supergroup, South Africa Harald Strauss a, Nicolas J. Beukes

b

a Ruhr-Universiti~t Bochum, lnstitutfftr Geologie, D-44780 Bochum, Germany t, Rand Afrikaans University, Department of Geology, Auckland Park 2006, South Africa

Received 31 January 1994; revised version accepted 1 December 1995

Abstract

Sedimentary organic matter and pyrite from parts of the lower Transvaal Supergroup and few samples from the Olifantshoek Group have been studied for their carbon and sulfur isotopic compositions, respectively. Carbon isotopic compositions for minimally altered organic matter range from - 4 3 to -31%o. Sulfur isotope values for sedimentary pyrite display values between -1.3 to + 23.6%o. Results from this and previously published work are being interpreted with respect to past biochemical pathways. Carbon isotope data indicate autotrophic carbon fixation as the process of primary production. In addition, products of secondary biological reworking contributed to the total organic carbon content. Sulfur isotope data indicate bacterial sulfate reduction as the main process. However, changes in sulfate availability and/or the rate of sulfate reduction may have caused the observed range in isotopic compositions.

1. I n t r o d u c t i o n

The study of carbon and sulfur isotopes in sedimentary organic matter and pyrite is believed to reveal insights into ancient biochemical pathways, point to differences in depositional environments, and enhance our understanding of the sedimentary cycling through time. Coupling of the carbon and sulfur geochemical cycles bear information on the exogenic cycle of past sedimentary systems. Furthermore, pyrite genesis and past depositional environments can be assessed through evaluation of the system Corg-Spyr-FeR (reactive iron). All three parameters strongly affect the biological formation of sedimentary pyrite, with one of them--depending on

the sedimentological a n d / o r geochemical conditions - - b e i n g the limiting factor. Evaluation of their relationships has been used successfully in characterizing geochemical conditions in depositional environments of Phanerozoic age (i.e., Raiswell and Berner, 1986). Previous studies on the paleobiology of the Precambrian, specifically the geochemical aspects (Hayes et al., 1983, 1992; Schidlowski et al., 1983; Strauss et al., 1992a, b, c), have clearly demonstrated the amount of information concerning ancient biochemistries obtainable from studying Precambrian sediments. The aspect of preservation, in particular, appears to be limiting in this regard. Taking this into consideration, it appears feasible to interpret geo-

0301-9268/96/$15.00 © 1996 Elsevier Science B.V. All rights reserved SSDI 0301-9268(95)00088-7

58

H. Strauss, N.J. Beukes / Precamhrian Research 79 (1996) 57-71

chemical and isotopic data obtained from Precambrian sediments in a way similar to younger ones. In this regard, abundances and isotopic compositions of organic carbon and sulfide sulfur (pyrite) and abundances of reactive iron species have been determined for sediments from the Transvaal Supergroup and the Olifantshoek Group. In addition, relevant literature data were screened. The objective of this study was to present a comprehensive view of the biogeochemistry of carbon and sulfur and their characteristics at the time of deposition. Specific questions include the prominent way of carbon fixation, the process of sulfate reduction and the mode of pyrite formation.

in numerous studies, and the reader is referred to more recent reviews by Beukes (1986), Eriksson and Clendenin (1990), and references therein• Similarly, associated banded iron-formations have been studied by Beukes (1983, 1984), Beukes et al. (1990) or Beukes and Klein (1992). Samples analysed during this study were taken from the lower part of the stratigraphic sequence, namely the Schmidtsdrif and Campellrand Subgroups (Ghaap Group), and the Mapedi Shale, which is part of the Olifantshoek Group and unconformably overlies the VoiSlwater Subgroup. A representative stratigraphic column is shown in Fig. 1. This study focuses on the shales and siltstones which are interbedded with stromatolitic and laminated carbonates and contain mostly disseminated pyrite. Samples containing concretions have been avoided during this study• A Paleoproterozoic age has been assigned to the Transvaal Supergroup. Published geochronological

2. Geological setting and age Aspects of the geology and stratigraphy of the Transvaal Supergroup have been treated extensively

Griqualand

Eastern

West i. •

Olifantshoek G r o u p L2 I 0 0 - 1900 Ma ] Hartley Lava / .i Mapedl Shale ~_ Postmasburg Group Voelwater S u b g r o u p O n g e l u k Lava M a k g a n y e n e Diamlctite

Ghaap G r o u p

Transvaal

Pretoria G r o u p

i

Magaliesburg Formation Silverton F o r m a t i o n Daspoort Formation

[ 2 2 2 4 ± 21 Ma [ - ~3ncOtago~X~}t2/~ . _ ~ . . / . . . . . . . . . . .

-

Strubenkop Formation Hekpoort Lava Boshoek F o r m a t i o n --Tlmeb--all-I-~ill-Fo1"naati%n _Ro_oihoo__gteF _ o r m a U o n

Chunlespoort Group Dultschland Formation

Koegas S u b g r o u p Asbestos Subgroup Danielskull Formation K u r u m a n Iron Formation f

2 4 3 2 ± 3 1 Ma

Penge Iron F o r m a t i o n Malmanl Subgroup

Campbellrand Subgroup - -

[ 2552

Gamohaan Formation ~ [ Kllpfontetnheuwel F o r m a t i o n Relvilo F o r m a t i o n Montevllle F o r m a t i o n Schmldtsdrif Subgroup

±

11 Ma

Eccles F o r m a t i o n Lyttelton F o r m a t i o n Monte Crlsto F o r m a U o n Oak Tree Formation

.¢~ ~

[ 2 5 5 7 x 49 Ma I

Lokamonna FormaUon/ Boomplas Formation -Vryburg Formation -I

Frisco F o r m a t i o n

2 5 2 4 Ma

• _

Black _Reef. . . . . . .

/. .

2 6 4 0 Ma

]

Fig. I. S t r a t i g r a p h y o f the T r a n s v a a l S u p e r g r o u p in the G r i q u a l a n d - W e s t and E a s t e r n T r a n s v a a l a r e a s ( r e v i s e d after S.A.C.S., 1980)

H. Strauss, N.I. Beukes / Precambrian Research 79 (1996) 57-71

dates for the lower part of the sequence are 2432 ± 31 Ma for the Kuruman Iron Formation (Trendall et al., 1990) and 2557 ± 49 Ma for the stromatolitic carbonates of the Boomplas Formation (Jahn et al., 1990). The Hekpoort/Ongeluk Lava higher up in the sequence has been dated at 2224 ± 21 Ma (Hamilton, 1977). Most recent studies suggest a late Archean age with new dates of 2640 Ma for a tuffaceous bed within the Vryburg Formation (F. Walraven et al., pers. commun., 1994), 2552 ± 11 Ma for a tuffaceous bed in the Campbellrand Subgroup (Barton et al., 1994) and 2524 Ma for a tuffaceous bed in the Gamohaan Formation (Sumner and Bowring, 1996). Beukes and Smit (1987) presented stratigraphic and structural evidence for the position of the Mapedi Shale (Griqualand Wes0 in relation to the rest of the Transvaal sequence. In general, the Mapedi Shale is a red-bed sequence which rests with angular unconformity on the

I 24* ~ -~ -~

/~)

59

Transvaal Supergroup. It contains black shales, some of which have been studied here. An age range of 2100-1900 Ma is assigned to this unit (cited in Holland and Beukes, 1990).

