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Earth and Planetary Science Letters 317-318 (2012) 96–110

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Earth and Planetary Science Letters

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The origin of decoupled carbonate and organic carbon isotope signatures in the early Cambrian (ca. 542–520 Ma) Yangtze platform

Ganqing Jiang a,⁎, Xinqiang Wang b, Xiaoying Shi b, Shuhai Xiao c, Shihong Zhang b, Jin Dong d a Department of Geoscience, University of Nevada, Las Vegas, NV 89154-4010, USA b State Key Laboratory of Biogeology and Environmental Geology, China University of Geosciences, Beijing, 100083, China c Department of Geosciences, Virginia Polytechnic Institute and State University, Blacksburg, VA 24061, USA d Institute of Geology, Chinese Academy of Geological Sciences, Beijing 100037, China article info abstract

Article history: The early Cambrian (ca. 542–520 Ma) strata in South China record two prominent negative carbonate carbon iso- Received 14 April 2011 13 tope (δ Ccarb)excursionsofearlyNemakit–Daldynian (N–D) and early Tommotian ages. Across each of these Received in revised form 12 September 2011 13 13 excursions, carbonate and organic carbon isotopes (δ Ccarb and δ Corg) are strongly decoupled. Regional corre- Accepted 18 November 2011 13 lation across a shelf-to-basin transect shows lateral heterogeneity of δ Corg during the early-middle N–Dbut Available online xxxx 13 more homogenized δ Corg values across the basin during the late N–D and Tommotian. The temporal and lateral δ13 δ13 δ13 – δ13 Editor: P. DeMenocal variations in Corg suggest that decoupled Ccarb and Corg across the N D Ccarb excursion were possibly caused by diagenetic alteration of organic matter and/or ampliﬁcation of detrital organic carbon isotope signa- δ13 δ13 – Keywords: ture in low-TOC carbonates. In contrast, decoupled Ccarb and Corg of the upper N D and Tommotian were Ediacaran likely resulted from chemoautotrophic–methanotrophic biomass contribution to TOC in organic-rich black 13 13 early Cambrian shale and carbonates. The decoupled δ Ccarb–δ Corg pattern from the lower N–D strata (ca. 542 Ma) shows organic carbon isotopes striking similarities with those from the basal (ca. 635 Ma) and upper (ca. 551 Ma) Doushantuo Formation. In ocean anoxia 13 13 all three cases, decoupled δ Ccarb–δ Corg are seen in organic-poor carbonates (TOC≤ 0.1‰)andcoupled Yangtze platform 13 13 13 δ Ccarb–δ Corg occur in organic-rich black shale and carbonates at the end of the negative δ Ccarb excur- South China 13 13 sion. These similarities suggest that the shift from decoupled to coupled δ Ccarb–δ Corg has no causal link with the terminal oxidation of a large oceanic DOC reservoir. Given the pervasive anoxia/euxinia in Ediacaran– early Cambrian oceans, local DOC-rich environments may have been common, but a large oceanic DOC reservoir capable of buffering the δ13C of marine organic matter requires independent evidence. © 2011 Elsevier B.V. All rights reserved.

1. Introduction for millions of years from the onset of the Cryogenian glaciations (ca. 720 Ma; Swanson-Hysell et al., 2010) to the latest Ediacaran Period 13 Decoupled carbonate and organic carbon isotopes (δ Ccarb and (b560 Ma; Fike et al., 2006; McFadden et al., 2008; Rothman et 13 δ Corg) from some late Neoproterozoic successions have been inter- al., 2003). These uncertainties question the casual link between 13 13 preted as evidence for the existence of a large oceanic dissolved organic δ Ccarb–δ Corg decoupling and an inferred large oceanic DOC reser- 13 13 carbon (DOC) reservoir capable of buffering carbon isotopes of organic voir, and require tests on the validity of taking decoupled δ Ccarb–δ Corg matter (Fike et al., 2006; McFadden et al., 2008; Rothman et al., 2003; from single sections as globally consistent isotope signature. Swanson-Hysell et al., 2010). However, existing isotope record from Similar to the late Neoproterozoic successions, the early Cambrian 13 roughly time-equivalent Ediacaran sections in Oman (Fike et al., (ca. 540–520 Ma) strata record large δ Ccarb excursions (e.g., Brasier 2006), South China (McFadden et al., 2008), northwest Canada et al., 1990, 1994, 1996; Ishikawa et al., 2008; Kouchinsky et al., 2007; 13 (Kaufman et al., 1997), and Australia (Calver, 2000)showlargevaria- Maloof et al., 2005, 2010a, 2010b), but existing δ Corg data show a 13 13 tions in absolute δ Corg values from −23‰ to −37‰, which is inconsis- rather complex picture. Earlier studies reported a negative δ Corg ex- 13 13 tent with an isotopically homogenized ocean. Coupled δ Ccarb–δ Corg cursion at the Precambrian–Cambrian transition (Banerjee et al., data from the entire Cryogenian–Ediacaran interval in northwest Canada 1997; Kimura and Watanabe, 2001; Knoll et al., 1995; Shen and 13 13 (Kaufman et al., 1997) and at the Cryogenian–Ediacaran transition in Schidlowski, 2000) and a weak co-varying δ Ccarb–δ Corg trend of South China (Jiang et al., 2010) also contradict with the isotopic buffering Tommotian age (Goldberg et al., 2007; Guo et al., 2007a). Most of from a long-lasting DOC reservoir that has been inferred to have lasted these data, however, were obtained from organic-rich shales with 13 13 only a few paired δ Ccarb–δ Corg analyses. More recent studies of the carbonate-rich succession in Morocco showed that decoupled δ13 δ13 ⁎ Corresponding author. Tel.: +1 702 895 2708; fax: +1 702 895 4064. Ccarb and Corg are common in early Cambrian successions 13 E-mail address: [email protected] (G. Jiang). (Maloof et al., 2010b), but detailed δ Corg proﬁles across the

