Pyrite morphology and redox fluctuations recorded in the Ediacaran Doushantuo Formation

Palaeogeography, Palaeoclimatology, Palaeoecology 333 (2012) 218–227

Contents lists available at SciVerse ScienceDirect

Palaeogeography, Palaeoclimatology, Palaeoecology journal homepage: www.elsevier.com/locate/palaeo

Pyrite morphology and redox fluctuations recorded in the Ediacaran Doushantuo Formation Lin Wang a, Xiaoying Shi a, b,⁎, Ganqing Jiang c a b c

School of Earth Sciences and Resources, China University of Geosciences (Beijing), Beijing 100083, China State Key Laboratory of Biogeology and Environmental Geology, Beijing 100083, China Department of Geoscience, University of Nevada, Las Vegas, NV 89154-4010, USA

a r t i c l e

i n f o

Article history: Received 28 December 2011 Received in revised form 20 March 2012 Accepted 23 March 2012 Available online 30 March 2012 Keywords: Ediacaran Euxinic ocean Doushantuo Formation Framboidal pyrites South China

a b s t r a c t Recent geochemical studies of the Doushantuo Formation (ca. 635–551 Ma) in South China suggested that the Ediacaran ocean was strongly stratified, with an oxic surface layer above a euxinic wedge that was sandwiched within ferruginous deep waters. This ocean redox model, however, was derived largely from the data obtained from stratigraphic sections in the Yangtze Gorges area that, according to recent paleogeographic reconstruction, were deposited in a restricted intrashelf lagoon. In order to test the redox conditions in open-ocean, deep-water environments, we have conducted a detailed morphological analysis of authigenic pyrites from the upper to lower slope sections of the Doushantuo Formation. In analogy to modern euxinic basins such as the Black Sea, framboidal pyrites with smaller and less variable size distribution are taken as evidence for sulfide precipitation in euxinic water column, while larger and more variable framboidal and euhedral pyrites are formed or diagenetically altered in sediments with an overlying oxic/ dysoxic water column. Except for a few brief intervals, pyrites from the Doushantuo Formation in upper slope sections (Siduping and Taoying sections) are mainly of early diagenetic origin and do not record water column euxinia during deposition. In contrast, pyrites from the Doushantuo Formation in the lower slope section (Wuhe section) are dominated by fine-grained framboids indicative of pervasive water column euxinia below chemocline. In both upper and lower slope sections, temporal changes in genetic pyrite types occur at centimeter to decameter scales, suggesting frequent chemocline fluctuations. The overall decrease in abundance of framboidal pyrites toward the upper Doushantuo Formation in upper slope sections suggests increasing water column oxygenation and deepening of the chemocline. Macroalgae and metazoan fossils are found mainly from shale intervals without syngenetic pyrites in upper slope sections, indicating the sensitivity of macroscopic eukaryotes to the ocean chemocline. The redox fluctuations recorded by the Doushantuo pyrites compels for comprehensive geochemical data in deep-water successions to further test the existing paleoceanographic models of the Ediacaran ocean. © 2012 Elsevier B.V. All rights reserved.

1. Introduction The Ediacaran Period (635–542 Ma) was a critical time interval in earth history that witnessed significant changes in biotic evolution, global climate and ocean chemistry. Large negative carbon isotope anomalies with magnitudes greater than 10‰ have been reported from its sedimentary successions, indicating major perturbations of the global carbon cycle (e.g., Kaufman and Knoll, 1995; Calver, 2000; Halverson et al., 2005, 2010; Fike et al., 2006; Jiang et al., 2007, 2008, 2011; Zhou and Xiao, 2007; Zhu et al., 2007, in press; McFadden et al., 2008; Wang and Shi, 2009). Highly variable sulfur isotopes and low sulfate–sulfide sulfur isotope fractionation (Δδ 34S) suggest that

⁎ Corresponding author at: School of Earth Sciences and Resources, China University of Geosciences (Beijing), Beijing 100083, China. E-mail address: [email protected] (X. Shi). 0031-0182/$ – see front matter © 2012 Elsevier B.V. All rights reserved. doi:10.1016/j.palaeo.2012.03.033

sulfate concentration in Ediacaran ocean seawater may have been much lower than in the modern ocean (~ 28 mmol/L or mM) (Anbar and Konll, 2002; Kah et al., 2004; Hurtgen et al., 2005, 2006; Fike et al., 2006; Halverson and Hurtgen, 2007; McFadden et al., 2008; Halverson et al., 2010), although the marine sulfate reservoir may have increased significantly during the latest Ediacaran Period. This period also records the earliest macroscopic metazoan fossils in earth history (Xiao et al., 1998; Xiao and Knoll, 2000; Knoll et al., 2006; Yin et al., 2007; Yuan et al., 2009, 2011). It is generally believed that Ediacaran biological innovations and ocean chemical changes were closely related to the oxidation of earth's surface environments (Hurtgen et al., 2005; Fike et al., 2006; Canfield et al., 2007, 2008; Halverson and Hurtgen, 2007; Kaufman et al., 2007; McFadden et al., 2008; Scott et al., 2008; Shen et al., 2008; Lyons et al., 2009; Wang and Shi, 2009; Halverson et al., 2010; Jiang et al., 2010), but the casual link among these events remains controversial and requires additional data from individual sedimentary basins.