3. Sampling details and analytical methods All samples were collected from drillcores in order to avoid surficial weathering. Samples from the Mapedi Shale (DDH A) and from the Ghaap Group (DDH B) are from the Griqualand-West area. DDH C (Black Reef) is located in the central Transvaal basin, about 15 km east-southeast of Carletonville (Fig. 2). Standard analytical techniques have been applied. Organic carbon abundances and carbon dioxide for isotopic analyses were obtained through sealed tube combustion of carbonate-free rock powder with CuO

28* f ÷ + + + + + +

÷ + + + + + ~

N

++::iiiiiiii!iiiiii!iiiiiiiiii::,: :,

~

~

++ d- + + + + + + + + + + + + + + + • _A ~_=\DDH-A (~'-j~Vryburg + + -~KlerkSdOr LPZ':: •~ ~

~ ~

:

"

.

:

,

: 7

~

+

*

+

++++ +++ +++~

+

+

+

+

+

÷

+

TRANSVAAL

DDH- C

~:!:!: + Y'~

- - ] C o v e r rocks D

Bushveld Complex

(~) Drill cores ~

Olifantshoek Group

28 ° -

[ ] T r a n s v a a l Supergroup [ ~ Basement

GRIQUALAND W~EST :"~rie~

~ 240

Oi

1001

200[ km

28*1

Fig. 2. Simplified geological map of the area showing the locations of ddllholes DDH-A, DDH-B and DDH-C.

60

H. Strauss. N.J. Beukes / Precambrian Research 79 (1996) 57-71

Table 1 Abundances and isotopic compositions of organic carbon and pyrite sulfur, S / C ratios, iron abundances and degree of pyritization for Transvaal Supergroup sediments Sample

Depth (m)

Mapedi Shale, DDH A B 630 357.0 B 631 358.0 B 632 364.5 B 633 365.0 B 634 365.9 B 635 371.3

~345

T.O.C. (rag/g)

~13C (%c, PDB)

S (rag/g)

(%c, CDT)

0.1 0.7 41.6 10.6 22.3 20.1

-24.6

4.3 4.3 13.3 27.2 20.9 14.4

+22.6 +23.6 + 1.9 + 6.6 + 4.9 +0.7

-41.4 - 43.2 - 43.4

5.3 0.9 0.2 0.1

+ 13.0

-38.1 - 38.6 -38.1 -38.0 -40.1 - 38.3 - 40.2 -40.9 -39.5 -41.5 - 41.5 -41.9 -42.7 -41.7 - 37.1

0.6 1.8 2.5 2.8 0.7 1.0 2.7 1.9 1.8 2.9 6.6 2.8 3.9 11.2 245.0

-32.1 -33.2 -33.8 - 31.4

6.0 4.2 2.2 0.2

-37.7 -37.3 - 36.5

-41.0 - 39.0 - 39.7 - 40.4 -41.2 -43.4

Reivilo Formation, DDH B B 598 493.1 20.9 B 599 493.6 5.9 B 600 493.9 6.4 B 601 502.5 0.3 Monteoille Formation, DDH B B 602 514.5 17.4 B 603 515.3 25.0 B 604 518.7 6.0 B 605 522.5 18.9 B 606 523.6 28.8 B 607 532.7 20.7 B 608 543.8 25.3 B 609 546.9 23.6 B 610 549.5 11.7 B 611 553.4 24.1 B 612 558.4 19.7 B 613 560.8 23.7 B 614 563.7 24.6 B615 610.7 30.7 B 616 615.7 10.9 Lokamonna Formation, DDH B B 617 635.5 1.0 B 618 651.3 10.3 B 619 652.5 3.8 B 620 665.1 6.2 Boomplas Formation, DDH B B 621 737.5 4.3 B 622 792.4 1.6 B 623 796.0 0.7 Vryburg Formation, DDH B B 624 820.5 1.3 B 625 850.4 1.9 B 626 921.2 3.6 B 627 940.8 4.0 B 628 969.4 3.4 B 629 1012.8 4.3

- 34.2 - 34.0 - 34.4 -35.1

Fepy r (%)

Fenc I (%)

33.08 6.40 0.32 0.57 0,94 0.72

0.37 0.37 1.16 2.36 1.82 1.25

1.95 2.60 0.40 0.37 0.52 0.17

2.32 2.97 1.56 2.73 2.34 1,42

0.16 0.12 0.97 0.86 0.78 0.88

0.25 0.15 0.03 0.38

0.46 0.08 0.02 0.01

1.96 6.39 2.20 2.18

2.42 6.47 2,22 2,19

0.19 0.01 0.03 0.01

0.04 0.07 0.42 0.15 0.02 0.05 0.11 0.08 0.16 0.12 0.33 0.12 0.16 0.37 22.5

0.06 0.15 0.22 0.24 0.06 0.09 0.24 0.17 0.16 0.25 0.57 0.24 0.34 0.98 21.3

4.10 1.87 6.29 0.47 0.32 0.69

4,16 2,02 6,51 0.71 0.38 0.78

0.01 0.07 0.03 0.34 0.16 0.12

0.98 1.57 1.13 0.70 0.53 0.67 2.00

1.14 1.82 1.70 0.94 0.87 1.65 23.3

0.14 0.15 0.34 0.26 0.39 0.59 0.91

+2.4 + 1.3 + 1.4

6.00 0.41 0.59 0.03

0.52 0.37 0.16 0.02

3.25 4.39 3.23 10.9

3.77 4.76 3.39 10.9

0.14 0.08 0.05 0.01

2.2 3.9 1.5

+3.5 + 1.5 - 1.3

0.51 2.41 1.99

0.19 0.34 0.13

4.71 5.01 4.46

4.90 5.35 4.59

0.04 0.06 0.03

1.8 2.7 0.1 0.6 0.2 1.9

+ 0.4 - 0.1

1.37 1.43 0.04 0.14 0.06 0.43

0.16 0.23 0.01 0.05 0.02 0.16

3.55 3.23 1.99 2.69 3.54 2.31

3.71 3.46 2.00 2.74 3.56 2.47

0.04 0.07 0.01 0.02 0.01 0.06

+ 9.3 + 12.7 + 6.6 + 12.3 + 17.5 + 14.6 +7.1

+0.9

S/C

Fe R (%)

DOP

61

tl. Strauss, N.J. Beukes / Precambrian Research 79 (1996) 57-71

Table 1 (continued) Sample

Depth (m)

T.O.C. (mg/g)

~13C (%~, PDB)

S (mg/g)

0.6 14.3 0.8 14.3 14.9 13.7 3.4 17.0 2.6 27.3 12.9 22.5 23.7 1.0 8.3 0.4 0.4 0.1

- 21.3 -31.3 - 23.0 - 32.6 -36.7 - 36.8 - 36.7 - 37.8 - 34.3 - 36.4 - 35.2 -36.4 - 36.2 - 32.4 -33.1 - 26.4 - 26.5 -26.0

0.1 9.2 0.1 26.7 9.3 0.5 4.6 8.8 2.9 2.3 2.9 2.5 2.1 3.5 4.8 0.1 0.1 1.2

~34S (%,~,CDT)

S/C

Fepyr (%)

FeHo (%)

FeR (%)

DOP

0.17 0.64 0.13 1.86 0.62 0.04 1.35 0.52 1.12 0.08 0.23 0.11 0.09 3.50 0.58 0.25 0.25 12.0

0.01 0.81 0.01 2.33 0.81 0.04 0.40 0.75 0.25 0.20 0.26 0.22 0.18 0.31 0.42 0.01 0.01 0.10

1.29 0.30 1.70 3.03 2.81 2.96 1.10 2.20 0.58 1.49 3.64 0.94 2.77 0.42 6.90 0.09 3.21 0.17