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13 individual δ Ccarb excursions and their spatial variability are not the base of the Yanjiahe Formation. The presence of small shelly fossils yet available. To better understand the temporal and spatial vari- Anabarites trisulcantus and Protohertzina unguliformis of Assemblage 13 13 ability of δ Ccarb–δ Corg patterns, in this study we present paired Zone 1 (SS1; Qian et al., 2001; Steiner et al., 2007) in the lower Yanjiahe 13 13 δ Ccarb–δ Corg data of the early Cambrian (ca. 542–520 Ma) strata FormationsuggestsanageofNemakit–Daldynian or lower-middle from the Yangtze platform in South China and compare the early Cam- Meishucunian. The uppermost Yanjiahe Formation and lower Shuijingtuo 13 13 brian δ Ccarb–δ Corg patterns with those from the Ediacaran strata in Formation contain typical small shelly fossil Aldanella yanjiahensis of As- the same region. semblage Zone 3 (SS3; Steiner et al., 2007), which is of Tommotian or upper Meishucunian age. The first appearance of trilobites in the lower 2. Paleogeography and stratigraphy to middle Shuijingtuo Formation markedthebaseoftheAtdabanianor Qiongzhusian (Qian et al., 2001; Steiner et al., 2007; Zhu et al., 2003). 2.1. Paleogeography Other microfossils including organic-walled microfossils, macroscopic carbonaceous compressions, and sponge spicules (e.g., Guo et al., 2008, The Ediacaran–Cambrian Yangtze platform developed over a late 2010; Zhao et al., 1988) have also been reported, but they have limited Neoproterozoic rifted continental margin that is inferred to have initiat- biostratigraphic significance. ed at the southeastern side of the Yangtze block (Fig. 1A) at ca. 830–820 Ma (Wang and Li, 2003; Zhao et al., 2011). A passive continen- 2.3. Regional stratigraphic correlation tal margin setting has been inferred for the Ediacaran–Cambrian strata on the basis of stratigraphic pattern, thickness, and the lack of signifi- Regional correlation of the early Cambrian (ca. 542–520 Ma) strata cant igneous activity (Jiang et al., 2003, 2011; Wang and Li, 2003). in South China is mainly based on small shelly fossils and distinctive During the latest Ediacaran and earliest Cambrian (lower-middle marker beds including phosphorite-rich and highly metalliferous Meishucunian or Nemakit–Daldynian), the Yangtze platform was black shales (Fig. 1C). Small Shelly Fossil Assemblage Zone 1 (SS1) characterized by predominately carbonate deposits in the northwest represented by A. trisulcantus and P. unguliformis and Assemblage and mudstone/shale deposits in the southeast (Fig. 1B; Steiner et al., Zone 3 (SS3) represented by Watsonella crosbyi are widespread in 2007; Wang, 1985; Zhu et al., 2007). A ‘transitional zone’ character- shelffaciesoftheYangtzeplatform(Steiner et al., 2007; Zhu et al., ized by interbedded shale and carbonate deposits is located in north- 2003). Traditionally their appearance was taken as the base of the N– eastern Guizhou, northern Hunan and Hubei provinces (Fig. 1B), D and Tommotian stages, respectively, but integrated biostratigraphy representing outer shelf environments of the Yangtze platform. At and carbon isotope chemostratigraphy suggested that the base of the the end of the middle Meishucunian (after small shelly fossil assem- N–D stage (i.e., the Precambrian–Cambrian boundary) would be lower blage 3), the shelf regions of the Yangtze platform was subaerially ex- than SS1, most likely coincident with the regional stratigraphic discon- posed (Steiner et al., 2007; Zhu et al., 2003), resulting in partial tinuity at the top of the Dengying Formation (Fig. 1C; Brasier et al., erosion of the earliest Cambrian strata, particularly in the inner 1990; Goldberg et al., 2007; Ishikawa et al., 2008; Li et al., 2009; Shen shelf regions west of the Yangtze Gorges area (Fig. 1C). Subsequent and Schidlowski, 2000; Steiner et al., 2007; Zhou et al., 1997; Zhu et transgression during the upper Meishucunian (Tommotian) led to al., 2003). Small Shelly Fossil Assemblage Zone 2 (SS2) represented by deposition of deep-water black shales across the entire Yangtze plat- Paragloborilus subglobosus–Purella squamulosa is more restricted to the form (Fig. 1C). In comparison with the other early Cambrian succes- proximal regions of the Yangtze platform (e.g., eastern Yunnan and sions such as those in Morocco (Maloof et al., 2005, 2010a, 2010b) Sichuan; Sections 1 and 2 in Fig. 1C) and thus its biostratigraphic signif- and Siberia (e.g., Kouchinsky et al., 2007), the early Cambrian (ca. 542 icance is limited (Steiner et al., 2007). In the slope and basin facies, the to 520 Ma) succession in the Yangtze platform are highly condensed Precambrian–Cambrian boundary is traditionally defined by the first and possibly less complete (Fig. 1C). occurrence of sponge spicules, aided by the presence of phosphorite layers/nodules that are close to the Precambrian–Cambrian boundary. 2.2. Early Cambrian stratigraphy in the Yangtze Gorges area The highly metalliferous black shales with high Ni–Mo concentrations from the Niutitang and Xiaoyanxi formations (Fig. 1C; Lehmann et al., In the vicinity of the Yangtze Gorges area (Fig. 2A), the late Ediacaran 2007; Wille et al., 2008; Yang et al., 2004) was once considered as Dengying Formation is overlain by the Yanjiahe Formation (Fig. 2B). The close to the Precambrian–Cambrian boundary, but recent radiometric Yanjiahe Formation is composed of sandy/cherty dolostones in the ages of 539.4±2.9 Ma (Compston et al., 2008), 532.3±0.7 Ma (Jiang lower part and interbedded limestone and black shale in the upper et al., 2009)and536.3±5.5Ma(Chen et al., 2009)fromashbeds part. Thin phosphorite layers (mostly chertified) are present in the below the highly metalliferous black shales indicate that the metallifer- lower part and at the top of this unit. The Yanjiahe Formation is overlain ous beds are significantly younger than the Precambrian–Cambrian by the black shales of the Shuijingtuo Formation, the lower part of boundary, likely having an early Tommotian age (Fig. 1C). which contains large carbonate concretions (with diameters up to 30 cm) and thin (b10 cm thick) carbonate layers. The contact between 3. Isotope analysis the Dengying and Yanjiahe Formations marks a regional discontinuity, which in places truncates the uppermost Dengying Formation Three well-exposed outcrop sections near Yichang and Changyang (Ishikawa et al., 2008; Zhu et al., 2003). The Yanjiahe–Shuijingtuo (Fig. 2A) were sampled at an average spacing of 10–20 cm. Collected boundary has also been considered as a stratigraphic discontinuity samples were cut into two or more chips so that visual and petro- that in places truncates a portion or all of the Yanjiahe Formation and graphic observations can match the drilling spots. Sample powders its time-equivalent units, particularly in the western part of the Yangtze (20–50 mg) were drilled from relatively fine-grained dark laminae platform (e.g., loc. 1, 2, and 3 in Fig. 1BandC;Zhu et al., 2003). However, in cleaned slabs and were then split for carbonate and organic isotope in the measured sections of this study, physical evidence for strati- analyses. All isotope values are reported as δ values with reference to graphic truncation at this level is insignificant. the Vienna Pee Dee Belemnite standard (VPDB). Sample preparation The Ediacaran–Cambrian boundary in this region is traditionally de- and analyses were performed in the Las Vegas Isotope Science fined by the first appearance of small shelly fossils (Chen, 1984; Qian, (LVIS) Laboratory at the University of Nevada Las Vegas. 13 18 1999)andMicrhystridium-like acritarchs (Ding et al., 1992; Dong et al., For δ Ccarb and δ Oanalyses,about50–200 μgofcarbonatepowders 2009; Yin, 1997), which is located 5–15 m above the Dengying–Yanjiahe were reacted with orthophosphoric acid for 10 minutes at 70 °C, using a boundary (Fig. 2B). However, considering the lithologically dependent Kiel IV carbonate device connected to a Finnigan Delta V Plus mass spec- fossil preservation, the Ediacaran–Cambrian boundary may be down to trometer via dual-inlet. Analytical reproducibility was better than 0.1‰ Author's personal copy

98 G. Jiang et al. / Earth and Planetary Science Letters 317-318 (2012) 96–110

Fig. 1. (A) Tectonic outline of the Yangtze Block. (B) Paleoenvironmental reconstruction of the Yangtze platform during the early Cambrian Nemakit–Daldynian (lower-middle Meishucunian) stage. Modiﬁed from Wang (1985) and Steiner et al. (2007). Locations 1–7 mark the areas where reference sections are used for regional correlation. (C) Strati- graphic correlation of the early Cambrian (ca. 542–520 Ma) strata, with biostratigraphic and U–Pb age data marked in sections. Data source: 1—Meishucun (Compston et al., 2008; but see Jenkins et al., 2002; Zhu et al., 2009, for slightly different ages from the same section); 2—Shatan (Goldberg et al., 2007; Guo et al., 2007a); 3—Songlin (Jiang et al., 2009); 4—Yangtze Gorges (this study); 5—Ganziping (Chen et al., 2009); 7—Yuanling (Guo et al., 2007a). Small shelly fossil assemblage zones are adopted from Steiner et al. (2007) and Micrhystridium-like acritarchs are from Dong et al. (2009). Molybdenum isotopes from black shales of high Ni–Mo contents in Songlin (loc. 3), Ganziping (loc. 5), and Yuanling (loc. 7) were reported as evidence for pervasive euxinia in early Cambrian ocean (Lehmann et al., 2007; Wille et al., 2008).

13 18 for both δ Ccarb and δ O values, as monitored by NBS-19 and an internal for 4 h. This wash and dry process was repeated up to 3 times until pH standard. tests gave a near neutral value (≥6.0). Samples were then dried and For organic carbon isotope analysis, aliquots of powdered samples be- wrappedintincapsulesandstoredinanovenat105°Cpriortoanalysis. tween 5 and 30 mg (depending on TOC content) were decarbonated in Organic carbon isotopes were analyzed using an elemental analyzer (EA) silver capsules overnight through acid fumigation (Harris et al., 2001) coupled with a conﬂow interface that automatically transfers carbon di- with concentrated HCL (12 M). After drying, one or more drops of 1 M oxide gas into a Finnigan Delta Plus mass spectrometer. USGS-24 (graph- HCLwereaddedtothecapsulestoensurecompletecarbonateremoval. ite), IAEA-600 (caffeine), and acetanilide (Costech Analytical The silver capsules with carbonate-free residue were then neutralized Technologies) standards were used to monitor the external and internal 13 using a stepwise washing process, by which each capsule was immersed uncertainty that was better than 0.2‰ for δ Corg and 0.1% for TOC. To in DI water (using 20 mL weighing dishes) for 6 h and then dried at 70 °C monitor the reproducibility of organic-rich and organic-poor samples, Author's personal copy

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Fig. 2. (A) Simpliﬁed geological map of the Yangtze Gorges area in Hubei Province, South China, showing location of measured sections. 1—Jiuqunao; 2—Jijiapo; 3—Hezi'ao. (B) Gen- eralized stratigraphic column of the early Cambrian (ca. 542–520 Ma) succession in the Yangtze Gorges area. Biostratigraphic data are summarized from Chen (1984), Ding et al. (1992), Dong et al. (2009), Qian (1999), Steiner et al. (2007), Yin (1997), and Zhu et al. (2003).

