Post-depositional origin of highly 13C-depleted carbonate in the Doushantuo cap dolostone in South China: Insights from petrography and stable carbon isotopes

Sedimentary Geology 242 (2011) 71–79

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Post-depositional origin of highly 13C-depleted carbonate in the Doushantuo cap dolostone in South China: Insights from petrography and stable carbon isotopes Zhijia Lin a, b, 1, Qinxian Wang a, b,⁎, 1, Dong Feng a, Qian Liu a, b, Duofu Chen a,⁎ a b

CAS Key Laboratory of Marginal Sea Geology, Guangzhou Institute of Geochemistry, Chinese Academy of Sciences, Guangzhou 510640, China Graduate University of Chinese Academy of Sciences, Beijing 100049, China

a r t i c l e

i n f o

Article history: Received 23 April 2011 Received in revised form 10 October 2011 Accepted 16 October 2011 Available online 20 October 2011 Editor: B. Jones Keywords: Doushantuo cap dolostone Carbon isotopes Methane oxidation South China Neoproterozoic Marinoan

a b s t r a c t Extremely negative carbon isotope values from the Doushantuo cap dolostone succession in the Yangtze Gorges, South China were interpreted as direct evidence of the hypothesis of methane hydrate destabilization during the Marinoan deglaciation. However, this suggestion remains uncertain due to the obscurity of the paragenetic sequence of carbonate minerals with diverse carbon isotopic compositions. Here, we conducted macroscopic and microscopic petrographic and carbon isotopic investigations of the cap dolostone succession at the Jiulongwan section in the Yangtze Gorges. Our results show that extreme δ13C values down to − 44‰ exclusively occur in calcite, whereas δ13C values for the host dolomite range mostly between − 4‰ and − 2‰, and clearly reveal that the host dolomite were locally dissolved and/or replaced by the calcite bearing highly 13C-depleted signals. This finding suggests that methane-oxidation activity as indicated by the strongly depleted δ13C signals did not occur until after the cap dolostone deposition. Thus, the presence of the extremely negative δ13C values could not be taken as evidence for the proposed methane hypothesis. Nevertheless, this conclusion is not contradictory to the scenario that massive releases of methane from permafrost hydrates during deglaciation, much prior to methane oxidation documented here, could drive substantial negative δ13C shifts well exhibited by the Doushantuo cap dolostone and others around the world. © 2011 Elsevier B.V. All rights reserved.

1. Introduction The late Neoproterozoic era, about 750 to 542 million years (Ma) ago, is punctuated by at least two global glaciations, generally termed Sturtian and Marinoan, respectively, each extending to very low paleolatitudes (Kennedy et al., 1998; Fairchild and Kennedy, 2007; Hoffman and Li, 2009; Macdonald et al., 2010). The glacial deposits are commonly capped sharply by ‘cap dolostones’ on almost every continent (e.g. Kennedy, 1996; Kaufman et al., 1997; Hoffman et al., 1998; Jiang et al., 2003a), implying severe and rapid climatic oscillations thought to serve as an ‘environmental filter’ for biological evolution (Hoffman et al., 1998; Runnegar, 2000; Hoffman and Schrag, 2002). These postglacial cap dolostones have been of especial interest, understandably, as they may provide important hints about the nature of the environmental fluctuations and the evolution of life in the aftermath of global glacial ages. The most compelling candidate for such carbonate rocks is the well-developed Marinoan cap dolostones at the base of the Ediacaran, dated as ~ 635 Ma (Condon et al., 2005). Marinoan cap dolostones are thin (an average thickness of 18.5 m),

⁎ Corresponding authors. E-mail addresses: [email protected] (Q. Wang), [email protected] (D. Chen). 1 Both authors should be regarded as joint first authors. 0037-0738/$ – see front matter © 2011 Elsevier B.V. All rights reserved. doi:10.1016/j.sedgeo.2011.10.009

regionally consistent carbonate rocks composed predominantly of microcrystalline dolomite, and are characterized by a curious suite of sedimentary features including tepee-shaped structures, stromatactislike cavities, tube-like structures, sheet cracks, peloids, isopachous cements, stromatolites, and local barite fans (e.g. Kennedy, 1996; James et al., 2001; Kennedy et al., 2001; Hoffman and Schrag, 2002; Jiang et al., 2003a, 2006; Nogueira et al., 2003; Gammon et al., 2005; Hoffman et al., 2007; Shields et al., 2007; Wang et al., 2008; Hoffman, 2011; Peng et al., 2011). In addition, they are geochemically distinctive, with unusually negative carbon isotope shifts of about 4‰ to 7‰ (e.g. Kennedy, 1996; Hoffman et al., 1998; James et al., 2001; Jiang et al., 2003a; Font et al., 2006; Shen et al., 2008). To explain the genesis of the enigmatic Marinoan cap dolostones, several alternative models have been proposed. One theory is that the cap dolostones result from postglacial upwelling of deep water rich in 13C-depleted alkalinity in a stratified ocean (the upwelling hypothesis, Grotzinger and Knoll, 1995; Knoll et al., 1996; Shen et al., 2008). Another one ascribes the cap dolostones to the aftermath of postglacial enhanced weathering of both carbonate and silicate rocks (the snowball Earth hypothesis, Hoffman et al., 1998; Hoffman and Schrag, 2002; Higgins and Schrag, 2003). In contrast to above models, a plausible explanation by Kennedy et al. (2001) invokes destabilization of methane hydrates in terrestrial permafrost following postglacial warming and marine transgression as the mechanism for

