Spatial and temporal evolution of Ediacaran carbon and sulfur cycles in the Lower Yangtze Block, South China

Journal Pre-proof Spatial and temporal evolution of Ediacaran carbon and sulfur cycles in the Lower Yangtze Block, South China Wei Wang, Chengguo Guan, Yongliang Hu, Huan Cui, A.D. Muscente, Lei Chen, Chuanming Zhou PII:

S0031-0182(19)30536-X

DOI:

https://doi.org/10.1016/j.palaeo.2019.109417

Reference:

PALAEO 109417

To appear in:

Palaeogeography, Palaeoclimatology, Palaeoecology

Received Date: 7 June 2019 Revised Date:

21 October 2019

Accepted Date: 21 October 2019

Please cite this article as: Wang, W., Guan, C., Hu, Y., Cui, H., Muscente, A.D., Chen, L., Zhou, C., Spatial and temporal evolution of Ediacaran carbon and sulfur cycles in the Lower Yangtze Block, South China, Palaeogeography, Palaeoclimatology, Palaeoecology, https://doi.org/10.1016/ j.palaeo.2019.109417. This is a PDF file of an article that has undergone enhancements after acceptance, such as the addition of a cover page and metadata, and formatting for readability, but it is not yet the definitive version of record. This version will undergo additional copyediting, typesetting and review before it is published in its final form, but we are providing this version to give early visibility of the article. Please note that, during the production process, errors may be discovered which could affect the content, and all legal disclaimers that apply to the journal pertain. © 2019 Published by Elsevier B.V.

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Spatial and temporal evolution of Ediacaran carbon and sulfur cycles in the Lower

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Yangtze Block, South China

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Wei Wang 1*, Chengguo Guan 1, Yongliang Hu 1, Huan Cui 1,2,3, A. D. Muscente 4, Lei Chen5,

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Chuanming Zhou 1

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Palaeontology and Center for Excellence in Life and Palaeoenvironment, Chinese Academy

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of Sciences, 39 East Beijing Road, Nanjing 210008, China

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State Key Laboratory of Palaeobiology and Stratigraphy, Nanjing Institute of Geology and

Research Group of Analytical, Environmental and Geo- Chemistry, Division of Earth

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System Science, Vrije Universiteit Brussel, Brussels 1050, Belgium

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Research Consortium, Belgium

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Minerals, Shandong University of Science and Technology, Qingdao, Shandong 266590,

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China

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Corresponding authors: W. Wang, Tel: 86 25 83284311; fax: +86 25 83357026; e-mail:

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[email protected]

ET-HOME (Evolution and Tracers of the Habitability of Mars and Earth) Astrobiology

Department of Geology, Cornell College, Mount Vernon, Iowa, 52314, USA. Shandong Provincial Key Laboratory of Depositional Mineralization and Sedimentary

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Abstract:

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Major radiations of microscopic and macroscopic eukaryotes occurred respectively in the

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early and middle Ediacaran Period. Various hypotheses have been proposed to attribute these 1

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evolutionary events to changes in ocean redox conditions. To date, published models of the

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Ediacaran ocean in South China have largely been focused on the Upper and Middle Yangtze

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blocks, as opposed to the slope and basin sections of the Lower Yangtze Block, where the

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oldest complex macrofossils of the Ediacaran (i.e. the ~600 Ma Lantian Biota) are preserved.

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To achieve a holistic understanding of the record of ocean chemistry in the entire Yangtze

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Block, we carried out carbon (δ13Ccarb and δ13Corg) and sulfur (δ34SCAS and δ34Spyr) isotopic

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investigations of four new sections (i.e., Wujialing, Meishuxia, Wangfu and Zhushuwu) in

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the Lower Yangtze Block, in addition to one previously reported drill core at Lantian. The

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Doushantuo and Lantian formations at these sections represent shelf margin, upper slope,

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lower slope, and deep basin environments. In conjunction with previous work on sections of

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the Upper and Middle Yangtze blocks, our study reconstructs the redox structure of the entire

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Ediacaran ocean in South China. The new results show profound δ13Ccarb negative excursions

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in the slope and deep basinal sections, which can be correlated to the EN3 and Shuram

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excursions at a global scale. Notably, the δ13Ccarb and δ34Spyr profiles from shelf to basinal

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sections show distinct spatial and temporal patterns, which can be explained by model of a

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redox stratified ocean during the early and middle Ediacaran Period. The finding of

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isotopically light (34S-depleted) pyrite from this time suggests that the oceanic chemocline

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extended to intermediate depth. Integrating this study with existing ecological and

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taphonomic data, it is likely that marine redox stratification in the early Ediacaran limited the

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spatial distribution of early metazoans in the Lower Yangtze Block.

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Key words: Ediacaran ocean, carbon isotopes, sulfur isotopes, redox stratification, spatial

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variation, Lantian Biota 2

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1. Introduction

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The biogeochemical evolution of the Precambrian ocean involved three major transitions

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with respect to oxygen availability. After a ‘whiff’ of oxygenation in the Archean Eon (Anbar

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et al., 2007; Kaufman et al., 2007; Ostrander et al., 2019) and Great Oxidation Event around

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2.4–2.1 Ga (Holland, 1992, 2006; Kasting, 1993; Lyons et al., 2014), deep ocean water

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ultimately became oxygenated over a prolonged period of time in the Neoproterozoic (1–0.54

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Ga), such that, the oceanic chemocline probably reached the sediment-water interface at the

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end of the era (Canfield and Teske, 1996; Shields-Zhou and Och, 2011; Och and

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Shields-Zhou, 2012). This prolonged phase of deep ocean oxygenation coincided with the

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radiation and extinction of eukaryotes (Riedman et al., 2014; Muscente et al., 2018, 2019;

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Darroch, 2018), including the rapid rise and diversification of complex eukaryotes in the

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Ediacaran (Yuan et al., 2011; Droser and Gehling, 2015; Xiao et al., 2016). For this reason, a

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number of hypotheses attribute the radiation and extinction to transient and secular changes

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in the redox architecture of the Ediacaran ocean (Canfield et al., 2007; Wood et al., 2015; Cui

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et al., 2016; Bowyer et al., 2017; Muscente et al., 2018; Zhang et al., 2018).