duplicates were analyzed for every sixth sample and their reproducibility in the lower part to −8±2‰ in the upper part of the section, but they 13 13 in δ Corg and TOC was better than 0.5‰ and 0.2%, respectively. do not show covariation with δ Ccarb (Fig. 3). Organic carbon isotope 13 13 (δ Corg) values are decoupled from δ Ccarb values. At the base of the 13 4. Results Yanjiahe Formation, a 2‰ shift in δ Corg from −26‰ to −28‰ occurs 13 at the beginning of the negative δ Ccarb excursion (CN1), but for the ma- 13 4.1. Jiuqunao section jority of CN1, δ Corg values remain at −27‰ to −28‰. A large negative 13 shift in δ Corg from −27‰ to −35‰ occurs at the end of CN1, coinci- In the Jiuqunao section (loc. 1 in Fig. 2A), a total of 109 samples dent with the lithological change from cherty dolostone to fossiliferous, δ13 were analyzed for the 62-m-thick interval that covers stratigraphic phosphorous-cherty shale/limestone. Very stable Corg values around units from the uppermost Dengying Formation to the lowermost −33‰ span the rest of the section (CP1 and CN2) and they do not 13 show a covariation with δ13C or TOC. The carbonate and organic car- Shuijingtuo Formation (Fig. 3; Table A1). A negative δ Ccarb excursion carb Δδ13 Δδ13 δ13 −δ13 (CN1) with minimum values down to −4‰ covers the lowermost 12 m bon isotope difference ( C; C= Ccarb Corg)changesfrom 13 28‰ to 24‰ before the top of CN1, to 33±2‰ across CP1, and to ~28‰ of the Yanjiahe Formation. The negative to positive shift in δ Ccarb occurs right after the phosphorous-cherty shale/limestone that contains at the basal Shuijingtuo Formation (CN2). A clear covariation between δ13 Δδ13 abundant Micrhystridium-like acritarchs such as Heliosphaeridium Ccarb and Cisobservedthroughoutthesection(Fig. 3). ampliatum (cf. Dong et al., 2009; Yao et al., 2005) and small shelly fossil Anabarites trisulcatus (cf. Steiner et al., 2007). The rest of the Yanjiahe 4.2. Jijiapo section 13 Formation (CP1; Fig. 3)hasδ Ccarb values mostly within the range of 0‰ to +2‰, with a few down to −2‰. The second negative shift in Samples (n=131) from the 88-m-thick Jijiapo section (loc. 2 in 13 δ Ccarb (CN2) occurs at the Yanjiahe–Shuijingtuo transition, with min- Fig. 2A) covers the uppermost Dengying Formation, the entire Yanjiahe imum values down to ≤−4‰ in limestone layers and concretions of the Formation and the lower Shuijingtuo Formation (Fig. 4; Table A1). Two 13 basal Shuijingtuo Formation. Oxygen isotope values change from −6‰ negative δ Ccarb excursions (CN1 and CN2) are present at the lower Author's personal copy

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13 13 18 13 13 13 Fig. 3. Paired carbonate and organic δ C and cross plots of δ Ccarb–δ O, δ Ccarb–δ Corg, TOC–δ Corg of the early Cambrian strata from the Jiuqunao section (loc. 1 in Fig. 2A). Abundant Micrhystridium-like acritarchs such as Heliosphaeridium ampliatum and small shelly fossil Anabarites trisulcatus are found from 19.9 m to 20.5 m in the section, coincident with the organic carbon isotope minimum. See Fig. 2B for explanation of symbols.

Yanjiahe Formation and at the Yanjiahe–Shuijingtuo transition, respec- 4.3. Hezi'ao section tively, consistent with those documented from a drilled core adjacent to 13 the section (Ishikawa et al., 2008). However, strongly negative δ Ccarb Carbon isotope data (n=133) from the Hezi'ao section in Changyang 13 values down to −9‰ reported in Ishikawa et al. (2008) from a 10- (loc. 3 in Fig. 2A) show two negative δ Ccarb excursions at the lower cm-thick interval below the Yanjiahe–Shuijingtuo boundary were not Yanjiahe and Shuijingtuo formations, respectively (Fig. 5;TableA1). 13 13 18 observed in this study. Between the two negative δ Ccarb excursions The δ Ccarb and δ O values and trends are similar to those of the Jiuqu- 13 13 are positive δ Ccarb values mostly from +2‰ to +4‰ (CP1). Oxygen nao and Jijiapo sections (Figs. 3 and 4). Again, δ Corg is decoupled from 13 isotope values are mostly around −6‰ in the lower part of the section δ Ccarb, with a minor negative shift at the Dengying–Yanjiahe boundary, (EP and CN1) and are around −8‰ in the upper part of the section. The a major negative shift down to −36‰ associated with the fossiliferous, −2‰ shift in δ18O coincides with the lithological change from dolos- phosphorous-cherty shale/limestone that contains small shelly fossil A. 13 18 13 tone to limestone. No δ C–δ O covariance is observed for the entire trisulcatus,verystableδ Corg values around −33‰ in the upper Yan- 13 dataset and across each individual δ Ccarb excursions. jiahe Formation, and a 2‰ increase from −34‰ to −32‰ in the lower Organic carbon isotopes are quite variable in the lower part of the Shuijingtuo Formation. Similar to the other two sections, the Δδ13C 13 13 section. A negative shift in δ Corg from −27‰ to −31‰ occurs near trend tracks that of δ Ccarb, with lowest values down to 21‰ associated 13 13 the Dengying–Yanjiahe boundary. Up section, δ Corg values increase with the CN1 δ Ccarb minimum, maximum values up to 36‰ across CP1 13 from −31‰ to −27‰ as δ Ccarb shifts toward the minimum of CN1. in the upper Yanjiahe Formation, and moderate values around 30‰ 13 An abrupt negative shift in δ Corg from −27‰ to −36‰ is present at across CN2 in the lower Shuijingtuo Formation. the end of CN1, concomitant with a lithological change from cherty 13 13 dolostone to phosphorous-cherty shale/limestone that contains 4.4. Composite δ Ccarb–δ Corg patterns abundant small shelly fossil A. trisulcatus (Ding et al., 1992; Guo et 13 13 13 al., 2008). For the rest of the Yanjiahe Formation (CP1), δ Corg values The overall δ Ccarb and δ Corg trends are persistent in the three 13 13 remain very stable around −33‰. Across the δ Ccarb excursion at the measured sections, although absolute δ C values slightly differ due 13 Yanjiahe–Shuijingtuo transition (CN2), δ Corg show a minor change to variable sampling resolution among sections. Composite strati- from −33‰ to −34‰, followed by a gradual increase from −34‰ graphic and carbon isotope data from these sections, normalized to to −31‰. Except for the intervals at the base and top of CN1, the the average thickness of the Yanjiahe and Shuijingtuo formations, 13 13 Δδ C trend tracks that of δ Ccarb (Fig. 4). provide the following isotope feature (Figs. 6 and 7). Author's personal copy

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13 13 18 13 13 13 Fig. 4. Paired carbonate and organic δ C and cross plots of δ Ccarb–δ O, δ Ccarb–δ Corg,TOC–δ Corg of the early Cambrian strata from the Jijiapo section (loc. 2 in Fig. 2A). Locations of small shelly fossils are from Chen (1984) and Qian (1999).SeeFig. 2B for explanation of symbols.

13 13 18 (1) Two negative δ Ccarb excursions are present in the lower Yanjiahe dolostone to limestone. There is no δ C–δ O covariance across 13 and lower Shuijingtuo formations, respectively (Fig. 6). The negative each of the negative δ Ccarb excursions and for the entire stratigraph- 13 δ Ccarb excursion from the lower Yanjiahe Formation (CN1) has the ic interval (Fig. 7A). 13 13 nadir slightly below the Small Shelly Fossil Assemblage Zone 1 (SS1) (3) Both CN1 and CN2 show decoupled δ Ccarb and δ Corg, but 13 13 that has an early Nemakit–Daldynian (N–D) age. The negative δ Ccarb ex- their expressions differ (Fig. 6). The δ Corg values across CN1 show cursion from the lower Shuijingtuo Formation (CN2) is associated with a negative shift at the beginning of CN1 (CN1a), an upward increase 13 Small Shelly Fossil Assemblage Zone 3 (SS3) of early Tommotian age. towards the CN1 δ Ccarb minimum (CN1b), and a large negative 13 These δ Ccarb excursions have been documented from other sections in shift from −27‰ to −36‰ followed by a positive shift from −36‰ the Yangtze platform (e.g., Brasier et al., 1990; Goldberg et al., 2007; to −33‰ at the end of CN1 (CN1c). In contrast, no significant 13 Ishikawa et al., 2008; Li et al., 2009; Shen and Schidlowski, 2000; Zhou (b2‰) δ Corg change is observed across CN2. The most significant 13 et al., 1997) and have been correlated with those from global successions change in δ Corg at the end of CN1 (CN1c) coincides with an increase in Morocco (Maloof et al., 2005, 2010a, 2010b)andSiberia(Brasier et al., in TOC and a lithological change from cherty dolostone to 1994; Kaufman et al., 1996; Kouchinsky et al., 2007). The sub-million- phosphorous-cherty shale/limestone (Figs. 6 and 7B). The strati- 13 year-scale δ Ccarb excursions documented from Morocco (Maloof et al., graphic interval between CN1 and CN2 (interval CP1) has very stable 13 13 2005, 2010a, 2010b), however, are not well displayed in the studied sec- δ Corg values around −33‰ and high Δδ C(>32‰). 13 13 tions. The upper N–DintervalbetweenCN1andCN2(CP1inFig. 6)has (4) There is no covariance between δ Ccarb and δ Corg for each of 13 13 δ Ccarb values scattered between 0‰ and 4‰,butitisdifficult to define the δ Ccarb excursions and for the entire stratigraphic interval 13 13 13 13 discrete δ Ccarb excursions. The lack of sub-million-year-scale δ Ccarb ex- (Fig. 7C). The Δδ C profile tracks that of δ Ccarb (Fig. 6) and there 13 13 cursions in the Yangtze platform is likely due to highly condensed depo- is a clear covariance between δ Ccarb and Δδ C for each of the iso- sition that has obscured the record of short-term carbon cycles. tope excursions (Fig. 7E). The data points, however, fall into two dis- (2) Oxygen isotopes from the lower part of the sections (EP and tinct groups with different slopes and y-axis intercepts in the 13 13 CN1) are mostly within the range of −5‰ to −7‰, while for the δ Ccarb–Δδ Cplot(Fig. 7F). Data points from the lower part of the sec- upper part of the sections (CP1 and CN2), δ18O values are 2–3‰ tions (EP+CN1a+CN1b in Fig. 6) have a slope of 0.87 and a y-intercept lower, mostly within the range of −7‰ to −10‰ (Fig. 7A). The 2– of −23.7‰, while those from the upper part (CN1c+CP1+CN2 in 3‰ decrease in δ18O coincides with a lithological change from Fig. 6) show a slope of 0.99 and a y-intercept of −33.3‰ (Fig. 7E). Author's personal copy

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13 13 18 13 13 13 Fig. 5. Paired carbonate and organic δ C and cross plots of δ Ccarb–δ O, δ Ccarb–δ Corg, TOC–δ Corg of the early Cambrian strata from the Hezi'ao section in Changyang (loc. 3 in Fig. 2A). Abundant Micrhystridium-like acritarchs such as Heliosphaeridium ampliatum and small shelly fossil Anabarites trisulcatus are found from 23.5 m to 25.5 m in the section, coincident with the organic carbon isotope minimum. See Fig. 2B for explanation of symbols.