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the formation of Marinoan cap dolostones. Subsequently, extremely negative δ 13C values as low as −48‰ have been found by Jiang et al. (2003a, 2006) and Wang et al. (2008) from the Ediacaran Doushantuo cap dolostone succession in the Yangtze Gorges area, South China, and have been taken as direct evidence to support the proposed methane hydrate destabilization event during deglaciation. Conversely, Zhou et al. (2010), on the basis of macroscopic petrographic observations of the Doushantuo cap dolostone, found that highly negative δ13C values occur in calcite that precipitated after the cap dolostone deposition, and thus argued that the exceptionally negative δ13C signals were not recorded until after the deposition of the cap dolostone, and accordingly concluded that the presence of the extremely negative δ13C values does not support the ‘methane hydrate’ hypothesis. Similarly, Bristow et al. (2011) presented strong evidence from carbonate clumped isotope thermometry, 87Sr/ 86Sr ratios, trace element content and clay mineral from the Doushantuo cap dolostone that calcite bearing highly negative δ 13C values formed more than 1.6 Ma after the cap dolostone deposition, and suggested that their finding denies exceptional carbon isotope evidence for the methane release event during Marinoan deglaciation. However, there is still dearth of systematic studies of petrography and stable isotopes for the Doushantuo cap dolostone, which would likely place important constraints on the relationship between extremely negative carbon isotope signals registered in the cap dolostone and the proposed methane hydrate destabilization. In this paper, we carried out macroscopic and microscopic petrographic and bulk and micro-drilled carbon isotopic investigations of the Doushantuo cap dolostone succession (Jiulongwan section) in the Yangtze Gorges, South China. Our own data clearly reveal that highly 13C-depleted signals are unrelated to the proposed methane release during deglaciation. Consequently, we consider that the presence of extremely low δ 13C values could not be used as evidence for the inferred methane release hypothesis. Nonetheless, our proposal is not contradictory to the scenario that methane hydrate destabilization during deglaciation could contribute to the global negative δ13C anomalies of cap dolostones (Kaufman et al., 1997; Kennedy et al., 2001, 2008). 2. Geological background The late Cryogenian Nantuo Formation and the early Ediacaran Doushantuo Formation were deposited in a southeast-facing rift to passive margin that developed on the Yangtze Block of South China after the breakup of the supercontinent Rodinia (Fig. 1; ~ 850– 750 Ma) (Jiang et al., 2003b; Wang and Li, 2003). The Nantuo Formation mainly comprises glacial diamictite (Fig. 1B), equivalent to glacial deposits of the Marinoan glaciation in the Neoproterozoic, and is radiometrically constrained between 654.5 ± 3.8 Ma and 635.2 ± 0.6 Ma (Condon et al., 2005; Zhang et al., 2008). It is conformably overlain by the Doushantuo Formation which is mostly composed of carbonates, black shales, and phosphatic deposits (Fig. 1B). The Doushantuo Formation is richly fossiliferous, containing fossil embryos, acanthomorphic acritarchs, multicellular algae, and purported bilaterians that are purportedly the earliest animals (Xiao et al., 1998; Chen et al., 2004; Yin et al., 2007). In addition, U–Pb Zircon dates indicate that the depostion of the Doushantuo Formation occurred between 635.2 ± 0.6 Ma and 551.1 ± 0.7 Ma (Condon et al., 2005), presenting most of the Ediacaran Period. The base of the Doushantuo Formation on the Yangtze Block consists of the cap dolostone succession, typically b10 m thick and in sharp contact with the underlying Nantuo diamictite without evidence of reworking or hiatus (Jiang et al., 2003a; Shen et al., 2005). The Doushantuo cap dolostone exhibits peculiar sedimentological and isotopic characteristics mentioned above, similar to equivalent cap dolostones in Brazil (Font et al., 2006), Namibia (Kennedy et al., 2001; Hoffman et al., 2007), Australia (Kennedy, 1996; Rose and Maloof, 2010) and elsewhere. A detailed lithologic study reveals that the