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Redox reconstructions of the Ediacaran ocean have emerged from data on numerous

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different geochemical proxies. Most models suggest that ferruginous conditions (anoxia with

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free iron) were widespread in the deep ocean water (Canfield et al., 2007, 2008; Li et al.,

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2010; 2015a; Jin et al., 2018; Wang et al., 2018). This widespread ferruginous deep water

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likely persisted until the early Paleozoic (Dahl et al. 2010; Sperling et al. 2015; Jin et al.,

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2016), when the rise of land plants may have significantly increased atmospheric pO2 and

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contributed to oxygenation of the deep ocean (Bergman et al. 2004). However, such redox 3

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model is often inconsistent with the occurrence of benthic macroscopic and morphologically

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complex fossils that are preserved in situ within Ediacaran black shales, such as the Lantian,

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Miaohe, and Wenghui biotas sampled from the Lantian and Doushantuo formations of South

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China (Xiao et al., 2002; Yuan et al., 2011; Wang Y. et al., 2014). This inconsistency may

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indicate that iron speciation records long-term trends in ocean chemistry rather than

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short-term redox fluctuations affecting habitat viability and fossil preservation (Guan et al.,

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2014; Cheng et al., 2017; Sperling et al., 2016, 2018; Muscente et al., in press). Indeed, other

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geochemical proxies provide evidence for, at least, partial oxygenation of deep water during

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the Ediacaran Period, including basinal enrichment of sedimentary redox-sensitive trace

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metals like Mo and V (Sahoo et al., 2012), rising Ediacaran δ98/95Mo in euxinic shales (Chen

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et al., 2015), widespread higher δ15N values from deep water facies (Wang et al., 2017, 2018;

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Chen et al., 2019), and oscillating pyrite δ34S signals in both slope and deep-water facies

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(Sahoo et al., 2016; Wang et al., 2017; Shi et al., 2018). In this context, efforts to reconstruct

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ancient ocean redox conditions should utilize multiple geochemical proxies, as they reflect

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different reduction systems and timescales of response to redox oscillation (Cheng et al.,

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2017).

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The Yangtze Block contains fossiliferous successions of Ediacaran strata representing

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shelf to basinal facies (Zhou and Xiao, 2007; Zhu et al., 2007; Jiang et al., 2011; Zhou et al.,

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2019). Thus, it provides an opportunity to reconstruct the spatial redox conditions of the

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Ediacaran ocean. For instance, an episodic ocean oxygenation process has been proposed

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based on carbon and sulfur isotopic variations through a shallow water Ediacaran succession

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in the Yangtze Gorges area (McFadden et al., 2008). After that, oceanic redox heterogeneity 4

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was demonstrated in detail on the basis of facies-dependent variations of iron speciations,

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carbon and sulfur isotopic compositions across the Yangtze Block (e.g., Li et al., 2010, 2015a,

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b; Shi et al., 2018).

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Geographically, the Yangtze Block of South China is divided into the Upper, Middle,

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and Lower Yangtze blocks. The areas of each part are subdivided on basis of the upper,

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middle and lower reaches of the modern Yangtze River. The lower block represents the

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eastern part of the Yangtze Block (Fig. 1a) (Cao et al., 1989). Historically, most studies have

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focused on sections in the Upper and Middle blocks (McFadden et al., 2008; Li et al., 2010,

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2015a, b; Sahoo et al., 2016), while the Ediacaran successions in the Lower Block have

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received less attention. Accordingly, developing a complete model of the redox architecture

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of the Ediacaran ocean in South China requires analysis of sections in the Lower Yangtze

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Block. In this context, we present data on four new sections, including the Zhushuwu section

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in southern Anhui, and Wangfu, Meishuxia and Wujialing sections in western Zhejiang, as

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well as one previously reported drill core section at Lantian in southern Anhui (Wang et al.,

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2017). All of these sections are located in the Lower Yangtze Block, and collectively

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represent early-to-middle Ediacaran depositional environments spanning a variety of water

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depths. Thus, they can provide insights into redox conditions across a shallow-to-deep water

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transect. Temporal profiles of high-resolution carbon (δ13Ccarb and δ13Corg) and sulfur (δ34SCAS

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and δ34Spyr) isotopic values complied with isotopic comparison amongst different

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depositional facies are presented in this paper. Together, they provide empirical support for a

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redox stratified ocean during the early and middle Ediacaran.

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2. Geological background

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2.1. Facies change across the Lower Yangtze Block

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The Ediacaran successions of the Lower Yangtze Block (Fig. 1a) were deposited on the

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southeastern side of rifted, passive continental margin (Jiang et al., 2011). These successions

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exhibit significant changes in both lithology and thickness, which are exemplified by classic

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sections across the shelf-to-basinal transect (Cao et al., 1989, Fig. 1b). In the northwest, the

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successions are dominated by black carbonaceous shale interbedded with grey argillaceous

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dolostone, limestone, and black chert, whereas in the southeast region, they are characterized

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by siltstone, limestone, and dolostone, with dolomitic stromatolites occurring in the Jiangshan

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area during the late Ediacaran (Fig. 1b). These changes in lithology indicate that the

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environments became progressively deeper from southeast to northwest (Cao et al., 1989).

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2.2. New sections in this study All samples in this study were collected from four new outcrop sections (Fig. 1, Table 1),

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including the Wujialing (shelf margin; 28.848217°N, 118.652500°E ) (Fig. 2a, b, c),

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Meishuxia (upper slope facies; 30.103800°N, 119.644667°E) (Fig. 2d, e, f), Wangfu (lower

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slope facies; 29.952233°N, 118.897783°E) (Fig. 2g, h, i), and Zhushuwu sections (deep

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basinal facies; 29.991567°N, 118.817100°E) (Fig. 2j, k, l). In these sections, the Ediacaran

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strata belong to the Lantian/Doushantuo and Piyuancun/Dengying formations. The

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Doushantuo and Dengying formations change toward northwest to the stratigraphically

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equivalent Lantian and Piyuancun formations, respectively. This study focuses on the

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Doushantuo Formation and equivalent Lantian Formation. At Wujialing (Fig. 1c), the 6

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Doushantuo Formation consists of medium-bedded dolostone (Figs. 2a, 3a), limestone (Fig.

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3b), and dolomitic siltstone (Fig. 2b), representing shallow-water platform facies. At

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Zhushuwu (Fig. 1c), the Lantian Formation mainly consists of organic rich black shales (Figs.

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2j, 3c), thin-bedded dolostones (Figs. 2k, 3d), and dolomitic siltstones, which were deposited

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in the deep basinal facies. At Meishuxia (Fig. 1c), the Doushantuo Formation includes

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limestone interbedded with dolostone (Fig. 3e), argillaceous dolostone (Fig. 3f), and

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limestone. The overlying Dengying Formation at Meishuxia consists of sandstones (Fig. 1c)

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and carbonate rocks, indicating relatively shallow-water deposits. At Wangfu (Fig. 1c), it

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consists mainly of light grey siltstone in the lower part, dark grey mudstone and limestone

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(Fig. 3g, h) in the middle part, and dolostone (Fig. 3i) and mudstone in the upper part of the

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Lantian Formation. The predominance of chert in the overlying Piyuancun Formation

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resembles that at the Zhushuwu and Lantian sections (Figs. 2l, 4), indicating deposition in

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relatively deep-water environment. Therefore, we tend to interpret that the Ediacaran

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successions at Meishuxia and Wangfu were deposited in upper and lower slope facies,

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respectively, though the slumping structures that are common in the slope facies of the Upper

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Yangtze Block have not been discovered in these two sections. The ages of the strata in these

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new sections are supported by lithostratigraphic and carbon isotope (δ13C) chemostratigraphic

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correlations with the Lantian Formation in the Lantian drill core, and the Doushantuo

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Formation at Jiulongwan in the Yangtze Gorges area (Fig. 4) (Yuan et al., 2011).

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3. Analytical methods Prior to geochemical analysis, sample surfaces were removed to limit measurements of 7

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contaminations introduced by drilling fluid and/or weathering. Geochemical analyses of

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δ34SCAS, δ34Spyr, δ13Ccarb, δ13Corg, δ18Ocarb and total organic carbon content (TOC) were

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conducted at the Nanjing Institute of Geology and Palaeontology, Chinese Academy of

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Sciences.