These features are typical of late Neoproterozoic carbonates but they are to be explained by the isotopic buffering of a large and persistent DOC 13 13 different from those of the Cenozoic δ Ccarb–Δδ C plot that has a slope reservoir in the ocean (cf. Jiang et al., 2010), in which temporal changes 13 of 0.2–0.3 and a y-intercept of −6‰ (Rothman et al., 2003). in δ Corg would be much less signiﬁcant or invariant (Fike et al., 2006; McFadden et al., 2008; Rothman et al., 2003; Swanson-Hysell et al., 5. The origin of decoupled carbonate and organic carbon isotopes 2010).

13 13 The δ C patterns from the early Cambrian strata in the Yangtze 5.1. Regional δ Corg variability Gorges area are similar to those of the late Neoproterozoic in that 13 (1) there is no comparable δ Corg excursions associated with the Compilation of existing organic carbon isotope data (Fig. 8) across 13 δ Ccarb excursions (Fig. 6) and (2) there is a strong correlation be- the Ediacaran–Cambrian transition in South China provides insights 13 13 13 tween δ Ccarb and Δδ C(Fig. 7F). These features have been inter- on the regional variability of δ Corg in the Yangtze platform. In the 13 preted as resulting from burial diagenetic alteration of δ Ccarb inner shelf section of eastern Yunnan Province (Section 1; Fig. 8), 13 13 (Derry, 2010a, 2010b) or as evidence for the existence of a large dis- co-varying δ Ccarb and δ Corg were reported from the N–D strata, al- 13 13 solved organic carbon (DOC) reservoir capable of buffering carbon though no paired δ Ccarb–δ Corg data are available (Shen and isotopes of oceanic organic matter (Rothman et al., 2003). The diage- Schidlowski, 2000). Distal inner shelf (intrashelf lagoon) section in netic interpretation emphasizes the δ13C–δ18O covariance through Shatan (Section 2; Fig. 8) has limited N–D strata due to stratigraphic 13 13 negative δ Ccarb excursions and implies that the negative δ Ccarb ex- truncation by unconformities and/or stratigraphic condensation. A few 13 cursions from late Neoproterozoic successions are not necessarily δ Corg values reported from the Kuanchuanpu Formation of N–Dage globally synchronous. The lack of δ13C–δ18O covariance across the are ≤−34‰ (Goldberg et al., 2007; Guo et al., 2007a). It is uncertain if 13 N–D and Tommotian negative δ Ccarb excursions (Fig. 7A) and these values are correlatable with those from the upper N–Dinthe 13 more importantly, the global occurrence of these δ Ccarb excursions Yangtze Gorges area (Fig. 6). Tommotian strata in this section show 13 13 13 in biostratigraphically correlatable units (Ishikawa et al., 2008; co-varying δ Ccarb–δ Corg trends, but Δδ Cvaluesaremostly≥32‰, Kouchinsky et al., 2007; Li et al., 2009; Maloof et al., 2010a, 2010b) implying chemoautotrophic biomass contribution (Goldberg et al., 13 argue against substantial diagenetic reset of δ Ccarb values. Similarly, 2007). From the slope section in Ganziping (Section 5 in Fig. 8; Chen 13 13 13 the large temporal variations in δ Corg across the N–D δ Ccarb excur- et al., 2009), a few δ Corg values reported from the Dengying Formation 13 13 sion, particularly the co-varying positive shift of δ Ccarb and δ Corg are ≥−28‰, similar to those from the Dengying Formation in the Yang- 13 13 at the end of the N–D δ Ccarb excursion (CN1c in Fig. 6), are difﬁcult tze Gorges area. Up to 10‰ shift in δ Corg occurs at the Ediacaran– Author's personal copy

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Fig. 6. Composite stratigraphy and carbon isotope data of the early Cambrian strata from all three sections in the Yangtze Gorges area, normalized to the average thickness of the Yanjiahe and Shujingtuo formations. The black curves fitted to the data represent a five point floating average. See Fig. 2B for explanation of symbols.

13 Cambrian transition, coincident with a lithological change from dolos- in δ Corg across the shelf-to-basin transect (Fig. 8) is inconsistent tone to siliceous shale and chert. The N–D and Tommotian strata in with an early Cambrian ocean that was isotopically homogeneous and 13 this section are highly condensed, but δ Corg values display up to 5‰ buffered by a large DOC reservoir. Instead, it indicates that the 13 13 temporal variations within N–D strata and remain very stable around decoupled δ Ccarb–δ Corg pattern seen in the Yangtze Gorges area −32‰ after the N–D and Tommotian boundary. It is uncertain whether (Figs. 3–6) is a localized phenomenon that may change with paleogeo- 13 the N–D δ Corg shift seen in the Yangtze Gorges areas (CN1c in Figs. 6 graphic settings. In addition, the data also show that during the latest 13 and 8) should be correlated with the base or the top of the Liuchapo For- Ediacaran and earliest Cambrian (basal N–D), δ Corg values from mation in this section. In the basinal sections at Songtao and Yuanling shale- and chert-dominated basinal sections are 8–10‰ lower than (Sections 6 and 7 in Fig. 8; Guo et al., 2007a), the uppermost Ediacaran those from the carbonate-dominated shelf sections, but more homoge- 13 13 and basal Cambrian (N–D) strata have most δ Corg values around neous δ Corg values around −33‰ characterize the lower Tommotian 13 −34‰, but a small (~2‰)negativeδ Corg shift is seen at the upper- strata across the basin. most Liuchapo Formation. Given the existing age constraints and sam- 13 ple resolution, it is difficult to determine if this negative shift in 5.2. Controls on δ Corg 13 13 δ Corg is time equivalent to the entire N–D δ Ccarb excursion (CN1 in Fig. 6) or the uppermost portion of the N–D excursion (CN1c in Fig. 6) As have been reviewed in numerous publications (e.g., Ader et al., in the Yangtze Gorges area. The lower Tommotian strata in the basinal 2009; Des Marais et al., 1992; Goldberg et al., 2007; Hayes et al., 1983, 13 sections have δ Corg values around −32±1‰, which are similar to 1999; Jiang et al., 2010; Kaufman and Knoll, 1995; Maloof et al., those from the outer shelf sections in the Yangtze Gorges area, but 2010b; Young et al., 2008), carbon isotopes of bulk organic matter can they are 1–2‰ higher than those from the inner shelf section (Section be influenced by many factors including carbon isotope fractionation 2; Fig. 8). In the Yuanling section (Section 7; Fig. 8), a 5‰ positive during primary and secondary production, detrital organic carbon 13 δ Corg ‘excursion’ is present in the upper Tommotian–Atdabanian stra- input, DOC contribution, post-depositional alteration through diagene- ta, but its regional correlation and significance need to be clarified in the sis/metamorphism (e.g., methanogenesis, bio- and thermal degrada- future because (1) the available age constraints are insufficient to pre- tion), and hydrocarbon contamination (Fig. 9A). While co-varying 13 13 cisely correlate the hosting stratigraphic unit across the basin, and (2) δ Ccarb and δ Corg imply that organic carbon in marine sediments (sed- 13 no Tommotian δ Corg excursion is seen in the other basinal section imentary rocks) was mainly derived from primary (photosynthetic) pro- (Section 6; Fig. 8). duction without substantial post-depositional alteration, decoupled 13 13 13 In spite of uncertainties on δ Corg excursions in deep-water sections δ Ccarb and δ Corg have a much less definitive meaning. If organic car- that require clarification with additional isotope data of higher strati- bon contribution from one or more of the non-primary sources (second- graphic resolution and better age constraints, the regional variability ary production, detrital organic carbon, DOC, post-depositional Author's personal copy

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13 18 Fig. 7. Cross plots of isotope data from the Yangtze Gorges area. Intervals EP, CN1, CP1, and CN2 match with those in Fig. 6. (A) δ Ccarb versus δ Ocarb. Neither the individual in- 13 18 13 13 tervals nor the entire data show δ C–δ O covariance. (B) Total organic carbon (TOC) versus δ Corg. No clear TOC–δ Corg covariance is observed, except for interval CN1 2 13 13 13 13 13 13 13 (R =0.53). (C) δ Ccarb versus δ Corg. There is no δ Ccarb–δ Corg covariance for the individual intervals and for the entire data. (D) δ Corg versus Δδ (Δδ=δ Ccarb −δ Corg). 2 13 13 13 Except for CN1 (R =0.73), no δ Corg–Δδ covariance is observed for other intervals. (E) Cross plot of δ Ccarb versus Δδ showing clear δ Ccarb–Δδ covariance for each interval. 13 (F) δ Ccarb versus Δδ for intervals EP+CN1a +CN1b and CN1c+CP1 +CN2 (Fig. 6), with a linear ﬁt computed using the reduced major axis method. Author's personal copy

