Raman geothermometry of carbonaceous material in the basal Ediacaran Doushantuo cap dolostone: The thermal history of extremely negative δ13C signatures in the aftermath of the terminal Cryogenian snowball Earth glaciation

Accepted Manuscript Raman geothermometry of carbonaceous material in the basal Ediacaran Doushantuo cap dolostone: the thermal history of extremely negative δ 13C signatures in the aftermath of the terminal Cryogenian snowball Earth glaciation Zhou Wang, Jiasheng Wang, Yui Kouketsu, Robert J. Bodnar, Benjamin C. Gill, Shuhai Xiao PII: DOI: Reference:

S0301-9268(17)30073-6 http://dx.doi.org/10.1016/j.precamres.2017.06.013 PRECAM 4795

To appear in:

Precambrian Research

Received Date: Revised Date: Accepted Date:

11 February 2017 5 June 2017 12 June 2017

Please cite this article as: Z. Wang, J. Wang, Y. Kouketsu, R.J. Bodnar, B.C. Gill, S. Xiao, Raman geothermometry of carbonaceous material in the basal Ediacaran Doushantuo cap dolostone: the thermal history of extremely negative δ 13C signatures in the aftermath of the terminal Cryogenian snowball Earth glaciation, Precambrian Research (2017), doi: http://dx.doi.org/10.1016/j.precamres.2017.06.013

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Raman geothermometry of carbonaceous material in the basal Ediacaran Doushantuo cap dolostone: the thermal history of extremely negative δ13C signatures in the aftermath of the terminal Cryogenian snowball Earth glaciation

Zhou Wang a,b, Jiasheng Wang a,*, Yui Kouketsu c, Robert J. Bodnar b, Benjamin C. Gill b, Shuhai Xiao b,* a

State Key Laboratory of Biogeology and Environmental Geology, School of Earth Sciences, China University of

Geosciences, Wuhan 430074, P. R. China b c

Department of Geosciences, Virginia Tech, Blacksburg, VA 24061, USA Department of Earth and Planetary Sciences, Nagoya University, Nagoya, 4648602, Japan

* Corresponding authors: Jiasheng Wang ([email protected]) and Shuhai Xiao ([email protected])

ABSTRACT Extremely negative δ13C values (as low as –48‰) of blocky calcite cements filling cavities, sheet-cracks, and fractures in the basal Ediacaran cap dolostone of the Doushantuo Formation in South China have been considered as evidence for methane release possibly associated with gas hydrate destabilization in the aftermath of the terminal Cryogenian snowball Earth glaciation. Alternatively, these extremely negative δ13C values may have resulted from oxidation of thermogenic methane derived from localized hydrothermal activity long after the deposition of the Doushantuo cap dolostone. The latter interpretation was based on carbonate clumped isotope thermometry suggesting that the precipitation of extremely

13

C-depleted calcite occurred at temperatures up to 1

476 °C (mean=378 °C, n=4) whereas carbonate fabrics with moderately negative δ 13C values (generally > –10‰) were formed at much lower temperatures (mean=112 °C, n=4). To assess these competing hypotheses, we systematically investigated the petrography, carbonate carbon and oxygen isotopes, and carbonaceous material (CM) Raman spectroscopy of the Doushantuo cap carbonate. Raman geothermometry suggests that CM associated with different microfabrics and δ13C signatures—including sedimentary fabrics with moderately negative δ13C values (e.g., peloids, stromatolite laminae, microbial mat fragments), diagenetic fabrics with moderately negative δ 13C values (e.g., some void-filling or crack-filling cements), and diagenetic fabrics with extremely negative

δ13C

values

(e.g.,

void-filling

or

crack-filling

blocky and

bladed

calcites)—experienced temperatures generally <300 °C (largely 200–300 °C). Contrary to temperature estimates based on carbonate clumped isotopes, CM Raman geothermometry reveals that different microfabrics show largely overlapping paleotemperatures and that there is no apparent relationship between δ13C signatures and paleotemperatures. Whereas the origins of these extremely negative δ13C values remain unresolved, the Raman geothermometric data challenge the previous interpretation that they were derived from thermogenic methane related to localized hydrothermal activity at temperatures up to almost 500 °C that only affected the thermal history of the calcite with extremely negative δ13C values. Instead, the new data suggest that the analyzed microfabrics experienced the same peak temperatures <300 °C, probably related to the burial history of the host rocks or pervasive hydrothermal activity that affected all microfabrics in the 2

Doushantuo cap dolostone.

Keywords: Raman geothermometry; carbonaceous material; methane; Ediacaran System; Doushantuo cap dolostone; South China

1. Introduction

The terminal Cryogenian snowball Earth glaciation had a catastrophic ending: virtually every succession of terminal Cryogenian glacial diamictite is overlain by a cap dolostone with negative δ13C values (around –5‰) and enigmatic sedimentary-diagenetic structures such as seafloor precipitates, tubestones, sheet-cracks, and tepee-like structures (Hoffman et al., 1998; Jiang et al., 2006). Various hypotheses have been proposed to explain the carbon isotopic signatures found in the cap dolostone. Leading hypotheses include (1) rapid chemical weathering coupled with subdued bioproductivity in a post-glacial Earth with high pCO2 levels (Hoffman et al., 1998; Higgins and Schrag, 2003; Huang et al., 2016), (2) oceanic upwelling of

12

C-enriched alkalinity from the deep ocean to shallow

carbonate platforms (Grotzinger and Knoll, 1995; James et al. 2001), and (3) post-glacial destabilization of gas hydrates in permafrost (Kennedy et al., 2001) or marine methane seepages (Jiang et al., 2003; Wang et al., 2008). The highly negative carbon isotopic compositions and a broad range of δ18O values from –25 ‰ to 12 ‰ in the cap dolostone (Kennedy et al., 2008) have been interpreted as evidence for mixing between 3