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3.1. Carbonate carbon and oxygen isotope analysis An aliquot of 80~100 µg powder was allowed to react with orthophosphoric acid for

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150~200 s at 72 °C in a Kiel IV carbonate device connected to a MAT 253 mass

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spectrometer. CO2 generated from this reaction was introduced to the mass spectrometer for

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δ13Ccarb and δ18Ocarb analysis. The standard GBW-04405, with a δ13Ccarb value of 0.57 ± 0.03‰

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and a δ18Ocarb value of ‒8.49 ± 0.13‰ (VPDB), was used for analytical calibration.

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Analytical precision was better than 0.08‰ for δ13Ccarb (VPDB) and 0.1‰ for δ18Ocarb

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(VPDB).

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3.2. Organic carbon isotope analysis An aliquot of ~3 g sample powder was decarbonated overnight using 5% (V/V) HCl in a

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polypropylene centrifuge tube. After 24 h, the acid solution was centrifugated, and the

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supernatant was removed. This process was repeated at least three times to ensure complete

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removal of carbonates. The carbonate-free residue was then neutralized using deionized water

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with a stepwise washing procedure, and then totally dried at 70 °C in an oven. Samples were

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loaded into capsules and flash combusted at 1060 °C in an Organic Elemental Analyzer

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(FLASH 2000) fitted with zero blank autosampler. The resulting CO2 gas was measured by 8

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thermal conductivity detector to obtain TOC content (wt%), then transferred by continuous

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flow for isotopic ratio determination on a Delta V Advantage Isotope Ratio Mass

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Spectrometer. 13C/12C ratios were reported in the conventional delta notation as per mil

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deviation from the VPDB standard. Calibration of δ13C values was accomplished using a

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working standard B2151 (δ13Corg = –26.27‰), GBW04407 (black carbon, δ13Corg = –22.43‰)

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and UREA (δ13Corg = –39.79‰). Sample δ13Corg values were repeatable to a standard error of

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< 0.1‰.

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3.3. Sulfur isotope analysis

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To remove soluble sulfates (e.g., from sulfide oxidation) and water-soluble sulfate

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minerals (e.g., anhydrite), carbonate powder was rinsed with 10% NaCl solutions. This

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process was repeated at least three times until no sulfate was detected in filtered solutions.

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After dissolving the NaCl-rinsed powders in HCl for 2 hours, acid solution was filtered.

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BaCl2 solution was added into the supernatant to precipitate BaSO4 overnight for δ34SCAS

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analysis.

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About 3 g sample powder was prepared for pyrite extraction. Chromium reduction

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method was used for pyrite extraction to finally obtain Ag2S precipitate (Canfield et al.,

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1986).

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Dried BaSO4 or Ag2S precipitates were combined with an excess of V2O5 and then

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analyzed on a Delta V Advantage Isotope Ratio Gas Source Mass Spectrometer fitted with a

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peripheral Organic Elemental Analyzer (FLASH 2000) for on-line sample combustion. Sulfur

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isotope compositions were expressed in standard δ-notation as per mil (‰) deviation from 9

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Vienna Canyon Diablo Troilite (VCDT), with an analytical error of < 0.2‰, calculated from

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replicate analyses of samples and laboratory standards. Analyses were calibrated using four

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international standards including NBS-127 (BaSO4 = 20.3‰), IAEA-S-1 (Ag2S, δ34S = –

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0.3‰), IAEA-S-2 (Ag2S, δ34S = 22.7‰) and IAEA-S-3 (Ag2S, δ34S = –32.3‰).

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4. Results

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4.1. Carbon isotope chemostratigraphy

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The δ13Ccarb values of cap carbonates (dolostones overlying the Cryogenian tillite) in the

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Wujialing, Meishuxia, Wangfu, and Zhushuwu sections are around –5‰. The stratigraphic

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profiles of the sections at Meishuxia and Wangfu contain a profound δ13Ccarb negative

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excursion in the middle part of the Doushantuo/Lantian formation, wherein values drop to –

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11.2‰ and –17.7‰, respectively. Above these excursions, the δ13Ccarb values of the

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Doushantuo/Lantian formation in these two sections rise and stabilize at 0‰ (Fig. 5a, b).

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Although the δ13Ccarb values in the Zhushuwu section are almost entirely negative, their

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stratigraphic profile contains a similar trend with negative shift in δ13Ccarb (low value of –

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10.2‰) in the middle part of the Lantian Formation (Fig. 5c). Likewise, most of the δ13Ccarb

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values in the Wujialing section are positive, but their stratigraphic profile shows a mild

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negative shift from 4.6‰ to 0‰ in the middle part of the Doushantuo Formation (Fig. 5d).

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Thus, despite different δ13Ccarb values, all of the sections record the existence of negative

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δ13Ccarb excursions.

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With the exception of the Wujialing section, the δ13Corg variation in all the sections is

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largely decoupled from the change in δ13Ccarb (Table S1, Fig. 5). The δ13Corg values in the

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Meishuxia section range from –31.5‰ to –21.1‰ with an average value of –26.1 ± 2.1‰ 10

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(Fig. 4a). At Wangfu, the δ13Corg values exhibit greater variation ranging from –27.6‰ to –

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10.7‰ with an average value of –22.9 ± 4.7‰. The stratigraphic profile of this section

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provides a positive shift in δ13Corg coupled to the δ13Ccarb positive excursion in the upper of

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the Doushantuo/Lantian formation (Fig. 5b). Compared with the δ13Corg values at Wangfu, the

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δ13Corg values in the Zhushuwu and Wujialing sections have relatively narrow ranges. The

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δ13Corg values range from –31.0‰ to –25.3‰ with an average value of –28.1 ± 1.3‰ for the

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Zhushuwu section, and from –33.3‰ to –24.0‰ with an average value of –28.3 ± 2.3‰ for

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the Wujialing section, respectively (Fig. 5c, d).

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4.2. Sulfur isotope chemostratigraphy The stratigraphic δ34Spyr profiles of the Wujialing, Meishuxia and Wangfu sections show

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broad similarities (Fig. 5). Above the highly positive values of the cap carbonate, the δ34Spyr

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values significantly decrease from 25.0‰ to –21.0‰ at Meishuxia, 37.6‰ to –20.8‰ at

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Wangfu, and 24.2‰ to –29.0‰ at Wujialing. Further up in these sections, the δ34Spyr values

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return to positive values, which are recorded up to the top of the Doushantuo/Lantian

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formation (Fig. 5a, b, d). This negative shift in δ34Spyr was not observed above the cap

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carbonate at Zhushuwu, where the exposure of the lower part of the section is poor. Instead,

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all measurements of δ34Spyr returned to positive values, ranging from 1.5‰ to 36.4‰, for the

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Zhushuwu section (Fig. 5c).

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The four outcrop sections differ with regard to their stratigraphic δ34SCAS profiles. In

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general, the profiles exhibit average values of 21.3 ± 4.0‰ and 22.4 ± 7.4‰ for the

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Zhushuwu and Meishuxia sections, respectively (Fig. 5a, c). These δ34SCAS values are

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consistent with the δ34S value of modern ocean water (about 21.0‰), and are also similar to

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those reported from the Lantian drill-core section, where the average δ34SCAS value is 20.4 ±

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6.8‰ (Wang et al., 2017). The δ34SCAS profiles at Wangfu and Wujialing have lower (10.5 ±

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8.0‰) and higher (27.8 ± 7.3‰) mean values, respectively, with regard to Zhushuwu and

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Meishuxia (Fig. 5b, d).

