Anomalous hopane distributions at the Permian–Triassic boundary, Meishan, China – Evidence for the end-Permian marine ecosystem collapse

Organic Geochemistry Organic Geochemistry 38 (2007) 52–66 www.elsevier.com/locate/orggeochem

Anomalous hopane distributions at the Permian–Triassic boundary, Meishan, China – Evidence for the end-Permian marine ecosystem collapse Chunjiang Wang

*

Laboratory of Geochemistry and Environmental Sciences, Faculty of Geosciences, China University of Petroleum, Beijing 102249, China Received 5 June 2006; received in revised form 16 August 2006; accepted 31 August 2006 Available online 27 October 2006

Abstract Molecular and carbon isotopic geochemistry of hopanes in marine sediments of the Meishan section in Changxing County, Zhejiang Province of China, were studied to reveal biotic and environmental changes across the Permian–Triassic boundary (PTB) and the Wuchiapingian–Changhsingian boundary (WCB). The hopane distribution at the PTB is characterized by high ratios of moretane/hopane, Tm/Ts, Tm/C30 hopane, C31/C32 hopane and hopane/sterane. This hopane distribution is anomalous for these marine sediments, but it is typical of coal measure strata and soils. Two possible genetic mechanisms for such an anomaly are suggested: (1) these hopanes were terrigenous and most probably originated from acidified soil and peat. This mechanism implies that the marine primary production and aerobic bacteria production at the PTB were extremely low. (2) This hopane distribution was possibly caused by freshening and acidification of the upper water column during the end-Permian to Early Triassic marine stagnation and stratification. This hopane anomaly, which coincides with the end-Permian mass extinction and carbon-isotope excursion, may signal the end-Permian mass extinctions and marine and terrigenous ecosystem collapse. The 13C-enriched hopanes with the fingerprint of typical anoxic marine shales, coupled with an n-C15–n-C17-dominated distribution of n-alkanes, strongly indicate that the organic matter around the WCB at Meishan originated mainly from cyanobacteria. The corresponding positive d13Ccarbonate, and higher total organic carbon and hydrocarbon index (HI) values, strongly suggest that high marine primary production and marine anoxia caused by intermittent cyanobacterial blooms, resulted in the local mass extinction at Meishan.  2006 Elsevier Ltd. All rights reserved.

1. Introduction The end-Permian extinction devastated both marine and terrigenous ecosystems (Erwin, 1994; Hallam and Wignall, 1999; Erwin et al., 2002; *

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Wignall and Twitchett, 2002). A recent overview of the end-Permian mass extinction was provided by Racki and Wignall (2005). Numerous mechanisms for the end-Permian extinction have been discussed. Among them, environmental stress and ecological changes are mostly emphasized. Anoxia has been proposed to play a major role in driving the extinction (Wignall and Twitchett, 1996; Isozaki, 1997;

0146-6380/$ - see front matter  2006 Elsevier Ltd. All rights reserved. doi:10.1016/j.orggeochem.2006.08.014

C. Wang / Organic Geochemistry 38 (2007) 52–66

Grice et al., 2005a). Widespread sulfidic conditions could have caused the biotic crisis in the Late Permian (Nielsen and Shen, 2004; Grice et al., 2005a; Kump et al., 2005). Massive release of hydrogen sulfide to the surface ocean and atmosphere during intervals of oceanic anoxia was also suggested as a killing mechanism that might account for the terrigenous and marine extinction (Grice et al., 2005a; Kump et al., 2005). Decreased marine primary production was suggested as a cause of the distinctively negative carbon-isotope shift that coincides with the end-Permian extinction (Wang et al., 1994; Bowring et al., 1998; Rampino et al., 2002). Marine environmental changes, ascribed to effects of terrestrial ecosystem crisis, are notably related to elevated freshwater input to shallow seas (Kozur, 1998; Korte et al., 2003). Comparisons of the timing of biotic crises on land and in the sea suggested that they were coincident events (Twitchett et al., 2001; Steiner et al., 2003). Abundant fungal remains in Permian–Triassic transition sections from various parts of the world may present evidence for terrigenous ecosystem collapse (Eshet et al., 1995; Visscher et al., 1996; Steiner et al., 2003). Other evidence of terrigenous crisis includes worldwide occurrences of abnormal morphotypes of spores and pollen, which may reflect increased mutagenesis during a period of excessive environmental stress (Visscher et al., 2004; Foster and Afonin, 2005). Looy et al. (2001) proposed that terrigenous ecosystem collapse precedes the negative d13C excursion, whereas plant extinctions take place long after the onset of the isotope event. In recent years, molecular organic geochemical analyses on the biotic and environmental changes during the end-Permian biotic crisis have been carried out (Sephton et al., 1999, 2002, 2005; Schwab and Spangenberg, 2004; Grice et al., 2005a,b,c; Summons et al., 2005; Wang et al., 2005a,b; Watson et al., 2005; Xie et al., 2005; Wang and Visscher, 2006). Anomalous distributions of isorenieratane and aryl isoprenoids (Grice et al., 2005a), 2-methylhopane (Xie et al., 2005) and pristane (Wang et al., 2005a) near the PTB at the Meishan section, China, the Global Stratotype Section and Point (GSSP) for the Permian–Triassic boundary (PTB), reflect dramatic changes in marine ecosystems. An unusual abundance of oxygen-containing aromatic compounds (such as dibenzofuran) was discovered close to the end-Permian extinction interval in shallow-marine sections in northern Italy (Sephton et al., 1999,

53

2002) and in slope facies at the Meishan section (Wang and Visscher, 2006). These compounds represent land-derived diagenetic products of polysaccharides, and may be indicative of massive soil erosion (Sephton et al., 2005). Hopanoids are a complex group of biomarkers that occur ubiquitously in sediments (Rohmer et al., 1984). Derivatives of C35 bacteriohopanepolyol (BHP) are exclusively synthesized by prokaryotic organisms and are used as membrane rigidifiers. They occur predominantly in aerobic bacteria (i.e., methanotrophs, heterotrophs and cyanobacteria) (Rohmer et al., 1992; Sinninghe Damste´ et al., 2004). As biomarkers of bacteria, the distributions of hopanoids and their carbon-isotopic compositions play an important role in paleoenvironmental reconstruction. In order to obtain further information about the marine and terrigenous ecological crises around the PTB, this paper interprets hopane biomarkers around the PTB at the Meishan section. Apart from the PTB, the Meishan section also includes the GSSP for the Wuchiapingian–Changhsingian boundary (WCB) (Jin et al., 2003). The WCB is characterized by the extinction of a significant number of marine fauna (Jin et al., 1995; Knoll et al., 1996; Bowring et al., 1998). Based on molecular and carbon-isotopic geochemistry of hopanes, the biotic and environmental changes around the WCB at Meishan section will also be discussed briefly. 2. Samples and experimental 2.1. Description of the key beds of the PTB and WCB and the samples The PTB section at Meishan is exposed in closely-spaced limestone quarries (abandoned) near Meishan in Changxing County of Zhejiang Province, China (Fig. 1). The quarries are labeled A–E and Z from west to east, among which D serves as the GSSP of the PTB. Detailed descriptions of the PTB and WCB are given by Yin et al. (1996, 2001) and Jin et al. (2003), respectively. Here, a brief description with some basic geochemical data (Table 1), is given for the samples from Beds 24– 30 and Beds 1–6, which are the key beds of the PTB and WCB, respectively. A simplified lithological column is given in Fig. 2. Bed 30 and Bed 29 share similar lithological characters: light grey marl and grayish black mudstone