Doushantuo cap dolostone was deposited below or near storm wave base (Jiang et al., 2006) and records postglacial transgression following the Marinoan glaciation. Although the Doushantuo cap dolostone crops out widely, this paper focuses on the Jiulongwan section of the Yangtze platform (Fig. 1A), situated at the southern limb of the Huangling anticline in the Yangtze Gorges area (Fig. 1C). At the Jiulongwan section (Fig. 1D), the cap dolostone succession is approximately 3.4 m in thickness and consists primarily of microcrystalline dolomite. The succession can be divided into three lithostratigraphic sequences. The lower unit C1 is presented by 0.9-mthick disrupted and cemented microcrystalline dolomite that contains abundant fractures, sheet cracks, localized breccias and cavities commonly filled with multiple generations of mineral cements. The middle unit C2 is characterized by ~1.5-m-thick laminated microcrystalline dolomite and contains dolomitic limestone, with low angle cross-bedding. The upper unit C3 consists of ~1.0-m-thick silty dolomite and limestone with mm-scale laminae. 3. Sampling and methods Forty-nine large rock blocks were continuously collected from fresh outcrops of the Jiulongwan section. Each of them was cut perpendicular to the bedding surface into two pieces in the laboratory. One was used to prepare a thin section and a polished slab for petrographic observations. The other was further cut transversely into smaller pieces, each 1–4 cm thick, totally producing 110 bulk samples. These samples were crushed into chips, which were subsequently picked for clean fragments, avoiding visible veining and siliciclastic components. The selected chips were sonicated in double-distilled deionized water, then air-dried at room temperature. After drying, the chips were pulverized into powder (less than 200 mesh) for X-ray diffraction (XRD), carbon and oxygen isotope analyses. Thin sections and polished slabs of the samples were stained with Alizarin Red-S and potassium ferricyanide solutions before petrographic observations to differentiate between calcite and dolomite (Dickson, 1966). Stained thin sections were observed using a LEICADMRX optical microscope equipped with a LEICA DC500 digital camera. Stained polished slabs were photographed with a Nikon D300 digital camera. Besides, powdered samples of dolomite, calcite and their mixture were micro-drilled, respectively, from some stained slabs (which were deliberately selected, according to the mineralogical and isotopic compositions of bulk samples) for carbon and oxygen isotope measurements. For XRD analyses, a set of powdered samples were mounted on glass slides with ethanol. Then prepared slides were scanned using a Rigaku D/MAX 2500 diffractometer with CuKα radiation operating at 40 kV and 150 mA at the First Institute of Oceanography, State Oceanic Administration, China. Scans were run from 3 to 80° 2θ, at a speed of 2° 2θ/min. The divergence, scattering and receiving slits were 1°, 1° and 0.3 mm, respectively. The XRD patterns were subsequently analyzed using the program SIROQUANT, which is based upon the Rietveld method to quantify mineralogical content (Taylor and Matulis, 1991). Both bulk and micro-drilled powdered samples were reacted with 100% phosphoric acid to liberate CO2 gas for carbon and oxygen isotope analyses. The prepared gas from bulk samples was measured at the State Key Laboratory of Environmental Geochemistry, Chinese Academy of Sciences, using a GV IsoPrime I stable isotope ratio mass spectrometer. The resulting gas from micro-drilled samples was analyzed at the Key Laboratory of Isotope Geochronology and Geochemistry, Chinese Academy of Sciences, using a GV IsoPrime II stable isotope ratio mass spectrometer. All carbon and oxygen isotope data are reported in the conventional delta notation in per mil (‰) relative to the Vienna PeeDee Belemnite (VPDB) international standard. The standard deviations are less than 0.1‰ (2σ) for both δ13C and δ18O values.

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Fig. 1. Geological background for the Doushantuo cap dolostone. A: Paleogeographic reconstruction of the southeast-facing rift to passive margin on the Yangtze block during late Neoproterozoic (Jiang et al., 2003b; Wang and Li, 2003; Zhou et al., 2004), showing the location of the Jiulongwan section collected and analyzed in the Yangtze Gorges area (square). B: A conceptual transect (a–a′ in A; not to scale) across the Yangtze block showing the stratigraphic occurrence of the Doushantuo cap dolostone (Jiang et al., 2003a, b). Age data are from Condon et al. (2005) and Zhang et al. (2008). C: Simplified geological map of the Yangtze Gorges area (Zhao et al., 1985), showing locations (triangles) of occurrence of extreme δ13C signals registered in the Doushantuo cap dolostone succession (Jiang et al., 2003a; Wang et al., 2008; Zhou et al., 2010; Bristow et al., 2011; this paper). D: Field photograph of the Doushantuo cap dolostone at the Jiulongwan section in the Yangtze Gorges area, showing lithostratigraphic sequences.

4. Petrography Recent sedimentological and petrographic investigations of the Doushantuo cap dolostone succession in South China have been carried out by some authors (Jiang et al., 2003a, 2006; Wang et al., 2008). However, none of these investigations refer to the paragenetic sequence of minerals (dolomite and calcite) in the cap dolostone. Here, we carefully conducted petrographic observations on polished slabs and thin sections stained with Alizarin Red-S and potassium ferricyanide, and present the mineral paragenetic sequence in the cap dolostone at the Jiulongwan section. Our petrographic analyses show that dolomite, which is stained light gray to blue on the polished slabs due to the presence of ferroan

dolomite (Fig. 2), often occurs as microcrystalline forms and is the principal component of the host matrix in the cap dolostone. The microcrystalline dolomite matrix is commonly crosscut by veins filled with microsparry to sparry calcite which is stained violet to red or occasionally with silica, but generally retains most of original petrographic textures, such as laminations (Fig. 2). Two different types of the calcite veins are identified. The first (Type I) are irregular in shape, with jagged margins (Fig. 2). The Type I veins can be seen to locally interrupt and dislocate the primary stromatolitic laminae that are made up of microcrystalline dolomite matrix (Fig. 2B, C and D), and sometimes dolomite residues are found within the veins on the slabs (Fig. 2C). These findings strongly suggest that calcite veins of Type I precipitated later than the host