1.30 1.11 1.71 5.36 3.62 3.00 1.50 2.95 0.83 1.69 3.90 1.16 2.95 0.73 7.32 0.10 3.22 0.61

0.01 0.73 0.01 0.43 0.22 0.02 0.27 0.25 0.30 0.12 0.07 0.19 0.06 0.42 0.06 0.09 0.03 0.17

Black Reef, DDH C

B 484 B 485 B 486 B 487 B 488 B 489 B 490 B 491 B 492 B 493 B 494 B 495 B 496 B 497 B 498 B 499 B 500 B 501

1351.5 1355.1 1356.7 1358.9 1364.5 1365.2 1365.7 1366.7 1367.8 1368.4 1368.9 1369.9 1370.5 1370.7 1371.4 1371.8 1373.1 1375.8

at 850°C. Sulfide sulfur abundances were determined via coulometric titration following combustion in an oxygen stream at 1000°C (Lange and Brumsack, 1977). Sample preparation for sulfur isotopic analyses involved the reaction with chromium chloride (Canfield et al., 1986) and subsequent combustion of the resulting Ag2S with V20 5 at 1100°C on-line. All isotopic analyses were performed on a Finnigan M A T 251 mass-spectrometer, with a multicollector. Isotope values are presented in the standard 8-notation. Organic carbon isotopic compositions are reported as per mil deviations from the PDB-standard, while sulfur isotopic compositions are reported as per mil deviations from the CDT-standard. Reproducibility, as determined through replicate analyses, was generally better t h a n _ 0.2%° for carbon and + 0.5%0 for sulfur. Abundances of HCl-leachable iron were determined following the method outlined by Berner (1970) and Leventhal and Taylor (1990). Pyritebound iron was calculated from sulfide-sulfur abundances. The degree of pyritization (DOP) was calculated according to Raiswell et al. (1988) as the ratio of pyrite-bound iron and reactive iron (pyrite-bound iron + H C l - l e a c h a b l e iron). Analytical results are presented in Table 1, and a summary of all data,

+4.1

+4.7

+ 3.6

newly determined and screened from the literature, is given in Table 2.

4. Results and discussion 4.1. O r g a n i c carbon

Abundances and isotopic compositions of organic carbon (T.O.C. and ~13Corg.) display large variations ranging between 0.1 and 41.6 m g / g and between - 4 3 . 4 and -21.3%o, respectively. Two distinct populations of carbon isotope data can be distinguished, a smaller group of samples with ~ 3 C - v a l u e s between - 2 6 . 5 and - 2 1 . 3 % o and the main population with values more negative than - 3 1 % o (Fig. 3). The carbon isotopic composition of sedimentary organic matter records both, primary and secondary processes within the biogeochemical carbon cycle. Thereby, the isotopic composition of TOC is largely governed by the isotopic fractionation associated with primary production. However, subsequem biological reworking involving additional fractionation will alter this primary signal once the residues of these organisms become incorporated in the sediments. The carbon isotope record of sedimentary organic

H. Strauss, N.J. Beukes / Precambrian Research 79 (1996) 57-71

62

Table 2 Summary of carbon and sulfur isotopic compositions in organic carbon and sedimentary sulfides from the Transvaal Supergroup Stratigraphic unit

Lithology

813C (%~)

Mapedi Shale Daspoort Formation Silverton Formation Strubenkoop Formation Timeball Hill Formation Penge Iron-Formation

shale

- 32.5

dolomite

- 19.4

BIF BIF carbonate + shale carbonate chert chert dolomite carbonate shale chert carbonate shale siderite-BIF oxide-BIF shale shale shale shale shale shale

-

chert carbonate limestone dolomite dolomite carbonate shale chert

-

Asbestos Hills Formation Kuruman Iron-Formation Malmani Subgroup Gamohaan Formation

Reivilo Formation Monteville Formation Lokamonna Formation Boomplas Formation Vryburg Formation Black Reef Transvaal Supergroup (undifferentiated)

n

Reference

Mineral

834S (~)

n

Reference

5

12

1

2

pyrite sulfide sulfide sulfide sulfide sulfide pyrite

+ 9.8 + 8.5 - 9.7 + 8.6 - 18.5 - 5.5 - 4.9

3 1 6 4 14 10 10

12 1 1 1 1 1 4

24.6 26.0 32.9 30.7 13.0 37.4 36.1 33.8 39.1 33.2 37.0 37.0 36.1 26.7 42.7 39.9 32.7 37.2 40.8 32.2

1 8 15 21 1 1 1 11 3 4 14 8 8 2 3 15 4 3 6 18

3 5 5 3 3 7 7 8 8 8 5 5 5 5 12 12 12 12 12 12

sulfide

- 0.9

3

5

sulfide sulfide pyrite

- 2.3 + 2.5 - 2.5

25 11 5

1 6 8

pyrite pyrite pyrite pyrite pyrite pyrite sulfide

+ 13.0 + 11.4 + 1.3 + 3.5 + 0.4 + 4.4 + 4.2

1 7 1 1 1 2 1

12 12 12 12 12 12 1

28.0 26.9 37.1 28.9 18.7 25.5 37.0 35.4

1 15 l 1 1 3 4 4

9 2 10 11 3 8 8 8

References: 1 = Cameron (1982); 2 -~ Eichmann and Schidlowski (1975); 3 = Hayes et al. (1983); 4 = Cameron and Garrels (1980); 5 = Beukes et al. (1990); 6 = Bottomley et al. (1992); 7 = Klein et al. (1987); 8 = Strauss and Moore (1992); 9 = Oehler et al. (1972); 10 = Hoering (1967); 11 = Hoering (1962); 12 = this study.

m a t t e r h a s b e e n u s e d s u c c e s s f u l l y in o r d e r to c h a r a c -

severe alteration or even complete loss of primary

terize ancient biochemical pathways of carbon fixa-

signals.

t i o n , e s t i m a t e r a t e s o f c a r b o n b u r i a l in p a s t s e d i m e n -

D a t a r e f l e c t i n g t h e state o f p r e s e r v a t i o n o f s e d i -

t a r y s y s t e m s a n d e v a l u a t e t h e r m a l m a t u r i t y o r alter-

m e n t a r y o r g a n i c matter ( k e r o g e n ) f r o m the Paleopro-

ation (Summons

terozoic Transvaal Supergroup have been presented

a n d H a y e s , 1992).