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C. Wang / Organic Geochemistry 38 (2007) 52–66

Fig. 1. Geographic setting of the Meishan section in South China (modified after Yin et al., 1996).

in the lower and upper parts, respectively. Total organic carbon (TOC) values range from 0.09% to 0.64% for Bed 30, and only 0.06% to 0.16% for Bed 29. Bed 28 is composed of two layers: a lower yellowish white illite-montmorillonite claystone and an upper yellowish grey to grayish black shale of locally variable thickness. The average TOC of this bed in Section B is about 0.08%. Beds 28–30 are regarded as a secondary extinction horizon (Yin and Tong, 1998; Jin et al., 2000). Bed 27 is lithologically homogeneous and described as light grey medium-bedded limestone or as grey dolomitic marl. It can be divided into four intervals: 27a–d based on fossil assemblages, of which the 27b/27c division is the Permian/Triassic boundary (Yin et al., 1996). TOC ranges from 0.08% to 0.10% in Bed 27. Bed 26 is black grey mudstone with TOC ranging from 0.82% to 0.89%. This bed has the most negative stable carbon isotopic composition in the section, with d13Ckero of 31.8& to 31.4&. Bed 25 is bluish grey illite-montmorillonite claystone and weathered outcrops usually are milky white. The thickness of this layer is locally variable, with TOC of grey claystone at Section C up to 0.22%. This layer of claystone is regarded as being of volcanic origin (Yin et al., 1996). The base of this layer is regarded as the main mass extinction level (Jin et al., 2000). Bed 24 is dark grey medium-bedded micrite,

with Bed 24e about 10 cm thick at the top, which is composed of black grey thin-bedded micrite. TOC in this interval ranges from 0.22% to 0.79%. Beds 1–6 are mainly composed of dark grey thinto medium-bedded bioclastic micritic limestone, with siliceous bandings or intercalating light grey thin-bedded calcareous mudstone. At Section C, Beds 1–6 are richer in silty or argillaceous limestone than those in Section D, and the top of the upper Longtan formation (Bed 1) is mainly composed of silty and sandy limestone. The Wuchiapingian– Changhsingian boundary is located in the lower part of Bed 4 (Jin et al., 2003). The vitrinite reflectance (Ro %) of the samples from Sections D and B (Table 1) ranges from 0.62% to 0.75% for Beds 2–26 (except for Sample D 9-3), with an average of 0.67% (n = 10); but varies from 1.07% to 1.12% in Beds 29–30. Because the organic matter (OM) in the Meishan section is mainly of medium maturity, thermal destruction of biomarkers should be limited, and molecular geochemical information should be well preserved. 2.2. Experimental Eighty-five whole rock outcrop samples were collected from the Meishan Sections B–D, of which 67 samples were selected for molecular geochemical and 21 samples for compound-specific carbon-isoto-

C. Wang / Organic Geochemistry 38 (2007) 52–66

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Table 1 Basic geochemical data for the samples Sample C32 B30-4a C30-2 C30-1 B30-1 B29-2 C29-3 B29-1 C29-1 B28 B27d B27c B27b B27a B26-2 C26 B26-1 C25 B24e-3 B24e-2 B24e-1 B24d-2 B24d-1 D24d-2 D24d-1 C24d-1 D23-2 D23-1 C22 D22-1 D21 C21 C20 D20 C19 D19-2 D19-1 D18 D16 D15-1 D14-1 D13-1 D12-1 D11-2 D10-2 D10-1 D9-3 D9-1 D8 D8-1 D7 D6-2 C6-4-3 C6-4-2 C6-1 D5/6 C5-10 C5-3

Lithology Grayish black mudst. Grayish black mudst. Grayish black mudst. Bluish grey marl Bluish grey marl Black grey argil. marl Black grey marl Light grey marl Light grey marl Dark grey shale + pale clay Light grey limestone Light grey limest. Light grey limest. Light grey limest. Black grey mudst. Black grey mudst. Black grey mudst. Bluish grey clayst. Grey thin-bedded micrite Grey thin-bedded micrite Dark grey wackst. Dark grey wackst. Dark grey wackst. Grey dolomatic micrite Dark grey dolomatic micrite Dark grey dolomatic micrite Dark grey micrite Dark grey micrite Dark grey micrite Black grey micrite Dark grey micrite Dark grey micrite Black grey micrite Black grey micrite Black grey micrite Black siliceous limest. Black laminar shale Black siliceous limest. Black grey micritic limest. Black grey micritic limest. Black siliceous limest. Black grey micritic limest. Black grey micritic limest. Black siliceous limest. Black limest. Black siliceous limest. Black siliceous laminar shale Black limest. Black limest. Black siliceous shale Black siliceous limest. Black grey limest. Black grey laminar limest. Dark grey shale Dark grey siliceous shale Black laminar shale Black grey micritic limest. Grey black marl

Syst./Fm. b

Tr/Yk Tr/Yk Tr/Yk Tr/Yk Tr/Yk Tr/Yk Tr/Yk Tr/Yk Tr/Yk Tr/Yk Tr/Yk Tr/Yk P2/Yk P2/Yk P2/Yk P2/Yk P2/Yk P2/Yk P2/Chxc P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx

Ro (%)

Carbonate (%)

TOC (%)

HI (mg HC/g TOC)

– 1.07 – – – 1.12 – – – – – – – – – – 0.75 – 0.67 – 0.66 – – 0.64 – – 0.62 – – – – – – 0.62 – – 0.74 – – – – – – – – – 1.05 – – – – – – – – 0.7 – –

– 5.2 – – 29.8 44.6 – 51.3 – 17.2 – – – – 9.7 – 7.2 2.0 79.0 90.4 78.1 – 62.7 – – – – – – – – – – – – – 24.4 – – – 68.5 – 87.4 – – – 21.9 76.9 – 38.4 – – – – – 20.0 – –

