Newly discovered Sturtian cap carbonate in the Nanhua Basin, South China

Precambrian Research 293 (2017) 112–130

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Newly discovered Sturtian cap carbonate in the Nanhua Basin, South China Wenchao Yu a,b,d, Thomas J. Algeo a,c,d, Yuansheng Du a,b,⇑, Qi Zhou e, Ping Wang a,b, Yuan Xu a,b, Liangjun Yuan f, Wen Pan f a

State Key Laboratory of Biogeology and Environmental Geology, China University of Geosciences-Wuhan, Wuhan 430074, China School of Earth Sciences, China University of Geosciences, Wuhan 430074, China State Key Laboratory of Geological Processes and Mineral Resources, China University of Geosciences-Wuhan, Wuhan 430074, China d Department of Geology, University of Cincinnati, Cincinnati, OH 45221-0013, USA e Guizhou Bureau of Geology and Mineral Exploration and Development, Guiyang 550004, China f 103 Geological Party, Guizhou Bureau of Geology and Mineral Exploration and Development, Tongren 554300, China b c

a r t i c l e

i n f o

Article history: Received 30 June 2016 Revised 7 March 2017 Accepted 9 March 2017 Available online 10 March 2017 Keywords: Cryogenian Glacial Dolomite U-Pb dating Carbon isotopes Manganese

a b s t r a c t Sturtian cap carbonate deposits, previously unknown in South China, are documented in this study from the base of the Datangpo Formation in the Nanhua Basin (Guizhou Province). The age of these deposits is constrained by an LA-ICP-MS U-Pb zircon age of 662.7 ± 6.2 Ma from the overlying Mn-shale layer, which is approximately coeval with the termination of the Sturtian glaciation globally. The newly discovered cap carbonates are composed of massive, laminated, pisolitic, and sandy dolomites, and formed subtidally on the summits of horsts within the Nanhua Basin. Two d13Ccarb profiles (from the JJS-1 and ZK01 sections) show positive shifts of 4–5‰ from the base to the top of the cap carbonates with a
2‰ offset between the profiles. These features suggest both a large global perturbation as well as local influences on carbonate carbon isotopes during the Sturtian deglaciation, which is consistent with aspects of both the Snowball Earth and gas hydrate destabilization hypotheses for the origin of cap carbonates. The newly discovered cap carbonates are inferred to be correlative with rhodochrosite (Mn-carbonate)bearing deposits of the basal Datangpo Formation in adjacent grabens of the Nanhua Basin. This correlation is based on (1) the equivalent stratigraphic positions of the cap carbonates and Mn-deposits, both of which are found at the base of the Datangpo Formation, and (2) consistent U-Pb dates of 667–663 Ma for tuff layers interbedded with these deposits at two locales. The Mn-deposits show significantly higher TOC content (1–3% vs <0.7%) and more negative d13Ccarb values (10.5 to 5.5‰ vs 4.0 to +2.5‰) than the cap carbonates. The upsection trend toward higher d13Ccarb in the cap carbonates may reflect increases in marine productivity and the burial rate of organic carbon following the Sturtian glaciation. Despite coeval formation, the dolomitic cap carbonates deposited on horsts and the Mn-carbonate sediments deposited in grabens are heterogeneous facies with different formation mechanisms. Ó 2017 Elsevier B.V. All rights reserved.

1. Introduction Cap carbonates are continuous layers of limestone and/or dolomite that sharply overlie Neoproterozoic glacial deposits, formed during deglacial transgressions, and exhibit thicknesses ranging from 0.5 to 30 m (Shields, 2005; Hoffman et al., 2011). Two main glacial events are recognized in Cryogenian successions globally, the Sturtian (
717–660 Ma) and the Marinoan (
650–635 Ma) ⇑ Corresponding author at: State Key Laboratory of Biogeology and Environmental Geology, China University of Geosciences-Wuhan, Wuhan 430074, China. Tel.: +86 13971241916, fax: +86 27 87481365. E-mail address: [email protected] (Y. Du). http://dx.doi.org/10.1016/j.precamres.2017.03.011 0301-9268/Ó 2017 Elsevier B.V. All rights reserved.

(Hoffman and Li, 2009; Rooney et al., 2015). Different types of cap carbonates overlie the Sturtian and Marinoan glacial deposits (Corsetti and Lorentz, 2006). Sturtian cap carbonates tend to consist of dark, finely laminated carbonates, whereas Marinoan cap carbonates are characterized by a lighter color (whitish or pinkish) and exhibit features not seen in the Sturtian cap carbonates, e.g., prismatic pseudomorphs after barite or aragonite (Hoffman and Schrag, 2002; Shields et al., 2007), tubestone (Kennedy et al., 1998) and pseudo-tepee structures (Kennedy, 1996). Also, different patterns of carbonate carbon isotopic variation are observed: Sturtian cap carbonates yield negative d13C values at their base, increasing to weakly positive values upsection over a few meters or tens of meters, whereas Marinoan cap carbonates occasionally