F o r t h e f i r s t t w o , r e a s o n a b l e p r e s e r v a t i o n is t h e

b y H a y e s et al. ( 1 9 8 3 ) a n d S t r a u s s et al. ( 1 9 9 2 a ) . A

most important requirement, as postdepositional pro-

r e l i a b l e c r i t e r i o n u s e d to a s s e s s t h e p r e s e r v a t i o n o f

cesses

k e r o g e n is its e l e m e n t a l c o m p o s i t i o n ( C , H , N ) , a n d

strongly

affect

the

organic

matter

causing

H. Strauss, N.J. Beukes / Precambrian Research 79 (1996) 57-71

-

This study Ishale

I _

63

-~o~,

Literature [ ] oxide - BIF Ill siderite- BIF r-1 chert

10-

[] carbonate

n

6 t3C -30

[ ] shale

°~a/!he..~J a"e,atmn

Io

-35]C

0

_~1o

o~

IU o O 0

N

o

~

o O0

00

I

I

-40

I

-30

-20

~

I

-10

451 q

6taCo~ ~,~, POel Fig. 3. S u m m a r y of newly determined and previously published organic carbon isotope data (source of literature data: Beukes et al., 1990; Strauss and Moore, 1992).

in particular the H/C-ratio (Durand, 1980). Relevant data are shown in Table 3. Reasonably well preserved organic matter as characterized by an H / C ratio >/0.2 (Strauss et al., 1992b) for this unit is associated with ~t3C-values around -35%o. Organic carbon isotopic compositions substantially heavier likely reflect the effects of diagenesis and thermal degradation, and no information concerning primary signals and related biogeochemical pathways can be obtained from these samples. Based on a global survey, well preserved early Paleoproterozoic (2500-2300 Ma) organic matter ( H / C >/0.2) seems to be characterized by 813C-val ues between - 40 and - 30%o (Strauss et al., 1992b). Primary isotopic compositions might have been even more 13C-depleted as evident from 'recalculated'

o

o

o

0 0

0 o q30

o

o oo

0

o 0

oo ~o

o °

o

TOC ling/g)

Fig. 4. Cross-plot of organic carbon abundances and isotopic compositions determined during this study indicating the effect of postdepositional alteration.

~13Corg" values which take postdepositional thermal alteration into account (Des Marais et al., 1992). Carbon isotope data obtained during this study (predominantly from shales) display a bimodal distribution. The main population of carbon isotope values between - 4 3 . 4 and -31.4%o is comparable to the range of values previously published for the Transvaal Supergroup (Fig. 3) and to other isotope values of Paleoproterozoic age assessed as 'least altered' by Strauss et al. (1992b, c). The smaller population with values between - 2 6 . 5 and -21.3%o is likely a result of postdepositional thermal alteration with samples low in TOC trending to more positive ~13C-values (see Fig. 4). These values characterize samples from the Black

Table 3 Elemental and carbon isotopic compositions of kerogen samples from the Transvaal Supergroup Lithology Carbonate Carbonate Carbonate Carbonate Carbonate Carbonate Carbonate Shale Shale Shale

C

H

N

(%)

(%)

(%)

69.0 52.0 75.7 73.0 78.4 86.0 68.9 71.7 74.5 81.8

0.81 0.86 1.71 1.47 1.71 1.98 1.49 0.77 2.14 2.10

0.23 0.40 0.60 0.63 0.78 0.95 0.74 0.67 0.46 0.40

References: 1 = Hayes et al. (1983); 2 = Strauss et al. (1992c).

H/C

~ 13C

Reference

(%°) 0.14 0.21 0.27 0.25 0.26 0.28 0.26 0.13 0.35 0.31

-

35.0 32.5 32.1 32.7 34.6 34.5 34.6 34.5 36.4 36.7

1 I 2 2 2 2 2 2 2 2

64

H. Strauss, NJ. Beukes / Precambrian Research 79 (1996) 57-71

813Corg"

~13Cearb. e--

¢ • = • I - 40

i

I

Oxide-BIF Siderite-BIF Shale Dolomite Limestone

'

- 30

I

o o o o o i

- 20 ~13C

(°/oo,

Fig. 5. Facies-dependent variations (modified after Beukes et al., 1990).

I

i

- 10

I 0

PDB)

of 13Corg., 13Ccarb" and

A8

Reef, derived from an area known for tectonic deformation and some associated metamorphic overprint. Beukes et al. (1990) presented a most detailed geochemical study of Transvaal Supergroup sediments illustrating the facies dependence of carbon isotopic compositions of carbonate and organic carbon in the transition from limestone to iron-formation deposition. Different depositional environments, a stratified water column and differences in productivity as well as biological a n d / o r inorganic degradation (reworking) were processes used to interpret various degrees of 13C-depletion. Organic matter associated with limestones, shales and siderite-rich banded iron-formations were most depleted in 13C with ~13C-values between - 4 4 and -30%0, with a slight trend of decreasing depletion in the order: limestone ~ shale ~ siderite-rich BIF (see Fig. 3). The most lac-enriched organic carbon samples discussed by Beukes et al. (1990) were from oxide-rich BIF, where the oxidation of labile organic matter by ferric iron most likely resulted in the concentration of refractory organic matter. Even less depleted carbon isotope values between - 2 0 and -13%o (as illustrated in Fig. 4), were interpreted as a result of postdepositional thermal alteration. Transvaal carbonates for comparison, both limestones and dolomites, display carbon isotopic compositions of - 1 . 2 ___1.1%o (Gamohaan Formation, Beukes et al., 1990) and - 0 . 9 ___0.7%0 (Malmani Formation, Veizer et al., 1992). This results in a A~, the isotopic difference between the carbonate and organic carbon isotopic compositions, of around 35%0, close to the maximal isotopic fractionation

(J3C-depletion) assumed possible during production of primary photosynthetic organic carbon at that time (for discussion see Strauss et al., 1992c). Facies-dependent variations in 8'3Ccarb, 5~3Corg" or A~ for sediments from the Transvaal Supergroup have been discussed in detail by Beukes et al. (1990) and are illustrated in Fig. 5. Partially, these are interpreted to reflect primary geochemical and physical variations within the depositional environment (i.e. low vs. high productivity, shallow vs. deep water, well mixed vs. stratified, hydrothermal input, Fe-content and O2-content). Superimposed on primary variations are isotopic variations related to different biological (i.e., sulfate reduction, denitrification) and abiological (thermal alteration, metamorphism) reworking of organic matter. In summary, carbon isotope data obtained for TOC from Transvaal Supergroup sediments encode different lines of information. Considering only reasonably well preserved organic matter ( H / C ~> 0.2), ~'3Corg" values ranging from - 43.4 to - 31.4%o can be interpreted primarily as a result of biochemical processes. As pointed out above, the TOC-based isotope values record isotope fractionations associated with primary production and secondary reworking. In particular, a resulting A8 as large as 43%o, thus exceeding the maximum fractionation involved with primary production of biomass via the C 3Calvin-cycle, requires some additional process which increases this parameter. While processes involving respiratory metabolism tend to decrease the A8 (Hayes et al., 1989), fermentative processes increase this isotopic difference between carbonates and organic carbon (Hayes et al., 1987; Boreham et al., 1989). A variety of reactions including methanogenesis, sulfide- or ammonia-oxidation result in biomass depleted in 13C. These processes require distinct redox conditions achieved for example through water column stratification. The presence of such biochemical pathways and resulting products in sedimentary organic matter of Phanerozoic age has been demonstrated (i.e., Kepkay et al., 1979; Summons and Powell, 1986, 1987; Hayes et al., 1987; Freeman et al., 1990). Accumulation of these materials in Transvaal Supergroup sediments might explain the observed organic carbon isotope values. Unfortunately, organic geochemical techniques (i.e., biomarker studies) yielding clear evidence for the

H. Strauss, N J . Beukes / Precambrian Research 79 (1996) 57-71

operation of these processes are inhibited due to the state of preservation of the organic material.