0.34 0.64 0.43 0.13 0.09 0.16 0.18 0.06 0.06 0.08 0.09 0.10 0.08 0.09 0.82 0.76 0.89 0.22 0.79 0.24 0.22 0.63 0.84 0.24 0.34 0.28 0.48 0.47 0.08 0.11 0.47 0.25 0.08 0.85 0.19 1.26 2.78 0.17 0.36 0.33 1.64 0.30 0.39 0.21 0.23 1.35 2.61 0.85 0.21 3.05 0.49 0.30 0.70 2.21 0.77 4.39 0.21 2.65

55.9 45.6 64.8 23.5 16.3 31.5 33.4 80.5 32.1 52.8 10.2 10.6 11.4 12.0 36.6 80.0 44.9 49.5 77.3 50.0 54.8 67.8 98.5 66.3 97.8 47.2 73.6 97.3 36.7 36.9 131.8 67.5 26.6 125.3 72.0 139.3 102.1 77.2 82.4 91.9 214.4 23.6 176.0 22.3 123.4 130.0 139.2 63.3 60.7 57.2 19.8 89.8 210.7 151.6 123.8 42.0 19.3 56.9

d13Ckero (&) 28.2 29.3 29.2 25.1 26.0 27.6 28.1 25.8 –

26.2 27.1 26.8 26.8 27.2 31.8 30.7 31.4 29.4 29.0 28.7 28.9 28.8 28.7 28.7 29.9 30.0 29.7 29.2 27.3 28.0 28.2 27.8 27.5 28.0 28.9 29.6 30.0 28.0 27.3 27.8 28.0 28.0 26.8 26.8 27.9 27.9 29.0 29.5 27.4 27.7 27.0 26.7 26.7 27.0 27.1 26.2 27.4 26.8 (continued on next page)

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C. Wang / Organic Geochemistry 38 (2007) 52–66

Table 1 (continued) Sample

Lithology

Syst./Fm.

Ro (%)

Carbonate (%)

TOC (%)

HI (mg HC/g TOC)

D4-3A D4-1 D4-2A D4-2 C4-4 D3 D2-2A D2 C1-2

Black grey micritic limest. Dark grey siliceous shale Dark grey shale Dark grey limest. Black siliceous limest. Dark grey thin-bedded shale Dark grey limest. Dark grey limest. Black thin-bedded marl

P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2/Chx P2 /Ltd

– – – – – 0.62 – 0.63 –

– 29.3 – – – 53.4 – 89.8 –

0.25 3.03 1.97 0.43 1.88 3.34 0.45 0.34 0.87

200.7 227.0 251.3 83.9 291.1 348.8 64.6 52.6 200.1

d13Ckero (&) 26.6 27.2 27.5 27.2 28.0 26.7 25.8 26.3 27.1

Beds

Fm.

Syst.

mudst. = mudstone; limest. = limestone; wackst. = wackestone. d13Ckero = stable carbon isotopic composition of kerogen. a Section B, Bed 30, sample No. 4. b Yinkeng Formation. c Changxing Formation. d Longtan Formation.

Lith./Age(Ma)

Moretane/Hopane 0.0 0.1 0.2 0.3 0.4

13

Tm αβC 29

αβ C 31 homohopane

C30-2

δ C (per mil) -36

-32

-28

-24

-20

Yinkeng

L.Triassic

250 36

250.2 ± 0.2

33

250.4 ± 0.5

31 29-30 28 27

Ts

áâ C 35

C 29Hop C 30Hop

βαC 30 moretane βα C 31 moretane

250.7 ± 0.3

C26

Tm n-alk

251

26

22

Changxing

Upper

Permian

25 24 23

21 20 18-19 17

C25

252 B24e-3 252.3 ± 0.3

15-16 12-14 10-11 8-9 7 2-6

L

251.4 ± 0.3

253

C 29Ts βαC 30 moretane

C5-3

αβC31 homohopane 253.4 ± 0.2

Tm Ts

WCB

1

Fig. 2. Variations in moretane/hopane ratios and carbon-isotope compositions of biomarkers (n-alkanes and hopanes) across the section, with m/z 191 mass chromatograms showing hopane distributions in some representative beds. The hopane distribution and carbon-isotope composition in Beds 24–32 are different from those below Bed 23. The column and age are after Jin et al. (2000). The base of Bed 25 is the level of the main mass extinction (Jin et al., 2000); while the Permian–Triassic biological boundary is located in the middle of Bed 27 with the marker of the first appearance of the H. parvus Zone (Yin et al., 1996). The WCB is located in the lower part of Bed 4 (Jin et al., 2003). ‘‘L’’ on the base of the column is the abbreviation of ‘‘Longtan Formation’’. Arrows indicate location of each m/z 191 chromatogram in the stratigraphic column.

pic measurements. TOC for each whole rock sample was measured by using Leco-WR-112 Carbon Detector. The precision for the TOC measurement is ±1% of the carbon present. Hydrogen index

(HI) was determined by using OGE-II Rock-Eval System made by the Institute of Petroleum Exploration and Exploitation of China National Petroleum Corporation. The precision for the HI measurement

C. Wang / Organic Geochemistry 38 (2007) 52–66

ranges from 610% for S2 values >3 mg/g rock to as much as 50% for S2 values of <0.1 mg/g rock. The samples (100–750 g samples were used depending on TOC) were extracted using a Soxhlet apparatus with dichloromethane/methanol solvent (97:3 v:v). About 150 ml solvent per 100 g sample was used for extraction. The extracts were separated into maltene and asphaltene fractions by precipitation in heptane. The maltenes were fractionated by silica–alumina column (450 mm · 8 mm) chromatography into aliphatic, aromatic and polar fractions by using hexane, benzene, and ethanol elution solvents, respectively. The aliphatic and aromatic hydrocarbons were analysed on a Finnigan SSQ710 GC–MS equipped with a HP-5MS (30 m · 0.25 mm · 0.25 lm) fused silica capillary column. Analyses were conducted in electron impact mode at 70 eV and helium was used as a carrier gas. After an initial period of 1 min at 80 C, the oven was heated to 250 C at 7 C/min, and then to 300 C at 2 C/min, followed by an isothermal period of 15 min. The saturated fraction was further separated into n-alkane and branched/cyclic compounds by adduction using 5A molecular sieve. Carbon isotopic compositions of individual compounds were measured with an Agilent 6890-Isoprime GC–IR–MS system equipped with an HP5 column (30 m · 0.32 mm · 0.25 lm). GC conditions were: injection port temperature was 290 C; initial oven temperature was 60 C, the oven temperature increased at 10 C/min to 120 C, then at 5 C/min to 290 C, followed by an isothermal period of 20 min; oxidation tube temperature was 850 C; splitless injection was used and helium was the carrier gas. Stable carbon isotopic compositions are reported in the delta notation against the VPDB standard. The values reported were determined by duplicate analyses and averaged ±0.5& error. 3. Results and discussion 3.1. Hopane distribution anomaly and its genesis around the PTB 3.1.1. Hopane distribution anomaly around the PTB Fig. 2 gives m/z 191 mass chromatograms showing hopane distributions in the representative beds around the PTB and WCB as well as the variation in moretane/hopane ratios and d13CHop (carbon-isotopic composition of hopane) values across the section. The hopane distribution around the PTB (Beds