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4.3 Other results

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The TOC values range from 0.0% to 0.5% at Meishuxia, 0.0% to 4.0% at Wangfu, 0.0%

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to 0.6% at Wujialing, and 0.0% to 5.8% at Zhushuwu. Total pyrite content ranges from 0.0%

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to 1.2% at Meishuxia, 0.0% to 13.5% at Wangfu, 0.0% to 1.5% at Wujialing, and 0.0% to 6.2%

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at Zhushuwu. Statistical analysis shows that the shallow-water facies (Wujialing and

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Meishuxia) contain significantly less organic carbon and pyrite than the relatively deep-water

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facies (Wangfu and Zhushuwu) (Mann-Whitney test, p < 0.01) (Figs. 5, 6).

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Except for the Wujialing section, the δ18O values in the sections are, by and large, lower

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than −10‰ (Fig. 5). They covary to a degree with the δ13Ccarb values in the Meishuxia section

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(Fig. 5a). Nonetheless, the variables are decoupled in the other three sections (Fig. 5b, c, d).

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5. Discussion

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5.1 Evaluation of post-depositional alteration

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5.1.1. Evaluation of the δ34Spyr, δ13Ccarb, and δ13Corg data

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Several lines of petrographic and geochemical evidence support the interpretation that the measured δ34Spyr, δ13Ccarb, and δ13Corg values represent primary signals (Figs. 3, 7). First, 12

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the carbonate rocks in most samples consist of fine-grained calcite (calcite crystals with

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diameters between 4 µm and 64 µm), argillaceous limestone, and dolomicrite that exhibit no

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obvious features of diagenetic alteration, such as dissolution and overgrowth (Fig. 3).

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Dolosparite (dolomite crystals diameter >30 µm) interbedded with fine-grained calcite occurs

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in the middle part of the Doushantuo Formation at the Meishuxia section, but δ13Ccarb values

271

do not show much difference between these two lithologies (e.g., Figs. 2f, 3e). Indeed, the

272

δ13Ccarb value of the dolosparite sample is −9.1‰ (WY1507-50 black), only 2‰ higher than

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the fine-grained calcite (WY1507-50 white) (Table S1).

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Second, sedimentary pyrites in the samples are mostly preserved as disseminated

275

microcrystals or framboidal pyrites with diameters between 5 µm and 50 µm but mostly less

276

than 10 µm (Fig. 3d, h, i). These grains are interpreted to have formed in an anoxic water

277

column (syngenetic pyrites) and/or within sediments immediately below the redox boundary

278

between oxic and anoxic conditions (Berner, 1982). Their δ34S values, therefore, may record

279

the primary isotopic composition of seawater, or alternatively, porewater that developed

280

beneath bottom seawater (Machel et al., 2001). Although coarse pyrites (visible to the naked

281

eye) have also been observed in some samples (e.g., Fig. 3c), they were carefully removed

282

during crushing and geochemical analysis so that they would not affect bulk δ34Spyr values.

283

The lack of a clear correlation between δ34Spyr values and total pyrite (wt %) indicate that the

284

addition of late-digenetic pyrite did not have a strong influence on δ34Spyr values (Fig. 7d, h, l,

285

p).

286 287

Third, although the low δ18O values (< −10.0‰) in some sections indicate late digenetic alteration (Kaufman and Knoll, 1995), the correlation between δ13Ccarb and δ18O values is 13

288

negligible (Fig. 7a, e, i, m). Ergo, the δ13Ccarb values most likely reflect original seawater

289

signals due to the buffering effect of carbon in carbonate-dominated rocks.

290

Finally, δ13Corg data of the measured four sections, exhibiting no correlation with TOC

291

(Fig. 7c, g, k, o), are relatively consistent, except for a few outliers with values higher than

292

−15‰ at the Wangfu section. These outliers may result from incomplete carbonate removal

293

during acidification. Based on the above reasoning, we interpret the trends of δ34Spyr, δ13Ccarb,

294

and δ13Corg data as mostly primary signals, which can be used for redox reconstruction.

295 296 297

5.1.2. Evaluation of the δ34SCAS data Due to the large reservoir size and a long residence time of close to 10 Myr of seawater

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sulfate (Chiba and Sakai, 1985), the isotopic composition of seawater sulfate is expected to

299

be homogeneous. δ34SCAS values of the middle Ediacaran (EN3 interval), however, exhibit

300

spatial variation, with a decrease from inner-shelf to upper-slope sections, which may suggest

301

a small sulfate reservoir during that period (Shi et al., 2018). In contrast, the δ34SCAS values in

302

present study show no spatial gradient from shelf to basinal facies, and most of the δ34SCAS

303

values (around 20‰) of the four new sections are statistically lower than the previously

304

reported data (Shields et al., 2004; McFadden et al., 2008; Li et al., 2010; Xiao et al., 2012).

305

In particular, the average δ34SCAS value in the Wangfu section is only 10.5‰, significantly

306

lower than the regular value (around 25‰). Therefore, we speculate that the carbonate

307

associated sulfate (CAS) of our samples has been contaminated by pyrite oxidation or

308

incomplete NaCl rinsing during the extraction process, as both processes could drive the

309

δ34SCAS towards lower values (Marenco et al., 2008; Peng et al., 2014). 14

310 311 312

5.2． ．δ13Ccarb and δ34Spyr chemostratigraphic correlation in the Lower Yangtze Block Although the sections exhibit variation in δ34Spyr and δ13Ccarb, both support a degree of

313

stratigraphic correlations. In general, the cap dolostone intervals are characterized by positive

314

δ34Spyr values, which increase slightly up section (highlighted by green arrows in Fig. 5).

315

Above the cap dolostones, the δ34Spyr profiles at Wangfu, Meishuxia and Wujialing contain

316

short intervals of negative excursion (down to ‒29.0‰) below positive values towards the

317

tops of the sections (highlighted by red arrows in Fig. 5).

318

Previous δ13Ccarb chemostratigraphic studies on the Doushantuo Formation from shallow

319

water platform facies have shown that there are three marked negative excursions, namely

320

EN1 (Ediacaran δ13Ccarb negative excursion 1), EN2, and EN3, occurring at the base, middle,

321

and upper Doushantuo Formation, respectively (Zhou et al., 2007; McFadden et al., 2008).

322

Except for EN2, which is sometimes absent in successions from slope and basinal facies (Zhu

323

et al., 2007; Jiang et al., 2007, Ouyang et al., 2017; Lan et al., 2019), EN1 and EN3 have

324

been widely reported in the Doushantuo Formation across the platform, slope, and basinal

325

facies in the Upper Yangtze Block (Zhu et al., 2007, 2013). Similarly, two δ13Ccarb negative

326

excursions have been found in the Doushantuo/Lantian formation at the Meishuxia, Wangfu

327

and Zhushuwu sections (Figs. 4, 5a, b, c). The absence of EN2 may be due to the coverage

328

that limits sampling and unavailability of lithology suitable for δ13Ccarb analysis in the lower

329

parts of these sections (Fig. 5).