L. Wang et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 333 (2012) 218–227

The Doushantuo Formation (ca. 635–551 Ma) in South China plays important roles in understanding the Ediacaran ocean redox evolution and its relationship with biotic changes owing to its wellpreserved fossils, relatively good age constraints, and preservation of shelf-basin sedimentary sections (Jiang et al., 2011; and references cited therein). Recent geochemical studies from this formation suggested that the Ediacaran ocean was strongly stratified, with an oxic surface layer resting above a euxinic wedge that was sandwiched within ferruginous deep waters (Li et al., 2010). This model provides a conceptual framework for the Ediacaran ocean redox conditions, but existing data were mainly obtained from the inner and outer shelf sections that, according to the new paleogeographic reconstruction, may have been deposited in an intrashelf lagoon (Jiang et al., 2011; Xiao et al., 2012; Zhu et al., in press). Thus, further study is needed to test this model in the open ocean, deep-water successions. The morphology and size distribution of pyrites have proven to be an effective tool for understanding the water-column redox conditions in modern and ancient sedimentary basins (e.g., Raiswell, 1982; Wilkin et al., 1996, 1997; Wignall and Newton, 1998; Wilkin and Arthur, 2001). In modern euxinic sediments, less than 4% of pyrite framboids are >10 μm; while in non-euxinic sediments, more than 10% of pyrite framboids are >10 μm (Wilkin et al., 1996). Diagenetic processes within sediments would increase the proportion of euhedral/amorphous pyrites (Wilkin et al., 1996, 1997; Wignall and Newton, 1998; Wilkin and Arthur, 2001; Wignall et al., 2005, 2010), but their formation may not significantly change the attributes of the well-preserved portion of framboidal pyrites (e.g., Wignall and Newton, 1998). The maximum framboid diameter of pyrites (MFD) has been also used as a paleoredox proxy. Sediments/sedimentary rocks deposited from oxic/dysoxic environments usually have greater MFD than those deposited from euxinic environments (Wilkin et al., 1996, 1997; Wignall and Newton, 1998; Wilkin and Arthur, 2001; Nielsen and Shen, 2004; Chang et al., 2009). The validity of using the size and morphology of pyrites as a paleoredox indicator has been confirmed by a number of independent tests, including sedimentological, paleontological and geochemical methods (Wignall and Newton, 1998; Raiswell et al., 2001), and has been successfully applied in paleoenvironmental analysis (e.g. Nielsen and Shen, 2004; Wignall et al., 2005, 2010; Shen et al., 2007; Chang et al., 2009; Zhou and Jiang, 2009; Liao et al., 2010). In this paper, we report a detailed pyrite morphological study of the Doushantuo Formation from the open ocean, deep-water (slope) sections of the Ediacaran Yangtze platform. Using pyrite morphology and size distribution as redox indicators, we discuss the spatial and temporal redox variations during the Doushantuo deposition and their potential influences on the evolution of early metazoans. 2. Geological setting The Ediacaran Doushantuo Formation in South China was deposited on a passive continental margin that faced southeast (Fig. 1A) (Jiang et al., 2003, 2011; Wang and Li, 2003). It overlies the upper Cryogenian Nantuo Formation and underlies the Ediacaran Dengying Formation or its chert-dominated correlative units (Liuchapo or Laobao Formation) in deep-water sections (Fig. 1B). The base and top of the Doushantuo Formation have been dated as ca. 635 Ma and ca. 551 Ma, respectively (Condon et al., 2005; Zhang et al., 2005). The Doushantuo Formation started with a widespread, 3–6 m thick cap carbonate across the basin (Jiang et al., 2006). Shortly after this marker bed, a shoal complex was developed at the shelf-margin, which separated the intrashelf lagoon from the open ocean until the deposition of the lower Dengying Formation (Fig. 1C; Jiang et al., 2011). In this study, three sections in the open ocean side of the Ediacaran Yangtze platform were sampled for pyrite morphological investigations. Their locations are shown in Fig. 1A and C.

219

Among the studied sections, Siduping and Taoying sections are from the upper slope setting of the Yangtze platform. The Doushantuo Formation at Siduping mainly consists of carbonates, slump blocks and olistostromes, while at Taoying it is dominated by dark and black shales, with subordinate carbonates and olistostromes. The lowerslope Wuhe section consists of mainly black shales with thin carbonate interbeds and olistostrome layers. It is worth noting that abundant multicellular algae and Ediacaran type metazoan fossils have been reported from the top 25 m black shale interval of the Doushantuo Formation in the Taoying section (Zhao et al., 2004, 2010; Wang and Wang, 2006; Wang et al., 2007; Tang et al., 2008; Zhu et al., 2008). The algal fossils share great similarity with those found in the Miaohe member of the Doushantuo Formation in the Yangtze Gorges area (Xiao et al., 2002; Zhao et al., 2010). In addition, Horodyskia (the so called “strings of beads”) and Palaeopascichnus have been found in the Liuchapo Formation at Taoying and Siduping in this study, which are identical with those described from the same formation in eastern Guizhou (Dong et al., 2008). 3. Samples and methods For this study, 175 samples were collected from three sections (Fig. 1A and C). Sampling spacing is decimeter to meter level, which varies according to the exposure condition and thickness of studied sections (Table S1). Weathered intervals were avoided during sampling. Pyrite morphology was examined in thin sections using optical microscope and scanning electron microscope (SEM). Diameters of 10,205 framboids in thin sections were measured by eyepiece micrometer under reflected light mode of an Axio Scope A1 optical microscope made by Carl Zeiss, Inc. The eyepiece micrometer has a minimum scale of 1 μm under maximum magnification. Although the resolution of the microscope is slightly lower than the backscattered electron microscope used in other related researches, it still meets the requirement for size distribution analysis of pyrite framboids (down to 0.1 to 1 μm level). Thin sections used for framboidal pyrite analysis are mostly 2 × 3 cm in size. Diameters of framboidal pyrites were measured along several parallel lines across each thin section until a size population more than 100 is achieved (less than 100 in several samples due to sparse framboids or smaller thin section size). The minimal, mean and maximum diameter, the standard deviation of diameters, the percentage of pyrite framboids with diameter ≥10 μm, and the proportion of pyrite framboids with infilling feature (microcrystals welded together by secondary pyrites) were calculated from measured framboidal pyrite population in each sample. We also estimated the area ratios of framboidal pyrites to all pyrites in our samples. The analytic results are listed in Table 1. 4. Results In most samples, disseminated pyrites are randomly distributed in thin sections (Fig. 2A to C) and occasionally, pyrites form clusters (of single crystal or framboids) or lenses. Some samples contain crosscutting secondary veins or cracks that are filled with calcite or quartz, but the distribution of pyrites shows little relationship with them. In general, carbonates contain less pyrites than shale and mudstone (Fig. 2C), but no other differences (i.e., pyrite size and morphology) are recognizable among lithologies. Genetic types of pyrites are evaluated and classified by the morphology and size distribution of framboids. Four genetic types have been identified and they reflect different forming conditions (Table 2). A sample could contain more than one pyrite type, but euhedral and amorphous (Type A) are the most common morphology

220

L. Wang et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 333 (2012) 218–227

L. Wang et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 333 (2012) 218–227

221

Table 1 Size distribution of pyrite framboids in studied sections. Section

Samples

Height

n

dMin

dMax

dAvg

S.D.

RD

RI

RF

Siduping

08SDP-1-6 08SDP-4-1 08SDP-5-1 08SDP-6-1 08SDP-7-1 08SDP-8-1 08SDP-9-1 08SDP-10-1 08SDP-10-2 08SDP-10-4 08SDP-11-1 08SDP-12-1 08SDP-12-3 08SDP-29-4 08SDP-29-5 08SDP-29-7 08SDP-30-2 08SDP-30-7 08SDP-30-9 08SDP-31-2 08SDP-36-3 WG8.7 09TY-17.8 WG11.7 08WG-3r-2 08WG-4a WG15.8 WG16.8 08WG-10-2 WH09-3.2 WH09-6.0 WH09-8.6 WH09-9.7 WH09-11.9 WH09-13.0 WH09-24.4 WH09-26.0 WH09-28.0 WH09-28.6 WH09-30.9 WH09-31.1 WH09-35.1 WH09-37.1 WH09-38.0 WH09-40.2 WH09-44.7 WH09-45.2 WH09-47.2 WH09-50.1 WH09-53.4 WH09-56.7 WH09-60.7 WH09-85.7 WH09-86.3 WH09-88.9 WH09-90.1 WH09-92.0 WH09-95.1 WH09-98.3 WH09-100.0 WH09-105.1 WH09-110.7 WH09-118.0 WH09-119.1 WH09-120.0