+ • 0 o

4.2. Sulfide sulfur

Reivilo Formation__.~.~3

MontevilleFormation Lokamonna Formation. BoomplasFormation/2

x Vryburg Formation O Black Reef

10--

Variable contents of sulfide sulfur between 0.1 and 27.2 m g / g have been determined for sediments analyzed here (a single value of 245.0 m g / g has been measured for a sample with pyrite concretions). Most S / C ratios vary between 0 and 2, with an average value of 0.4. Reactive iron contents, the sum of pyrite-bound and HCl-leachable iron, range from 0.3 to 10.9%. The majority of samples studied here display DOPvalues < 0.5, with only a few Mapedi Shale samples showing values between 0.8 and 0.9. Finally, a range of sulfur isotopic compositions from - 1 . 3 to + 23.6%0 has been measured for sedimentary pyrite from the Transvaal Supergroup and the Mapedi Shale, Olifantshoek Group. Bacterial sulfate reduction is one of the key processes in the biological sulfur cycle. In modem marine sediments it is responsible for the reworking of a major part of the organic matter (Jorgensen, 1982), thus resulting in a coupling of both geochemical cycles. The evolution of bacterial sulfate reduction through time and particularly the timing of its onset in Earth's history has been discussed and is largely controversial (for reviews see Schidlowski, 1989; Lambert and Donnelly, 1990; Hayes et al., 1992; Ohmoto, 1992). The discussion in all cases was chiefly based upon the sulfur isotopic composition of sedimentary (and presumably biogenic) pyrite. In principle, the process of bacterial sulfate reduction and resulting formation of sedimentary pyrite involves three basic constituents: organic carbon to fuel the reduction process, hydrogen sulfide as the result of sulfate reduction, and reactive iron which fixes the H2S as iron sulfide. The interpretation of the abundances of each of these reactants and some derivative parameters (i.e., DOP) has been used successfully to characterize the geochemical and depositional environment for clastic sediments of both Phanerozoic (i.e., Bemer and Raiswell, 1983, 1984; Raiswell and Bemer, 1985, 1986; Raiswell et al., 1988) and Proterozoic age (i.e., Donnelly and Crick, 1988; Donnelly and Jackson, 1988; Carrigan and Cameron, 1991; Donnelly and Crick, 1992). Abun-

65

S (rag/g)-

1

/ /

~

0

~IG_~ •

0

2 0

0

,d/+ 0

3

, 10

o

"

, 20

,

"~}

,

30

Corg" (rag/g)

Fig. 6. Sulfur/carbon cross-plot for sediments from the lower Transvaal Supergroup.

dances of sulfur, pyrite-bound iron (Fepyr), reactive iron (Fe R) and degree of pyritization (DOP), determined for Transvaal Supergroup sediments during this study, show substantial variations (Table 1). The principle information potential of C, S, Fe, DOP data lies in a clear evidence for biogenicity of the sedimentary sulfide and a characterization of the geochemical framework of its formation (i.e., limiting factors). The degree of information decreases through effects of postdepositional alteration on one or all of these parameters. Organic carbon isotope data discussed above already indicated variable, sometimes even severe thermal alteration for some of the Transvaal Supergroup sediments studied. Still, some relevant information can be obtained from S/C-ratios and DOP-values. Organic carbon and sulfide sulfur data (Fig. 6) are quite variable, but sediments show an increase in sulfur abundance with increasing organic carbon contents. This indicates the coupling of both cycles which, analogues to Phanerozoic examples, can be interpreted as evidence for biogenic pyrite formation. An S/C-ratio of 0.36 is commonly used as a reference value for Phanerozoic normal marine shales (Berner, 1984). Despite the scatter of the Transvaal Supergroup data, differences exist for individual stratigraphic units. Samples from the Boomplas and Lokamonna formations display S/C-ratios slightly higher than this

H. Strauss, N J . Beukes / Precambrian Research 79 (1996) 57-71

66

01 ==5 ~12 1~6 1;33 FR7 10-

~4

n

0 0

0.5 DOP

1.0

Fig. 7. Degree of pyritization (DOP) for sediments from the lower Transvaal Supergroup and Olifantshoek Group. 1 = Black Reef; 2 = Vryburg Formation; 3 = Boomplas Formation; 4 = Lokamonna Formation; 5 = Monteville Formation; 6 = Reivilo Formation; 7 = Mapedi Shale.

reference value. This could indicate an enhanced amount of pyrite formed per unit organic carbon, resulting from an enhanced preservation of metabolizable organic material, or the additional presence of pyrite which was formed in the water column (Raiswell and Bemer, 1985). Both processes would require oxygen-free or at least oxygen-deficient bottom-waters. In contrast, S/C-ratios for the overlying Monteville and Reivilo formations are in general substantially lower than 0.36. Analogous to Phanerozoic conditions, this could be indicative of a lowsulfate environment (Berner, 1984), where a proprotionally higher percentage of organic carbon is preserved in the sediment instead of being consumed during sulfate reduction. Finally, Black Reef samples show quite variable S/C-ratios between < 0.1 and 3.5. This variability might be in part due to different degrees of postdepositional alteration as indicated by organic carbon results. Samples B0484, 486, 499,

t

ri o

500, 501, for example, are characterized by comparably low organic carbon abundances and relatively heavy (less U~C_depleted) carbon isotopic compositions. This, however, indicates fairly severe thermal alteration of the sedimentary organic matter (see above) and, thus, the preferential loss of a single parameter. The degree of pyritzation (DOP), the ratio of pyrite-bound and reactive iron, has been used as indicator of bottom-water oxygenation and the position of the H2S/O2-interface (Raiswell and Bemer, 1985). It is a measure, whether sufficient iron was available to capture all the H2S produced during bacterial sulfate reduction. In principle, DOP-values are considerably lower for normal marine than for euxinic or semi-euxinic sediments. Low oxygen to oxygen-free bottom-water conditions reduce the amount of organic carbon recycling in the water and, thus, enhance the burial of more metabolizable organic material and increase the rate of sulfate reduction. In addition, pyrite formed in such an anoxic water column is being buried, together with pyrite formed in the sediment (Raiswell and Bemer, 1985). Both processes account for a higher sulfur content. Generally, sediments from euxinic and semi-euxinic environments show DOP-values well above 0.5. Reasonably high contents of reactive iron and low to moderate DOP-values < 0.5 for most samples analyzed during this study suggest that iron has not been the limiting factor in sedimentary pyrite formation. Only a few Mapedi Shale samples suggest a deficiency in reactive iron (Fig. 7). In terms of bottom-water oxygenation, this leaves the site of bacterial sulfate reduction below the sediment-water interface. The range of sulfur isotope values measured here essentially represents the positive end of the total

[] Literature data

Fin i

-30

n

b.,-m,.

,,,,,n,7!,,,,n i

-20

,

i

-10

1

0 634S (%o,CDT)

i

+10

,,,, i

,

+20

Fig. 8. Summary of newly determined and previously published sulfide sulfur isotope data for sediments from the Transvaal Supergroup. Source of literature data: Beukes et al. (1990); Bottomley et al. (1992), Strauss and Moore (1992).