18

O-enriched clathrate-derived fluids and

18

O-depleted glacial meltwaters as would be

expected during the decomposition of permafrost methane clathrates (Matsumoto, 1989; Bowen and Wilkinson, 2002; Kennedy et al., 2008). The most convincing evidence for a methane contribution to the precipitation of the cap dolostone, however, comes from the basal Ediacaran Doushantuo cap dolostone where microfabrics such as blocky calcite spars, clotted dolomicrite, and peloids have been reported to have extremely negative δ13C values as low as –48‰ (Jiang et al., 2003, 2006; Wang et al., 2008). The precipitation of these microfabrics was likely influenced by dissolved inorganic carbon (DIC) derived from anaerobic oxidation of methane (AOM) that likely occurred within the sediment. An important premise of the gas hydrate hypothesis is that methane release and AOM occurred during sedimentation or authigenesis of the cap dolostone, contributing to the negative δ13C values of the bulk cap dolostone (typically around –5‰). Thus, vigorous methane release, AOM, and precipitation of extremely

13

C-depleted

microfabrics must have occurred at low temperatures in sedimentary environments, although these microfabrics could later experience burial, hydrothermal, and metamorphic temperatures. This prediction can be tested against geothermometric data. Bristow et al. (2011) provided a test of the gas hydrate hypothesis using clumped isotope geothermometry. These authors analyzed the carbonate clumped isotope composition (Δ47) of and calculated the precipitation temperatures for different microfabrics of the Doushantuo cap dolostone in the Yangtze Gorges area, including dolomicrite with moderately negative δ13C values (around –5‰) and void-filling gray 4

calcite with extremely negative δ13C values (–26‰ to –41‰). The dolomicritic matrix yielded lower temperatures (112 ± 31 °C, n = 4), whereas extremely 13C-depleted calcite cements yielded Δ47 values indicating much higher temperatures (378 ± 91 °C, n = 4; up to 476 °C). The authors thus concluded that extremely 13C-depleted calcite cements were influenced by oxidation of thermogenic methane related to brief and localized hydrothermal activity, which could have occurred in the Cambrian or millions of years after the deposition of the Doushantuo cap dolostone. However, clumped isotope geothermometry can sometimes give spurious results, particularly when disequilibrium behavior occurs due to kinetic isotopic effects from biologically-influenced processes (Yeung et al., 2015; Wang et al., 2015). Likewise, clumped isotope geothermometry could overestimate the formation temperature of biogenic methane and the precipitation temperature of methane-derived carbonate, due to the isotopic disequilibrium produced during AOM coupled to microbial sulfate reduction (Wang et al., 2015; Loyd et al., 2016). As sulfate-driven AOM likely played an important role in the precipitation of extremely

13

C-depleted calcite in the Doushantuo cap

dolostone (Peng et al., 2015), there are reasons to be skeptical about the Δ47 geothermometry of the extremely

13

C-depleted calcite. Furthermore, kinetic effects

related to solid-state C-O bond reordering limits the application of clumped isotope analysis of carbonates for rocks with deep and long burial histories (Henkes et al., 2014). Therefore, it is important to further test the results of Bristow et al. (2011) using an independent geothermometer in order to better constrain the thermal history of the 5

Doushantuo cap dolostone. Two recent studies add further insight into the debate over the origin of the extremely negative δ13C values in the Doushantuo cap dolostone. Zhou et al. (2016) reported early authigenic calcite with extremely negative δ 13C values from a 20-cm-thick horizon immediately above the cap dolostone. This horizon was regarded as the top unit of the cap dolostone (i.e., unit C3 of Jiang et al., 2003, 2006) but has since been removed from the cap dolostone, therefore Zhou et al.’s (2016) study confirms previous report of methane signal in C3 (Wang et al., 2008). Zhou et al. (2016) argue that this horizon is genetically related to extremely negative δ13C values previously reported from the cap dolostone. They also argue that the extremely negative δ13C values in both units had a low-temperature, albeit post-depositional, origin related to AOM. On the other hand, Zhou et al. (2017) reported fluid inclusions in quartz crystals filling sheet-cracks in the Doushantuo cap dolostone, and calculated homogenization temperatures of 160–220 ºC (mean, 192 ºC, n = 31); these quartz crystals demonstrably predate calcite spars with extremely negative δ13C values; the latter fabrics often fill the residual cavity in sheet-cracks and represent the final stage of void-filling minerals. Despite the fact that these homogenization temperatures are much lower than those derived from the Δ47 data (378 ± 91°C, Bristow et al., 2011), Zhou et al. (2017) nonetheless concluded that the Doushantuo cap dolostone experienced hydrothermal temperatures compatible with the Δ47 data. Given these ongoing debates, it is imperative to explore independent methods to constrain the thermal history of the Doushantuo cap dolostone to test previous 6

paleotemperature estimates. In this study, we analyzed the carbonaceous material (CM) in the Doushantuo cap dolostone using Raman spectroscopy in order to provide an independent estimate of the thermal history of this unit. Carbonaceous material is ubiquitous in sediments and, with increasing temperature, can be irreversibly transformed from disordered organic matter to partially crystallized organic matter and eventually to fully crystallized graphite. The degree of CM crystallinity can be quantified with Raman spectroscopy. Thus, Raman spectroscopic analysis of CM offers a geothermometer to estimate the peak burial or metamorphic temperature of the host sediments (Yui et al., 1996; Beyssac et al., 2002a, b; 2003a, b; Schiffbauer et al., 2007; Buseck and Beyssac, 2014; Kouketsu et al., 2014). Importantly, because it has been shown experimentally that Raman spectra of CM from Precambrian organic-walled microfossils are affected even during heating for less than a year (Schiffbauer et al., 2012), the CM Raman geothermometer could also be used to detect localized hydrothermal events that lasted a relatively short interval of geological time. If, as Bristow et al. (2011) argued, extremely 13C-depleted calcite in the Doushantuo cap dolostone was precipitated from hydrothermal fluids at nearly 500 °C, then CM Raman geothermometry of the extremely 13C-depleted calcite should record temperatures significantly higher than those of host rock, considering that the host rock has not experienced background burial temperatures greater than 300 °C (Loyd et al., 2015) and the overlying units of the Doushantuo Formation contain exceptionally preserved microfossils (Xiao, 2004, Xiao et al., 2014; Liu et al., 2014). Thus, CM Raman 7

geothermometry offers a potential test of the hydrothermal hypothesis that was based on clumped isotope and fluid inclusion geothermometries (Bristow et al., 2011; Zhou et al., 2017).