7.1 15.2 19.1 24.9 26.7 29.5 32.1 35.1 39.2 42.7 45.7 51.7 53.9 115.8 116.9 118.1 125.7 131.1 133.9 138.3 174.2 10.7 12.2 16.7 16.7 17.4 20.8 21.8 38.3 3.2 6.0 8.6 9.7 11.9 13.0 24.4 26.0 28.0 28.6 30.9 31.1 35.1 37.1 38.0 40.2 44.7 45.2 47.2 50.1 53.4 56.7 60.7 85.7 86.3 88.9 90.1 92.0 95.1 98.3 100.0 105.1 110.7 118.0 119.1 120.0

89 332 136 160 193 242 218 156 215 102 181 115 171 203 122 88 143 103 182 109 187 215 203 143 217 237 168 80 221 162 120 151 172 121 169 153 154 158 163 152 151 148 153 196 160 159 159 160 138 151 136 128 136 152 131 109 142 150 145 135 151 139 108 132 130

3.1 1.3 1.8 3.0 1.8 1.9 1.3 1.6 2.0 1.8 2.2 1.6 1.8 2.0 2.1 3.0 1.5 1.6 2.2 2.3 2.2 2.2 3.2 1.8 2.0 1.5 6.0 6.3 2.0 3.0 2.5 2.2 2.7 2.4 2.5 1.9 2.3 2.0 2.4 1.9 1.9 2.3 2.0 2.0 1.9 2.0 1.5 2.2 2.2 1.5 1.6 2.5 2.3 2.0 2.2 3.0 1.9 2.0 2.2 2.1 2.1 2.1 2.0 2.1 2.2

16.5 18.8 20.5 19.0 11.0 19.0 7.8 12.0 13.0 13.3 8.5 9.0 9.1 10.3 11.0 11.8 12.0 21.9 11.0 13.2 21.3 13.0 16.0 34.1 (17.9)a 14.5 13.5 50.0 41.5 14.0 10.7 16.9 10.3 10.6 8.3 8.2 11.1 20.9 14.2 17.6 13.5 14.0 16.0 14.7 16.0 18.0 15.1 17.0 11.6 15.1 15.5 18.5 15.8 18.2 9.0 11.0 17.1 15.0 13.0 13.0 12.0 14.0 16.0 9.0 9.0 10.0

6.5 4.6 5.9 7.8 4.8 5.2 3.4 3.7 4.8 4.6 5.1 4.2 3.7 6.1 5.1 5.8 7.2 8.5 5.8 4.8 6.7 5.5 6.5 4.4 4.8 5.6 20.4 15.5 5.6 5.7 5.8 5.0 6.0 5.2 5.6 5.8 7.4 5.7 6.0 5.3 6.1 5.8 5.0 5.3 5.4 4.9 5.4 4.8 5.1 5.3 5.1 5.4 5.6 5.0 5.2 6.9 5.5 5.1 4.9 4.7 4.8 5.5 5.0 4.7 4.6

2.5 1.7 2.7 2.9 1.6 2.4 1.0 1.3 1.9 2.0 1.3 1.5 1.2 1.9 1.7 1.8 2.2 2.3 1.7 1.6 3.0 1.9 1.7 3.1(1.9)a 1.6 1.9 9.4 4.1 1.9 1.3 2.1 1.3 1.5 1.1 1.2 1.6 3.1 2.2 2.6 2.1 2.6 2.3 1.8 2.4 2.4 2.0 2.2 1.7 1.9 2.4 1.8 2.0 2.5 1.5 1.7 3.0 2.4 1.8 1.5 1.2 1.9 2.3 1.3 1.5 1.4

11.2% 0.6% 8.8% 18.1% 0.5% 4.6% 0.0% 0.6% 3.7% 2.9% 0.0% 0.0% 0.0% 3.5% 2.5% 3.4% 14.7% 33.0% 1.7% 1.8% 11.2% 2.8% 3.5% 2.8% 1.4% 3.0% 94.1% 96.3% 3.6% 1.2% 5.0% 0.7% 0.6% 0.0% 0.0% 2.0% 20.1% 5.7% 11.7% 5.3% 11.3% 6.8% 1.3% 6.1% 6.3% 3.8% 1.9% 2.5% 2.2% 3.3% 0.7% 3.9% 5.9% 0.0% 1.5% 19.3% 6.3% 1.3% 0.7% 0.7% 3.3% 3.6% 0.0% 0.0% 0.8%

67.4% 100.0% 68.9% 81.3% 81.9% 92.1% 74.4% 61.3% 91.6% 82.6% 72.4% 41.8% 40.9% 92.1% 95.1% 87.5% 97.9% 99.0% 100.0% 94.5% – 64.7% 17.7% 93.7% 77.4% 54.8% 99.4% 100.0% 48.4% 9.3% 50.0% 22.5% 12.2% 15.7% 18.3% 68.0% 26.6% 32.9% 30.1% 71.7% 48.3% 47.3% 46.4% 48.0% 43.1% 57.9% 58.5% 59.4% 54.4% 30.5% 71.3% 60.2% 61.0% 82.9% 62.6% 70.6% 49.3% 50.0% 66.2% 61.9% 51.7% 71.9% 25.0% 39.4% 93.1%

3.6% 15.0% 23.3% 3.6% 4.2% 37.3% 25.8% 14.0% 29.6% 4.4% 9.9% 16.1% 9.1% 11.0% 18.2% 13.8% 11.8% 3.7% 36.4% 30.3% 72.5% 12.7% 43.3% 0.9% 21.8% 52.9% 46.6% 49.4% 40.4% 24.9% 24.5% 32.5% 61.5% 31.1% 36.0% 46.4% 43.5% 31.2% 41.0% 16.6% 28.4% 31.6% 25.9% 28.1% 37.1% 38.2% 29.5% 32.9% 39.9% 37.1% 18.0% 32.3% 35.3% 50.0% 33.4% 13.4% 26.0% 41.1% 30.0% 24.9% 34.7% 17.0% 5.4% 6.5% 2.5%

Taoying

Wuhe

Notes: 1. “Height” in the table refers to the distance of the sampling level to the base of the Doushantuo Formation, measured in meters. 2. “n” refers to the number of framboidal pyrites analyzed. 3. “dMin”, “dMax” and “dAvg” refer to the minimal, maximal and average value of the diameter of framboidal pyrites, measured in μm. 4. S.D. refers to standard deviation of framboid diameters. 5. “RD” refers to the percentage of pyrite framboids with diameter ≥10 μm in the total pyrite framboids. “RI” refers to the proportion of infilled pyrite framboids to total framboids, “–” means that RI can not be obtained from the sample due to oxidation of pyrite. “RF” refers to the area ratio of pyrite framboids to all pyrites. a The number in brackets is the calculated results when excluded the exceptionally large diagenetic framboidal pyrites in WG11.7 (see supplementary Table S1).

Fig. 1. (A) Paleogeographic reconstruction during the deposition of the Doushantuo Formation, with study sections marked (after Jiang et al., 2011). (B) Summary of the late Neoproterozoic stratigraphic units in the Nanhua basin in South China, with major age constraints and marker beds for shelf-to-basin correlation (after Jiang et al., 2007, 2011). (C) Shelf-to-basin transect with section numbers matching those in (A). Modified from Jiang et al. (2007, 2011) and Ader et al. (2009). Filled circles and triangles in (A) and (C) indicate sections for this study.