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Fig. 2. A–D: Representative polished slab images of the Doushantuo cap dolostone at the Jiulongwan section showing the diverse carbonate components and their carbon isotope values (‰, VPDB). Polished slabs were stained with Alizarin Red-S and potassium ferricyanide. Matrix dolomite is stained light gray to blue, and is generally crosscut by two types of veins (I and II) filled with calcite which is stained violet to red. Type I veins are irregular in shape with jagged margins, and locally disrupted the primary laminae with dolomite residues (arrow); Type II veins are simple with smooth margins, generally lying sub-vertical to the bedding plane, and sometimes cut across Type I veins. δ13C values for dolomite (circles) range from − 6.2‰ to + 1.4‰, from − 44.4‰ to − 26.2‰ for calcite (Type I) (triangles), from − 5.5‰ to − 1.3‰ for calcite (Type II) (diamond), and from − 28.3‰ to − 4.5‰ for mixtures of calcite and dolomite (squares). A–D are from polished slabs JLW-007, JLW-44, JLW-45 and JLW-48, respectively.

dolomite matrix. The second (Type II) are comparatively simple with smooth margins, and generally lie sub-vertical to the bedding plane (Fig. 2A). In some cases, it can be seen that Type II veins cut sharply across Type I veins (Fig. 2A); therefore, there is no doubt that the calcite veins of Type II formed well after those of Type I. On the thin sections which are mirror images of polished slabs, we observed that: 1) the matrix dolomite was dissolved along the veins of Type I and the resultant pore spaces were filled by calcite (Fig. 3); 2) dark organic-rich laminae appear to be crosscut by calcite veins of Type I, with remnants of corroded dolomite within these veins (Fig. 3B), suggesting dedolomitization; and 3) some dolomite rhombs within or near the calcite veins of Type I are found to be partially or completely replaced by calcite (Fig. 3D), showing calcitization of the dolomite matrix (dedolomitization). These microscopic observations unambiguously reveal that the calcite veins of Type I did not occur

until after the formation of the host dolomite, which is consistent with the macroscopic observations. 5. Mineralogy and Isotope geochemistry 5.1. Mineralogical and isotopic compositions of bulk samples The XRD analyses of the bulk samples show that dolomite and calcite are the most abundant minerals (average 48.2 wt.% and 39.6 wt.%, respectively) in the cap dolostone succession, with less amounts of quartz (average 10.0 wt.%) and clay minerals (illite and chlorite; average b2.0 wt.%) (Table 1). The abundance of calcite changes variably along the profile, with relative high contents roughly in units C1 and C3 (Fig. 4B), whereas the abundance of dolomite clearly exhibits a reverse trend, with high contents in unit C2 (Fig. 4C).

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Fig. 3. Representative thin-section photomicrographs of the Doushantuo cap dolostone at the Jiulongwan section, showing dedolomitization of the host dolomite. A: Photomicrograph showing dolomite crystals in which the cloudy, inclusion-rich cores were dissolved along calcite veins of Type I (see text for explanations of the veins) and the resultant pore spaces were filled by calcite (arrows). B: Photomicrograph that shows dark organic-rich laminae interrupted by the Type I veins with dolomite residues. C: Photomicrograph showing that the matrix dolomite was corroded along the Type I veins. Note the linear dolomitic structure (arrow). D: Photomicrograph showing partial or complete replacement of the dolomite rhombs by calcite (arrows). E: Photomicrograph showing island-like remnants of corroded dolomite matrix. Thin sections were stained with Alizarin Red-S and potassium ferricyanide. A is from thin section JLW-41; B–D are from thin section JLW-44; E is from thin section JLW-48. Thin sections JLW-44 and JLW-48 are mirror images of polished slabs JLW-44 and JLW-48 (as indicated in Fig. 2), respectively.

The high resolution δ 13C data are presented in Table 1, and the δ C profile accompanies the stratigraphic section in Fig. 4D. δ 13C values in units C1 and C2 range from −4.8‰ to − 2.6‰ (mean = −3.3‰, SD = 0.4‰, n = 79), showing little variation. In contrast, δ 13C values in unit C3 show wide variations between − 22.7‰ and −3.1‰ (mean = −9.1‰, SD = 4.8‰, n = 31), consistent with high variability of δ 13C obtained from two other sections in the Yangtze Gorges (Huajipo and Wangzishi sections, Fig. 1C) (Jiang et al., 2003a; Wang et al., 2008). In addition, δ 13C values show a moderate negative correlation with calcite contents (r = −0.6) (Fig. 5). Thus, it is conceivable that calcite may be isotopically much lighter than dolomite.

δ 18O values for matrix dolomite range widely between − 12.9‰ and −1.1‰ (mean = −8.1‰, SD = 3.2‰, n = 30), while calcite within Type I veins displays a narrow range of δ 18O values from −9.3‰ to −6.1‰ (mean = −7.8‰, SD = 1.2‰, n = 8). We also note that calcite within veins of Type II is characterized by lower δ 18O values ranging from −15.1‰ to − 11.9‰ (mean = −13.6‰, SD = 1.1‰, n = 7).