2. Materials and methods

Samples of the basal Ediacaran Doushantuo cap dolostone were collected from the Jiulongwan (30°49′14″N, 111°05′14″E) and Wangzishi (30°32′46″N, 111°10′04″E) sections in the Yangtze Gorges and its adjacent area, Hubei Province of South China (Fig. 1A–B). Both sections were paleogeographically located in the intrashelf lagoon in the Yangtze Block (Jiang et al., 2011). The Doushantuo cap dolostone is geochronologically constrained at 635.26 ± 1.07 Ma (Fig. 1C, U-Pb zircon age from an ash bed immediately above the cap dolostone; Schmitz, 2012). This age is similar to a U-Pb zircon age of 635.21 ± 0.59 Ma from an ash bed within the Marinoan-age diamictite in Namibia, which is believed to be equivalent to the Nantuo diamictite underlying the Doushantuo cap dolostone (Prave et al., 2016). Given that these two ages are indistinguishable within analytical error, the deposition of the Doushantuo cap dolostone has been suggested to occur over a rather short geological time interval likely less than one million years. The Doushantuo cap dolostone in South China was previously divided into three units (Jiang et al., 2003, 2006): in ascending stratigraphic order, C1 consisting of disrupted massive dolostone with abundant sheet-cracks and irregular cavities filled with 8

chalcedony and calcite (Xiao et al., 2012); C2 characterized by laminated dolostone with microbial fabrics, tepee-like structures, and sheet-cracks; and C3 consisting of thin-bedded marly dolostone and silty limestone. Subsequent studies (e.g., Zhou et al., 2016, 2017) restricted the Doushantuo cap dolostone to include only C1 and C2, because C3 is more akin to the overlying thin-bedded strata whereas C1 + C2 are distinct owing to their strong resistance to weathering. This redefinition has resulted in some confusion. For example, the 20-cm-thick horizon of early authigenic calcite at the Jiulongwan section in the Yangtze Gorges area with extremely negative δ13C values was said to occur stratigraphically immediately above the cap dolostone (Zhou et al., 2016). However, this horizon is actually in the basal C3 that has been previously known to contain intervals of extremely negative δ13C values at the Jiulongwan and Wangzishi sections (Wang et al., 2008). To avoid further confusion, we choose to adopt the definition of Zhou et al. (2016, 2017) and exclude C3 from the Doushantuo cap dolostone. In so doing, the Doushantuo cap dolostone is comprised of a lower and an upper unit (Fig. 1C), equivalent to C1 and C2 of Jiang et al. (2003, 2006). The lower unit of Doushantuo cap dolostone in the study area is characterized by massive dolomicrite containing abundant breccias, cavities (e.g. stromatactis), sheet-cracks, irregular fractures, and tepee-like structures (Jiang et al., 2006). In contrast, the upper unit is dominated by laminated dolomicrite with microbial fabrics and structures (e.g. stromatolites). Multiple generations of void-filling and crack-filling calcite cements occur in both the lower and upper units, and some of these cements are 9

characterized by extremely negative δ13C values indicative of methane (Jiang et al., 2003, 2006; Wang et al., 2008; Zhou et al., 2010; Lin et al., 2011; Zhou et al., 2016). Samples were split to make thin sections and polished slabs. Thin sections were examined under a transmitted light microscope for petrographic analysis. Guided by petrographic analysis, powders were microdrilled from specific microfabrics (e.g., dolomicritic matrix, stromatolite, blocky calcite spar, isopachous cement) on corresponding polished slabs. These powders were used for carbonate carbon and oxygen isotope analysis, which was carried out on an Isoprime 100 isotope ratio mass spectrometer interfaced with a MultiFlow-Geo headspace sampler in the Department of Geosciences at Virginia Tech. About ~ 350 µg of carbonate powder was allowed to react with 80% ortho-phosphoric acid (H3PO4), and CO2 evolved from this reaction was collected for isotopic analysis. Several standards, including IAEA CO-1 (CaCO3 marble), IAEA CO-9 (BaCO3), and NBS 18 (CaCO3 calcite), were used for quality control and calibration. Isotopic data are reported as δ values relative to the Vienna Pee Dee Belemnite (VPDB). The uncertainties based on repeated standard measurements are better than 0.10‰ (1σ) for carbon isotope and 0.15‰ (1σ) for oxygen isotope analyses. Raman spectra of CM were obtained in the Department of Geosciences at Virginia Tech, using a high-resolution 800 mm focal length spectrometer (JY Horiba LabRam HR800) with a 514 nm laser source with maximum power of 50 mW at the source. Laser power at the sample was typically ca. 10 mW and could be reduced further using a series of filters placed in the laser path. The scattered Raman light was collected through a 400 10

μm pinhole and dispersed with a 600 gr/mm grating. The detector consisted of an electronically cooled charge-coupled device (CCD) of 1024×256 pixels (Andor Technology). CM targets in standard petrological thin-sections were positioned at the location of the laser spot for Raman analysis using a microscope (Olympus BX-41) equipped with ×40 (Olympus SLCPlanFl 40×, NA = 0.55) and ×100 objectives (Olympus MSPlan 100×, NA = 0.95). The laser beam was focused on CM that was beneath the exposed surface of the petrographic thin sections, in order to minimize artifacts related to mechanical damage and structural re-organization of CM during thin section preparation (Beyssac et al., 2003b). Acquisition of Raman spectra was conducted with the software Labspec5.0 and acquisition time for each run was confined to less than 45 seconds to limit laser-induced heating. Spectral corrections for fluorescence background were achieved by subtracting a linear baseline in the spectral range of 1000–1750 cm–1 (Kouketsu et al., 2014). Reference spectra from the Crystal Sleuth Raman library (http://rruff.info/) were used for mineralogical identification. Several parameters, including peak center position, peak intensity ratio, peak area ratio, and peak full width at half maximum (FWHM hereinafter) (Fig. 2A), were acquired after peak decomposition and fitting of baseline-corrected CM Raman spectra. A four-band peak fitting procedure as summarized in Kouketsu et al (2014) was followed to decompose each baseline-corrected spectrum into four CM bands, including the D1-band (~ 1350 cm–1), D2-band (~ 1600 cm–1), D3-band (fixed at 1510 cm–1), and D4-band (fixed at 1245 cm–1). Each band was fitted as a Gaussian-Lorentzian Sum function 11

(Voigt-like lineshape) (Fig. 2B; Beyssac et al., 2003b). The software PeakFit 4.12 was used in peak decomposition and fitting. Estimated paleotemperatures were calculated using FWHM-based CM Raman geothermometers (as Eq. 1 and Eq. 2) originally calibrated with linear regressions using CM from Paleozoic and Mesozoic metasediments (Kouketsu et al., 2014).