222

L. Wang et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 333 (2012) 218–227

Fig. 2. Microscopic features of pyrites from the Doushantuo Formation. (A) Black shale sample from the lower slope Wuhe section, rich in framboidal pyrites. (B) Black shale sample from the upper slope Taoying section. Framboidal pyrites are less abundant and usually less well-preserved compared to the lower slope section. (C) Carbonate sample from the upper slope Siduping section. Both framboidal and non-framboidal pyrites are less abundant than in the lower slope shale samples. (D) Black shale sample from the Wuhe section. A euhedral pyrite grows on a framboidal pyrite, but the later can still be identified. The framboidal pyrite also show secondary infilling feature, which occurs at the surface of microcrystals and welds them together. The secondary growth, however, did not change the overall size of the framboids significantly. (E) A ~ 3 mm layer dominated by large euhedral and amorphous pyrites in a black shale sample from the Taoying section. Pyrite-rich and pyrite-poor layers alternate at millimeter scale, possibly recording frequent redox fluctuations. (F) SEM image of a euhedral pyrite crystal from the Taoying section. Abbreviations: N-non-framboidal pyrite, F-Framboidal pyrite.

across the sections (Table S1). Framboidal pyrites (Type B, C and D) commonly coexist with non-framboidal pyrites (Table S1) and often show variable levels of infilling (Table 1; Fig. 2D). The size distribution of type C and D pyrites resembles syngenetic pyrite precipitated from euxinic water column. Samples

with these pyrite types were likely deposited under euxinic water column (or at least periodic euxinic bottom water), they were found in dispersed intervals separated by shale and carbonates (vary from b1 m to > 10 m in thickness) that contain only diagenetic pyrites (type A and B) or negligible pyrites. The thickness of a single

Table 2 Genetic types of pyrites of the Doushantuo Formation. Genetic Morphology types

Proportion of framboids Mean diameter S.D.a (≥ 10 μm) (μm)

Type A

Euhedral and amorphous –

–

Type B

Framboidal

>10%

6.0–20.4

Type C

Framboidal

4.6%–6.8%

5.2–5.8

Type D

Framboidal

b4%

3.4–6.5

–

MFDb (μm)

Environmental interpretation

–

Diagenetic pyrites formed within sediments. Forms slowly at saturation levels below those of Fe monosulfides 2.2–9.4 12–50 Diagenetic pyrites formed within sediments, with overlying oxic/dysoxic water column. Form in pore waters highly supersaturated with respect to both Fe monosulfide and pyrite 2.1–2.4 13.5–19 Mixture of syngenetic/diagenetic pyrites. Overlying water column likely shift ephemerally between dysoxic and euxinic. 1.0–2.4 7.8–18.8 Syngenetic pyrites precipitated from euxinic water column

Note: There are two samples (WG11.7 and 08SDP-5-1) contain framboidal pyrites that can not be simply classified by the standard above. Their genetic types are decided separately (see supplementary Table S1). a S.D. = Standard deviation of framboid diameters. b MFD = Maximum framboid diameter of pyrites.

L. Wang et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 333 (2012) 218–227

syngenetic pyrite interval varies from less than 1 meter to tens of meters (Fig. 3). The temporal distribution of syngenetic pyrites varies among sections. In the upper-slope Siduping section (Fig. 3A), syngenetic pyrites mainly appear in two discrete intervals: the first 50 m following the Doushantuo cap carbonate and a 30-m-thick interval in the upper Doushantuo Formation. In the upper-slope Taoying section (Fig. 3B), syngenetic pyrites appear mostly in the lower half of the Doushantuo Formation, and their abundance decreases significantly in the upper half of the formation. Black shales that host multicellular algae and metazoan fossils in this section contain only diagenetic pyrites or no pyrites (Fig. 3B). In the lower-slope Wuhe section (Fig. 3C), the majority of the Doushantuo Formation samples contain syngenetic pyrites. Discrete intervals include the basal Doushantuo Formation (5–15 m), the lower-middle Doushantuo Formation (23–64 m), and the upper Doushantuo Formation (85–120 m). In general, framboidal pyrites indicative of euxinic water column pyrite precipitation in the upper slope sections are much less abundant than those in the lower slope section, indicating more pervasive euxinia in the deep-water section.

223

5. Discussion 5.1. Framboidal pyrites as a paleoredox indicator The common occurrence of euhedral/amorphous pyrites (type A in Table 2) in most Doushantuo samples indicates diagenetic alteration and challenges the reliability of using pyrite morphology and size distribution as a paleoredox indicator. In modern euxinic basins such as the Black Sea, the majority of pyrites are fine-grained (b10 μm) framboids precipitated from the euxinic water column and they have undergone negligible secondary growth within sediments (Calvert et al., 1996; Wilkin et al., 1996; Wilkin and Arthur, 2001). In ancient sedimentary successions, however, pyrites inevitably underwent some degree of diagenetic alteration and the coexistence of primary and secondary pyrites would be expected. To prove the validity of using the preserved portion of framboidal pyrites as a paleoredox indicator, Wignall and Newton (1998) studied the British Jurassic black shales and found that, although framboidal pyrites coexist with a large proportion of non-framboidal pyrites, the size distribution of framboidal pyrites matches well with paleoecologically-determined

Fig. 3. Stratigraphic distribution of pyrite morphology and mean diameter of framboidal pyrites in the Doushantuo Formation. (A) Siduping section (upper slope). (B) Taoying section (upper slope). (C) Wuhe section (lower slope). Scale in X-axis of pyrite morphology diagram indicates the combinations of pyrite types (Table 2). 0 — samples with negligible pyrites; 1 — samples with type A pyrites only; 2 — samples with type A and B pyrites; 3 — samples with type A and type C/D pyrites. Type C and D pyrites are taken as evidence for euxinic bottom waters during the Doushantuo deposition. Dark gray band in each of the mean diameter diagrams marks diameter range of the syngenetic framboidal pyrites. The black dots adjacent to the lithological column show the sampling horizons.