330 331

It is notable that the δ13Ccarb profiles in the upper part of the Doushantuo Formation at Meishuxia and Wangfu sections show consistently positive values. This trend contrasts with 15

332

the δ13Ccarb profiles in the Upper Yangtze Block, where the termination of EN3 typically

333

occurs right near the Doushantuo/Dengying boundary. Considering that a lithostratigraphic

334

boundary can be diachronous among sections, it is possible that the different δ13Ccarb profiles

335

reflect the diachronous nature of the Doushantuo/Dengying boundaries between the Upper

336

and Lower Yangtze blocks.

337

Unlike the Meishuxia, Zhushuwu and Wangfu sections, the shallow-water Wujialing

338

section only shows positive δ13Ccarb values above the cap dolostone (Fig. 5d). The most likely

339

explanation for the absence of an EN3-equivalent excursion in the Wujialing section is its

340

poor exposure; a 20 m-thick covered interval in the upper part of the section may encompass

341

the excursion. Regardless, even if the EN3 interval is present at Wujialing, the negative

342

δ13Ccarb shift in the stratigraphic profile could be small (< 5‰), given the previously reported

343

δ13Ccarb data from the nearby (5–7 km away in distance) Jiangshan (Peng et al., 2006),

344

Meihuishan (Ling et al., 2007), and Jiujiawu (Liu et al., 1992) sections, where the excursions

345

are less than 5‰ (Fig. 4). Thus, it appears unlikely that the Wujialing section records the

346

large magnitude negative δ13Ccarb excursion (down to –17.7‰) that has been observed in the

347

other three sections. We interpret the δ13Ccarb negative excursions at the Meishuxia, Wangfu,

348

and Zhushuwu sections as the expression of Shuram/EN3 anomaly in slope and basinal

349

sections, which has also been identified in the Lantian drill core in southern Anhui Province

350

and the Jiulongwan section in the Yangtze Gorges area (Fig. 4).

351 352 353

5.3. Origin of a large δ13Ccarb gradient in the Yangtze Block Statistical evaluation of the δ13Ccarb data from different depositional facies reveals 16

354

significant variation across the platform, slope, and basinal sections in the Lower Yangtze

355

Block with the exception of the cap dolostone, which does not vary much in its chemical

356

makeup (Fig. 6a). The average δ13Ccarb value obtained from shelf facies sections (0.8 ± 2.9‰),

357

including the Wujialing section and previously reported Jiangshan, Meihuishan and Jiujiawu

358

sections, is significantly higher than those measured from the Meishuxia (–4.6 ± 3.8‰),

359

Wangfu (–7.1 ± 4.8‰), Lantian drill cores (–6.0 ± 3.6‰), and Zhushuwu sections (–7.6 ±

360

1.9‰) (Fig. 6a). This facies-dependent variation in δ13Ccarb may corroborate the existence of

361

a dissolved organic carbon (DOC) reservoir in the deep ocean (Jiang et al., 2007). Anaerobic

362

oxidation of the DOC reservoir may have contributed to the low δ13Ccarb values recorded in

363

the deep water sections. Decoupling of δ13Ccarb and δ13Corg, as observed in the EN3/Shuram

364

intervals of the new sections (Fig. 5), was previously shown in the Doushantuo Formation in

365

the Upper and Middle Yangtze blocks at the inner shelf Northern Xiaofenghe (Xiao et al.,

366

2012), intra-shelf Jiulongwan (McFadden et al., 2008), outer shelf Zhongling (Li et al., 2010;

367

Cui et al., 2015), and slope Siduping sections in South China (Wang et al., 2016), which may

368

indicate the purported DOC reservoir spanned the entire Yangtze Block. If so, the

369

depth-dependent δ13Ccarb gradient in the Yangtze Block may reflect the existence of a

370

widespread stratified ocean during the Ediacaran.

371

Alternatively, the gradient in δ13Ccarb may reflect variation in the input of 13C-depleted

372

authigenic carbonates, wherein, deep water environments received the largest input. Mixing

373

of sedimentary carbonates with authigenic carbonates, which precipitate within sediment, can

374

affect the δ13Ccarb values of sedimentary rocks (Schrag et al., 2013; Cui et al., 2017; Jiang et

375

al., 2019). Of course, unlike outer shelf sections of the Doushantuo Formation (Ader et al., 17

376

2009; Furuyama et al., 2016; Cui et al., 2017; Cui et al., 2019), the slope and basin sections

377

(e.g. the Zhushuwu and Lantian drill-core sections) have not yet yielded any evidence of

378

methane-derived authigenic carbonates. Therefore, future work should investigate whether

379

13

380

facies.

C-depleted authigenic carbonates contribute to the low δ13Ccarb values in the deep-water

381 382 383

5.4. Origin of the δ34Spyr gradient across the Lower Yangtze Block Secular δ34Spyr variation in the geological record provides a basis for reconstructing the

384

redox state of the ocean (Canfield and Teske, 1996). However, the δ34Spyr signals in an

385

individual section (or basin) could also reflect overprints associated with changes of seawater

386

sulfate concentrations and sedimentation rate (Fike et al., 2015; Pasquier et al., 2017; Liu et

387

al., 2019) and/or the primary production and availability of organic matter (Leavitt et al.,

388

2013). Accordingly, efforts to reconstruct spatial patterns in redox state should utilize δ34Spyr

389

data from multiple sections and depositional facies rather than the δ34Spyr profile of a single

390

section. In this context, the δ34Spyr data in the present study provide a framework for

391

reconstructing the oxidation state of the Ediacaran ocean in South China. With the exception

392

of cap carbonate and negative δ34Spyr values intervals (highlighted by green and red arrows in

393

Fig. 5), statistical evaluation of the new δ34Spyr data, along with data compiled from the

394

Lantian drill-core, reveal a depth-dependent gradient between shallow and deep-water facies

395

of the Lower Yangtze Block in the early and middle Ediacaran Period (Fig. 6). The δ34Spyr

396

values of the deep basinal facies Zhushuwu section (16.2 ± 5.3‰), the upper slope facies

397

Meishuxia section (12.3 ± 11.0‰) and the shelf marginal facies Wujialing section (18.8 ± 18

398

10.5‰) are relatively higher than the lower slope facies Wangfu (4.7 ± 4.3‰) and Lantian

399

(2.4 ± 11.0‰) sections (Mann-Whitney test, p < 0.01) (Fig. 6b). The lowest δ34Spyr values

400

occur in the moderately deep-water facies (Wangfu and Lantian sections, Fig. 6b). The

401

observed δ34Spyr pattern may reflect multiple factors affecting sulfur isotope fractionation in

402

marine environments. Oceanic sulfate mainly originates from the terrigenous weathering of

403

pyrite and the dissolution of calcium sulfates (Fig. 8). In marine sediment, sulfate reduction

404

occurs in the anoxic zone, below the depth at which oxygen has been depleted in response to

405

aerobic respiration, and generally speaking, below the depth at which metal phases, such as

406

Mn4+ and Fe3+ oxides, are reduced by anaerobic microorganisms (Fike et al., 2015). The