In equations (1) and (2), FWHM-D1 and FWHM-D2 are FWHM of D1-band and D2-band, respectively. The calibrated Raman geothermometers given by Eq. 1 and Eq. 2 are applicable to temperatures between 150 °C and 400 °C (Kouketsu et al., 2014).

3. Results

3.1. Petrography and stable isotopic signatures The Doushantuo cap dolostone contains CM preserved in sedimentary, authigenic (syndepositional to early diagenetic mineralization within uncompacted sediments), and late diagenetic microfabrics (Figs. 3–5). Primary sedimentary fabrics include dolomicritic matrix, stromatolites, microbial laminae, and peloids. Dolomicrite forms the matrix of the cap dolostone and contains diffuse CM currently filling intercrystalline space between microspar (Fig. 3A–C). Stromatolites occur in the upper unit in the form of laterally linked, dome-shaped, organosedimentary laminae (Fig. 3D–F; Fig. 5A), which represent 12

vertically accreted organic-rich microbial mats. CM-rich crinkled laminae that do not occur in domal stromatolitic structures are also interpreted as microbial mats. Fragments of microbial mats are common in micritic matrix and are sometimes cemented by blocky sparitic calcite (Fig. 3J; Fig. 5B). These fragments are also rich in CM and often contain appreciable amounts of authigenic pyrite (Fig. 3J). Finally, as is common in basal Ediacaran cap dolostones (Xiao et al., 2004; Hoffman, 2011), peloids 50–100 μm in diameter (Fig. 3G–H) tend to occur in the dolomicritic matrix and low-lying shoulders of domal stromatolites. The peloids consist of patches of dolomicrite with trapped CM (Fig. 3H–I). Authigenic and diagenetic microfabrics include multiple generations of cements with different mineralogical compositions. Patchy cements (‘PC’ in Fig. 3D) of calcium carbonate, silica, and apatite occur within CM-rich stromatolitic laminae, whereas layers of microsparry calcite cements (‘LC’ in Fig. 3D) occur between and conform to these laminae. Both are interpreted as authigenic cements perhaps filling early voids created by gas bubbles in growing or degrading microbial mats (Bosak et al. 2010). Randomly oriented calcite needles (Fig. 3E–F) are abundant in stromatolites, and their occurrence is often confined between, but can also cross-cut, stromatolitic laminae (Fig. 3E). This cross-cutting relationship establishes the calcite needles as post-sedimentary minerals, probably formed during authigenesis. Diagenetic recrystallization of the dolomicritic matrix forms aggregates of dolomicrospars and segregates CM into space between these dolomicrospars (Fig. 3A–C). Post-sedimentary voids, cavities, sheet-cracks, fractures, 13

and veins are abundant in the Doushantuo cap dolostone. These are filled with CM and multiple generations of isopachous cements and blocky calcite. For example, some lamina-cutting fractures are filled almost entirely with CM possibly derived from late bituminous fluids (Fig. 3K). Isopachous cements consisting of conformable layers of calcite, silica, and CM are found in many sheet-cracks (Fig. 4B and E–F; Fig. 5D) (Xiao et al., 2012). Dark gray-colored blocky and bladed calcite cements tend to occur in the center of these cracks/cavities and are subtended by isopachous cements (Fig. 4), thus postdating the isopachous cements based on cement stratigraphy. Occasionally, bladed calcite fills irregular fractures that cross-cut sheet-cracks (Fig. 4B), establishing their relatively late diagenetic origin. It is these blocky and bladed calcite cements (Fig. 5C–D; Wang et al., 2008), along with some peloids (Jiang et al., 2003), that typically have extremely negative δ13C values, and they also contain a fair amount of CM (Fig. 4A and C–D). Finally, still later diagenetic veins and fractures are filled with light-colored calcite and cross-cut various microfabrics of the cap dolostone, however CM is usually scarce in these late veins. Given our goal to constrain the thermal history of methane-related fabrics using CM Raman geothermometry, we categorized CM into two major groups based on associated microfabrics:

sedimentary

CM

(CM

in

dolomicritic

matrix,

stromatolitic

laminae/microbial mat, and peloid) and methane signal-related CM (CM in blocky calcite and bladed calcite typically characterized by extremely negative δ 13C values, as well as CM in isopachous silica cement in contact with blocky and bladed calcites). 14

Different microfabrics of the Doushantuo cap dolostone show a wide range of δ13C values (Supplementary Data Table S1). Dolomicritic matrix has δ13C values from –0.7 to –10.4‰ (Fig. 5). Patchy and layered cements within stromatolitic laminae as well as calcite needles in stromatolite yield a relatively narrow range of δ13C values from –2.3 to –4.0‰ (Fig. 5A). Calcite spars cementing microbial mat fragments and filling blocky voids of stromatolite have more negative δ13C values from –2.3 to –13.7‰ (Fig. 5A–B), and isopachous calcite cements also have negative δ13C values from –2.5 to –10.2‰ (Fig. 5D). While it is possible that the most negative δ13C values in these fabrics may reflect the mixing with microsparry calcite cements characterized by methane signals (e.g., Cui et al., 2017), none of these microfabrics carry a diagnostic and definitive δ13C signature of methane. δ18O values of these microfabrics range from –7.2‰ to –15.2‰. In sharp contrast, dark gray-colored blocky and bladed calcite cements are characterized by profoundly different δ13C values, generally less than –20‰ and as low as –48.3‰ (Fig. 5C–D). Such extremely negative δ13C values are diagnostic signals of inorganic carbon derived from the anaerobic oxidation of methane (AOM). δ18O values of the blocky and bladed calcites range from –6.5‰ to –13.5‰, comparable to those of sedimentary microfabrics (Fig. 5E). δ13C and δ18O values show no significant covariation (Fig. 5E).