224

L. Wang et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 333 (2012) 218–227

biofacies. They concluded that pyrite framboids, when they are present, could preserve information of original depositional environments through diagenetic processes. In the Doushantuo Formation, a large proportion of framboidal pyrites exhibit the so called “infilling feature” as described in previous study (e.g., Wilkin et al., 1996), which may indicate diagenetic alteration of these framboids. In order to evaluate its effect on the size distribution analysis, we calculate the mean diameter of framboids with and without secondary infilling feature separately. For Type C and D pyrites, the mean diameter of all infilled framboids (5.2 μm) is a little larger than unfilled ones (5.1 μm), but the difference is negligible and would not have significant effect on the size distribution analysis. Therefore, we think that it is valid to use the size and morphology of well-preserved framboidal pyrites as a paleoredox indicator for the Doushantuo Formation (cf., Wilkin et al., 1996; Wignall and Newton, 1998). The coexistence of primary and secondary pyrites in the Doushantuo samples, however, indicates that cautions should be taken when geochemical analyses such as iron speciation (DOP and FePY/FeHR) are used to evaluate the depositional redox conditions of sedimentary rocks. Because intensive diagenetic pyrite formation within sediments under non-euxinic water column may significantly increase DOP and FePY/FeHR values, it may mistakenly identify euxinic water column conditions. For example, in the modern Achterwasser lagoon in southwest Baltic Sea (Neumann et al., 2005) and oxic shelf region of the Black Sea (Wijsman et al., 2001), high DOP values reflect pore water rather than water column redox conditions because of the formation of secondary pyrites. While FePY/FeHR ratio in sediments remains as an important proxy for water-column euxinia, it is better to be used in combination with other methods such as the analysis of pyrite size and morphology. 5.2. Deep-water euxinia in the Ediacaran Nanhua basin If framboidal pyrites from the Doushantuo Formation record water column euxinia, their spatial distribution (Fig. 3) provides insights to the redox structure of the Ediacaran Nanhua basin (and the Ediacaran ocean). In a recent redox model, Li et al. (2010) proposed that the deep-water environments in the Nanhua basin were largely ferruginous, with a dynamic wedge of euxinic water resting on the shelf (Fig. 4A), which is similar to the redox structures characteristic of late Archean to early Neoproterozoic oceans (Reinhard et al., 2009; Johnston et al., 2010; Kendall et al., 2010; Poulton et al., 2010; Poulton and Canfield, 2011). This model was derived from the geochemical data from the inner shelf section in the Yangtze Gorges area (i.e., the Jiulongwan section), the outer shelf section (Zhongling section), and two poorly sampled deep-water sections (Mingle and Longe sections). The dominance of framboidal pyrites in the lower-slope Wuhe section, however, suggests that the deep-water environments were largely euxinic. Does the deep-water euxinia represented by the Wuhe section record the extension of the euxinic wedge to the lower slope environment? In Li et al. (2010), euxinic intervals were mainly documented from the inner shelf section in the Yangtze Gorges area (Fig. 5) and the outer shelf Zhongling section (and basinward) recorded largely ferruginous water-column conditions. Thus the proposed euxinic wedge did not reach the shelf margin (Fig. 4A). The spatial distribution of framboidal pyrites among the upper and lower slope sections (Fig. 3) does not support a continuous shelf-to-slope euxinic wedge as well. A more plausible interpretation, which is consistent with the recent paleogeographic reconstruction of the Nanhua basin, is that the euxinic intervals from the Yangtze Gorges area (Li et al., 2010) recorded the redox structure of a restricted intrashelf lagoon (Jiang et al., 2011; Xiao et al., 2012), while the upper-lower slope sections (Fig. 3) document the redox changes in the open-ocean side of the Nanhua basin (Fig. 4B and C). The spatial distribution of deep-water

Fig. 4. Alternative redox models for the Ediacaran Nanhua basin. (A) The stratified redox model in Li et al. (2010). In this model, a euxinic wedge was thought to have rested on the shelf north of the Zhongling section and the deep basin (ocean) was largely ferruginous. (B) Preferred redox model for the Ediacaran Nanhua basin based on the recent paleogeographic reconstruction of the Nanhua basin (Jiang et al., 2011; Xiao et al., 2012; Zhu et al., in press) and the distribution of framboidal pyrites indicative of euxinic bottom waters (Fig. 3). (C) Alternative redox model for the Nanhua basin assuming that the lower slope Wuhe section recorded a deep-water euxinic wedge.

euxinia toward the basin, however, is uncertain. One possibility is that the deep ocean below the lower-slope environment was largely euxinic, with a chemocline located between the upper and lower slope (Fig. 4B). Episodic expansion of euxinic water to the upper slope environments resulted in deposition of framboidal pyrites in shallower water, upper slope sections. Another scenario could be a euxinic wedge covering the lower slope water depth sandwiched in deep ferruginous waters, which is similar to the redox framework proposed by Li et al. (2010) but occurred in the open-ocean side of the basin (Fig. 4C). The euxinic wedge in this model (Fig. 4C), however, may be difficult to maintain due to the lack of sulfate supply from the continental source, unless the water column above the euxinic wedge was largely oxic rather than ferruginous. A more comprehensive geochemical study in deep-water sections is required to test these hypotheses. 5.3. Redox fluctuations Another important aspect of the syngenetic framboidal pyrites from the Doushantuo Formation is their temporal variation (Fig. 3). In each section, euxinia-dominated intervals are separated by submeterto decameter-scale non-euxinic carbonate and shales. These temporal variations indicate that temporal redox fluctuations in the Nanhua basin and potentially, in the Ediacaran ocean may have been much more frequent than previously thought (e.g., Jiang et al., 2007; Li et al., 2010). In the lower slope Wuhe section (Fig. 3C), framboidal pyrites appear mainly at three discrete intervals at the basal, lower, and upper Doushantuo Formation. Previous geochemical studies from the Jiulongwan sections in the Yangtze Gorges area have shown three euxinic intervals at approximately the same stratigraphic level

L. Wang et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 333 (2012) 218–227

Fig. 5. Stratigraphic distribution of major fossil groups and geochemical changes of the Doushantuo Formation in the Yangtze Gorges area. (A) Stratal column of the Doushantuo Formation at the Jiulongwan section, showing the four members (marked with I, II, III, IV) and major age constraints (after Li et al., 2010). (B) δ13 Ccarb excursion (after McFadden et al., 2008). EN1, 2, 3 mark the negative δ13 C excursions and EP1, 2 mark the intervals with highly positive δ13 C. (C) Iron speciation FeHR/FeT and FePY/FeHR data (after Li et al., 2010). (D) Variation of Mo and Mo/TOC (White and gray squares are from Scott et al., 2008. Cycles and triangles are from Li et al., 2010). (E) Variation of Δδ34SCAS-Py values (after McFadden et al., 2008). (F) Integrated temporal distribution of major fossil groups of the Doushantuo Formation. I — Acanthomorphic acritarchs (compiled from the data of Yin et al., 2007; McFadden et al., 2008, 2009). II — Multicellular algae (compiled from the data of Zhao et al., 2004; Wang and Wang, 2006; Wang et al., 2007; McFadden et al., 2008; Yuan et al., 2011). III — Metazoan embryos (compiled from the data of Xiao et al., 1998; Xiao and Knoll, 2000; Yin et al., 2007; McFadden et al., 2008, 2009). IV — Macroscopic metazoan (compiled from the data of Tang et al., 2008; Zhu et al., 2008; Liu et al., 2010; Zhao et al., 2010). Sections represented by numbers in (F): 1 — Xiaofeng section, Hubei; 2 — Sixi section, Hubei; 3 — Wangfenggang section, Hubei; 4 — Tianjiayuanzi section, Hubei; 5 — Jiulongwan section, Hubei; 6 — Zhancunping section, Hubei; 7 — Miaohe section, Hubei; 8 — Niuping section, Hubei; 9 — Weng'an section, Guizhou; 10 — Taoying section, Guizhou; 11 — Lantian section, Anhui. Gray bands show the proposed euxinic intervals by Li et al. (2010).