407

pyrite forms as a consequence of sulfate reduction, and its δ34Spyr value depends on the

408

conditions of the local depositional environment, including the supply of decomposable

409

organic matter, sulfate reduction rate, iron availability, dissolved sulfate concentration and

410

sediment deposition rate (Berner et al., 1982; Canfield, 2001; Lyons and Gill, 2010). Large

411

sulfur isotopic fractionation can occur during sulfate reduction, when sulfate concentrations

412

exceed 200 µM (Habicht et al., 2002), while no obvious sulfur isotopic fractionation (< 6‰)

413

between pyrite and sulfate occurs if sulfate concentration is < 200 µM (e.g., Habicht et al.,

414

2002). However, a recent study of Lake Matano (Indonesia) suggests that sulfate levels below

415

200 µM may, in some cases, lead to large (> 20‰) fractionation (Crowe et al., 2014). In

416

addition to sulfate concentration, organic matter can also affect the fractionation of sulfur

417

isotopes (Fike et al., 2015). Organic molecules represent the major electron donor in

418

reduction of seawater sulfate. Methane generally produces the highest δ34Spyr values among

419

organic substrates, due to the higher reaction rate between sulfate and hydrogen in a process 19

420

known as sulfate-driven anaerobic oxidation of methane (Borowski et al., 1996, 2000; Fike et

421

al., 2015; Lin et al., 2016).

422

It is notable that the δ34Spyr values do not vary substantially in the cap dolostone interval

423

of the sections (Figs. 5, 9a; Table S1). Similar δ34Spyr values have been reported from the cap

424

dolostone intervals in Hubei and Guizhou provinces (e.g., McFadden et al., 2008; Xiao et al.,

425

2012; Sahoo et al., 2016). This consistency among sections likely resulted from the unique

426

depositional environments at the end of the Marinoan glaciation (Zhou et al., 2010; Creveling

427

et al., 2016). During deglaciation, a massive amount of glacial drifts may have been deposited

428

in a short period of time in the platform, slope, and basinal environments. Those deposits

429

likely filled up the basin and flattened preexisting topography across the Yangtze Block, and

430

as a result, the cap dolostones may have been deposited simultaneously across the entire

431

shallow water environment on the Yangtze Block (Zhou et al., 2010; Creveling et al., 2016).

432

In such a scenario, pyrites formed in the loose carbonate mud during early diagenesis and

433

grew into euhedral or subhedral crystals. The positive δ34Spyr values are therefore indicative

434

of low concentrations of pore water sulfate (Fig. 9a) (Machel, 2001).

435

Going stratigraphically upward, the middle and upper parts of the Doushantuo/Lantian

436

formation show an unusual δ34Spyr gradient from shallow to deep water depth (Fig. 6b, 9b).

437

Higher δ34Spyr values occur in both shallow and deep-water facies, while relative lower δ34Spyr

438

values occur in slope facies. It is notable that some modern tropical deltaic muds exhibit

439

similar albeit larger magnitude δ34Spyr depth gradients; δ34Spyr values gradually shift from 36‰

440

in cores of shallow water mud (water depths 8–18 m) to around –25‰ in muds of deep water

441

(depths up to 50 m) (Gao et al., 2013). Because sedimentary pyrites typically form right near 20

442

the oxic-anoxic interface, the spatial gradient of δ34Spyr may indicate a change in the depth of

443

the redox chemoclines among different localities (Sahoo et al., 2016; Wang et al., 2017).

444

Two explanations may be able to account for the observation of higher δ34Spyr values in

445

basinal facies than slope facies. First, sulfate availability may have significantly declined

446

from the shore to the basin, given that the oxidation of terrestrial pyrite represents the primary

447

source of sulfate along continental margins (Li et al., 2010; Shi et al., 2018). A limited sulfate

448

supply, ergo, may have caused higher δ34Spyr values in deep water environments (Habicht et

449

al., 2002; Fike et al., 2015). Second, sulfate availability in the deep ocean may become

450

limited due to excessive consumption by sulfate-reducing microorganisms in the wake of the

451

Marinoan glaciation and development of the Ediacaran deep-ocean DOC reservoir (Rothman

452

et al., 2003; Fike et al., 2006; Jiang et al., 2010). In both cases, lower sulfate concentration in

453

deep waters would have resulted in higher δ34Spyr values due to limited sulfur isotope

454

fractionation between sulfate and sulfide.

455

Alternatively, the δ34Spyr gradient from shallow shelf to slope facies can also be

456

explained by biological sulfide oxidation. It has been shown that partial oxidation of sulfide

457

elevates δ34Spyr values in shelf environments (Fig. 9b) (Fry et al., 1988; Fike et al., 2015).

458

Sulfide oxidation reverses the process of sulfate reduction via the conversion of hydrogen

459

sulfide or sulfide minerals to more oxidized sulfur species. Although biological sulfide

460

oxidation has only a limited effect (~5‰) on isotopic fractionation (Fry et al., 1988; Canfield,

461

2001; Zerkle et al., 2009), this process may have contributed to the overall high δ34Spyr values

462

in shelf margin section if the recycling flux was great in shallow-water environments.

463

Multiple sulfur isotopes (δ34S, ∆33S), which could track the disproportionation reactions, may 21

464 465

help to test this hypothesis in the future work (Fike et al., 2015). Overall, the gradient suggests the existence of a redox stratified Ediacaran ocean,

466

including oxic conditions in shallow-water environments, unstable euxinic conditions in

467

environments of intermediate depth, and relatively constant anoxic conditions in the deep

468

environments (Li et al., 2010; Wang et al., 2017). Pyrites that precipitate in an euxinic water

469

column typically have low δ34Spyr values, because sulfate reducing microbes have full access

470

to sulfate in seawater. In contrast, pyrites that precipitate within surface sediments of shallow

471

water environments with high sedimentation rates often have high δ34Spyr values due to

472

limitations on the diffusion of sulfate from seawater to porewater (Pasquier et al., 2017; Liu

473

et al., 2019). In this context, the lower slope sections were mainly deposited near oxic-anoxic

474

interface, where pyrites may have formed in euxinic water columns (with full access to

475

sulfate supply). On the contrary, shelf sections were mostly deposited in shallow oxic

476

conditions, in which pyrites were formed in anoxic porewaters, and deep basin sections were

477

probably deposited beneath the oxic-anoxic interface in sulfate-limited environments (Fig.

478

9b). Following this simplified model, depositional depth left a signal in the δ34Spyr values of

479

each section, assumed the relatively uniform sedimentary facies during the early-middle

480

Ediacaran Period in the Lower Yangtze Block.