3.2. Raman spectra of carbonaceous matter Representative raw spectra, baseline-corrected spectra, and the results of peak fitting 15

are presented in Fig. 6. The raw spectra unavoidably include peaks of various minerals (e.g., calcite, dolomite, and quartz) found in close association with CM; these mineral peaks can sometimes overwhelm the CM peaks (e.g., gray lines in Fig. 6, left column). Additionally, almost all raw spectra show various levels of fluorescent background (Fig. 6), which is common in samples containing calcite and/or organic material. However, after baseline correction (which was carried out in the range of 1000–1750 cm–1 because our focus was on carbonaceous material), characteristic CM peaks can be readily identified. These peaks include a broad peak at ~ 1350 cm–1 (identified as the D1-band), a minor shoulder at ~ 1250 cm–1 (interpreted as the D4-band), and a relatively sharp and prominent peak at ~ 1600 cm–1 (interpreted to be the D2-band). The presence of a minor broad peak at ~ 1500 cm–1 (possibly representing the D3-band) can be further recognized after baseline correction (Fig. 6).

3.3. Peak fitting and paleotemperature estimates Kouketsu et al. (2014) summarized different schemes for peak decomposition and fitting of CM Raman spectra that are required to implement CM Raman geothermometers. Especially for low crystallinity CM (amorphous carbon), initial assessment of the GL/D1 peak intensity ratio, where GL (subscript ‘L’ denoting ‘low’) represents the integrated peak at ~ 1600 cm–1 (probably including the largely overlapping D2 and G bands) and D1 represents the D1-band at ~ 1350 cm–1, is required to determine which peak fitting scheme is most appropriate. If the GL/D1 ratio is greater than 1.5, the GL peak is 16

dominated by the D2-band. As such, the low-grade peak fitting scheme is adopted (‘fitting G’ in Kouketsu et al., 2014), and the Raman spectrum can be decomposed into four bands represented by the D1-band at ~ 1345 cm–1, D2-band at ~ 1605 cm–1, D3-band fixed at 1510 cm–1, and D4-band fixed at 1245 cm–1 (Fig. 6). If the GL/D1 ratio is less than 1.5, the GL peak is decomposed into G- and D2-bands (‘fitting F and E’ in Kouketsu et al., 2014). All CM Raman spectra from the Doushantuo cap dolostone are characterized by GL/D1 intensity ratios > 1.5, and are accordingly decomposed into the four D bands as described above (Figs. 6–7). Sedimentary CM and methane signal-related CM show largely overlapping ranges of GL/D1 intensity ratios and FWHM of GL peak (Fig. 7). These spectral peak parameters suggest

preliminary

paleotemperatures from ~ 230 to ~ 270 °C according to the Raman geothermometers calibrated by Kouketsu et al. (2014) (Fig. 7). Paleotemperatures calculated using decomposed peaks and FWHM-based Raman geothermometers described in Eqs. 1 and 2 are denoted as T-D1, and T-D2, respectively (Fig. 8; Supplementary Data Table S2). Several CM Raman spectra were beyond the calibration range of Kouketsu et al. (2014), and they were excluded from the calculations of T-D1 and T-D2; specifically, spectra with fitted peaks FWHM-D1 > 152.6 cm–1 or FWHM-D2 > 56.8 cm–1 were excluded in the calculations of T-D1 and T-D2. Sedimentary CM and methane signal-related CM exhibit comparable paleotemperature ranges, with small differences (< 20 °C) in mean temperatures (compare Fig. 8A vs. Fig. 8C, or Fig. 8B vs. Fig. 8D). Statistical tests show that there is no significant difference in 17

T-D2 (245 ± 33 °C for sedimentary CM, 259 ± 32 °C for methane signal-related CM; Mann-Whitney test, Z = –1.763, p = 0.0783), although T-D1 does show a small but significant difference (215 ± 13 °C for sedimentary CM, 233 ± 18 °C for methane signal-related CM; Mann-Whitney test, Z = –6.087, p < 0.001). Cross-plots of Raman spectral parameters show various degrees of correlation (Fig. 9A–H; Supplementary Data Table S2). For example, FWHM of D2-band shows good correlation with D2 peak position and various peak intensity ratios (Fig. 9E–H). In comparison, the correlation between FWHM of D1-band and other spectral parameters is weaker (Fig. 9A–D). The difference in band behavior might be attributed to different kinetic effects of these two bands in response to changing temperatures. These effects may also be responsible for the systematic differences between T-D1 and T-D2 (ΔT in Fig. 9I–O). Regardless, the estimated peak paleotemperatures of the Doushantuo cap dolostone-hosted CM are mostly < 300 °C based on the Raman geothermometers of Kouketsu et al. (2014) (Figs. 7–8).

4. Discussion

4.1. Thermal transformation and Raman spectroscopy of CM Organic material buried in sediments is subjected to progressive modifications in composition and structure in response to changes in thermal conditions throughout the diagenetic history (Beyssac et al., 2003a; Buseck and Beyssac, 2014). As temperature 18

increases, an irreversible transformation occurs, converting organic compounds to kerogens and eventually to graphite. The progressive transformations can be subdivided into two stages: carbonization and graphitization (Watanabe et al., 2009; Buseck and Beyssac, 2014 and references therein). Carbonization is characterized by the loss of volatile elements, such as nitrogen, hydrogen, and oxygen, and the decomposition of organic matter into liquid/gaseous hydrocarbons and solid residues or kerogens (Vendenbroucke and Largeau, 2007). Subsequent graphitization occurs in the presence of short-lived or persistent heating from regional or contact metamorphism, hydrothermal fluids, or coseismic fault friction (Beyssac et al., 2004; Marshall and Marshall, 2011; Bower et al., 2013; Buseck and Beyssac, 2014; Furuichi et al., 2015). In the graphitization process, temperature plays a crucial role in determining the ultimate structure and crystallinity of the CM, although this process can be accelerated at high pressures (Beyssac et al., 2002b; 2003a). As the crystallinity of CM increases, disordered CM becomes more organized structurally. This transformation has been directly observed by means of high-resolution transmission electron microscopy (HRTEM) on a nanoscale (Beyssac et al., 2002b; Buseck and Beyssac, 2014). The structural transformation of CM in response to temperature increase can also be analyzed using Raman spectroscopy, which can quantitatively assess the degree of CM organization or crystallinity (Yui et al., 1996; Beyssac et al., 2002a; 2003b; Buseck and Beyssac, 2014). In addition, Raman spectroscopy is capable of in situ analysis of CM with minimum sample preparation. Raman spectroscopy and geothermometry of CM has 19