67

H. Strauss, N.J. Beukes / Precambrian Research 79 (1996) 57-71

in Eastern Transvaal. In contrast, the 34S-depleted values obtained by Cameron (1982) and shown in Fig. 8 derive from units stratigraphically higher, a level which was not sampled for this study. Considering the entire set of available sulfur isotope values for pyrite from the Transvaal Supergroup, stratigraphic differences are indicated (Fig. 9). 834S-values for the Ghaap/Chuniespoort Group ( > 2400 Ma) range from - 1 2 to +22%0, with an average isotopic composition of + 0.7 + 7.5%0 (n = 68). In principle, this range of 834S-values can be caused by bacterial and abiological (thermochemical) sulfate reduction. Considering the observed S / C ratio pattern, even the sedimentary sulfides from the lower part of the Transvaal Supergroup could be interpreted as being of biological origin. In contrast, sedimentary pyrite from the Olifantshoek and Pretoria Groups ( < 2400 Ma) displays sulfur isotope values between - 31 and + 22%0 (avg.: - 8.4 + 13.5%o, n = 27). Apparently, sulfur isotopic compositions in the upper part of the Transvaal Supergroup (i.e. Timeball Hill Formation) are considerably more 34S-depleted than in the lower part and thus more

range of isotope data observed for the Transvaal Supergroup (Fig. 8). No obvious facies-dependency exists for the new data. Samples were taken from shales interbedded with carbonates, which is essentially the same lithology sampled by Cameron (1982) in a detailed sulfur isotope study of Archean and Early Proterozoic sediments from South Africa. Furthermore, no correlation exists between the sulfur isotopic composition and the S/C-ratio or the DOPvalues, which are both parameters related to the process of sulfate reduction and subsequent pyrite formation. The apparent discrepancy between our new data and those published previously can best be explained by looking at the stratigraphic variation of the sulfur isotope data. Sulfur isotope data for Black Reef sediments between + 3.6 and + 4.7%0 compare well with results obtained from the same unit by Cameron (1982). Sulfur isotopic compositions ranging from - 1 . 3 to + 17.5%o from the Schmidtsdrif and Campbellrand Subgroups display a range fairly similar to the one observed by Cameron (1982) and Bottomley et al. (1992) for the lower part of the Malmani Subgroup

Griqualand West

Eastern Transvaal /

Bm

, B_~

Pretoria Group ,

m

~

B

i

Postmasburg Group Ghaap Group i

I t;I

Chuniespoort Group ,

Asbestos Subgroup [-]

,

_ ¢,,vtxtitY I II

~i

i

O!ffantshoekGroup . • ,

I

Penge Iron Formation ~f

I--] l:l

~, L

m

i

[] ~

Campbellrand Subgroup

i

Malmani Subgroup ,

III

I I,I I II I I I I I I I I I r

]

i

i

?<-~ ~ _B!aek _Reef. . . . Sehmtdtsdrff S u b g r o u p . . /

-lo

o

+to +20

~34 s (Oloo,CDT)

f

i

-3o-;o-;o

i

o

+io +2o

~34S (Oloo,CDT) Fig. 9. Stratigraphicvariationsof 34S-values within the TransvaalSupergroup.

68

H. Strauss, N.J. Beukes / Precambrian Research 79 (1996) 57-71

easily accepted as being of a biological origin. This observation has been noted by Cameron (1982) and interpreted as a clear indication for the activity of sulfate-reducing bacteria in a sulfate-rich ocean at that time. He placed the 'isotopic transition' from a sulfate-poor to a sulfate-rich ocean within the lower part of the Malmani Subgroup which is a correlative to the Campbellrand Subgroup in Griqualand West. The suggested concentration of sulfate in the seawater prior to the proposed isotopic transition was < 0.001 mol SO 2-, a level below which the kinetic isotope effect associated with sulfate reduction decreases significantly (Harrison and Thode, 1958). However, other parameters (i.e., rate of sulfate reduction, open vs. closed system conditions with respect to sulfate availability) than the concentration of seawater sulfate also affect the kinetic isotope effect (for a review see Ohmoto, 1992). The sulfur isotopic composition of contemporaneous seawater sulfate has been estimated to be at + 15%o vs CDT, as proposed in a study of trace sulfate (anhydrite) in carbonates from the Malmani Subgroup (Buchanan and Rouse, 1982). This value is quite comparable to other trace sulfate data for Paleoproterozoic carbonates studied by Bottomley et al. (1992) and could well reflect the true seawater signature at that time. Comparing this isotopic composition for seawater sulfate to the entire dataset of sulfide sulfur isotope values, it would result in an isotopic difference between these two parameters (AS) of 27%0 for the most 3as-depleted sulfides. At the same time, interpretation of the 34S-enriched values of up to + 20%0 would indicate, at first glance, that the kinetic isotope effect was negligible or even slightly positive. While the first observation is generally accepted as clear evidence for bacterial sulfate reduction, interpretation of an apparently greatly reduced or even reversed isotopic fractionation requires additional explanation. The most important parameter involved in sulfate reduction is the availability of sulfate at the site of reduction. Its supply depends on the overall transport from the water column into the sediment, which, as a result of missing burrowing organisms during most of the Proterozoic, was merely based on diffusion. In addition, this initial concentration at the site of reduction is continuously decreasing if the replenish-

ment through diffusion is less than the consumption during reduction. The speed of this decrease in sulfate availability depends on the rate of sulfate reduction. This is largely dependent on the availability of metabolizable organic matter. Processes affecting the sulfate concentration during bacterial sulfate reduction result in a decrease of the kinetic isotope effect as well as in a unidirectional evolution of the isotope system towards heavier, 3as-enriched values for the residual sulfate and the resulting sulfide. Thus, a scenario involving an initially low sulfate concentration in the pore waters being rapidly exhausted by an increased rate of sulfate reduction due to an enhanced availability of metabolizable organic matter could explain the total range of sulfur isotope values observed for the Transvaal Supergroup sediments. The question remaining relates to the significance of the observed stratigraphic variation in ~34S. An increase in the concentration of seawater sulfate and, thus, an increase in the kinetic isotope effect associated with bacterial sulfate reduction, during the time of deposition of the Transvaal Supergroup sediments, as suggested by Cameron (1982), would agree well with independent data from paleosols or organic carbon burial (Des Marais et al., 1992) for a rise in atmospheric oxygen at that time. It contradicts, however, with an alternative interpretation of the Archean sulfur isotope record by Ohmoto (1992) suggesting a seawater sulfate concentration of more than 1 / 3 of the modern ocean already as far back as 3.5 Ga ago. Accordingly, a variable and reduced isotopic difference between parental sulfate and resulting biogenic sulfide would have been caused by variable but higher rates of sulfate reduction. The proposed 'isotopic transition' (Cameron, 1982) could, thus, indicate a net increase in the kinetic isotope effect as a result of an increase in seawater sulfate concentration o r / a n d a decrease in the rate of bacterial sulfate reduction as a result of less available metabolizable organic matter. A final explanation might be derived through a detailed isotope study of individual pyrite grains instead of bulk rock analyses, as suggested by Ohmoto (1992). It should be noted, however, that the proposed principle process of pyrite formation for these sulfides is bacterial sulfate reduction. Variable biogeochemicai conditions at the site of reduction and during the continuous evolution of this process in the sediment (i.e., availability of sulfate and me-

H. Strauss, N,I. Beukes / Precambrian Research 79 (1996) 57-71

tabolizable organic matter, rate o f sulfate reduction) h a v e caused the isotopic variability o b s e r v e d .

5. Conclusions G e o c h e m i c a l e v i d e n c e f r o m w e l l p r e s e r v e d sediments f r o m the T r a n s v a a l S u p e r g r o u p results in the f o l l o w i n g c o n c l u s i o n s r e g a r d i n g the b i o g e o c h e m i c a l carbon and sulfur cycles. C a r b o n isotopic c o m p o s i tions for m i n i m a l l y altered organic m a t t e r ( k e r o g e n ) ranging f r o m - 4 3 to - 3 1 % o indicate p r i m a r y production via autotrophic carbon fixation, i n v o l v i n g close to m a x i m a l isotopic fractionation. In addition, ~3C-depleted m e t a b o l i c products f r o m subsequent c o n s u m e r s contributed to the total organic carbon pool. F a c i e s - d e p e n d e n t variations in the isotopic c o m p o s i t i o n s o f organic and carbonate carbon are caused by both primary (depositional) and secondary (postdepositional) processes. S e d i m e n t a r y pyrites are thought be a result o f bacterial sulfate reduction. Variable ~34S-values b e t w e e n - 31 and + 22%o for the entire dataset indicate differences in sulfate supply a n d / o r c h a n g e s in the rate o f sulfate reduction. The latter is likely a result o f c h a n g i n g proportions o f p r e s e r v e d m e t a b o l i z a b l e organic matter. T h e ultimate significance o f an apparent t e m p o r a l trend in the sulfur isotopic c o m p o s i t i o n r e m a i n s controversial and awaits detailed analysis with h i g h e r spatial resolution (i.e., individual pyrite grains).