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Fig. 8. Carbonate and organic carbon isotope profiles of the early Cambrian strata from sections across a shelf to basin transect (see Fig. 1A and C for location of sections and age constraints). Data source: 1—Laolin (Shen and Schidlowski, 2000); 2—Shatan (Goldberg et al., 2007; Guo et al., 2007a); 4—Yangtze Gorges area (this study; Fig. 6); 5—Ganziping (Chen et al., 2009); 6—Songtao (Goldberg et al., 2007; Guo et al., 2007a); 7—Yuanling (Goldberg et al., 2007; Guo et al., 2007a). overprints) constitutes a significant portion of TOC in sedimentary rocks, Similarly, methanogenesis releases 13C-depleted methane and results 13 13 13 decoupled δ Ccarb and δ Corg values would be expected (Fig. 9A), at in C-enrichment in residual TOC (Irwin et al., 1977; Summons et al., least locally. 1998). The degree of 13C-enrichment in TOC through methanogenesis Carbon isotope signature of non-primary organic matter is strong- depends on the amount of substrate consumption and can be highly ly dependent on geological settings and plaeoenvironmental condi- variable in different microbial ecosystems (Bradley et al., 2009; tions, and thus there is no definitive δ13C value for each category of Summons et al., 1998). non-primary organic carbon. However, numerous studies on modern It has been frequently reviewed and widely cited that diagenesis and ancient sedimentary organic matter have provided guidance for and thermal maturation do not substantially alter carbon isotope qualitatively evaluating the isotope differences between primary values of organic matter in sedimentary rocks (e.g., Ader et al., and non-primary organic matter. For secondary production, chemo- 2009; Bekker et al., 2008; Chen et al., 2009; Maloof et al., 2010b; autotrophic organisms use recycled carbon during carbon fixation Young et al., 2008). For rocks metamorphically lower than greens- and their biomass is up to 15‰ depleted in 13C relative to primary chist facies, which is most likely the case for the late Neoprotero- photosynthate (e.g., Conway et al., 1994; Hollander and Smith, zoic–early Cambrian strata in South China (Wang et al., 1993, 2001). Methanotrophic bacteria assimilate carbon from 13C-depleted 2008), 13C-enrichment through bio- and thermal degradation would methane and their biomass can be 13C-depleted relative to photosyn- be in the order of 2–3‰ (e.g., Clayton, 1991; Hayes et al., 1983; thetic organic matter by −15‰ to −40‰ (Conway et al., 1994; Tocqué et al., 2005). In addition, it has been argued that thermal deg- Summons et al., 1994, 1998). Statistic review of paired carbonate radation relates to the burial history of sedimentary basins; it may 13 13 and organic carbon isotopes from the geological record suggests change the absolute δ Corg values but not the temporal δ Corg that organic matter with Δδ13C values ≥32‰ must have involved trend (Des Marais et al., 1992). These statements may be true for chemoautotrophic biomass contribution (Hayes et al., 1999). Carbon organic-rich sedimentary rocks, but for organic-poor carbonates, 13 isotope values of detrital organic carbon should reflect those of the δ Corg may be more sensitive to diagenesis and thermal maturation. weathering source rocks and, if not following a long-lasting unconfor- First, assuming that the same amount of 12C-rich carbon were re- mity, may be similar with those of the underlying stratigraphic units. moved during methanogenesis and thermal maturation, the change Dissolved organic carbon (DOC) may escape degradation under anox- in δ13C of residual organic matter in organic-poor carbonates may ic conditions during which small molecules remain as highly aged be significantly larger than that in organic-rich shale and carbonates. and slowly cycling components (e.g., Loh et al., 2004). Because aged Second, bio- and thermal degradation of organic matter may lead to DOC may have experienced multiple recycling, their carbon isotope the amplification of isotope signature of detrital organic carbon. For values may be more negative than those of their particulate and dis- both organic-rich and organic-poor sedimentary rocks, there may be solved organic carbon precursors (Hayes, 2001). Migrating hydrocarbons a certain amount of detrital organic carbon. Because detrital fossil or- have highly variable δ13Cvaluesfrom−21‰ to −33‰ (Whelan and ganic matter is much less reactive compared to fresh photosynthetic Thompson-Rizer, 1993), but they should be localized along lithologically organic carbon, diagenesis (e.g., biodegradation and methanogenesis) or fracturally controlled fluid pathways. Thermal maturation of organic and thermal maturation would preferentially remove the latter. In matter removes 13C-depleted gases and shifts TOC towards higher δ13C organic-rich rocks, removal or altered portion of the primary organic 13 values (e.g., Clayton, 1991; Hayes et al., 1983; Tocqué et al., 2005). carbon may not substantially change the δ Corg values, but for Author's personal copy

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13 13 Fig. 9. (A) Schematic diagram showing the potential controls of organic carbon isotope values in sedimentary rocks. While co-varying δ Ccarb and δ Corg imply that organic carbon 13 13 in sedimentary rocks was mainly derived from primary (photosynthetic) production without substantial post-depositional alteration, decoupled δ Ccarb and δ Corg have a much 13 13 less deﬁnitive meaning. (B) A tentative interpretation for the δ Ccarb–δ Corg patterns observed from the early Cambrian strata in the Yangtze Gorges area.

organic-poor carbonates, organic carbon from primary production also inconsistent with substantial organic carbon contribution from may be largely or completely removed, leaving the less reactive detri- an oceanic DOC reservoir. Hydrocarbon contamination may cause 13 13 tal organic carbon as the only organic remains. In this case, δ Corg local positive shift in δ Corg, but it is inconsistent with low TOC con- values of organic-poor rocks may record the isotope signature of de- tents at the same stratigraphic level in multiple sections (Figs. 3–5), trital organic matter rather than photosynthetic organic carbon the lack of physical evidence for such an event, and invariable δ18O from the surface ocean. values across CN1. The more likely cause of the decoupled δ13C- 13 carb–δ Corg across CN1 may be the combination of diagenetic alter- 13 13 5.3. The origin of decoupled δ Ccarb–δ Corg pattern ation and detrital organic carbon contribution. In organic-poor carbonates, substantial or complete removal of primary organic car- 13 13 13 ﬁ The decoupled δ Ccarb–δ Corg across the N–D δ Ccarb excursion bon through bio- and thermal degradation may result in ampli cation (CN1) occurs in organic-poor dolostones that have TOC contents of δ13C signature of less reactive detrital organic carbon (interval A; 13 13 mostly b0.1% (Fig. 6). As summarized in Fig. 9A, such a carbon isotope Fig. 9B). In contrast, the co-varying δ Ccarb–δ Corg trends at the 13 – δ13 pattern implies that the δ Corg values across the majority of CN1 may end of the N D Ccarb excursion (interval B; Fig. 9B) indicate that record isotope signature of non-photosynthetic origin. Among the the majority of TOC was derived and isotopically fractionalized from possible non-photosynthetic organic carbon sources (Fig. 9A), sec- contemporaneous DIC by primary producers. 13 ondary production through chemoautotrophic/methanotrophic bio- The invariable δ Corg valuesofthemiddleN–D through lower Tom- 13 mass contribution commonly results in lower δ Corg and higher motian (intervals C and D; Fig. 9B) occur in organic-rich shale and car- 13 13 Δδ13 ≥ ‰ Δδ C. This is inconsistent with the slight increase of δ Corg and de- bonates (Fig. 6)withhigh C values ( 32 in interval C). The 13 13 13 crease of Δδ C towards the minimum of CN1 (Fig. 9B). The large decoupled δ Ccarb–δ Corg from this interval may have resulted from 13 ﬁ – temporal and spatial variations in δ Corg across CN1 (Fig. 8) are signi cant chemoautotrophic methanotrophic biomass contribution Author's personal copy