481 482

5.5. Paired spatial depth-gradient of carbon and sulfur isotope compositions in the

483

Ediacaran ocean

484

The paired spatial depth-gradient of carbon and sulfur isotopic values may have been

485

caused by coupled oceanic carbon and sulfur cycling (Fig. 8). In moderate temperatures (< 22

486

100°C), the formation of sedimentary pyrite begins with the activity of prokaryotic

487

microorganisms that reduce sulfate while oxidizing organic matter or hydrogen (Machel,

488

2001; Lyons and Gill, 2010). This anaerobic pathway of organic degradation likely

489

dominated anoxic Ediacaran deep-water columns and marine sediments (Canfield et al., 2007;

490

Lyons and Gill, 2010; Fike et al., 2015). During this process, dissolved inorganic carbon with

491

the depleted 13C/12C was likely released into the water, creating the shallow-to-deep water

492

δ13Ccarb depth-gradient, while the product of hydrogen sulfide reacting with iron oxides or

493

other iron phase yielded sedimentary pyrites with depleted 34S/32S patterns relative to their

494

precursor sulfate. Along with low oxygen and sulfate concentrations, the accumulation of the

495

deep-water organic matter would foster the marine redox heterogeneity of water columns,

496

leading to δ34Spyr depth-gradient in the Ediacaran ocean. The depth-gradients in the δ13Ccarb

497

and δ34Spyr data reflect the redox heterogeneity of the Ediacaran ocean water. Additional data

498

may help to test this hypothesis by tracing the spatial δ13Ccarb and δ34Spyr pattern from the

499

Yangtze Block to other Ediacaran depositional basins.

500 501

5.6. Redox architecture in the Ediacaran ocean and its effect on spatial distribution of

502

the Lantian Biota

503

Evidence from the Lower Yangtze Block of South China robustly supports the

504

interpretation that the early-to-middle Ediacaran ocean (> 550 Ma) was characterized by

505

widespread deep water anoxia (Canfield et al., 2008). Most black shale samples from the

506

lower slope Lantian outcrop section have high FeHR/FeT ratios ( > 0.38) (Shen et al., 2008),

507

and positive δ15N and δ34Spyr values are observed in the Lantian drill-core (Wang et al., 2017), 23

508

both supporting reconstructions of anoxic conditions in the deep water environments of the

509

Ediacaran ocean.

510

The spatial gradient of δ34Spyr values across the shelf-to-basin transect suggests that the

511

ocean chemocline may have dropped, at least episodically, to the intermediate depth. The

512

high frequency of low δ34Spyr values observed in the moderately deep marine facies may

513

indicate that the early-to-middle Ediacaran chemocline, at least the sulfate reduction zone,

514

extended to intermediate depths (Fig. 9). Sustaining the large sulfur isotopic fractionation

515

effects recorded near the oxic-anoxic interface requires sulfate diffusion from shallow water

516

and an ample supply of organic material (Fig.8). Oceanic sulfate reduction zonation in the

517

intermediate water column might have guaranteed adequate bacterial sulfate reduction, for

518

high burial rate of sedimentary pyrite with low δ34Spyr values at the intermediate water-depth

519

environment (Figs. 6, 9b). On the contrary, the shallower or deeper depth of the sulfate

520

reduction zonation mismatched the relatively higher δ34Spyr values in both shallow and deep

521

water facies, as well as the lower pyrite content in shallow water facies. Indeed, the δ34Spyr

522

values in this study corroborate this interpretation, indicating the chemocline was located a

523

moderate depth in the ocean. Additionally, pervasive positive δ15N values in the early

524

Ediacaran slope sections (both in upper and lower slopes), which are interpreted as the result

525

of partial denitrification and anammox, also support this interpretation (Wang et al., 2018). To

526

maintain the persistent release of the isotopically light N2 in slope facies, a relatively

527

moderately deep chemocline was needed (Ader et al., 2016; Wang et al., 2018). This

528

moderately deep chemocline is also consistent with the enrichment of Mo, V, and U only

529

observed in relatively deep-water facies of the early Ediacaran ocean (Sahoo et al., 2012). 24

530

Admittedly, the intermediate chemocline is unstable, which finally goes a step deeper,

531

removes redox crisis, and facilitates the occurrence of early benthos due to the increase of the

532

atmospheric oxygen at the end of Ediacaran Period (Canfield et al., 2007; Chen et al., 2015;

533

Cheng et al., 2017).

534

An ocean chemocline may help to account for the distributions of the Lantian Biota in

535

black shales of slope facies. The Lantian Biota includes an array of macroscopic and

536

morphologically complex life forms (about 600 million years old) preserved in situ as

537

carbonaceous compressions or pyritized fossils in the slope Lantian and Shiyu sections of the

538

Lower Yangtze Block (Yuan et al., 1999, 2011; Wang W. et al., 2014). So far, no Lantian

539

fossils has been discovered from other areas. The absence of fossils in other sections may

540

indicate that their environments were unfavorable for benthic lifeforms and/or their

541

preservation. On one hand, the anoxic deep-water likely hampered the survival of the Lantian

542

Biota in the deep-water environments. On the other hand, a shelf margin environment with a

543

flourishing benthic community may have been unfavorable for the preservation of the

544

organisms, as soft tissue preservation generally benefits from proximity to anoxic conditions

545

(Muscente et al., 2017). In this context, slope facies were likely most favorable for the

546

survival and preservation of the Lantian Biota. Suboxic slope water near the anoxic and oxic

547

interface may have favored periodic flourishment of the macroscopic Lantian biota in oxic

548

conditions and their subsequent preservation after death during times and in areas of anoxic

549

conditions. Indeed, detailed analysis of the Lantian Formation black shales provides evidence

550

that varying redox conditions in the water column may have allowed for their preservation

551

(Guan et al., 2014). 25

552 553 554

6. Conclusions High-resolution carbon and sulfur isotopic profiles of the Ediacaran sequences deposited

555

on the shallow water platform, slope, and basinal facies of the Lower Yangtze Block shed

556

light on the spatial patterns of redox conditions in the Ediacaran ocean. The results provide

557

support for a stratified early-to-middle Ediacaran ocean model.

558

Negative δ13Ccarb shift with a nadir value of –17.7‰ occurs in the middle part of the

559

Doushantuo Formation. The δ13Ccarb shift could be equivalent to the previously reported EN3/

560

Shuram excursion. Despite mostly negative values, δ13Ccarb exhibits a shallow-to-deep ocean

561

gradient (~8‰) in the Lower Yangtze Block. The δ13Ccarb values in shallow water-facies are

562

significantly higher than those in deeper water facies.

563

The δ34Spyr profiles of the Lower Yangtze Block show broad similarities. While δ34Spyr

564

values are mostly positive with a maximum of 37.6‰ across sections, a remarkable negative

565

shift with a magnitude of > 40‰ appears in the lower part of the Doushantuo Formation (Fig.

566

5). This negative shift, which immediately followed the end of Marinoan glaciation might

567

have significance for stratigraphic correlation. Results based on spatial analysis of δ34Spyr

568

values demonstrate a depth-gradient between the shallow and deep-water facies in the Lower

569

Yangtze Block. Statistically higher δ34Spyr values occur in the shallow and deep water

570

sections, while lower δ34Spyr values occur in the moderately deep water section, probably

571

indicating that the oceanic chemocline at least reached to the mid-depth water column in the

572

early-middle Ediacaran ocean.

573

Spatial variation of both δ13Ccarb and δ34Spyr values could result from heterogenous redox 26

574

conditions of the Ediacaran ocean, which may have placed constraints on the spatial

575

distribution of the exceptionally-preserved early Ediacaran Lantian Biota.