Acknowledgements A c c e s s to core material and support during field w o r k in 1989 by G o l d Fields o f South A f r i c a Ltd. is gratefully a c k n o w l e d g e d . A n a l y t i c a l w o r k at the Ruhr-Universifftt B o c h u m , G e r m a n y , was financed by the D e u t s c h e F o r s c h u n g s g e m e i n s c h a f t (Grant V e 112/1-1/1-2).

References Barton, E.S., Altermann, W., Williams, I.S. and Smith, C.B., 1994. U-Pb zircon age for a tuff in the Campbell Group, Griqualand West Sequence, South Africa: implications for Early Proterozoic rock accumulation rates. Geology, 22: 343346.

69

Berner, R.A., 1970. Sedimentary pyrite formation. Am. J. Sci., 268: 1-23. Berner, R.A., 1984. Sedimentary pyrite formation: an update. Geochim. Cosmochim. Acta, 48: 605-615. Bemer, R.A. and Raiswell, R., 1983. Burial of organic carbon and pyrite sulfur in sediments over Phanerozoic time: a new theory. Geochim. Cosmochim. Acta, 47: 885-862. Berner, R.A. and Raiswell, R., 1984. C / S method for distinguishing freshwater from marine sedimentary rocks. Geology, 12: 365-368. Beukes, N.J., 1983. Palaeoenvironmental setting of iron-formations in the depositional basin of the Transvaal Supergroup, South Africa. In: A.F. Trendall and R.C. Morris (Editors), Iron-formations: Facts and Problems. Elsevier, Amsterdam, pp. 131-209. Beukes, N.J., 1984. Sedimentology of the Kuruman and Griquatown Iron-Formations, Transvaal Supergroup, Griqualand West, South Africa. Precambrian Res., 24: 47-84. Beukes, N.J., 1986. The Transvaal Sequence in Griqualand West. In: C.R. Anhaeusser and S. Maske (Editors), Mineral Deposits of Southern Africa. Geological Society of South Africa, Johannesburg, pp. 819-828. Beukes, N.J. and Klein, C., 1992. Models for iron-formation deposition. In: J.W. Schopf and C. Klein (Editors), The Proterozoic Biosphere: A Multidisciplinary Study. Cambridge University Press, Cambridge, pp. 147-151. Beukes, N.J. and Smit, C.A., 1987. New evidence for thrust faulting in Griqualand West, South Africa: implications for stratigraphy and the age of red beds. S. Afr. J. Geol., 90: 378-394. Beukes, N.J., Klein, C., Kaufman, A.J. and Hayes, J.M., 1990. Carbonate petrography, kerogen distribution, and carbon and oxygen isotope variations in an early Proterozoic transition from limestone to iron-formation deposition, Transvaal Supergroup, South Africa. Econ. Geol., 85: 663-690. Boreham, C.J., Fookes. C.J.R., Popp, B.N. and Hayes, J.M., 1989. Origins of etioporphyrins in sediments, evidence from stable isotopes. Geochim. Cosmochim. Acta, 53: 2451-2455. Bottomley, D.J., Veizer, J., Nielsen, H. and Moczydlowska, M., 1992. Isotopic composition of disseminated sulfur in Precambrian sedimentary rocks. Geochim. Cosmochim. Acta, 56: 3311-3322. Buchanan, D.L. and Rouse, J.E., 1982. Role of contamination in the precipitation of sulphides in the Platreef of the Bushveld complex. Proc. IGCP Projects 91 and 161, 3rd Nickel Sulphide Field Conference, Perth, W.A., pp. 141-146. Cameron, E.M., 1982. Sulphate and sulphate reduction in early Precambrian oceans. Nature, 296: 145-148. Cameron, E.M. and Garrels, R.M., 1980. Geochemical compositions of some Precambrian shales from the Canadian Shield. Geochim. Cosmochim. Acta, 47: 1069-1074. Canfield, D.E., Raiswell, R., Westrich, W.T., Reaves, C.M. and Bemer, R.A., 1986. The use of chromium reduction in the analysis of reduced inorganic sulfur in sediments and shales. Chem. Geol., 54: 149-155. Carrigan, W.J. and Cameron, E.M., 1991. Petrological and stable isotope studies of carbonate and sulfide minerals from the

70

H. Strauss, N.J. Beukes / Precambrian Research 79 (1996) 57-71

Gunflint Formation, Ontario: evidence for the origin of early Proterozoic iron-formation. Precambrian Res., 52: 347-380. Des Marais, D.J., Strauss, H., Summons, R.E. and Hayes, J.M., 1992. Carbon isotope evidence for the stepwise oxidation of the Proterozoic environment. Nature, 359: 605-609. Donnelly, T.H. and Crick, I.H., 1988. Depositional environment of the Middle Proterozoic Velkerri Formation in northern Australia: geochemical evidence. Precambrian Res., 42: 165172. Donnelly, T.H. and Crick, I.H., 1992. Biological and abiological sulfate reduction in two Northern Australian Proterozoic basins. In: M. Schidlowski, S. Golubic, M.M. Kimberley, D.M. McKirdy, P.A. Trudinger (Editors), Early Organic Evolution. Springer-Veflag, Berlin, pp. 398-407. Donnelly, T.H. and Jackson, M.J., 1988. Sedimentology and geochemistry of a Mid-Proterozoic lacustrine unit from northem Australia. Sediment. Geol., 58: 145-169. Durand, B., 1980. Kerogen. Editions Technip, Paris, 519 pp. Eichmann, R. and Schidlowski, M., 1975. Isotopic fracrionation between coexisting organic carbon-carbonate pairs in Precambrian sediments. Geochim. Cosmochim. Acta, 39: 585-595. Eriksson, K.A. and Clendenin, C.W., 1990. A review of the Transvaal Sequence, South Africa. J. Aft. Sci., 10:101-116. Freeman, K.H., Hayes, J.M., Trendal, J.M. and Albrecht, P., 1990. Evidence from carbon isotope measurements for diverse origins of sedimentary hydrocarbons. Nature, 343: 254-256. Hamilton, J., 1977. Sr isotope and trace element studies of the Great Dyke and Bushveld mafic phase, and their relations to early Proterozoic magma genesis in Southern Africa. J. Petrol., 18: 24-52. Harrison, A.G. and Thode, H.G., 1958. Mechanism of the bacterial reduction of sulphate from isotope fractionarion studies. Trans. Faraday Soc., 54: 84-92. Hayes, J.M., Kaplan, I.R. and Wedeking, K.W., 1983. Precambrian Organic Geochemistry, Preservation of the Record. In: J.W. Schopf (Editor), Earth's Earliest Biosphere: Its Origin and Evolution. Princeton Univ. Press, Princeton, N.J., pp. 93-135. Hayes, J.M., Takigiku, R., Ocampo, R., Calloth, H.J. and Albrecht, P., 1987. Isotopic compositions and probable origins of organic molecules in the Eocene Messel Shale. Nature, 329: 48-51. Hayes, J.M., Popp, B.N., Takigiku, R. and Johnson, M.W., 1989. An isotopic study of biogeoehemical relationships between carbonates and organic carbon in the Greenhorn Formation. Geochim. Cosmochim. Acta, 53: 2961-2972. Hayes, J.M., Lambert, I.B. and Strauss, H., 1992. The sulfur-isotopic record. In: J.W. Schopf and C. Klein (Editors), The Proterozoic Biosphere: A Multidisciplinary Study. Cambridge University Press, Cambridge, pp. 129-131. Hoering, T.C., 1962. The stable isotopes of carbon in the carbonate and reduced carbon of Precambrian sediments. Carnegie Inst. Washington, Yearb., 1961/1962: 190-191. Hoering, T.C., 1967. Criteria for suitable rocks in Precambrian organic geochemistry. Carnegie Inst. Washington, Yearb., 1967: 365-372.