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24–30) shows a very different pattern from other intervals below Bed 23. The former is characterized by higher relative abundances of moretanes (bahopanes), Tm and C31ab-hopanes while the latter shows a high relative abundance of C30ab-hopane, with low abundances of moretanes, Tm and C31ab-hopanes. It is interesting that the moretane/ hopane index varies little below Bed 23, but increases very quickly from the top of Bed 23 upward, reaching to a high value in Bed 26. It also shows high values in the upper parts of Bed 29 and Bed 30 as well as in Bed 32. In Fig. 3, the hopane distribution across the section is illustrated by various hopane indices. The variation of moretane/ hopane, Tm/Ts and C31/C32 hopane ratios are well correlated in the intervals from Bed 24 upward, but not below this bed. High moretane/hopane ratio commonly occurs in immature OM regardless of the lithology of sediments, but, when accompanied by high Tm/Ts (or Tm/C30 hopane) and C31/C32 hopane ratios (Fig. 3) in the mature or high maturity stage, it usually indicates terrigenous plant-derived organic matter, such as that in coaly strata (Wang, 1995; Chen et al., 2001). Here, this pattern of hopane distribution is called ‘‘coal-type’’ hopane. This type of hopane distribution occurs in these typical marine sediments (without coal) around the PTB, and can be regarded as a ‘‘hopane distribution anomaly’’. Considering the geochemical role of hopanoids as distinctive biomarkers of bacteria and the onset of this anomaly just below the main marine mass extinction horizon and prior to the end-Permian d13C excursion, the genesis of this anomaly and its relationship to the end-Permian biotic-environmental events must be investigated. As for Beds 12–18, the moretane/hopane ratio does not correlate with other indices (Fig. 3). Some thin-bedded siliceous micritic limestones from these intervals are characterized by high abundances of C2930-norhopane (with C2930-norhopane/C30 hopane ratio of ca. 0.7–2.5) and C28 29,30-bisnorhopane, indicating typical anoxic sedimentary conditions (see Peters et al., 2005, and references therein). 3.1.2. Genesis of the hopane anomaly around the PTB 3.1.2.1. Effect of sedimentary facies on hopane distribution. The variation in moretane/hopane across the section may demonstrate that no relationship exists between moretane enrichment and sedimentary facies (Fig. 2), even though it seems that

L it h .

B eds

C. Wang / Organic Geochemistry 38 (2007) 52–66

Fm.

58

Moretane/Hopane 0.0

0.1

0.2 0.3

0.4 1.2

C 31 /C 32Hopane 1.6

2.0

2.4

10

20

C312mH/CH30

Hopane/Sterane

Tm/Ts 2.8 0

30

1

10

100 0.00

0.03

0.06

0.09

36

Yi n k e n g

33 31 29-30 28 27

PTB

26 25 24 23

EB

C h an g x i n g

22 21 20 18-19 17

15-16 12-14

L

10-11 8-9 7 2-6

WCB

1

Fig. 3. Variations in various hopane geochemical indices across the section. The big difference in hopane distribution between the Beds 24–32 and the Beds below Bed 23 is illustrated clearly. Variations of various hopane indices (except C312mH/C30H) in Beds 24–32 or in Beds 1–6 are well correlated. The strong variations of these indices in Beds 11–15 are discussed in the text.

high moretane/hopane is closely related to the sedimentary facies in the interval of Beds 29 and 30, where moretanes are enriched only in the black grey laminated calcareous mudstone at the upper part of each bed. In addition, it seems that the effect of sedimentary facies does not occur in the intervals from Bed 23 to Bed 26, where moretane/hopane increases gradually from the upper Bed 23 to Bed 26. (Beds 23–24 are homogeneously composed of micrite, but Bed 26 consists of shale). Thus, the hopane anomaly should be related to the source input of organic matter in these marine sediments and/or some special sedimentary conditions. 3.1.2.2. Genesis mechanisms of the ‘‘coal-type’’ hopane. The ‘‘coal-type’’ hopane is characterized by higher relative abundance of moretane, Tm and C31 hopane. Its genesis mechanism needs to be discussed in detail: 1. High moretane/hopane ratio: the primary concentration of moretane precursors may be controlled by acidity and redox potential of depositional environments. The precursor for

17b(H),21a(H)-moretane is seldom discussed in the literature. 17b(H),21a(H)-moretan-29-ol was once reported together with ba- and ab-C32 hopanoids and alcohols in peat samples, and these compounds were suggested to be associated with bacterial activities in acidic, nutrient deficient environments (Quirk et al., 1984). A high abundance (compared with diploptene) of 17b(H)-moret-22(29)-ene was detected in the sediments from four acid lakes (pH ranging from 3.4 to 4.7) and was suggested to be closely associated to acidic conditions (Uemura and Ishiwatari, 1995). As no evidence for direct biosynthesis of these ba-precursors has been found up to now, the two compounds, 17b(H),21a(H)-moretan29-ol and 17b(H)-moret-22(29)-ene are considered to be chemical or microbial conversion products of hopanoids, just like hopanoid acids, in shallow sediments or the water column. Some evidence suggests that the moretane/hopane ratio depends partly on source input or sedimentary environments. For example, the ratio increases from transgressive to highstand system tracts in Lower-Middle Triassic mudstone from the