225

226

L. Wang et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 333 (2012) 218–227

(Fig. 5). Although a precise correlation of these euxinic intervals across the basin needs additional age constraints, it is possible that the occurrence and disappearance of euxinic intervals in Jiulongwan and Wuhe sections record at least basin-scale events, because these two sections were located at very different geographic position. Euxinic intervals from the Jiulongwan section roughly match the three negative δ 13C excursions (Fig. 5) that have been interpreted as resulting from pulsed oxidation of the Ediacaran ocean (McFadden et al., 2008). There is a potential casual link among euxinia, negative δ 13C excursion, and ocean oxidation, which requires further geochemical analyses in deep-water sections. The lack of three distinctive euxinic intervals in the upper slope sections (Fig. 3) suggests that these sections may have been located in transient redox conditions episodically emerged below the chemocline (e.g., Fig. 4B). These regions may have been affected by chemocline fluctuations more frequently than Jiulongwan and Wuhe sections (e.g., Fig. 4B). 5.4. Redox fluctuations and Ediacaran biotas In Jiulongwan section, most of the documented fossil groups are from non-euxinic intervals (Fig. 5; McFadden et al., 2008, 2009; Li et al., 2010), implying that water column euxinia had a strong influence on the organisms, especially for eukaryotes. However, fossils have been also reported from intervals with high FePY/FeHR values indicative of euxinic bottom waters. For example, in the first euxinic interval of the lower Doushantuo Formation at Jiulongwan, abundant acritarch fossils were documented (McFadden et al., 2008, 2009; Li et al., 2010). In the third euxinic interval, eukaryotic fossils of the Miaohe biota were found in nearby sections (Xiao et al., 2002). More recently, morphologically complex eukaryote fossils have been found from the lower Doushantuo Formation in the deep-water section in Lantian (Yuan et al., 2011), where iron speciation and sulfur isotope data suggested anoxic/euxinic conditions (Shen et al., 2008). The frequent redox fluctuations seen in the slope sections hint that the organisms may have lived during the non-euxinic short periods and were killed and preserved during the euxinic periods. This may imply that the existing geochemical studies have not yet reached the resolution required to reveal the submeter- to centimeter-scale redox fluctuations associated with the living and burial conditions of the Doushantuo organisms. Impact of euxinic waters may have had the regional control on the occurrence of the Doushantuo eukaryotic life. In the upper Doushantuo Formation of the Taoying section, various multicellular algal and metazoan fossils were found in the interval after syngenetic pyrites disappeared. While in Wuhe section, where syngenetic pyrites were well developed up to the upper Doushantuo Formation, no multicellular algal and metazoan fossils have been found so far. The differences between the two sections might, to some extent, indicate that macroscopic eukaryotes could only live in the region above chemocline where the water column was less frequently intervened with euxinic conditions. 6. Conclusions Framboidal pyrites with small and narrow size distribution from the Doushantuo Formation in South China are taken as evidence for water column euxinia in the Ediacaran Nanhua basin (and the Ediacaran ocean). In upper slope sections, framboidal pyrites indicative of euxinic bottom waters appear only in the lower and locally in the upper Doushantuo Formation. These euxinic intervals are separated by submeter to decameter thick, non-euxinic shale and carbonates. In the lower slope section, framboidal pyrites are much more abundant and euxinic intervals revealed by pyrite size and morphology are proportionally thicker than non-euxinic intervals. The spatial distribution of framboidal pyrites indicates pervasive deep-water euxinia in the Ediacaran Nanhua basin, in contrast with

the previous redox model that suggested predominantly ferruginous deep waters. Temporal variations in framboidal pyrites at submeter scales in individual sections suggest frequent redox fluctuations in local depositional environments that may have controlled the living and burial redox conditions for the Doushantuo biotas. Decameterscale euxinic and non-euxinic alternations in the deep-water, lower slope section are comparable with those from the shelf section in the Yangtze Gorges area, likely recording basin-wide redox events. The causal link among euxinia, carbon isotope excursions, and ocean oxidation, however, remains to be tested with high-resolution and comprehensive geochemical study in deep-water sections. Acknowledgments This study was supported by the National Natural Science Foundation of China (40972022, 40921062), and the Ministry of Science and Technology of China (2011CB808806). We are grateful to Dr. Finn Surlyk and two anonymous reviewers for their constructive comments and suggestions that have improved the paper greatly. Thanks are given to Hao Xinxin, Tang Dongjie, Wang Xinqiang and Zhang Mingyuan for their kind assistance in field work. Supplementary data Supplementary data associated with this article can be found in the online version, at doi:10.1016/j.palaeo.2012.03.033. These data include Google maps of the most important areas described in this article. Appendix A. Supplementary data Supplementary data to this article can be found online at doi:10. 1016/j.palaeo.2012.03.033. References Ader, M., Macouin, M., Trindae, R.I.F., Hadrien, M.H., Yang, Z., Sun, Z., Besse, J., 2009. A multilayered water column in the Ediacaran Yangtze platform? Insights from carbonate and organic matter paired δ13C. Earth and Planetary Science Letters 288, 213–227. Anbar, A.D., Konll, A.H., 2002. Proterozoic ocean chemisty and evolution: a bioinorganic bridge? Science 297, 1137–1142. Calver, C.R., 2000. Isotope stratigraphy of the Ediacarian (Neoproterozoic III) of the Adelaide Rift Complex, Australia, and the overprint of water column stratification. Precambrian Research 100, 121–150. Calvert, S.E., Thode, H.G., Yeung, D., Karlin, R.E., 1996. A stable isotope study of pyrite formation in the Late Pleistocene and Holocene sediments of the Black Sea. Geochimica et Cosmochimica Acta 60, 1261–1270. Canfield, D.E., Poulton, S.W., Narbonne, G.M., 2007. Late-Neoproterozoic deep-ocean oxygenation and the rise of animal life. Science 315, 92–95. Canfield, D.E., Poulton, S.W., Knoll, A.H., Narbonne, G.M., Ross, G., Goldberg, T., Strauss, H., 2008. Ferruginous conditions dominated later Neoproterozoic deep-water chemistry. Science 321, 949–952. Chang, H., Chu, X., Feng, L., Huang, J., 2009. Framboidal pyrites in cherts of the Laobao Formation, South China: evidence for anoxic deep ocean in the terminal Ediacaran. Acta Petrologica Sinica 25, 1001–1007. Condon, D., Zhu, M., Bowring, S., Wang, W., Yang, A., Jin, Y., 2005. U–Pb ages from the Neoproterozoic Doushantuo Formation, China. Science 308, 95–98. Dong, L., Xiao, S., Shen, B., Zhou, C., 2008. Silicified Horodyskia and Palaeopascichnus from upper Ediacaran cherts in South China: tentative phylogenetic interpretation and implications for evolutionary stasis. Journal of the Geological Society 165, 367–378. Fike, D.A., Grotzinger, J.P., Pratt, L.M., Summons, R.E., 2006. Oxidation of the Ediacaran ocean. Nature 444, 744–747. Halverson, G.P., Hurtgen, M.T., 2007. Ediacaran growth of the marine sulfate reservoir. Earth and Planetary Science Letters 263, 32–44. Halverson, G.P., Hoffman, P.F., Scharg, D.P., Maloof, A.C., 2005. Toward a Neoproterozoic composite carbon-isotope record. GSA Bulletin 117, 1181–1207. Halverson, G.P., Wade, B.P., Hurtgen, M.T., Barovich, K.M., 2010. Neoproterozoic chemostratigraphy. Precambrian Research 182, 337–350. Hurtgen, M.T., Arthur, M.A., Halverson, G.P., 2005. Neoproterozoic sulfur isotopes, the evolution of microbial sulfur species, and the burial efficiency of sulfide as sedimentary pyrite. Geology 33, 41–44.