5.2. Isotopic compositions of micro-drilled samples

As mentioned above, calcite within Type I veins is exclusively characterized by highly negative carbon isotope values down to −44‰. These values are too depleted in 13C to be ascribed to conventional mechanisms such as diagenetic alteration. Nevertheless, such low δ13C values in carbonate rocks have been known almost from modern and ancient seep carbonates which predominantly resulted from anaerobic oxidation of methane using sulfate in the cold-seep environments (Roberts and Aharon, 1994; Boetius et al., 2000; Peckmann et al., 2002; Feng et al., 2010). Accordingly, Jiang et al. (2003a, 2006) and Wang et al. (2008) contended that the extremely negative δ13C values, in the Doushantuo cap dolostone, are most likely associated with microbial oxidation of methane in cold-seep fluids. In contrast, Bristow et al. (2011), based on carbonate clumped isotope thermometry, Mn/Sr and 87 Sr/86Sr ratios and clay mineral evidence, inferred that these exceptionally 13C-depleted signals are products of thermogenic methane oxidation in hydrothermal fluids at depth. Understandably, although there is a controversy regarding the source of methane-rich fluids (cold-seep or hydrothermal fluids), it is conceivable that the calcite

13

In order to verify above speculation, we micro-drilled powder samples of dolomite, calcite (within veins of Types I and II), and their mixture, respectively, from selected stained slabs corresponding to some bulk samples characterized by high calcite contents or very negative δ 13C values, and analyzed carbon and oxygen isotopic compositions of diverse components, shown in Table 2, Figs. 2 and 4D. Matrix dolomite exhibits relatively negative δ 13C values ranging from −6.2‰ to + 1.4‰ (mean = − 3.0‰, SD = 1.7‰, n = 30), but mostly between − 4.8‰ and − 2.2‰. Similarly, calcite within veins of Type II has slightly depleted δ 13C values from −5.5‰ to − 1.3‰ (mean = − 4.2‰, SD = 1.3‰, n = 7). However, calcite within Type I veins records distinct δ 13C values between −44.4‰ and −26.2‰, with an average value of −36.6‰ (n = 8). Additionally, mixtures of calcite and dolomite record a noticeably large range of isotope values from −28.3‰ to −4.5‰ (mean =−11.6‰, SD =6.0‰, n = 28), consistent with δ13C data of the bulk samples.

6. Discussion 6.1. Methane oxidation after the cap dolostone deposition

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Table 1 Mineralogical and isotopic compositions of bulk samples from the Doushantuo cap dolostone succession at the Jiulongwan section. Sample number JLW-02-01 JLW-02-02 JLW-02-03 JLW-03-01 JLW-03-02 JLW-04-01 JLW-04-02 JLW-04-03 JLW-04-04 JLW-05-01 JLW-05-02 JLW-05-03 JLW-06-01 JLW-06-02 JLW-06-03 JLW-06-04 JLW-07 JLW-08-01 JLW-08-02 JLW-09-01 JLW-09-02 JLW-10-01 JLW-10-02 JLW-10-03 JLW-10-04 JLW-11-01 JLW-11-02 JLW-11-03 JLW-12 JLW-13-01 JLW-13-02 JLW-13-03 JLW-14-01 JLW-14-02 JLW-14-03 JLW-15-01 JLW-15-02 JLW-16-01 JLW-16-02 JLW-16-03 JLW-17 JLW-18 JLW-19-01 JLW-19-02 JLW-20-01 JLW-20-02 JLW-20-03 JLW-21-01 JLW-21-02 JLW-22-01 JLW-22-02 JLW-23-01 JLW-23-02 JLW-23-03 JLW-24-01 JLW-24-02 JLW-25 JLW-26-01 JLW-26-02 JLW-26-03 JLW-26-04 JLW-27 JLW-28-01 JLW-28-02 JLW-29 JLW-30 JLW-31-01 JLW-31-02 JLW-32-01 JLW-32-02 JLW-33 JLW-34-01 JLW-34-02

Height

δ13C

Calcite

Dolomite

Quartz

(cm)

(‰, VPDB)

(wt.%)

(wt.%)

(wt.%)

0.0 3.0 6.0 9.0 12.5 16.0 19.0 22.0 25.0 28.0 31.0 34.0 38.0 41.0 44.0 47.0 51.0 54.5 58.0 61.5 65.0 68.5 71.5 74.5 76.5 77.5 81.0 84.5 88.0 91.5 94.5 97.5 99.5 102.5 105.5 109.5 112.0 114.5 117.5 120.5 124.5 126.0 129.0 131.5 134.0 137.0 140.0 142.5 145.5 149.0 152.5 156.0 159.0 162.0 165.0 168.5 172.0 176.0 179.0 182.0 185.0 189.5 190.5 193.0 195.5 197.0 199.5 202.5 205.5 208.0 210.5 214.5 217.5

− 3.2 − 3.2 − 3.4 − 3.7 − 3.8 − 4.0 − 3.9 − 4.1 − 4.3 − 3.7 − 2.8 − 2.8 − 3.1 − 2.6 − 2.7 − 2.9 − 4.8 − 4.8 − 2.9 − 2.8 − 2.8 − 2.9 − 2.6 − 2.9 − 3.1 − 3.2 − 3.1 − 3.1 − 3.6 − 3.3 − 3.3 − 3.1 − 3.4 − 3.2 − 3.2 − 3.4 − 3.3 − 3.3 − 3.1 − 3.4 − 3.4 − 3.7 − 3.5 − 3.5 − 3.5 − 3.3 − 3.3 − 3.3 − 3.1 − 3.2 − 3.3 − 3.3 − 3.3 − 3.3 − 3.2 − 3.2 − 3.1 − 3.1 − 3.0 − 3.1 − 2.9 − 3.3 − 3.2 − 3.2 − 3.4 − 3.0 − 3.5 − 3.4 − 3.5 − 4.0 − 3.5 − 3.7 − 3.4