C. Wang / Organic Geochemistry 38 (2007) 52–66

Barents Sea, corresponding to increasing terrigenous higher-plant input (see Peters et al., 2005, and references therein). Based on a detailed investigation, it is suggested that high abundance of moretane is most probably related to acidic and oxic conditions in peat or soil. 2. High C31/C32 hopane ratio: the relative enrichment of C31 hopanes (high C31/C32 hopane ratio) is usually related to acidic and oxic conditions such as those in peat depositional environments where C32 hopanoid acids are easily formed and can be decarboxylated to form C31 hopanes during diagenesis. For example, Sinninghe Damste´ et al. (2005) suggested that bishomohopanoic acid forms as the dominant hopanoid acid after periodic acid treatment on the lipid extract of Sphagnum plants. The lipid extract contains C35 hopanetetrol derivatives produced by methanotrophic bacteria (including acidophilic methanotrophs) living in/on Sphagnum tissue. In Fig. 4, the C31/C32 hopane and moretane/hopane ratios of the samples from Bed 24 to Bed 32 show a good linear correlation (r2 = 0.74), implying that the ratios in these intervals must have been controlled by the same environmental condition or source input. 3. High Tm/C30 hopane and Tm/Ts ratio: very high Tm/C30 hopane and Tm/Ts (Ts is very low in abundance) are usually found in coals or coaly shales, even at the mature or high maturity stage. It has been demonstrated by using molecular mechanics that Tm (17a-22,29,30-trisnorhopane)

3.0 y = 3.35 x + 1.19

Beds 26-32 shales

2

R = 0.74

C 31/ C 32 h o p a n e

2.5 Beds 24-27 micrites/marls 2.0 Beds 24-32 Beds 19-23

1.5 Beds 28-30 marls

Beds 7-18 Beds 1-6

1.0 0.0

0.1

0.2

0.3

0.4

0.5

Moretane/Hopane Fig. 4. Correlation of C31/C32 hopane ratios with moretane/ hopane ratios, showing the important difference between various intervals and indicating the elevated acidic and oxic conditions in Beds 24–32, especially for mudstones or shales in Beds 26–32.

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is less stable than Ts (18a-22,29,30-trisnorhopane), but it is unknown whether conversion of Tm to Ts may also occur (see Peters et al., 2005). In our data, the very high Tm/Ts ratio at the mature stage suggests that the conversion of Tm to Ts is very limited, and the ratio cannot be explained by absence of acidic clay. Clay minerals are normally present in mudstone from coal measures, and samples like those in Bed 26, Bed 30 and Bed 32 contain clay minerals. Thus, it is suggested that high relative abundance of Tm is mainly a diagenetic product of peat-forming conditions (low pH and high Eh). In view of the genesis mechanisms of the ‘‘coaltype’’ hopane, the hopane anomaly around the PTB is most probably related to acidic and oxic conditions. This means that the OM in these intervals was either predominantly formed in such a depositional setting, or originated from peat, soil, or re-deposited coal. 3.1.2.3. C-isotopic evidence for the OM source in the intervals around PTB. No evidence of re-deposited OM sourced from coaly strata has been found for the marine sediments, which is also strongly supported by the systematic variation of carbon-isotopic composition of different organic molecules, kerogens and carbonates from Bed 24 to Bed 30 (Wang et al., 2005a,b). In Fig. 2, the carbon-isotope composition of hopane (d13CHop) co-varies with that of n-alkanes (and also with that of kerogen and pristane, see Tables 1 and 2, respectively), although d13CHop values of some beds were not obtained due to the very low concentrations of hopanes. The synchronously sharp negative shifts in d13Cmole, d13Ckero and d13Ccarb (carbon-isotopic composition of carbonate) from Beds 24 to 26 and the very 13C-depleted OM in Bed 26 all indicate that the terrigenous OM deposited in marine sediments did not originate from re-deposited coals (which usually have higher values of d13Cmole and d13Ckero, such as ca. 24  26&), but from newly-formed peat or soil resulting from changes of ocean–atmosphere CO2 system. This may be supported by the occurrence of the highly d13C-depleted palaeosols across the Permian–Triassic boundary section from the Antarctic region (Krull and Retallack, 2000). Based on the above discussion, one can infer that the hopane anomaly around the PTB at Meishan results from the contribution of acidified soil or peat to the OM in these marine sediments. This

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C. Wang / Organic Geochemistry 38 (2007) 52–66

Table 2 Molecular geochemical parameters of hopanes and d13Cmole values (&) of biomarkers Sample

Tm/Ts

Tm/C30 hopane

Moretane/ hopane

Gam/C30 hopane

C31/C32 hopane

C312mH/ C30H

Hopane/ sterane

Hopane/ n-alkane

C32 B30-4 C30-2 C30-1 B30-1 B29-2 C29-3 B29-1 C29-1 B28 B27D B27C B27B B27A B26-2 C26 B26-1 C25 B24e-5 B24e-4 B24e-3 B24e-2 B24e-1 D24-2 D24-1 C24-1 D23-2 D23-1 C22 D22-1 D21 C21 C20 D20 C19 D19-2 D19-1 D18 D16 D15-1 D14-1 D13-1 D12-2 D12-1 D11-2 D10-2 D10-1 D9-3 D9-1 D8 D8-1 D7 D6-2 C6-4-3 C6-4-2 C6-1 D5/6 C5-10 C5-3 D4-3A D4-2A C4-1 D4-2

24.94 18.47 19.40 1.66 1.21 19.76 20.14 1.49 1.48 1.03 2.06 3.31 2.74 2.98 19.35 14.46 17.98 7.16 15.90 7.52 3.33 7.40 8.36 2.64 2.56 2.61 2.97 2.09 1.21 1.48 1.68 1.29 0.92 0.76 1.12 0.60 0.65 0.64 0.75 0.74 1.36 2.43 3.29 3.80 3.85 0.64 0.86 0.93 0.87

0.81 0.46 0.59 0.22 0.23 0.67 0.61 0.27 0.19 0.23 0.32 0.38 0.33 0.40 0.54 0.59 0.50 0.53 0.61 0.56 0.36 0.45 0.45 0.32 0.42 0.44 0.34 0.41 0.32 0.31 0.37 0.28 0.30 0.20 0.40 0.22 0.25 0.30 0.31 0.35 0.28 0.45 0.40 0.35 0.39 0.29 0.30 0.26 0.27 0.27 0.24 0.28 0.20 0.20 0.17 0.14 0.14 0.27 0.16 0.19 0.24 0.12 0.17

0.417 0.358 0.409 0.158 0.143 0.319 0.353 0.153 0.149 0.124 0.177 0.235 0.206 0.224 0.320 0.396 0.305 0.308 0.266 0.230 0.178 0.207 0.220 0.142 0.161 0.161 0.097 0.098 0.085 0.070 0.072 0.091 0.088 0.066 0.075 0.067 0.053 0.055 0.060 0.069 0.064 0.032 0.068 0.061 0.058 0.035 0.054 0.094 0.061 0.066 0.072 0.077 0.060 0.085 0.083 0.093 0.063 0.076 0.091 0.074 0.078 0.089 0.083