G. Jiang et al. / Earth and Planetary Science Letters 317-318 (2012) 96–110 107 to bulk organic matter (Hayes et al., 1999) in anoxic–euxinic environ- through ocean circulation. The early-middle Ediacaran Doushantuo ments, which may have been pervasive across the Yangtze platform basin in the Yangtze Gorges area might well be one of such basins. during late N–D and Tommotian times (e.g., Goldberg et al., 2007; Here the lower-middle part of the Doushantuo Formation has high TOC 13 Guoetal.,2007b;Lehmannetal.,2007;Willeetal.,2008). Under perva- contents (TOC=1.2±1.0%) but invariable δ Corg values (−29.5± sive anoxia/euxinia, growth of aged DOC at regional or basinal scale, if 1.0‰; McFadden et al., 2008; Xiao et al., 2012), similar to those of the 13 not globally, would be an expected outcome, adding “buffering” effects upper N–DandTommotianstrata(Fig. 6). The invariable δ Corg values to the carbon isotopes of organic matter across the lower Tommotian from high-TOC black shale and carbonates of the lower-middle Doushan- 13 δ Ccarb excursion (interval D; Fig. 9B). tuo Formation could be caused by intensive organic carbon recycling by chemo- and methanotrophs under long-lasting anoxic conditions, but theycouldalsobecausedbyisotopic“buffering” from highly concentrat- 13 13 6. Implications for decoupled δ Ccarb–δ Corg from ed DOC in water column and porewater. In fact, the germanium/silica Ediacaran successions (Ge/Si) ratios from chert nodules of the lower Doushantuo strata in this area do suggest high DOC content in seawater and porewater 13 13 The lower N–D δ Ccarb–δ Corg pattern shows striking similarities to (Shen et al., 2011). However, the majority of the Ediacaran Doushantuo those from the basal (Jiang et al., 2010)andupper(McFadden et al., Formation in the Yangtze Gorges area was deposited from a restricted 2008) Doushantuo Formation of the Ediacaran Period (Fig. 10). In all intrashelf lagoon (Bristow et al., 2009; Jiang et al., 2011); to what extent 13 13 three cases, decoupled δ Ccarb–δ Corg are seen in organic-poor car- it records global ocean chemistry is questionable and requires tests in 13 13 bonates (TOC≤0.1‰) and coupled δ Ccarb–δ Corg occur in organic- more open-marine settings. 13 rich black shale and carbonates at the end of the δ Ccarb excursion. 13 13 The shift from decoupled to coupled δ Ccarb–δ Corg in the upper Doushantuo Formation (McFadden et al., 2008) and Shuram Formation 7. Conclusions (Fike et al., 2006) was taken as evidence for the terminal oxidation of a 13 13 unusually large oceanic DOC reservoir, which has been inferred to have Paired δ Ccarb–δ Corg analyses of the early Cambrian (ca. existed since the early Cryogenian (Swanson-Hysell et al., 2010). The 542–520 Ma) strata in the Yangtze Gorges area in South China reveal 13 13 13 13 repetitive occurrences of similar δ Ccarb–δ Corg patterns in Ediacar- strongly decoupled δ Ccarb and δ Corg across the lower Nemakit– 13 an–early Cambrian strata are apparently inconsistent with such a sim- Daldynian (N–D) and basal Tommotian negative δ Ccarb excursions. 13 13 13 plified oxidation model for late Neoproterozoic oceans. Instead, the Across the N–D δ Ccarb excursion, decoupled δ Ccarb–δ Corg are seen 13 13 13 strong correlation of δ Corg with lithology and TOC contents suggest in organic-poor carbonates (TOC≤0.1‰), but coupled δ Ccarb–δ Corg that the carbon isotope patterns from three distinctive stratigraphic in- occur in organic-rich black shale and carbonates at the end of the 13 13 13 tervals may share the same origin: δ Corg values from organic-poor δ Ccarb excursion. Across the basal Tommotian δ Ccarb excursion, 13 carbonates mainly record isotopic signature of diagenetically altered δ Corg values do not show significant change (b2‰). Regional correla- or detrital organic carbon, while those from organic-rich shale and car- tion across a shelf-to-basin transect reveals strong lateral heterogeneity 13 13 bonates record isotopic signature of primary organic carbon (Fig. 9B). It of δ Corg during the early-middle N–D but more homogenized δ Corg does not mean that low TOC carbonates had higher detrital organic car- values across the basin during the late N–D and Tommotian. The tempo- 13 13 bon content, but implies that organic-poor carbonates had originally ral and spatial variations in δ Corg suggest that the decoupled δ Ccarb 13 13 low primary/detrital organic carbon ratio, and diagenetic/thermal deg- and δ Corg across the N–D δ Ccarb excursion are localized features radation may have preferentially removed chemically more reactive resulting from diagenetically altered organic carbon and/or diagenetic photosynthetic organic carbon. If this is the case, the decoupled amplification of detrital organic carbon isotope signature in organic- 13 13 13 13 δ Ccarb and δ Corg patterns recorded in low-TOC carbonates of late poor carbonates. In contrast, decoupled δ Ccarb and δ Corg from the Neoproterozoic successions (e.g., Fike et al., 2006; McFadden et al., upper N–D and lower Tommotian strata were likely originated from 2008; Swanson-Hysell et al., 2010) should not be taken as evidence chemoautotrophic–methanotrophic biomass contribution to organic for the existence of a large oceanic DOC reservoir capable of buffering matter in strongly anoxic/euxinic environments. 13 13 13 the δ C of marine organic matter. The lower N–D δ Ccarb-δ Corg patterns show striking similarities The concept of a large DOC reservoir, however, still remains intrigu- to those from the basal (ca. 635 Ma) and upper (ca. 551 Ma) 13 13 13 ing, although its size and causal link with the negative δ Ccarb anoma- Doushantuo Formation. In all three cases, decoupled δ Ccarb–δ Corg lies require independent evidence. Given the pervasive anoxia in late are seen in organic-poor carbonates (TOC≤ 0.1‰)andcoupled 13 13 Neoproterozoic–early Cambrian oceans (Canfield et al., 2008; δ Ccarb–δ Corg occur in organic-rich black shale and carbonates at 13 Goldberg et al., 2007; Lehmann et al., 2007; Li et al., 2010; Wille et al., the end of the δ Ccarb excursion. These similarities suggest that the 13 13 2008), high DOC concentration in deep-ocean seawater and marine decoupled δ Ccarb–δ Corg from low-TOC carbonates of Ediacaran– pore waters would be an expected outcome. In modern anoxic basins early Cambrian successions recorded isotope signature of diagenetically such as the Black Sea, anoxic deep waters have an average DOC concen- altered or detrital organic carbon rather than the isotope buffering from tration of 120 μM(Ducklow et al., 2007), 2.6 times that of the deep open a large dissolved organic carbon (DOC) reservoir. Given the pervasive ocean (45 μM). In shallow water column close to the chemocline, DOC anoxia/euxinia in Ediacaran–early Cambrian oceans, DOC-rich environ- concentrations are as high as 280 μM, but most of it did not escape bac- ments may have been common, but the casual link between the 13 terial degradation within a few months (Cauwet et al., 2002; Ducklow negative δ Ccarb excursions and a large DOC reservoir requires inde- 13 13 et al., 2007). With much lower oxygen and sulfate level in late Neopro- pendent evidence other than decoupled δ Ccarb–δ Corg from single terozoic–early Cambrian ocean seawater, DOC enrichment of a few sections. times or even tens of times higher than that of the modern ocean is con- Supplementary materials related to this article can be found online ceivable, particularly in restricted basins with less frequent ventilation at doi:10.1016/j.epsl.2011.11.018.

13 13 Fig. 10. Comparison of the δ Ccarb–δ Corg pattern and TOC contents of the lower Yanjiahe Formation (A) with those from the upper Doushantuo Formation (B) and the basal Doushantuo 13 13 13 13 Formation (C). In all three cases, decoupled δ Ccarb–δ Corg are seen in organic-poor carbonates (TOC≤0.1‰) and coupled δ Ccarb–δ Corg occur in organic-rich black shale and carbon- 13 13 13 ates at the end of the δ Ccarb excursion. The repetitive occurrences of identical δ Ccarb–δ Corg patterns at ca. 635 Ma, ca. 551 Ma, and ca. 542 Ma are inconsistent with the isotopic buff- 13 ering from a large dissolved organic carbon reservoir that persisted since the Cryogenian time. Instead, the strong correlation of δ Corg with lithology and TOC contents suggest that the 13 carbon isotope patterns from three distinctive stratigraphic intervals share the same origin: δ Corg values from organic-poor carbonates mainly record isotopic signature of diagenetically altered or detrital organic carbon, while those from organic-rich shale and carbonates record isotopic signature of primary organic carbon. Author's personal copy