been discussed extensively in the literature (Fig. 7; Beyssac et al., 2002a, b; 2003a, b; Schopf et al., 2005; Wiederkehr et al., 2011; Kouketsu et al., 2014). Low-grade CM with peak temperatures < 300 °C is characterized by Raman spectra with the combination of a broad low-intensity D1-band (due to disordered sp2 carbon network) and a relatively narrow high-intensity D2-band (due to disordered sp2 lattice). The D2-band is at ~ 1600 cm–1 and completely overlaps the G-band (due to in-plane stretching of the C═C bond in an aromatic ring), and these two cannot be distinguished or deconvoluted; thus, the peak at ~ 1600 cm–1 is referred to as the GL peak (Kouketsu et al., 2014). With increasing temperature, the width of both the D1-band and D2-band (or GL peak) becomes systematically narrower (Schopf et al., 2005; Kouketsu et al., 2014) and the peak position of the D2-band gradually shifts toward higher wavenumbers (Fig. 9E; Beyssac et al., 2003b; Schopf et al., 2005; Kouketsu et al., 2014). For medium-to-high-grade CM that has experienced temperatures significantly exceeding 300 °C, the D1-band is characterized by two stages in spectral evolution (Schopf et al., 2005; Kouketsu et al., 2014). In the first stage, the D1-band gradually becomes the dominant peak with a decrease in GL/D1 intensity ratio (Fig. 7). In the meantime, the contribution of the G-band to the GL peak becomes more significant. Then, the D1-band exhibits a progressive weakening and is overwhelmed by the G-band with further increase in temperature. As the end state of the thermal evolution, well-crystallized graphite stabilizes at ~ 650 °C with a prominent and unique peak represented by the G-band. During this thermal maturation, the D3-band (amorphous 20

carbon) and D4-band (sp3 bonds or C-C and C═C stretching vibrations of polyene-like structures) continue to recede with increasing temperature until they become negligible in high-grade CM. Multiple CM Raman geothermometers have been empirically calibrated on the basis of linear or non-linear correlations between certain spectral band parameter(s) and independently estimated temperatures. These include CM Raman geothermometers based on FWHM-D1 or FWHM-D2 (Eqs. 1–2; Kouketsu et al., 2014), peak area ratios (Beyssac et al., 2002a; Aoya et al., 2010; Lahfid et al., 2010), and combined peak area and intensity ratios (Rahl et al., 2005). Among these, the FWHM-D1 and FWHM-D2 geothermometers (Eqs. 1–2; Kouketsu et al., 2014) are the most appropriate for CM in the Doushantuo cap dolostone because the calibrated ranges of CM Raman spectral features match our data very well (Figs. 6–7). Furthermore, the geothermometers of Kouketsu et al. (2014) were calibrated using CM hosted in Paleozoic–Mesozoic metasediments, which are more relevant to CM in the Doushantuo cap dolostone than CM in meteorites or in younger Cenozoic rocks (e.g., Lahfid et al., 2010). Thus, we choose to use the FWHM-based geothermometers of Kouketsu et al. (2014) to analyze CM in the Doushantuo cap dolostone. To

explore the relationships

between FWHM-D1,

FWHM-D2,

and other

paleotemperature proxies, such as D1 peak position, D2 peak position, and various D-band intensity ratios (Beyssac et al., 2002a; Beyssac et al., 2003b; Schopf et al., 2005; Kouketsu et al., 2014), a series of cross-plots are presented in Fig. 9 to show the 21

correlations of these parameters. These cross-plots show that, as FWHM-D1 and FWHM-D2 decrease (i.e., as temperature increases according to Eqs. 1–2), D1 peak position shifts to lower wavenumbers (Fig. 9A), D2 peak position shifts to higher wavenumbers (Fig. 9E), and D2/D1, D1/D3, D1/D4, D2/D3, and D2/D4 intensity ratios increase (Fig. 9B–D, F–H). Thus, there are generally good correlations among these paleotemperature proxies. Despite these generally good correlations, there are offsets in temperature estimates using different CM Raman geothermometers. For example, T-D1 and T-D2 show a difference (ΔT = T-D2 – T-D1) on the order of ±50 °C (Fig. 9I–O). Furthermore, there are some systematic correlations between ΔT and various D-band ratios (Fig. 9I–O). These correlations should be further investigated in future studies to explore secondary factors (e.g., intrinsic properties of organic precursors produced by various kinds of microbes; Qu et al., 2015, 2017) that have relatively minor impact on CM structures. Regardless, considering the offsets between T-D1 and T-D2, we estimate that the T-D1 and T-D2 geothermometers have an uncertainty on the order of ±50 °C.

4.2. Thermal history of the Doushantuo cap dolostone This study reveals two important discoveries about the thermal history of the Doushantuo cap dolostone. First, CM in different microfabrics of the Doushantuo cap dolostone has similar Raman features and shares a similar thermal history with largely overlapping peak paleotemperatures. Second, CM Raman geothermometry indicates that CM in the Doushantuo cap dolostone experienced a peak temperature of 200–300 °C 22

assuming that the CM has reached equilibrium with regard to the peak temperature. The CM Raman geothermometric data have several important implications for the thermal history of the Doushantuo cap dolostone. First, while there is overlap between paleotemperatures based on CM Raman and Δ47 data from the Doushantuo cap dolostone, the Raman data are apparently inconsistent with the carbonate Δ 47 data, which indicate divergent temperatures for dolomicritic matrix (112 ± 31 °C) and extremely 13C-depleted calcite cements (378 ± 91 °C; Bristow et al., 2011). Given the potential problems posed by disequilibrium on Δ47, particularly when microbial processes are involved (Wang et al., 2015; Peng et al., 2015; Loyd et al., 2016), there are reasons to question the paleotemperature estimates derived from Δ47 for the Doushantuo cap dolostone, particularly the calcite cements with extremely negative δ13C values. Indeed, fluid inclusions in quartz crystals, which fill sheet-cracks in the Doushantuo cap dolostone and are in close association with extremely