Holland, H.D. and Beukes, N.J., 1990. A paleoweathering profile from Griqualand West, South Africa: evidence for a dramatic rise in atmospheric oxygen between 2.2 and 1.8 b.y.b.p. Am. J. Sci., 290-A: 1-34. Jahn, B.-M., Bertrand-Sarfati, J., Morin, N. and Mace, J., 1990. Direct dating of stromatolitic carbonates from the Schmidtsdrif Formation (Transvaal Dolomite), South Africa, with implications on the age of the Ventersdorp Supergroup. Geology, 18: 1211-1214. Jorgensen, B.B., 1982. Mineralization of organic matter in the sea bed--the role of sulphate reduction. Nature, 296: 643-645. Kepkay, P.E., Cooke, R.C. and Novitsky, J.A., 1979. Microbial autotrophy: a primary source of organic carbon in marine sediments. Science, 204: 68-69. Klein, C., Beukes, N.J. and Schopf, J.W., 1987. Filamentous microfossils in the early Proterozoic Transvaal Supergroup: their morphology, significance and paleoenvironmental setring. Precambrian Res., 36: 81-94. Lambert, I.B. and Donnelly, T.H., 1990. The paleoenvironmental significance of trends in sulphur isotope compositions in the Precambrian: a critical review. In: H.K. Herbert and S.E. Ho (Editors), Stable Isotopes and Fluid Processes in Mineralization. Univ. Western Aust. Spec. Publ., 23: 260-268. Lange, J. and Brumsack, H.-J., 1977. Total sulphur analysis in geological and biological materials by coulometric titration following combustion. Fresenius Z. Anal. Chem., 286: 361366. Leventhal, J.S. and Taylor, C., 1990. Comparison of methods to determine degree of pyritizarion. Geochim. Cosmochim. Acta, 54: 2621-2625. Oehler, D.Z., Schopf, J.W. and Kvenvolden, K.A., 1972. Carbon isotopic studies of organic matter in Precambrian rocks. Science, 175: 1246-1248. Ohmoto, H., 1992. Biogeochemistry of sulfur and the mechanisms of sulfide-sulfate mineralization in Archean oceans. In: M. Schidlowski, S. Golubic, M.M. Kimberley, D.M. McKirdy, P.A. Trndinger (Editors), Early Organic Evolution. SpringerVerlag, Berlin, pp. 378-397. Raiswell, R. and Berner, R.A., 1985. Pyrite formation in euxinic and semi-euxinic sediments. Am. J. Sci., 285: 710-724. Raiswell, R. and Berner, R.A., 1986. Pyrite and organic matter in Phanerozoic normal marine shales. Geochim. Cosmochim. Acta, 50: 1967-1976. Raiswell, R., Buckley, F., Berner, R.A. and Anderson, T,F., 1988. Degree of pyritization of iron as a paleoenvironmental indicator of bottom-water oxygenation. J. Sediment. Petrol., 58: 812-819. S.A.C.S. (South African Committee for Stratigraphy), 1980. Stratigraphy of South Africa. Geological Survey of South Africa, Pretoria, Handbook, 8, 690 pp. Schidlowski, M., 1989. Evolution of the sulphur cycle in the Precambrian. In: P. Brimblecombe and A. Yu. Lein (Editors), Evolution of the Global Biogeochemical Sulphur Cycle. John Wiley, Chichester, U.K., pp. 3-19. Schidlowski, M., Hayes, J.M. and Kaplan, I.R., 1983. Isotopic inferences of ancient biochemistries: carbon, sulfur, nitrogen,

H. Strauss, N.J. Beukes / Precambrian Research 79 (1996) 57-71

and hydrogen. In: J.W. Schopf (Editor), Earth's Earliest Biosphere: Its Origin and Evolution. Princeton Univ. Press, Princeton, N.J., pp. 149-187. Strauss, H. and Moore, T.B., 1992. Abundances and isotopic compositions of carbon and sulfur species in whole rock and kerogen. In: J.W. Schopf and C. Klein (Editors), The Proterozoic Biosphere: A Multidisciplinary Study. Cambridge Univ. Press, Cambridge, pp. 709-797. Strauss, H., Des Marais, D.J., Summons, R.E. and Hayes J.M., 1992a. Concentrations of organic carbon and maturities and elemental compositions of kerogens. In: J.W. Schopf and C. Klein (Editors), The Proterozoic Biosphere: A Multidisciplinary Study. Cambridge Univ. Press, Cambridge, pp. 95-99. Strauss, H., Des Marais, D.J., Summons, R.E. and Hayes J.M., 1992b. The carbon-isotopic record. In: J.W. Schopf and C. Klein (Editors), The Proterozoic Biosphere: A Multidisciplinary Study. Cambridge Univ. Press, Cambridge, pp. 117127. Strauss, H., Des Marais, D.J., Hayes J.M. and Summons, R.E., 1992c. Proterozoic organic carbon--its preservation and isotopic record. In: M. Schidlowski, S. Golubic, M.M. Kimberley, D.M. McKirdy, P.A. Trudinger (Editors), Early Organic Evolution. Springer-Verlag, Berlin, pp. 203-211.

71

Trendall, A.F., Compston, W., Williams, I.S., Armstrong, R.A., Amdt, N.T., McNaughton, N.J., Nelson, D.R., Barley, M.E., Beukes, N.J., deLaeter, J.R., Retief, E.A. and Thome, A.M., 1990. Precise zircon U-Pb chronological comparison of the volcano-sedimentary sequences of the Kaapvaal and Pilbara cratons between about 3.1 and 2.4 Ga (abstr.). 3rd Int. Archaean Symp. Perth, Ext. Abstr. Vol., pp. 81-83. Summons, R.E. and Hayes, J.M., 1992. Principles of molecular and isotopic biogeochemistry. In: J.W. Schopf and C. Klein (Editors), The Proterozoic Biosphere: A Multidisciplinary Study. Cambridge Univ. Press, Cambridge, pp. 83-93. Summons, R.E. and Powell, T.G., 1986. Chlorobiaceae in Paleozoic seas revealed by biological markers, isotopes and geology. Nature, 319: 763-765. Summons, R.E. and Powell, T.G., 1987. Identification of aryl isoprenoids in source rocks and crude oils: biological markers for the green sulfur bacteria. Geochim. Cosmochim. Acta, 51: 557-566. Veizer, J., Clayton, R.N. and Hinton, R.W., 1992. Geochemistry of Precambrian carbonates: 4. Early Paleoproterozoic (2.25 ± 0.25 Ga) seawater. Geochim. Cosmochim. Acta, 56: 875-885.