576 577 578

Acknowledgments This study was supported by grants from the National Key Research and Development

579

Program of China (2017YFC0603101), National Natural Science Foundation of China

580

(U1562104), the Chinese Academy of Sciences (XDB18000000, XDB26000000), the

581

National Science and Technology Major Project (2017ZX05036-001-004), the Shale Gas

582

Resource Evaluation Method and Exploration Technology Research (2016ZX05034). This

583

paper is a contribution to the special issue “26 Vs1: South China Coevolution”. We thank

584

Xianguo Lang for helpful discussions; Li Peng and Jing Liu for technical assistance in the

585

laboratory; Thomas Algeo and Chao Li for handling this manuscript; Chengsheng Jin and the

586

other two anonymous reviewers for the constructive comments.

587 588

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Table 1. Facies and locations of the studied sections in the Lower Yangtze Block.

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Figure 1. Paleogeographic maps and lithostratigraphic columns of the studied sections. a,

882

Generalized paleogeographic map of the Yangtze Block during the Ediacaran Period,

883

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884

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LGW: Leigongwu Formation; DY: Dengying Formation; PYC: Piyuancun Formation.

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Figure 2. Field photographs showing lithothologies at Wujialing (a~c), Meishuxia (d~f),

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Wangfu (g~i), and Zhushuwu (j~l) sections. a, thin-bedded dolostone from the lower part of

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Formation, Wujialing section; d, contact between the Leigongwu Formation tillite and cap

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dolostone from the basal Doushantuo Formation, Meishuxia section; e, sandstone from the

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Formation, Wangfu section; j, black shale from the lower part of the Lantian Formation,

903

Zhushuwu section; k, thin-bedded dolostone from the middle part of the Lantian Formation,

904

Zhushuwu section; l, thin-bedded chert from the Piyuancun Formation, Zhushuwu section.

905

Hammers in (b), (d), (e), (h), and (i) are about 28 cm long; hammer heads in (g) and (k) are 41

906

about 18 cm long; hat in (f) is about 40 cm; thin-bedded cherts between two dashed red lines

907

is about 120 m in thickness.

908 909

Figure 3. Petrographic photomicrographs. a, Plane-polarized light photomicrograph of

910

dolostone, 12ZY-74, Doushantuo Formation, Wujialing. b, Plane-polarized light

911

photomicrograph of limestone, 12ZY-63, Doushantuo Formation, Wujialing. c, Reflected

912

light photomicrograph of black shales, 12ZSW-13, Lantian Formation, Zhushuwu. d,

913

Cross-polarized light of dolostone, 12ZSW-36, Lantian Formation, Zhushuwu. e,

914

Plane-polarized light photomicrograph of limestone dolostone beds, WY1507-50,

915

Doushantuo Formation, Meishuxia. f, Cross-polarized light photomicrograph of dolostone,

916

WY1507-53, Doushantuo Formation, Meishuxia. g, Cross-polarized light photomicrograph of

917

dolomitic limestone, CA1507-119, Lantian Formation, Wangfu. h, Cross -polarized light

918

photomicrograph of limestone, CA1507-116, Lantian Formation, Wangfu. i, Cross-polarized

919

light photomicrograph of calcareous dolostone, CA1507-48, Lantian Formation, Wangfu.

920

Thin sections were stained with Alizarin red S, the section name and stratigraphic height of

921

each sample are marked in the photomicrographs.

922 923

Figure 4. Spatial variation in δ13Ccarb of the Doushantuo/Lantian formations. Dashed lines

924

represent respective top and bottom boundaries of the Doushantuo/Lantian formations.

925

Yellow shade bar highlights the proposed negative δ13Ccarb excursion EN3 that is regionally

926

consistent and is widely regarded as equivalent to the Shuram excursion. EN3 is absent in the

927

Wujialing section. Data sources: the Jiulongwan section (Jiang et al., 2007), the Jiangshan 42

928

section (Peng et al., 2006), the Meihuishan section (Ling et al., 2007), the Jiujiawu section

929

(Liu et al., 1992), Lantian drill-core (Wang et al., 2017), and Meishuxia, Wangfu, Zhushuwu,

930

and Wujialing sections (this study). EN, Ediacaran δ13Ccarb negative excursion; EP, Ediacaran

931

δ13Ccarb positive excursion.

932 933

Figure 5. Integrated chemostratigraphic profiles of the studied sections. Green arrows mark

934

the increasing trend of δ34Spyr values occurring in the cap dolostone. Red arrow marks a

935

decreasing trend of δ34Spyr values following the cap dolostone.

936 937

Figure 6. Spatial comparison of carbon and sulfur isotopes, and pyrite and TOC contents of

938

the studied sections from different depositional facies. (a) The δ13Ccarb variation from shallow

939

to deep water facies; values in the shallow water facies are higher than those in the deep

940

water facies. The average δ13Ccarb value in the shallow water facies was derived from four

941

shelf marginal sections, including the Wujialing section (this study), the Jiujiawu section (Liu

942

et al., 1992), the Jiangshan section (Peng et al., 2006), and the Meihuishan section (Ling et al.,

943

2007). (b) The δ34Spyr variation from shallow to deep water facies; the lowest value occurs in

944

the moderately deep water facies (with the exception of cap carbonate and stratigraphic

945

intervals with negative δ34Spyr values, highlighted by green and red arrows in Fig. 4). (c) The

946

pyrite (FeS2) content variation. The highest content is found in deep water facies. (d) The

947

δ34SCAS variation. (e) The TOC contents variation. The highest TOC value is found in deep

948

water facies. (f) The δ13Corg variation.

949 43

950

Figure 7. Geochemical cross-plots. (a–d) for the Meishuxia section; (e–h) for the Wangfu

951

section; (i–l) for the Wujialing section; (m–p) for the Zhushuwu section.

952 953

Figure 8. Carbon and sulfur cycling model in the Ediacaran ocean (modified from Bottrell

954

and Newton, 2006). The inputs of oceanic carbon include the carbon from weathering of

955

carbonate and organic rich rocks, volcanic and metamorphic degassing, and the oxidation of

956

excess DOC reservoir. The oxidant used for DOC oxidation is mainly oceanic sulfate derived

957

from the terrigenous weathering of pyrite and sulfate.

958 959

Figure 9. Biogeochemical model showing the carbon and sulfur cycle and ocean redox

960

structure during the early-to-middle Ediacaran Period in South China. (a) Cap dolostone stage.

961

(b) Early-to-middle Ediacaran stage.

44

Table 1. Facies and locations of the studied sections in the Lower Yangtze Block. Section

Province

Latitude

Longitude

Facies

Wujialing

Western Zhejiang

28.848217°N

118.652500°E

Shelf margin

30.103800°N

119.644667°E

Upper slope

29.952233°N

118.897783°E

Lower slope

Province Meishuxia

Western Zhejiang Province

Wangfu

Western Zhejiang Province

Lantian drill-core

Southern Anhui Province

29.951567°N

118.037900°E

Lower slope

Zhushuwu

Southern Anhui Province

29.991567°N

118.817100°E

Deep basin

1. We report carbon and sulfur isotopes of four new sections in South China. 2. A profound negative δ13Ccarb excursion is found in the Lower Yangtze Block. 3. The δ13Ccarb and δ34Spyr profiles show distinct spatial and temporal patterns.

Conflict of interest We declare that we have no financial and personal relationships with other people or o rganizations that can inappropriately influence our work.