L. Wang et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 333 (2012) 218–227 Hurtgen, M.T., Halverson, G.P., Arthur, M.A., Hoffman, P.F., 2006. Sulfur cycling in the aftermath of a 635-Ma snowball glaciation: evidence for a syn-glacial sulfidic deep ocean. Earth and Planetary Letters 245, 551–570. Jiang, G., Sohl, L.E., Christie-blik, N., 2003. Neoproterozoic stratigraphic comparison of the Lesser Himalaya (India) and Yangtze Block (south China): paleogeographic implications. Geology 31, 917–920. Jiang, G., Shi, X., Zhang, S., 2006. Methane seeps, methane hydrate destabilization, and the late Neoproterozoic postglacial cap carbonates. Chinese Science Bulletin 51, 1121–1138. Jiang, G., Kaufman, A.J., Christie-blick, N., Zhang, S., Wu, H., 2007. Carbon isotope variability across the Ediacaran Yangtze platform in South China: implications for a large surface-to-deep ocean δ13C gradient. Earth and Planetary Science Letters 261, 303–320. Jiang, G., Zhang, S., Shi, X., Wang, X., 2008. Chemocline instability and isotope variations of the Ediacaran Doushantuo basin in south china. Science in China Series D-Earth Sciences 51, 1560–1569. Jiang, G., Wang, X., Shi, X., Zhang, S., Xiao, S., Dong, J., 2010. Organic carbon isotope constraints on the dissolved organic carbon (DOC) reservoir at the Cryogenian– Ediacaran transition. Earth and Planetary Science Letters 299, 159–168. Jiang, G., Shi, X., Zhang, S., Wang, Y., Xiao, S., 2011. Stratigraphy and paleogeography of the Ediacaran Doushantuo Formation (ca. 635–551 Ma) in South China. Gondwana Research 19, 831–849. Johnston, D.T., Poulton, S.W., Dehler, C., Porter, S., Husson, J., Canfield, D.E., Knoll, A.H., 2010. An emerging picture of Neoproterozoic ocean chemistry: insights from the Chuar Group, Grand Canyon, USA. Earth and Planetary Science Letters 209, 64–73. Kah, L.C., Lyons, T.W., Frank, T.D., 2004. Low marine sulphate and protracted oxygenation of the Proterozoic biosphere. Nature 431, 834–838. Kaufman, A.J., Knoll, A.H., 1995. Neoproterozoic variations in the C-isotopic composition of seawater: stratigraphic and biogeochemical implications. Precambrian Research 73, 27–49. Kaufman, A.J., Corsetti, F.A., Varni, M.A., 2007. The effect of rising atmospheric oxygen on carbon and sulfur isotope anomalies in the Neoproterozoic Johnnie Formation, Death Valley, USA. Chemical Geology 237, 47–63. Kendall, B., Reinhard, C.T., Lyons, T.W., Kaufman, A.J., Poulton, S.W., Anbar, A.D., 2010. Pervasive oxygenation along late Archaean ocean margins. Nature Geoscience 3, 647–652. Knoll, A.H., Walter, M.R., Narbonne, G.M., Christie-blick, N., 2006. The Ediacaran Period: a new addition to the geologic time scale. Lethaia 39, 13–30. Li, C., Love, G.D., Lyons, T.W., Fike, D.A., Sessions, A.L., Chu, X., 2010. A stratified redox model for the Ediacaran ocean. Science 328, 80–83. Liao, W., Wang, Y., Kershaw, S., Weng, Z., Yang, H., 2010. Shallow-marine dysoxia across the Permian–Triassic boundary: evidence from pyrite framboids in the microbialite in South China. Sedimentary Geology 232, 77–83. Liu, P., Yin, C., Chen, S., Tang, F., Gao, L., 2010. Affinity, distribution and staratigraphic significance of tubular microfossils from the Ediacaran Doushantuo Formation in South China. Acta Palaeontologica Sinica 49, 308–324. Lyons, T.W., Anbar, A.D., Severmann, S., Scott, C., Gill, B.C., 2009. Tracking euxinia in the ancient ocean: a multiproxy perspective and Proterozoic case study. Annual Review of Earth and Planetary Sciences 37, 507–534. McFadden, K.A., Huang, J., Chu, X., Jiang, G., Kaufman, A.J., Zhou, C., Yuan, X., Xiao, S., 2008. Pulsed oxidation and biological evolution in the Ediacaran Doushantuo Formation. Proceedings of the National Academy of Sciences of the United States of America 105, 3197–3202. McFadden, K.A., Xiao, S., Zhou, C., Kowalewski, M., 2009. Quantitative evaluation of the biostratigraphic distribution of acanthomorphic acritarchs in the Ediacaran Doushantuo Formation in the Yangtze Gorges area, South China. Precambrian Research 173, 170–190. Neumann, T., Rausch, N., Leipe, T., Dellwig, O., Berner, Z., Bottcher, M.E., 2005. Intense pyrite formation under low-sulfate conditions in the Achterwasser lagoon, SW Baltic Sea. Geochimica et Cosmochimica Acta 69, 3619–3630. Nielsen, J.K., Shen, Y., 2004. Evidence for sulfidic deep water during the late Permian in the East Greenland Basin. Geology 32, 1037–1040. Poulton, S.W., Canfield, D.E., 2011. Ferruginous conditions: a dominant feature of the ocean through earth's history. Elements 7, 107–112. Poulton, S.W., Fralick, P.W., Canfield, D.E., 2010. Spatial variability in oceanic redox structure 1.8 billion years ago. Nature Geoscience 3, 486–490. Raiswell, R., 1982. Pyrite texture, isotopic composition and the availability of iron. American Journal of Science 282, 1244–1263. Raiswell, R., Newton, R., Wignall, P.B., 2001. An indicator of water-column anoxia: resolution of biofacies variations in the Kimmeridge clay (Upper Jurassic, U.K.). Journal of Sediment Research 71, 286–294. Reinhard, C.T., Raiswell, R., Scott, C., Anbar, A.D., Lyons, T.W., 2009. A late Archean sulfidic sea stimulated by early oxidative weathering of the continents. Science 326, 713–716. Scott, C., Lyons, T.W., Bekker, A., Shen, Y., Poulton, S.W., Chu, X., Anbar, A.D., 2008. Tracing the stepwise oxygenation of the Proterozoic ocean. Nature 452, 456–459. Shen, W., Lin, Y., Xu, L., Li, J., Wu, Y., Sun, Y., 2007. Pyrite framboids in the Permian– Triassic boundary section at Meishan, China: evidence for dysoxic deposition. Palaeogeography, Palaeoclimatology, Palaeoecology 253, 323–331. Shen, Y., Zhang, T., Hoffman, P.F., 2008. On the coevolution of Ediacaran oceans and animals. Proceedings of the National Academy of Sciences of the United States of America 105, 7376–7381.