13

C-depleted calcite cements, suggest

temperatures of 160–220 ºC (Zhou et al. 2017). These temperatures based on fluid inclusion geothermometry are consistent with the paleotemperature estimate based on CM Raman geothermometry. The similar CM Raman paleotemperatures of the different microfabrics in the Doushantuo cap dolostone suggest that these microfabrics shared the same thermal history and experienced the same peak temperature, which could be related to either pervasive hydrothermal activity or burial history. The latter possibility is consistent with the cumulative thickness of the Ediacaran to Silurian-Devonian strata in western Hubei 23

Province; marine sedimentation ceased during the Silurian-Devonian Period when this area was uplifted. This thickness varies from ~ 5 km to ~ 10 km (Wang et al., 1993; Bureau of Geology and Mineral Resources of Hubei Province, 1988). Taking into consideration the potential thickness loss to denudation after uplift, we take the maximum estimate of 10 km as an estimated burial depth. Given a mean geothermal gradient of 24.1 ºC/km for the South China area (Yuan et al., 2006), this indicates a burial temperature of ~ 241 ºC. This estimate is supported by thermal maturity analyses of organic matter in Ediacaran-Cambrian strata in the Yangtze Gorges area that show the thermal maturation surpassed the gas window (Wang et al., 1993; Han et al., 2013). Both temperature estimates derived from the thickness of the stratigraphic overburden and thermal maturity of the organic matter are consistent with the peak temperature estimates based on CM Raman geothermometry. Thus, it is possible that both hydrothermal activity at 160–220ºC (Zhou et al., 2017) and the maximum burial depths of ~10 km were the main factors controlling the peak temperatures (~200–300 ºC) that Doushantuo cap dolostone experienced. Thus, on the basis of the largely overlapping and uniform CM Raman temperatures of different microfabrics in the Doushantuo cap dolostone, we regard it unlikely that the extremely

13

C-depleted calcite cements were derived from localized hydrothermal

activity up to nearly 500 ºC, even if the extremely negative δ13C values were derived from thermogenic methane (Bristow et al., 2011). Both the CM Raman data presented here and previously published fluid inclusion data (Zhou et al., 2017) suggest that the 24

Doushantuo cap dolostone experienced a peak temperature of ~200–300 ºC. However, we emphasize that, although our CM Raman data indicate a temperature much lower than 500 ºC, they do not unequivocally resolve the various interpretations about the origin of the methane signatures in the Doushantuo cap dolostone. It remains uncertain whether the extremely negative δ13C values in the Doushantuo cap dolostone were related to syndepositional methane release and oxidation (Jiang et al., 2003; Wang et al., 2008), authigenic carbonate precipitation in association with enhanced anaerobic oxidation of methane (Zhou et al., 2016), or the oxidation of thermogenic methane in association with pervasive hydrothermal activity during the early Cambrian Period (Bristow et al., 2011).

5. Conclusions

Integrated petrographic, carbonate δ13C, and Raman spectroscopic analysis suggests that carbonaceous material (CM) in different microfabrics of the basal Ediacaran Doushantuo cap dolostone, including calcite cements with methane-derived and extremely negative δ13C values, share a similar thermal history with a peak temperature in the range of 200–300 °C. These results are inconsistent with previously published Δ 47 data, which suggest drastically divergent paleotemperatures for the dolomicritic matrix (112 ± 31 °C) and the extremely 13C-depleted calcite cements (378 ± 91 °C) (Bristow et al., 2011). If the CM Raman geothermometric estimates obtained in this study are further confirmed in future studies, then the paleotemperatures of the Doushantuo cap dolostone 25

could reflect pervasive hydrothermal activity at temperatures of <300 °C or a maximum burial depth of about 10 km. It is important to note that these two scenarios are not mutually exclusive. Regardless, the CM Raman data presented here indicate that the methane-derived δ13C values in the Doushantuo cap dolostone were unlikely related to localized hydrothermal activity at temperatures up to nearly 500 °C.

Acknowledgements This research was supported by State Key R&D Project of China (2016YFA0601100), the Integrated Marine Biogeochemistry and Ecosystem Research (IMBER) project, the National Natural Science Foundation of China (41472085, 41272011), the U.S. National Science Foundation (EAR-1528553), NASA Exobiology Program (NNX15AL27G), and China Scholarship Council. We would like to thank Thomas F. Bristow for helpful discussions; Charles Farley for technical assistance with Raman analysis; and Yuangao Qu, and Ganqing Jiang for constructive reviews.

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Captions Fig. 1. Maps and stratigraphic column showing sampling locality and horizon. (A) Ediacaran paleogeography of the Yangtze Block (revised from Jiang et al., 2011). The Yangtze Gorges area is marked with a rectangle. (B) Simplified geological map of the Yangtze Gorges area and its vicinity (revised from Wang et al., 2008), showing the locations of the Jiulongwan section and Wangzishi section. (C) Generalized lithostratigraphic column of the Doushantuo cap dolostone in study area, with sample numbers and horizons (prefix JLW- refers to samples from the Jiulongwan section and prefix WZS- samples from the Wangzishi section). Lithostratigraphic thickness is based on measurements at the Jiulongwan section. Zircon U-Pb age from Schmitz (2012).

35

Fig. 2. Schematic diagram showing (A) peak center position, peak intensity, and peak full width at half maximum (FWHM); and (B) the shape of a Voigt profile (Gaussian-Lorentzian Sum; dashed line) with the limiting cases of Gaussian and Lorentzian profiles.