61.4

26.6

5.8

51.7

38.6

3.8

37.8

54.8

3.8

34.7

55.2

10.1

13.3

77.9

8.8

12.7 93.5 51 42.9

84.7 3.6 42.8 50

2.6 2.9 6.2 4.1

9.5

84.3

4.2

14.7

75.9

7.4

24.4

66.4

9.2

3.6

90.9

5.5

13.6

81.5

4.9

22.4

74.9

2.7

15.8

78.4

5.8

16.3

78.7

5

13.2

80.7

4.1

17

77.9

5.1

24.8

65.6

5.6

9.4

84.1

3.5

63.3

31.4

5.3

Table 1 (continued) Sample number JLW-35-01 JLW-35-02 JLW-35-03 JLW-36-01 JLW-36-02 JLW-36-03 JLW-37-01 JLW-37-02 JLW-38-01 JLW-38-02 JLW-38-03 JLW-38-04 JLW-39 JLW-40 JLW-41-01 JLW-41-02 JLW-41-03 JLW-42 JLW-43-01 JLW-43-02 JLW-44-01 JLW-44-02 JLW-45-01 JLW-45-02 JLW-45-03 JLW-45-04 JLW-46-01 JLW-46-02 JLW-47-01 JLW-47-02 JLW-47-03 JLW-48-01 JLW-48-02 JLW-48-03 JLW-49-01 JLW-49-02 JLW-50

Height

δ13C

Calcite

Dolomite

Quartz

(cm)

(‰, VPDB)

(wt.%)

(wt.%)

(wt.%)

220.5 223.5 226.5 229.5 232.5 235.5 238.5 241.5 245.0 248.0 251.0 254.0 258.0 262.0 265.0 268.5 272.0 276.0 280.0 282.5 285.0 287.5 290.5 293.5 296.5 299.5 303.5 306.5 309.5 312.5 315.5 318.5 321.5 324.5 327.5 330.5 334.0

− 3.0 − 3.0 − 3.8 − 2.9 − 3.5 − 3.6 − 3.9 − 3.6 − 3.4 − 3.4 − 3.6 − 3.6 − 6.2 − 8.3 − 5.8 − 3.1 − 11.4 − 15.3 − 3.3 − 9.5 − 14.0 − 11.8 − 14.5 − 14.0 − 6.7 − 11.6 − 11.8 − 14.5 − 11.1 − 12.3 − 11.9 − 12.3 − 9.6 − 8.3 − 22.7 − 7.0 − 3.3

3.6

87.3

6.1

23.8

73.3

2.9

9.1

88.9

2.0

10.8 35.4 43.8 46.6 57.0 67.1 43.1 64.6 62.4 68.7 81.4 83.4 47.2 48.0 40.8 40.7 40.0 43.3 32.3 55.7 49.8 62.5 70.4 48.8 51.5

82.9 55.2 29.3 17.9 26.3 7.5 30.3 7.4 27.0 19.8 9.6 11.8 38.0 43.0 48.7 33.1 48.5 46.2 33.3 34.3 40.1 20.8 15.2 27.5

6.3 2.8 15.3 27.0 13.0 20.1 24.6 25.0 9.6 9.5 7.9 3.8 11.0 7.9 8.9 19.0 10.5 9.6 27.1 8.8 8.5 16.7 13.3 22.5 45.5

extremely depleted in 13C was derived from methane oxidation in fluids enriched in methane. Furthermore, Zhou et al. (2010) and Bristow et al. (2011) suggested that the calcite bearing methane-oxidation signals formed later than the deposition of the Doushantuo cap dolostone. Here, our macroscopic and microscopic petrographic investigations of the Doushantuo cap dolostone succession (Jiulongwan section) show that the relatively 13 C-enriched host dolomite were often locally dissolved and/or replaced by the extremely 13C-depleted calcite. This finding undoubtedly reveals that the calcite bearing methane-oxidation signals precipitated after the cap dolostone deposition, which is consistent with the suggestion by Zhou et al. (2010) and Bristow et al. (2011). In combination with available data from other sections in South China (Jiang et al., 2003a; Wang et al., 2008; Zhou et al., 2010; Bristow et al., 2011), we deduce that methane-oxidation activity, indicated by the extreme δ13C values which have never been found outside the Yangtze Gorges area, occurred much later than the deposition of the cap dolostone.

6.2. Interpretation of carbon isotopic signatures of the host dolomite It is crucial to assess the potential degree of diagenetic alteration of the Doushantuo cap dolostone overlying Marinoan glacial deposits. Previous studies suggested that all carbon isotope data of Neoproterozoic marine carbonates were modified by diagenetic processes (Knauth and Kennedy, 2009). However, it is unlikely that diagenetic alteration played a significant role in the δ13C compositions (mostly between −4‰ and −2‰) for the host dolomite in the Doushantuo cap dolostone succession at the Jiulongwan section, as discussed below.

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Fig. 4. Lithostratigraphic column (A), changes in calcite (B) and dolomite contents (wt.%) (C), and carbon isotopic profile (bulk and micro-drilled samples) (D) of the Doushantuo cap dolosonte succession at the Jiulongwan section.