0.081 0.110 0.070 0.176 0.164 0.033 0.025 0.138 0.129 0.101 0.146 0.114 0.125 0.125 0.051 0.055 0.051 0.079 0.049 0.063 0.055 0.056 0.059 0.096 0.129 0.126 0.097 0.097 0.145 0.107 0.032 0.084 0.132 0.055 0.062 0.084 0.052 0.070 0.030 0.053 0.083 0.153 0.100 0.097 0.115 0.026 0.039 0.052 0.052 0.045 0.041 0.054 0.041 0.019 0.037 0.028 0.019 0.029 0.017 0.031 0.035 0.027 0.072

2.63 2.62 2.38 1.66 1.66 2.48 2.48 1.38 1.73 1.61 1.97 2.20 2.12 2.21 2.37 2.15 2.29 1.89 2.05 1.89 1.71 1.80 1.78 1.49 1.83 1.64 1.51 1.81 1.62 1.53 1.43 1.45 1.41 1.48 1.78 1.41 1.45 1.57 1.54 1.51 1.62 2.03 1.86 1.66 1.63 1.68 1.90 1.66 1.61 1.62 1.61 1.67 1.74 1.49 1.54 1.62 1.48 1.74 1.60 1.48 1.48 1.58 1.59

0.035 0.007 0.018 0.024 0.024 0.084 0.083 0.025 0.030 0.035 0.025 0.034 0.030 0.011 0.077 0.084 0.079 0.023 0.030 0.026 0.022 0.022 0.031 0.037 0.035 0.047 0.036 0.031 0.030 0.038 0.039 0.014 0.046 0.044 0.025 0.059 0.057 0.055 0.068 0.054 0.024 0.01 0.019 0.019 0.015 0.040 0.040 0.051 0.062 0.062 0.052 0.051 0.051 0.047 0.049 0.018 – 0.054 0.015 0.035 0.052 0.041 0.053

28.03 27.89 35.37 2.41 2.74 30.56 29.09 3.53 2.63 3.26 3.33 5.65 4.89 4.36 14.84 10.52 12.80 5.25 5.52 6.03 6.46 5.26 5.56 6.83 3.75 5.03 6.68 3.17 7.19 4.91 3.70 5.64 5.98 4.69 4.90 6.22 6.03 5.78 8.39 6.65 19.26 13.61 17.57 27.54 56.01 5.77 3.64 5.15 5.34 4.53 5.49 5.13 6.36 5.59 3.77 8.34 11.33 5.39 7.15 6.37 4.58 5.64 6.07

0.021 0.130 0.380 0.030 0.033 0.206 0.238 0.045 0.069 0.038 0.031 0.072 0.046 0.038 0.203 0.111 0.192 0.137 0.128 0.127 0.116 0.108 0.105 0.105 0.047 0.019 0.058 0.025 0.041 0.062 – 0.026 0.106 0.055 0.020 0.049 0.036 0.028 0.073 0.016 – 0.036 – – – 0.022 0.041 – –

0.75 0.98 0.46 0.78 1.28 1.18 0.76 1.93 1.66 1.49 1.28 1.07 1.21

d13Cn-alk 33.2 – 32.1 – – –

33. 8

32.0

30.8 –

33.2 33.5 –

34.6 35.0 –

30.1 29.7 – – – – – –

30.3 32.3 – – – – – –

30.6 –

32.3 –

29.8 – – –

31.7 – – –

29.0 – – 28.9 – –

30.5 – – – – –

28.2 –

31.0 –

28.3 – – – – –

29.6 – – – – –

29.2 – – – – – – – –

30.5 – – – – – – – – –

29.4 – – 29.4

30.2

30.8 – – – – – – – – – – – – – – – 24.8 – – – – – – – – – – – – – – – – – – – – – – –

22.7

22.1 – –

23.3 –

31.1 – –

31.7 – – – – – – – –

– – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – – –

30.7

30.0 – –

31.4

– –

–

33.6 – – –

– – – – – – – –

30.6 – –

–

32.6

30.5

30.1 –

32.6 –

– – –

– – – – –

d13CH30

31.5 –

– – –

– – – – –

d13 CH29

33.5 –

29.3 – 0.134 0.064 0.021 0.017 0.096 – 0.025 0.031 0.05 – 0.046 –

d13CPr

23.1 –

25.7 – –

27.1 – –

C. Wang / Organic Geochemistry 38 (2007) 52–66

61

Table 2 (continued) Sample

Tm/Ts

Tm/C30 hopane

Moretane/ hopane

Gam/C30 hopane

C31/C32 hopane

C312mH/ C30H

Hopane/ sterane

Hopane/ n-alkane

d13Cn-alk

d13CPr

d13 CH29

d13CH30

D4-1 D3 D2-2A D2 C1-2

0.77 1.38 1.29 1.25 1.11

0.10 0.17 0.17 0.14 0.16

0.083 0.101 0.072 0.097 0.117

0.024 0.046 0.032 0.021 0.041

1.83 1.60 1.64 1.62 1.57

0.036 0.021 0.019 0.044 0.042

6.46 8.20 5.97 5.64 4.88

0.231 – 0.119 0.176 0.023

– –

– –

– –

– –

28.3 –

29.2 –

28.9

21.7 –

30.7

21.6 –

23.9

22.9

Notes: Ts, 18a(H)-28,29,30-trisnorneohopane; Tm, 17a(H)-28,29,30-trisnorhopane; moretane/hopane, (C29 + C30)ba-/ab-hopanes; Gam/C30 hopane, P gammacerane/C30 hopane; C31/C32 hopane, C31/C32 hopanes (22S + 22R); C312mH/C30H, C312-methylhopane/C30 hopane; hopane/n-alkane, C27–C35 P P P hopanes/ n-C13–n-C35; hopane/sterane, C27–C35 hopanes/ C27–C29 steranes. (All biomarker indices were calculated by using peak area based on mass chromatograms: hopanes and gammacerane, m/z 191; 2-methylhopane, m/z 205; n-alkanes, m/z 85; and steranes, m/z 217.) d13Cmole, carbon isotopic composition of a single compound peak; n-Alk, normal alkane; Pr, pristane; d13Cn-alk is an average value of n-C15, n-C21 and n-C25.