Fig. 3. Transmitted plane-polarized light photomicrographs of petrographic thin sections showing microtextures and their Raman spectra. (A–C) CM (arrows) in intercrystalline pore space of the massive dolomicrite/dolomicrosparite, # JLW-2/2-A, # JLW-2/2-B, and # JLW-2/2-F, respectively. (D) Stromatolitic laminae with intercalated calcite cement layers (‘LC’) and authigenic patchy calcite, silica, and phosphatic cements (‘PC’), # JLW-2/2-D. (E) Randomly oriented needles of authigenic calcite that cross-cut microbial laminae (‘ML’), # JLW-2/2-D. (F) Magnified view of calcite needles in E. (G) CM-rich dolomitic peloidal aggregates (‘P’) in close association with fragments of microbial mats (‘FM’), # JLW-2/2-D. (H) Magnification of a peloid in G (red arrow). (I) Raman spectra of patchy cement (cross in D), needle-like calcite crystal (cross in F), dolomitic peloid (cross in H), and fine-grained pyrite (white dotted line in J), along with reference spectra of calcite, dolomite, quartz, apatite, and pyrite from the Crystal Sleuth database (http://rruff.info/). (J) Fragments of microbial mats (‘FM’, rich in organic matter, and perhaps pyrite) surrounded by blocky calcite spars (‘BCS’), # JLW-2/2-G. Note ghost structure (outlined by white dotted line) of mat debris with fine-grained authigenic pyrite. (K) CM filling irregular late-diagenetic cracks, # JLW-2/2-D. 36

Fig.

4.

Transmitted

non-polarized

(A–B)

and

plane-polarized

(C–F)

light

photomicrographs of petrographic thin sections with methane-derived calcite. (A) Void-filling, extremely 13C-depleted, blocky calcite spars (‘DC’) and dolomicrite (‘M’), # WZS-SX-1-25-07. (B) Crack- and fracture-filling, extremely 13C-depleted, bladed calcite (‘DBC’), isopachous silica cements (‘ISC’), isopachous calcite cements (‘IC’), and dolomicrite (‘M’), # JLW-F. Fracture-filling bladed calcite (‘DBC’) cross-cuts and thus postdates isopachous cements (‘ISC’ and ‘IC’) (red arrow). (C–D) Close-ups of CM in interstitial space in blocky calcite spars (‘DC’). (E–F) Magnified views of labeled rectangles in (B), showing extremely 13C-depleted bladed calcite (‘DBC’) and isopachous calcite (‘IC’) and silica (‘ISC’) cements with interlayered CM. Red arrow in E marks the boundary between bladed calcite and isopachous silica.

Fig. 5. Carbonate δ13C and δ18O of microsamples. (A) Thin section reflected light image of a stromatolite (# JLW-2/2-D) from the Jiulongwan section. Magnified views of microtextures are shown in Fig. 3D–H. (B) Thin section reflected light image of microbial mat fragments (# JLW-2/2-G) from the Jiulongwan section. Magnified view of microtextures is shown in Fig. 3J. (C) Reflected light image of a polished slab showing void-filling blocky calcite spars with extremely negative δ 13C values and silicified dolomicritic matrix with moderately negative δ 13C values. Sample from the Wangzishi section (# WZS-SX-1-25-07; see also Wang et al., 2008). (D) Reflected light image of a 37

polished slab (the same specimen from which the thin section in Fig. 4B was made), showing isopachous fibrous calcite cement with moderately negative δ 13C values, crackor fracture-filling bladed calcite with extremely negative δ 13C values, and brecciated dolomatrix with moderately negative δ 13C values. Sample from the Jiulongwan section (# JLW-F). (E) Crossplot of δ13C vs δ18O. See Figs. 3–4 for magnified views of different microfabrics in thin sections.

Fig. 6. Representative raw Raman spectra (at various levels of fluorescent background) of CM from different microfabrics (left row), corresponding spectra after baseline correction (right row), and resultant spectra (dotted red lines) after peak fitting with distinguishable fitted bands, including the D1-band (~ 1350 cm–1), D2-band (~ 1600 cm–1), D3-band (fixed at 1510 cm–1), and D4-band (fixed at 1245 cm–1) (blue lines). Gray lines in (4) and (5) are raw spectra where CM peaks are overwhelmed by mineral (calcite and quartz) peaks marked with dotted lines. ISC: isopachous silica cement.

Fig. 7. Crossplot of GL peak FWHM vs. GL/D1 intensity ratio. GL is the peak at ~ 1600 cm−1, which is dominated by the D2-band for lower-grade CM with GL/D1 intensity ratio > 1.5 and includes the G-band and D2-band for higher-grade CM with GL/D1 intensity ratio < 1.5. CM sample groups of Paleozoic and Mesozoic metasediments with calibrated temperatures in Kouketsu et al. (2014) are marked on the diagram as pale blue ovals. The Doushantuo data are categorized in two groups: sedimentary CM (n = 118) and methane 38

signal-related CM (n = 54).

Fig. 8. Histograms (blue bars), fitted normal distributions (red lines), and cumulative percentages (stepped orange lines) of temperatures calculated using Eq. 1 and Eq. 2, which are denoted as T-D1 and T-D2, respectively. Data are categorized into two groups: sedimentary CM (left column) and methane signal-related CM (right column). Spectra with FWHM-D1 > 152.6 cm–1 or FWHM-D2 > 56.8 cm–1 were excluded from the calculation of T-D1 and T-D2 because these FWHM values are outside the FWHM ranges calibrated in Kouketsu et al.’s (2014) Raman geothermometers. N in each histogram is the number of spectra used in the calculation.

Fig. 9. Crossplots of Raman band parameters including FWHM, peak center positions, intensity ratios, and temperature differences (ΔT = T-D2 ‒ T-D1) for sedimentary CM (n = 118) and methane signal-related CM (n = 54), which are color coded. ISC: isopachous silica cement.

Table S1. Carbonate carbon and oxygen isotope data.

Table S2. Carbonaceous material Raman spectroscopic and geothermometric data.

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Highlights 

Extremely negative δ13C in basal Ediacaran Doushantuo cap dolostone was enigmatic

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δ13C signature may be related to gas hydrate or localized hydrothermal activity

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Carbonaceous material Raman spectroscopy of different microfabrics was analyzed

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Different microfabrics with different δ13C show similar peak temperatures <300 °C

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Peak temperatures were related to burial history or pervasive hydrothermal activities

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Graphical abstract

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