Methane-rich fluids may alter primary δ13C compositions of carbonate rocks, due to the incorporation of 13C-depleted carbon from methane oxidation into the host rocks. In this study, however, differences of up to 20–38‰ in δ13C between the host dolomite and the methane-derived calcite on a centimeter scale (Table 2 and Fig. 2) suggest that the isotopic compositions of the host dolomite were not significantly altered by methane-enriched fluids. In addition, most of the dolomite samples have preserved original textures (laminations, dolomicrite), implying that they were not subjected to extensive alteration by the fluids, although local dissolution and/or replacement of dolomite by methane-derived calcite (Fig. 3) does indicate intense but very limited alteration of the host dolomite. Furthermore, there is no obvious covariation between δ13C and δ18O from the methane-derived calcite and the host dolomite (Fig. 6), arguing against significant modification of the isotopic signatures of the host dolomite. Consequently, we

consider that δ 13C compositions of the host dolomite are relatively immune to this type of isotopic alteration. In addition, diagenetic fluids (such as meteoric fluids, burial fluids) or dolomitization (if the dolomite was not primary) also did not severely alter the δ 13C compositions of the host dolomite in the Doushantuo cap dolostone succession, because: 1) the negative δ 13C values of the host dolomite are reproducible between all sections in the Yangtze Gorges area (Jiang et al., 2003a; Wang et al., 2008; this study), and are similar to data from other Marinoan cap dolostones around the world (e.g. Kennedy, 1996; Hoffman et al., 1998; James et al., 2001; Shen et al., 2008), 2) the original petrographic textures are usually preserved, as discussed above. In general, δ18O values more negative than −10‰ are used as an empirical index for alteration of Proterozoic carbonates by diagenetic fluids (e.g. Jacobsen and Kaufman, 1999). However, a majority of dolomite samples we analyzed have δ18O values greater than −10‰ (Table 2). Furthermore, a crossplot of δ 13C and δ18O from the host dolomite and the later-stage calcite (δ 18O values from −15.1‰ to −11.9‰) filling in Type II veins, shows little evidence for a linear correlation between δ13C and δ18O (R2 = 0.38, n = 37) (Fig. 6). In summary, our analyses suggest that the δ 13C compositions of the host dolomite were not significantly reset by diagenetic processes, and thus likely represent primary or near-primary isotopic signatures. 6.3. Extremely negative carbon isotope values: evidence for the hypothesis of methane hydrate destabilization?

Fig. 5. Cross plot of δ13C (bulk samples) and calcite contents (wt.%) for the Doushantuo cap dolostone succession. There is a moderate negative correlation between δ13C and calcite contents, suggesting calcite may be isotopically lighter than dolomite.

It is also a concern whether extremely negative carbon isotope values indicative of methane oxidation are related to a methane hydrate destabilization event proposed by Kennedy et al. (2001). In fact, some workers claimed that extreme δ 13C values, obtained from limestone clots, lenses and isopachous cements in the Doushantuo cap dolostone in the Yangtze Gorges (Fig. 1C), provide direct support to the inferred methane hypothesis (Jiang et al., 2003a, 2006; Wang et al., 2008). However, their suggestion remains some uncertainties

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Table 2 Isotopic compositions of micro-drilled samples from the Doushantuo cap dolostone succession at the Jiulongwan section. Sample number JLW-04-a JLW-04-b JLW-04-c JLW-007-a JLW-007-b JLW-007-c JLW-007-d JLW-07-a JLW-07-b JLW-07-c JLW-08-a JLW-08-b JLW-08-c JLW-08-d JLW-08-e JLW-32-a JLW-32-b JLW-32-c JLW-38-a JLW-38-b JLW-38-c JLW-39-a JLW-39-b JLW-39-c JLW-39-d JLW-40-a JLW-40-b JLW-40-c JLW-40-d JLW-41-a JLW-41-b JLW-41-c JLW-41-d JLW-41-e JLW-41-f JLW-41-g JLW-41-h JLW-42-a JLW-42-b JLW-42-c JLW-42-d JLW-43-a JLW-43-b JLW-43-c JLW-44-a JLW-44-b JLW-44-c JLW-44-d JLW-44-e JLW-44-f JLW-44-g JLW-45-a JLW-45-b JLW-45-c JLW-45-d JLW-46-a JLW-46-b JLW-46-c JLW-47-a JLW-47-b JLW-47-c JLW-47-d JLW-47-e JLW-48-a JLW-48-b JLW-48-c JLW-48-d JLW-49-a JLW-49-b JLW-49-c JLW-49-d JLW-49-e JLW-49-f

Height

Lithotype

(cm) 18.5 23.0 25.0 47.0 49.5 49.5 53.0 52.5 53.0 53.5 55.3 55.5 55.5 56.5 59.5 207.5 207.5 207.5 251.5 252.0 252.4 259.5 260.3 260.7 260.8 263.0 263.5 264.5 265.0 266.0 267.0 267.5 269.5 270.5 271.5 273.0 274.5 277.0 278.0 279.0 280.0 282.0 283.5 284.5 286.2 286.4 286.5 286.7 287.7 289.0 290.0 295.5 296.0 298.0 300.0 304.5 306.5 306.9 310.0 310.5 311.5 314.5 316.0 322.5 322.5 323.3 323.7 328.5 328.8 329.0 330.5 331.0 332.0