227

Tang, F., Yin, C., Bengtson, S., Liu, P., Wang, Z., Gao, L., 2008. Octoradiate spiral organisms in the Ediacaran of South China. Acta Geologica Sinica-English Edition 82, 27–34. Wang, J., Li, Z.X., 2003. History of Neoproterozoic rift basins in South China: implications for Rodinia break-up. Precambrian Research 122, 141–158. Wang, X., Shi, X., 2009. Spatio-temporal carbon isotope variation duing the Ediacaran period in South China and its impact on bio-evolution. Science in China Series D-Earth Sciences 52, 1520–1528. Wang, Y., Wang, X., 2006. The holdfasts of macroalgae in the Neoproterozoic Doushantuo formation in Northeastern Guizhou province and their environmental significance. Acta Micropalaeontologica Sinica 23, 154–164. Wang, Y., Wang, X., Huang, Y., 2007. Macroscopic algae from the Ediacaran Doushantuo formation in Northeast Guizhou, South China. Earth Science — Journal of China University of Geosciences 32, 828–844. Wignall, P.B., Newton, R., 1998. Pyrite framboid diameter as a measure of oxygen deficiency in ancient mudrocks. American Journal of Science 298, 537–552. Wignall, P.B., Newton, R., Brookfield, M.E., 2005. Pyrite framboid evidence for oxygenpoor deposition during the Permian–Triassic crisis in Kashmir. Palaeogeography, Palaeoclimatology, Palaeoecology 216, 183–188. Wignall, P.B., Bond, D.P.G., Kuwahara, K., Kakuwa, Y., Newton, R.J., Poulton, S.W., 2010. An 80 million year oceanic redox history from Permian to Jurassic pelagic sediments of the Mino-Tamba terrane, SW Japan, and the origin of four mass extinctions. Global and Planetary Change 71, 109–123. Wijsman, J.W.M., Middelburg, J.J., Herman, P.M.J., Böttcher, M.E., Heip, C.H.R., 2001. Sulfur and iron speciation in surface sediments along the northwestern margin of the Black Sea. Marine Chemistry 74, 167–180. Wilkin, R.T., Arthur, M.A., 2001. Variations in pyrite texture, sulfur isotope composition, and iron systematic in the Black Sea: evidence for Late Pleistocene to Holocene excursions of the O2–H2S redox transition. Geochimica et Cosmochimica Acta 65, 1399–1416. Wilkin, R.T., Barnes, H.L., Brantley, S.L., 1996. The size distribution of framboidal pyrite in modern sediments: an indicator of redox conditions. Geochimica et Cosmochimica Acta 60, 3897–3912. Wilkin, R.T., Arthur, M.A., Dean, W.E., 1997. History of water-column anoxia in the Black Sea indicated by pyrite framboid size distributions. Earth and Planetary Science Letters 148, 517–525. Xiao, S., Knoll, A.H., 2000. Phosphatized animal embryos from the Neoproterozoic Doushantuo Formation at Weng'an, Guizhou, South Chaina. Journal of Paleontology 74, 767–788. Xiao, S., Zhang, Y., Knoll, A.H., 1998. Three-dimensional preservation of algae and animal embryos in a Neoproterozoic phosphorite. Nature 391, 553–558. Xiao, S., Yuan, X., Steiner, M., Knoll, A.H., 2002. Macroscopic carbonaceous compressions in a terminal Proterozoic shale: a systematic reassessment of the Miaohe biota, South China. Journal of Paleontology 76, 347–376. Xiao, S., McFadden, K.A., Peek, S., Kaufman, A.J., Zhou, C., Jiang, G., Hu, J., 2012. Integrated chemostratigraphy of the Doushantuo Formation at the northern Xiaofenghe section (Yangtze Gorges, South China) and its implication for Ediacaran stratigraphic correlation and ocean redox models. Precambrian Research 192–195, 125–141. Yin, L., Zhu, M., Knoll, A.H., Yuan, X., Zhang, J., Hu, J., 2007. Doushantuo embryos preserved inside diapause egg cysts. Nature 446, 661–663. Yuan, X., Wang, D., Xiao, S., 2009. Animals in the Neoproterozoic Doushantuo epoch. Acta Palaeontologica Sinica 48, 375–389. Yuan, X., Chen, Z., Xiao, S., Zhou, C., Hua, H., 2011. An early Ediacaran assemblage of macroscopic and morphologically differentiated eukaryotes. Nature 470, 390–393. Zhang, S., Jiang, G., Zhang, J., Song, B., Kennedy, M.J., Christie-blick, N., 2005. U–Pb sensitive high-resolution ion microprobe ages from the Doushantuo Formation in south China: constraints on late Neoproterozoic galciations. Geology 33, 473–476. Zhao, Y., Chen, M., Peng, J., Yu, Y., He, M., Wang, Y., Yang, R., Wang, L., Zhang, Z., 2004. Discovery of a Miaohe-type biota from the Neoproterozoic Doushantuo formation in Jiangkou county, Guizhou province, China. Chinese Science Bulletin 49, 2224–2226. Zhao, Y., Wu, M., Peng, J., Yang, X., Yang, R., Yang, Y., 2010. Triridged lobe fossils from the Miaohe biota from the Ediacaran Doushantuo formation from Jiangkou county, Guizhou provience, SW China. Acta Micropalaeontologica Sinica 27, 305–314. Zhou, C., Xiao, S., 2007. Ediacaran δ13C chemostratigraphy of South China. Chemical Geology 237, 89–108. Zhou, C., Jiang, S., 2009. Palaeoceanographic redox environments for the lower Cambrian Hetang Formation in South China: evidence from pyrite framboids, redox sensitive trace elements, and sponge biota occurrence. Palaeogeography, Palaeoclimatology, Palaeoecology 271, 279–286. Zhu, M., Zhang, J., Yang, A., 2007. Integrated Ediacaran (Sinian) chronostratigraphy of South China. Palaeogeography, Palaeoclimatology, Palaeoecology 254, 7–61. Zhu, M., Gehling, J.G., Xiao, S., Zhao, Y., Droser, M.L., 2008. Eight-armed Ediacara fossil preserved in contrasting taphonomic windows from China and Australia. Geology 36, 867–870. Zhu, M., Lu, M., Zhang, J., Zhao, F., Li, G., Yang, A., Zhao, X., Zhao, M., in press. Carbon isotope chemostratigraphy and sedimentary facies evolution of the Ediacaran Doushantuo Formation in western Hubei, South China. Precambrian Research. doi:10.1016/j.precamres.2011.07.019.