interpretation is also supported by other evidence, such as the rapid increase in the abundance of dibenzofurans from the upper Bed 23 to very high levels above the Bed 24, suggesting terrigenous input of acidified soils or peat sediments (Wang and Visscher, 2006). 3.1.2.4. Collapse of marine primary production and autochthonous bacteria production. The mechanism of terrigenous OM input for the hopane anomaly requires that the relative contribution of marine autochthonous bacteria to the hopanes must have been limited during the deposition of this OM. This decreased marine autochthonous bacterial (heterotrophs, methanotrophs) production may be mainly caused by a sharp decrease of marine primary production or environmental mutagenesis. Moreover, the terrigenous OM from soil/peat was altered by oxidization and degradation during long-distance transport before entering the sea. Obviously, the terrigenous hopane predominance in marine sediments may indicate that the terrigenous and marine ecosystem collapsed synchronously. Did marine primary production greatly decrease in the east Tethys Sea from end-Permian to Early Triassic time? Decreased marine primary production was suggested as a cause of the distinctive negative carbon-isotope shift corresponding to the end-Permian extinction (Wang et al., 1994; Bowring et al., 1998; Rampino et al., 2002), although the Permian–Triassic section in Western Australia shows that marine primary production in Early Triassic was locally high (Thomas et al., 2004; Grice et al., 2005c). Much lower marine primary production at the Meishan section is supported by the following evidence: (1) Organic carbon content (TOC) and hydrogen index (HI) of the sediments around the PTB are very low. TOC values range from 0.06%

to 0.89% with an average of 0.31%; HI values range from 10.2 to 97.3 mg HC/g TOC with an average of 48 mg HC/g TOC for 33 samples from Bed 22 to Bed 32 (Table 1). These data may indicate that the original organic matter was hydrogen-lean when deposited. Since anoxic marine environments were extensively developed from the end Permian to Early Triassic (Isozaki, 1997; Wignall and Twitchett, 2002; Grice et al., 2005a) and the sediments are slope to basin facies at the Meishan section (Jin et al., 2000), the marine OM might be well preserved and hydrogen-rich. Thus, the hydrogen-lean character of the marine OM might be caused by mixing with much terrigenous OM. Microscopic analysis of the 27 samples shows that the kerogen is amorphous and heavily reworked. For example, only a few vitrinite fragments were detected in Bed 24. These results do not preclude significant terrigenous input to the sediments because of the extensive diagenetic alteration that they have undergone. (2) The hopane/sterane ratio is usually used to evaluate the relative input of bacteria versus phytoplankton in sediments. High hopane/sterane ratios usually occur in land-plant dominated sedimentary environments where bacteria flourish. Fig. 5 shows that extremely high hopane/sterane ratios occur in the shales or mudstones in Beds 24–32. The shales and mudstones (e.g., the upper parts of Bed 29 and Bed 30) are richer in TOC compared with marl or micrite (e.g., the lower parts of Bed 29 and Bed 30) (Table 1), possibly indicating that the enhanced TOC is due to terrigenous OM rather than marine primary production. The ‘‘coal type’’ hopane, low TOC and high hopane/sterane ratio in these marine sediments together support the interpretation that marine primary production greatly decreased or collapsed from end-Permian to Early Triassic. Strongly

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C. Wang / Organic Geochemistry 38 (2007) 52–66

100

Moretane/Ho pane

Bed 32 black shales

Beds 29-30 black shales

10 Beds 26 black shales Beds 24 black grey micrites Beds 27-30 marls 1

0.0

0.1

0.2

0.3

0.4

Hopane/n -Alkane Fig. 5. Variation in hopane/sterane and hopane/n-alkane ratios in Beds 24–32 at Meishan section.

decreased marine primary production must have led to decreased aerobic bacteria (i.e., methanotrophs, heterotrophs) production, which resulted in the terrigenous hopane fingerprint of the marine sediments. 3.1.2.5. Are cyanobacteria related to the hopane anomaly?. As photosynthetic primary producers, cyanobacteria are a major source of sedimentary hopanoids (Jahnke et al., 1999). Were cyanobacteria responsible for the hopane distribution anomaly around the PTB? It has been well known that 2methylhopane is a biomarker for cyanobacteria (Summons et al., 1999), while 3-methylhopanoids originate mainly from methanotrophic bacteria (Farrimond et al., 2004). In this study, predominantly 2-methylhopanes were detected in the samples from the Meishan section. Fig. 3 shows an obvious variation in the relative abundance of 2methylhopane (C312mH/C30H ratio) across the section. Four intervals in the section are relatively rich in 2-methylhopane, namely Beds 1–8, Beds 16–18, Bed 26 and the upper part of Bed 29. The distribution of 2-methylhopane in Beds 24–32 is similar to that previously described (Xie et al., 2005). Cyanobacteria hopanes are usually distinctively rich in 13C compared with other sedimentary lipids (Hayes, 1990; Kohnen et al., 1992; Sakata et al., 1997; Summons et al., 1998). In Bed 26, hopanes are rich in 13C relative to n-alkanes or pristane (i.e., d13CH30 = 30.8&, while d13Cn-alk and d13CPr = 33.5& and 35.0&, respectively, see Table 2). Coupled with the relative enrichment of 2-methylhopane, this indicates a significant contri-

bution of cyanobacteria to the hopanes in this bed. Moreover, as for the upper Bed 30 and Bed 32, hopanes are not enriched in 13C relative to n-alkanes or pristane (i.e., d13CH30 = 33.6 to 32.6&, while d13Cn-alk and d13CPr = 33.2 to 32.1& and 33.5 to 32.8&, respectively, see Table 2), and correspondingly, the relative abundance of 2-methylhopane in these beds is low. However, the variation of 2-methylhopane ratio (C312mH/C30H) does not correlate with other hopane indices (such as the moretane/hopane) across the section (Fig. 3), especially when comparing the very different hopane distribution around the PTB with that of the WCB (a typical hopane distribution of cyanobacteria in Beds 1–6, see the discussion in Section 3.2). The hopane distributions in these intervals around the PTB are nearly the same, regardless of the contribution of cyanobacteria, such that the hopane distributions in the upper Bed 30 and Bed 32 are very similar to that of the upper Bed 29 and Bed 26 (Figs. 2 and 3), but they are very different in C312mH/C30H ratio. These all indicate that the hopane anomaly does not originate from marine bacteria, such as cyanobacteria, but predominantly from a terrigenous or specialized environment like that in acidified soil or peat.