Dolomite Dolomite Dolomite Mixture Calcite (II) Dolomite Calcite (I) Mixture Mixture Mixture Dolomite Dolomite Dolomite Dolomite Calcite (II) Calcite (II) Calcite (II) Dolomite Dolomite Dolomite Dolomite Mixture Dolomite Dolomite Dolomite Mixture Calcite (II) Mixture Mixture Mixture Mixture Dolomite Mixture Calcite (I) Dolomite Calcite (I) Calcite (I) Calcite (II) Dolomite Mixture Calcite (II) Calcite (I) Mixture Dolomite Calcite (I) Mixture Mixture Dolomite Dolomite Mixture Mixture Mixture Calcite (I) Dolomite Dolomite Mixture Mixture Dolomite Mixture Mixture Mixture Mixture Dolomite Calcite (I) Mixture Mixture Dolomite Dolomite Dolomite Dolomite Dolomite Mixture Mixture

δ13C

δ18O

(‰, VPDB)

(‰, VPDB)

− 3.3 − 5.1 − 3.9 − 19.1 − 1.3 − 0.3 − 42.5 − 4.5 − 4.6 − 5.2 − 4.6 − 3.5 − 4.2 − 3.8 − 3.8 − 4.6 − 4.3 − 3.0 − 4.2 − 3.7 − 3.2 − 7.1 − 4.8 − 6.1 − 6.2 − 6.1 − 5.2 − 20.6 − 11.9 − 7.8 − 10.1 − 0.7 − 6.6 − 27.3 − 2.3 − 26.2 − 26.5 − 5.0 − 3.0 − 14.0 − 5.5 − 42.0 − 10.5 − 3.4 − 40.4 − 8.9 − 8.0 − 1.6 − 3.9 − 7.2 − 10.2 − 17.5 − 43.6 − 0.9 − 1.5 − 21.4 − 11.3 − 2.3 − 8.4 − 19.0 − 17.1 − 16.1 − 2.2 − 44.4 − 6.4 − 28.3 − 2.8 − 2.8 − 0.3 − 4.3 1.4 − 5.8 − 10.9

− 12.5 − 7.8 − 8.6 − 9.5 − 13.8 − 1.1 − 7.6 − 7.7 − 7.5 − 7.6 − 11.2 − 9.2 − 9.0 − 12.9 − 15.1 − 12.6 − 14.0 − 10.5 − 11.8 − 8.9 − 8.5 − 5.8 − 12.3 − 9.8 − 11.5 − 5.5 − 13.3 − 8.9 − 11.0 − 8.7 − 9.7 − 6.0 − 11.9 − 6.1 − 8.0 − 6.7 − 9.3 − 14.7 − 7.7 − 11.0 − 11.9 − 9.2 − 11.9 − 10.8 − 6.6 − 8.1 − 7.6 − 2.8 − 6.3 − 4.0 − 5.6 − 6.4 − 8.4 − 8.5 − 9.7 − 9.5 − 8.6 − 2.8 − 4.0 − 7.2 − 10.2 − 6.6 − 4.5 − 8.7 − 6.2 − 9.0 − 3.3 − 8.4 − 6.8 − 9.8 − 2.3 − 10.1 − 9.2

Fig. 6. Cross plot of δ13C and δ18O of micro-drilled samples from the Doushantuo cap dolostone succession.

because of the ambiguity of the diagenetic sequence of carbonate minerals with diverse carbon isotopic compositions in the cap dolostone succession. In this regard, our petrographic data from the cap dolostone succession (Jiulongwan section) unequivocally demonstrate that the calcite bearing methane-oxidation signals formed well after the host dolomite, indicating that methane-oxidation processes represented here occurred later than the deposition of the cap dolostone. Our inference is also supported by evidence from geochemical studies by Bristow et al. (2011). Accordingly, we consider that the presence of extremely negative δ13C values in the cap dolostone could not be taken as evidence to support the hypothesis of the methane hydrate destabilization. It also needs to be pointed out that, if it were true that the extremely negative δ13C signals resulted from thermogenic methane oxidation at depth (Bristow et al., 2011), it would not only refute the suggestion by Jiang et al. (2003a, 2006) and Wang et al. (2008), but also exclude the only known occurrence of a cold-seep carbonate derived from anaerobic methane oxidation in cap dolostones. Nevertheless, it is noteworthy that our proposal is not in conflict with the scenario that the primary or near-primary carbon isotopic signatures recorded well in the Doushantuo cap dolostone and others around the world could result from permafrost methane hydrate destabilization during deglaciation (Kaufman et al., 1997; Kennedy et al., 2001; 2008), much prior to methane oxidation described here. In other words, the presence of extremely negative δ13C values later than the Doushantuo cap dolostone does not witness a methane hydrate destabilization event, but at the same time, it does not negate possible contribution of methane hydrates to the global negative δ13C excursions of cap dolostones. 7. Conclusions By petrographic and carbon isotopic analyses of the Doushantuo cap dolostone succession at the Jiulongwan section in the Yangtze Gorges, we found that extreme δ13C values as low as −44‰ considered to be indicative of methane oxidation exclusively occur in calcite, whereas δ 13C values for the host dolomite range mostly from −4‰ to −2‰; we also observed the dolomite dissolution and/or replacement by the calcite bearing signals of methane oxidation. This finding implies that methane-related activity reported here did not occur until after the

Notes to Table 2: Footnote: Calcite (I): calcite within veins of Type I; Calcite (II): calcite within veins of Type II; Mixture: a mixture of dolomite and calcite; Polished slab JLW-007 is an additional sample from the Doushantuo cap dolostone succession at the Jiulongwan section.

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