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Acknowledgements Guo, Q., Shields, G.A., Liu, C., Strauss, H., Zhu, M., Pi, D., Goldberg, T., Yang, X., 2007b. Trace element chemostratigraphy of two Ediacaran–Cambrian successions in South China: implications for organosedimentary metal enrichment and silicification We thank Dr. M.J. Nicholl for proof reading the manuscript and in the Early Cambrian. Palaeogeogr. Palaeoclimatol. Palaeoecol. 254, 194–216. Drs. Peter deMenocal (Editor), David Fike, and an anonymous reviewer, Guo, J., Li, Y., Han, J., Zhang, X., Zhang, Z., Ou, Q., Liu, J.i., Shu, D., Maruyama, S., Komiya, fi T., 2008. Fossil association from the Lower Cambrian Yanjiahe Format ion in Yangtze for their careful review and constructive suggestions that signi cantly Gorges area, Hubei, South China. Acta Geol. Sin. (English ed.) 82, 1124–1132. improved the paper. This research was supported by the National Guo, J., Li, Y., Shu, D., 2010. Fossil macroscopic algae from the Yanjiahe Formation, Science Foundation (EAR-0745825 and EAR-0745827) and the National Terreneuvian of the Three Gorges area, South China. Acta Palaeontol. Sin. 49, – Natural Science Foundation of China (40621002). Jiang gratefully 336 342. Harris,D.,Horwath,W.R.,vanKessel,C.,2001.Acidfumigationofsoilstoremovecarbonates acknowledges UNLV Sabbatical Leave support. prior to total organic carbon or Carbon-13 isotopic analysis. Soil Sci. Soc. Am. J. 65, 853–1856. Hayes, J.M., 2001. Fractionation of carbon and hydrogen isotopes in biosynthetic processes. References Rev. Mineral. Geochem. 43, 225–277. Hayes, J.M., Kaplan, I.R., Wedeking, K.W., 1983. Precambrian organic geochemistry, Ader, M., Macouin, M., Trindade, R.I.F., Hadrien, M.H., Yang, Z., Sun, Z., Besse, J., 2009. A preservation of the record. In: Schopf, J.W. (Ed.), The Earth's Earliest Biosphere: multilayeredwatercolumnintheEdiacaranYangtzeplatform?Insightsfromcarbonate Its Origin and Evolution. Princeton University Press, Princeton, N.J., pp. 93–134. and organic matter paired δ13C. Earth Planet. Sci. Lett. 288, 213–227. Hayes, J.M., Strauss, H., Kaufman, A.J., 1999. The abundance of 13C in marine organic Banerjee, D.M., Schidlowski, M., Siebert, F., Brasier, M.D., 1997. Geochemical changes matter and isotopic fractionation in the global biogeochemical cycle of carbon across the Proterozoic–Cambrian transition in the Durmala phosphorite mine section, duringthepast800Ma.Chem.Geol.161,103–125. Mussoorie Hills, Garhwal Himalaya, India. Palaeogeogr. Palaeoclimatol. Palaeoecol. Hollander, D.J., Smith, M.A., 2001. Microbially mediated carbon cycling as a control on 132, 183–194. the δ13C of sedimentary carbon in eutrophic Lake Mendota (USA): new models for Bekker, A., Holmden, C., Beukes, N.J., Kenig, F., Eglingtone, B., Patterson, W.P., 2008. interpreting isotopic excursions in the sedimentary record. Geochim. Cosmochim. Fractionation between inorganic and organic carbon during the Lomagundi Acta 65, 4321–4337. (2.22–2.1 Ga) carbon isotope excursion. Earth Planet. Sci. Lett. 271, 278–291. Irwin, H., Curtis, C., Coleman, M., 1977. Isotopic evidence for source of diagenetic carbonates Bradley, A.S., Hayes, J.M., Summons, R.E., 2009. Extraordinary 13C enrichment of diether formed during burial of organic-rich sediments. Nature 269, 209–213. lipids at the Lost City Hydrothermal Field indicates a carbon-limited ecosystem. Ishikawa, T., Ueno, Y., Komiya, T., Sawaki, Y., Han, J., Shu, D., Li, Y., Maruyama, S., Yoshida, N., Geochim. Cosmochim. Acta 73, 102–118. 2008. Carbon isotope chemostratigraphy of a Precambrian/Cambrian boundary section Brasier, M.D., Magaritz, M., Corfield, R., Luo, H., Wu, X., Ouyang, L., Jiang, Z., Hamdi, B., He, T., in the Three Gorge area, South China: prominent global-scale isotope excursions just Fraser, A.G., 1990. The carbon- and oxygen-isotope record of the Precambrian–Cambri- before the Cambrian Explosion. Gondwana Res. 14, 193–208. an boundary interval in China and Iran and their correlation. Geol. Mag. 127, 319–332. Jenkins, R.J.F., Cooper, J.A., Compston, W., 2002. Age and biostratigraphy of Early Cambrian Brasier, M.D., Corfield, R.M., Derry, L.A., Rozanov, A.Y., Zhuravlev, A.Y., 1994. Multiple tuffs from SE Australia and southern China. J. Geol. Soc. London 159, 645–658. δ13C excursions spanning the Cambrian explosion to the Botomian crisis in Siberia. Jiang, G., Sohl, L.E., Christie-Blick, N., 2003. Neoproterozoic stratigraphic comparison of the Geology 22, 455–458. Lesser Himalaya (India) and Yangtze Block (South China): paleogeographic implications. Brasier, M.D., Shields, G., Kuleshov, V.N., Zhegallo, E.A., 1996. Integrated chemo- and Geology 31, 917–920. biostratigraphic calibration of early animal evolution; Neoproterozoic–Early Jiang, S.-Y., Pi, D.-H., Heubeck, C., Frimmel, H., Liu, Y.-P., Deng, H.-L., Ling, H.-F., Yang, J.-H., Cambrian of Southwest Mongolia. Geol. Mag. 133, 445–485. 2009. Early Cambrian ocean anoxia in South China. Nature 459, E5–E6. Bristow, T.F., Kennedy, M.J., Derkowski, A., Droser, M., Jiang, G., Creaser, R.A., 2009. Jiang,G.,Wang,X.,Shi,X.,Zhang,S.,Xiao,S.,Dong,J.,2010.Organiccarbonisotopeconstraints Mineralogical constraints on the paleoenvironments of the Ediacaran Doushantuo on the dissolved organic carbon (DOC) reservoir at the Cryogenian–Ediacaran transition. Formation. Proc. Natl. Acad. Sci. U. S. A. 106, 13190–13195. Earth Planet. Sci. Lett. 299, 159–168. Calver, C.R., 2000. Isotope stratigraphy of the Ediacarian (Neoproterozoic III) of the Jiang, G., Shi, X., Zhang, S., Wang, Y., Xiao, S., 2011. Stratigraphy and paleogeography of Adelaide rift complex, Australia, and the overprint of water column stratification. the Ediacaran Doushantuo Formation (ca. 635–551 Ma) in South China. Gondwana Precambrian Res. 100, 121–150. Res. 19, 831–849. Canfield, D.E., Poulton, S.W., Knoll, A.H., Narbonne, G.M., Ross, G., Goldberg, T., Strauss, H., Kaufman, A.J., Knoll, A.H., 1995. Neoproterozoic variations in the C-isotopic composition 2008. Ferruginous conditions dominated later Neoproterozoic deep-water chemistry. of seawater: stratigraphic and biogeochemical implications. Precambrian Res. 73, Science 321, 949–952. 27–49. Cauwet, G., Deliat, G., Krastev, A., Shtereva, G., Becquevort, S., Lancelot, C., Momzikoff, Kaufman, A.J., Knoll, A.H., Semikhatov, M.A., Grotzinger, J.P., Jacobsen, S.B., Adams, W., A., Saliot, A., Cociasu, A., Popa, L., 2002. Seasonal DOC accumulation in the Black 1996. Integrated chronostratigraphy of Proterozoic–Cambrian boundary beds in Sea: a regional explanation for a general mechanism. Mar. Chem. 79, 193–205. the western Anabar region, northern Siberia. Geol. Mag. 133, 509–533. Chen, P., 1984. Discovery of Lower Cambrian small shelly fossils from Jijiapo, Yichang, Kaufman, A.J., Knoll, A.H., Narbonne, G.M., 1997. Isotopes, ice ages, and terminal Proterozoic west Hubei and its significance. Prof. Pap. Stratigr. Palaeontol. 13, 49–66. earth history. Proc. Natl. Acad. Sci. U. S. A. 94, 6600–6605. Chen, D., Wang, J., Qing, H., Yan, D., Li, R., 2009. Hydrothermal venting activities in the Kimura, H., Watanabe, Y., 2001. Oceanic anoxia at the Precambrian–Cambrian boundary. Early Cambrian, South China: petrological, geochronological and stable isotopic Geology 29, 995–998. constraints. Chem. Geol. 258, 168–181. Knoll, A.H., Grotzinger, J.P., Kaufman, A.J., Kolosov, P., 1995. Integrated approaches to Clayton, C.J., 1991. Effect of maturity on carbon isotope ratios of oils and condensates. terminal Proterozoic stratigraphy: an example from the Olenek Uplift, northeastern Org. Geochem. 17, 887–899. Siberia. Precambrian Res. 73, 251–270. Compston, W., Zhang, Z., Cooper, J.A., Ma, G., Jenkins, R.J.F., 2008. Further SHRIMP Kouchinsky, A., Bengtson, S., Pavlov, V., Runnegar, B., Torssander, P., Young, E., Ziegler, K., geochronology on the early CambrianofSouthChina.Am.J.Sci.308,399–420. 2007. Carbon isotope stratigraphy of the Precambrian Cambrian Sukharikha River sec- Conway, N.M., Kennicutt, M.C., Van Dover, C.L., 1994. Stable isotopes in the study of tion, northwestern Siberian platform. Geol. Mag. 144, 609–618. marine chemosynthetic-based food webs. In: Lajtha, K., Michener, R. (Eds.), Stable Lehmann, B., Nägler, T.F., Holland, H.D., Wille, M., Mao, J., Pan, J., Ma, D., Dulski, P., 2007. Isotopes in Ecology and Environmental Science. Blackwell Scientific Publications, Highly metalliferous carbonaceous shale and Early Cambrian seawater. Geology Oxford, pp. 158–186. 35, 403–406. Derry, L.A., 2010a. A burial diagenesis origin for the Ediacaran Shuram-Wonoka carbon Li, D.A., Ling, H.-F., Jiang, S.-Y., Pan, J.-Y., Chen, Y.-Q., Cai, Y.-F., Feng, H.-Z., 2009. New isotope anomaly. Earth Planet. Sci. Lett. 294, 152–162. carbon isotope stratigraphy of the Ediacaran–Cambrian boundary interval from Derry, L.A., 2010b. On the significance of δ13C correlations in ancient sediments. Earth SW China: implications for global correlation. Geol. Mag. 146, 465–484. Planet. Sci. Lett. 296, 497–501. Li, C., Love, G.D., Lyons, T.W., Fike, D.A., Sessions, A.L., Chu, X., 2010. A stratified redox Des Marais, D.J., Strauss, H., Summons, R.E., Hayes, J.M., 1992. Carbon isotope evidence model for the Ediacaran Ocean. Science 328, 80–83. for the stepwise oxidation of the Proterozoic environment. Nature 359, 605–609. Loh, A.N., Bauer, J.E., Druffel, E.R.M., 2004. Variable ageing and storage of dissolved organic Ding, L., Li, Y., Chen, H., 1992. Discovery of Micrhystridium regulare from Sinian–Cambrian components in the open ocean. Nature 430, 877–881. boundary strata in Yichang, Hubei, and its stratigraphic significance. Acta Palaeontol. Maloof, A.C., Schrag, D.P., Crowley, J.L., Bowring, S.A., 2005. An expanded record of Early Sin. 9, 303–309. Cambrian carbon cycling from the Anti-Atlas Margin, Morocco. Can. J. Earth Sci. 42, Dong, L., Xiao, S., Shen, B., Zhou, C., Li, G., Yao, J., 2009. Basal Cambrian microfossils from 2195–2216. the Yangtze Gorges area (South China) and the Aksu area (Tarim Block, northwestern Maloof, A.C., Ramezani, J., Bowring, S.A., Fike, D.A., Porter, S.M., Mazouad, M., 2010a. China). J. Paleontol. 83, 30–44. Constraints on early Cambrian carbon cycling from the duration of the Nemakit– Ducklow, H.W., Hansell, D.A., Morgan, J.A., 2007. Dissolved organic carbon and nitrogen Daldynian–Tommotian boundary δ13C shift, Morocco. Geology 38, 623–626. in the Western Black Sea. Mar. Chem. 105, 140–150. Maloof, A.C., Porter, S.M., Moore, J.L., Dudas, F.O., Bowring, S.A., Higgins, J.A., Fike, D.A., Fike, D.A., Grotzinger, J.P., Pratt, L.M., Summons, R.E., 2006. Oxidation of the Ediacaran Eddy, M.P., 2010b. The earliest Cambrian record of animals and ocean geochemical Ocean. Nature 444, 744–747. change. Geol. Soc. Am. Bull. 122, 1731–1774. Goldberg, T., Strauss, H., Guo, Q., Liu, C., 2007. Reconstructing marine redox conditions McFadden, K.A., Huang, J., Chu, X., Jiang, G., Kaufman, A.J., Zhou, C., Yuan, X., Xiao, S., for the Early Cambrian Yangtze Platform: evidence from biogenic sulphur and organic 2008. Pulsed oxidation and biological evolution in the Ediacaran Doushantuo Formation. carbon isotopes. Palaeogeogr. Palaeoclimatol. Palaeoecol. 254, 175–193. Proc. Natl. Acad. Sci. U. S. A. 105, 3197–3202. Guo,Q.,Strauss,H.,Liu,C.,Goldberg,T.,Zhu,M.,Pi,D.,Heubeck,C.,Vernhet,E.,Yang,X.,Fu,P., Qian, Y., 1999. Taxonomy and Biostratigraphy of Small Shelly Fossils in China. Scientific 2007a. Carbon isotopic evolution of the terminal Neoproterozoic and early Cam- Publishing House, Beijing. brian: evidence from the Yangtze Platform, South China. Palaeogeogr. Palaeocli- Qian, Y., Li, G., Zhu, M., 2001. The Meishucunian stage and its small shelly fossil sequence matol. Palaeoecol. 254, 140–157. in China. Acta Palaeontol. Sin. 40, 54–62. Author's personal copy