3.1.2.6. Does a freshened and acidified upper water column occur around the PTB?. Although terrigenous OM input may explain the hopane anomaly around the PTB at Meishan, other factors should be considered. For example, the possibility of freshening and acidification of the upper water column in the Tethys Sea should be taken into account, because high moretane/hopane and C31/C32 hopane ratios are mostly related to acidic and oxic conditions. If these conditions occurred in the sea, there should be a salinity stratification in which the upper water column is fresh and oxic or acidified. Under these conditions biomass could be suspended and reworked by bacteria for a prolonged time, and the special hopane precursors might be formed. The high hopane/sterane ratios in Bed 26 and upper parts of Bed 29 and Bed 30 (Fig. 5) may be regarded as evidence that aerobic bacteria developed extensively. In addition, the very low sedimentation rate in these intervals (such as Beds 24–27) (about 0.3 m/Ma, Jin et al., 2000) might favor extensive reworking of the biomass. The possible evidence for marine freshening and acidification includes:

C. Wang / Organic Geochemistry 38 (2007) 52–66

1. Marine freshening: (1) ocean salinity was lowered by melting of the Gondwanan ice sheets and extensive Late Permian evaporite formation, and the whole-ocean salinity drop was probably 2 psu from Early Permian to Early Triassic (Kidder and Worsley, 2004, and references therein). (2) Kidder and Worsley (2004) further suggested that the frequency and intensity of tropical cyclonic storms increased by warming and the intense storm-generated rainfall in near coastal regions resulted in substantial flood runoff. Thus, strong rainfall should lead to water salinity stratification and freshening of the upper water column. (3) Marine stratification might be favored by extensive drainage of mid-latitude aquifers due to the expansion of desert belts and retreat of forests to high latitudes (Kidder and Worsley, 2004). (4) Enhanced erosion and freshwater input to shallow seas has been ascribed to effects of terrigenous ecosystem crisis which was revealed by the sharp rise in the 87Sr/86Sr ratio during the Late Permian–Early Triassic (Kozur, 1998; Korte et al., 2003). 2. Marine acidification: (1) acid rain caused by SO2 emissions during the Siberian Trap eruptions (Renne et al., 1995; Bowring et al., 1998) around the Permian–Triassic transitional period could acidify terrigenous depositional environments (Visscher et al., 1996, 2004) and lead to excessive soil erosion (Ward et al., 2000; Sephton et al., 2002, 2005; Visscher et al., 2004; Retallack, 2005; Watson et al., 2005). The enhanced terrigenous erosion and the acid precipitation probably affected the upper marine water column. (2) According to Knoll et al. (1996), from the latest Permian to earliest Triassic deep oceanic waters must have been isolated from the oxygen-rich surface ocean, accompanied by a rapid or transient turnover near the boundary. The rapid turnover of deep waters would lead to the release of CO2 and mixing of the ocean, which markedly increased the partial pressure of CO2 (P CO2 ) and elevated the acidity of the sea water. This was suggested to be the direct cause of the mass extinction, especially for the organisms producing CaCO3 skeletons (Knoll et al., 1996). Furthermore, methane hydrate release could be not only the main mechanism for the global negative d13C excursion, but also one of the main sources for the elevated CO2 concentration. Because the special hopane distribution occurs around the intervals

63

bearing the mass extinction events (Beds 24– 26, Beds 28–30) and coincides with the d13C excursion, the hopane anomaly around the PTB was probably related to the marine acidification. (3) Sulfidic (H2S-rich) bottom-water conditions were extremely extensive and intense in the late Permian, which might be the environmental stress that caused the biotic crisis in the Late Permian (Nielsen and Shen, 2004; Grice et al., 2005a; Kump et al., 2005). In addition, emissions of hydrogen sulfide to the atmosphere may have caused the terrigenous extinctions (Grice et al., 2005a; Kump et al., 2005). It was proposed that acid precipitation in atmosphere, on land, and in the ocean was accelerated by oxidation of massive H2S released from an anoxic ocean. Based on the above discussion, it can be concluded that marine freshening and acidification may have occurred and led to the formation of the hopane anomaly around the PTB. Certainly, this genesis mechanism could be secondary as compared with the first mechanism of terrigenous OM input. But, the two mechanisms may have occurred synchronously.

3.2. Is cyanobacterial bloom related to intermittent marine anoxia and local mass extinction around the WCB? 3.2.1. Hopane distribution around the WCB The hopane distribution seems to be very homogeneous in Beds 1–6 around the Wuchiapingian– Changhsingian boundary, although lithology varies. It is characterized by high abundance of C30abhopane, but low abundances of Tm and C31abhopanes, which is very different from Beds 24–32 (Figs. 2 and 3). This pattern of hopane distribution is common in anoxic marine organic-rich shales or argillaceous marls developed in slope or basin facies during transgression in various geologic eras (see Peters et al., 2005). It is also the typical hopane pattern for Cambrian organic-rich shale in which cyanobacteria are responsible for most primary production, such as in south China and the Tarim Basin in northwest China (our unpublished data). These intervals are characterized by high TOC and HI values (Table 1) that may be mainly due to favorable marine primary production and preservation. In addition, n-alkane distributions in these

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C. Wang / Organic Geochemistry 38 (2007) 52–66

intervals around WCB are very similar with n-C15–n-C17 as the main peaks, which also support the above concept that the OM is mainly sourced from cyanobacteria. 3.2.2. C-isotopic composition of hopanes around the WCB The carbon-isotopic composition of hopanes in Beds 1–6 are 13C-enriched, with the d13CHop values ranging from 21.5& to 26.0&, which are 13C enriched by ca. 4–8& compared with that of pristane (Table 2, Fig. 2). The 13C-enriched hopane indicates a cyanobacterial source (Hayes, 1990). Moreover, the heavier d13CHop value is accompanied by a positive d13Ccarb shift (Li, 1999; Jin et al., 2000, 2003; from 0.2& in Bed 1 to 3.8& in Bed 6, in our unpublished data) which coincides with high TOC and HI (Table 1), suggesting a major increase in primary production during marine transgression. Eutrophication and upwelling likely occurred intermittently. Thus, it is suggested that the local mass extinctions around the WCB (Jin et al., 1995; Knoll et al., 1996) are most probably related to the intermittent marine anoxia mainly caused by cyanobacterial blooms and the consequent poisoning of animals. 4. Conclusions 1. The hopane distribution anomaly in marine sediments at the PTB, characterized by high moretane/hopane, Tm/Ts, Tm/C30 hopane, C31/C32 hopane and hopane/sterane ratios, is most probably caused by: (1) input of terrigenous OM from acidified soil and peat; and (2) influence of freshening and acidification of the upper water column accompanied by stagnation and stratification of the marine waters during the Permian–Triassic transition. 2. The mechanism of terrigenous OM input for the hopane anomaly also implies global collapse of marine primary production and aerobic bacterial communities around the PTB, and that the terrigenous and marine ecosystem collapse occurred synchronously. 3. The hopane distribution anomaly associated with the end-Permian mass extinction and the typical carbon-isotope excursion around the PTB may indicate that the end-Permian mass extinctions were caused by marine and terrigenous ecosystem collapse triggered by the Siberian Trap eruptions.

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