110 G. Jiang et al. / Earth and Planetary Science Letters 317-318 (2012) 96–110

Rothman, D.H., Hayes, J.M., Summons, R.E., 2003. Dynamics of the Neoproterozoic carbon Xiao, S., McFadden, K.A., Peek, S., Kaufman, A.J., Zhou, C., Jiang, G., Hu, J., 2012. Integrated cycle. Proc. Natl. Acad. Sci. U. S. A. 100, 8124–8129. chemostratigraphy of the Doushantuo Formation at the northern Xiaofenghe Shen, Y., Schidlowski, M., 2000. New C isotope stratigraphy from Southwest China: section (Yangtze Gorges, South China) and its implication for Ediacaran implications for the placement of the Precambrian–Cambrian boundary on the stratigraphic correlation and ocean redox models. Precambrian Res. 192–195, Yangtze Platform and global correlations. Geology 28, 623–626. 125–141. Shen, B., Lee, C.-T.A., Xiao, S., 2011. Germanium/silica ratios in diagenetic chert nodules Yang, J., Jiang, S., Ling, H., Feng, H., Chen, Y., Chen, J., 2004. Paleoceanographic significance from the Ediacaran Doushantuo Formation, South China. Chem. Geol. 280, 323–335. of redox-sensitive metals of black shales in the basal Lower Cambrian Niutitang Steiner, M., Li, G., Qian, Y., Zhu, M., Erdtmann, B.-D., 2007. Neoproterozoic to Early Cambrian Formation in Guizhou Province, South China. Prog. Nat. Sci. 14, 152–157. small shelly fossil assemblages and a revised biostratigraphic correlation of the Yangtze Yao, J., Xiao, S., Yin, L., Li, G., Yuan, X., 2005. Basal Cambrian microfossils from the Yurtus Platform (China). Palaeogeogr. Palaeoclimatol. Palaeoecol. 254, 67–99. and Xishanblaq Formations (Tarim, north-west China): systematic revision and Summons, R.E., Jahnke, L.L., Roksandic, Z., 1994. Carbon isotope fractionation in lipids biostratigraphic correlation of Micrhystridium-like acritarchs. Palaeontology from methanotrophic bacteria: relevance for interpretations of the geochemical 48, 687–708. record of biomarkers. Geochim. Cosmochim. Acta 58, 2853–2863. Yin, L., 1997. Precambrian–Cambrian transitional acritarch biostratigraphy of the Yangtze Summons, R., Franzmann, P.D., Nichols, P.D., 1998. Carbon isotopic fractionation associated Platform. Bull. Natl. Mus. Nat. Sci. (Taipei) 10, 217–231. 13 with methylotrophic methanogenesis. Org. Geochem. 28, 465–475. Young, S.A., Saltzman, M.R., Bergström, S.M., Leslie, S.A., Xu, C., 2008. Paired δ Ccarb 13 Swanson-Hysell, N.L., Rose, C.V., Calmet, C.C., Halverson, G.P., Hurtgen, M.T., Maloof, and δ Corg records of Upper Ordovician (Sandbian-Katian) carbonates in North A.C., 2010. Cryogenian glaciation and the onset of carbon-isotope decoupling. Science America and China: implications for paleoceanographic change. Palaeogeogr. 328, 608–611. Palaeoclimatol. Palaeoecol. 270, 166–178. Tocqué, E., Behar, F., Budzinski, H., Lorant, F., 2005. Carbon isotopic balance of kerogen Zhao, Z., Xing, Y., Ding, Q., Liu, G., Zhao, Y., Zhang, S., Meng, X., Yin, C., Ning, B., Han, P., pyrolysis effluents in a closed system. Org. Geochem. 36, 893–905. 1988. The Sinian System of Hubei. China University of Geosciences Press, Wuhan. Wang, H.Z., 1985. Atlas of the Paleogeography of China. Cartographic Publishing House, Zhao, J.-H., Zhou, M.-F., Yan, D.-P., Zheng, J.-P., Li, J.-W., 2011. Reappraisal of the ages of Beijing. Neoproterozoic strata in South China: no connection with the Grenvillian orogeny. Wang, J., Li, Z.-X., 2003. History of Neoproterozoic rift basins in South China: implications Geology 39, 299–302. for Rodinia break-up. Precambrian Res. 122, 141–158. Zhou, C., Zhang, J., Li, G., Yu, Z., 1997. Carbon and oxygen isotopic record of the Early Wang, X., Hoffknecht, A., Xiao, J., Li, Z., 1993. Thermal maturity of the Sinian and Early Cambrian from the Xiaotan Section, Yunnan, South China. Sci. Geol. Sin. 32, Paleozoic in west Hubei, China, assessed by CAI, reflectance and geochemical studies. 201–211. Stratigr. Paleontol. China 2, 19–45. Zhu, M., Zhang, J., Steiner, M., Yang, A., Li, G., Erdtmann, B.D., 2003. Sinian-Cambrian Wang, T.G., Li, M., Wang, C., Wang, G., Zhang, W., Shi, Q., Zhu, L., 2008. Organic molecular stratigraphic framework for shallow- to deep-water environments of the Yangtze evidence in the Late Neoproterozoic Tillites for a palaeo-oceanic environment during platform: an integrated approach. Prog. Nat. Sci. 13, 951–960. the snowball Earth era in the Yangtze region, southern China. Precambrian Res. 162, Zhu, M., Zhang, J., Yang, A., 2007. Integrated Ediacaran (Sinian) chronostratigraphy of 317–326. South China. Palaeogeogr. Palaeoclimatol. Palaeoecol. 254, 7–61. Whelan, J.K., Thompson-Rizer, C.L., 1993. Chemical methods for assessing kerogen and proto- Zhu, R.X., Li, X.H., Hou, X.G., Pan, Y.X., Wang, F., Deng, C.L., He, H.Y., 2009. SIMS U–Pb zircon kerogen types and maturity. In: Engel, M.H., Macko, S.A. (Eds.), Organic Geochemistry — age of a tuff layer in the Meishucun section, Yunnan, southwest China: constraint on Principles and Applications. Plenum Press, New York, pp. 289–353. the age of the Precambrian–Cambrian boundary. Sci. China Ser. D 52, 1385–1392. Wille, M., Nagler, T.F., Lehmann, B., Schroder, S., Kramers, J.D., 2008. Hydrogen sulphide release to surface waters at the Precambrian/Cambrian boundary. Nature 453, 767–769.