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yield extremely negative d13C values (to 45‰) with little change upsection (Kennedy, 1996; Jiang et al., 2003). Neoproterozoic cap carbonates exist in South China, but the documented occurrences are all associated with the Marinoan glaciation. This is true of the cap carbonates at the base of the Ediacaran Doushantuo Formation, which overlie diamictites of the Nantuo Formation (Jiang et al., 2003; Wang et al., 2008). These carbonates have been dated by the TIMS zircon U-Pb method, yielding ages of 635.4 ± 0.6 Ma (Condon et al., 2005). Deposition of the Nantuo Formation was thus coeval with the Marinoan glaciation, and the overlying cap carbonates are earliest Ediacaran in age. Glacial sediments of Sturtian age have been identified in South China, e.g., diamictites of the Tiesi’ao Formation and correlative units, as shown by a zircon U-Pb date of 663 ± 4 Ma from a tuff bed directly overlying these units (Zhou et al., 2004; Corsetti and Lorentz, 2006; Chen et al., 2008). However, the existence of cap carbonates above these Sturtian glacial deposits is controversial. Some researchers believe that no cap carbonate was deposited (Dobrzinski and Bahlburg, 2007), whereas others regard a finely laminated, organic-rich manganese (Mn) carbonate layer at the base of the Datangpo Formation as a cap-carbonate equivalent (Zhou et al., 2004; Corsetti and Lorentz, 2006; Chen et al., 2008). These deposits contain considerable amounts of terrigenous detrital material (e.g., quartz, feldspar, illite), and the dominant carbonate mineral is rhodochrosite instead of calcite or dolomite. These features make this deposit unlike other Sturtian cap carbonates globally, e.g., the Scout Mountain Member of the Pocatello Formation, southeastern Idaho (Smith et al., 1994), the Ghubrah Member of the Ghadir Manqil Formation, Oman (Brasier et al., 2000), the Rasthof Formation in Namibia (Pruss et al., 2010), and the Tindelpina Formation in South Australia (Giddings and Wallace, 2009). In the present study, we identify for the first time shallowwater Sturtian cap carbonates from South China, which are located at the base of the Datangpo Formation in newly discovered sections in Guizhou Province. These cap carbonates are present on horst blocks of the Nanhua Basin, whereas the previously known Mn-carbonates of the basal Datangpo Formation represent deepwater deposits in adjacent grabens. The newly discovered cap carbonates are composed of massive or laminated, sandy to pisolitic dolomite, in which d13C values increase progressively from bottom to top. These deposits are thinner (<10 m) and more dolomitic than many Sturtian cap carbonates globally, although their laminated dolomitic composition and synsedimentary deformation structures are similar to those found in Marinoan cap carbonates (Hoffman and Schrag, 2002). However, we demonstrated the association of the newly discovered Nanhua Basin cap carbonates with the Sturtian glaciation by U-Pb dating of zircons in an overlying tuff layer to 662.7 ± 6.2 Ma (see Sections 5.1 and 6.1). In this contribution, we (1) document these new Sturtian cap-carbonate deposits at five locales in the Nanhua Basin, (2) develop a genetic model for the cap carbonates based on their petrographic and geochemical characteristics, and (3) contrast their occurrence and mode of formation with that of previously known Mn-carbonate deposits from the basal Datangpo Formation.

2. Geological background The study area is in Tongren City, northeastern Guizhou Province, South China (Fig. 1). Paleogeographically, it was located on the southeastern margin of the Yangtze Block, within the Nanhua Basin (Fig. 2A). Prior to the Cryogenian (>720 Ma), the Yangtze Block broke away from the supercontinent Rodinia, causing the Nanhua Basin to open as a failed rift basin between the Yangtze and Cathaysia blocks, accumulating a thick succession of continen-

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tal and volcaniclastic sediments (Wang and Li, 2003). Although the names of these deposits differ regionally across the Yangtze Block (e.g., Banxi Group in eastern Guizhou Province and western Hunan Province, Xiajiang Group in southern Guizhou Province, and Sibao Group in northern Guangxi Province; see Fig. 3), they are approximately coeval sedimentary successions with ages ranging from
820 Ma to 720 Ma (Wang et al., 2007, 2012; Zhao et al., 2011). During the Cryogenian (720–635 Ma), two main glacial events were recorded on the Yangtze Block. These events are known regionally as the Jiangkou and Nantuo glaciations, but they are in fact correlative with the globally known Sturtian and Marinoan glaciations, respectively (Zhang et al., 2011). The Jiangkou (Sturtian) glaciation on the Yangtze Block has been subdivided into two stages: the older Chang’an glaciation and later Gucheng glaciation (Lan et al., 2015a,b). The most complete Cryogenian sections are preserved in basinal facies of the Nanhua Basin, where the Chang’an Formation is composed of diamictite with a thickness of >1000 m, and the onset of the Chang’an glaciation has been dated to
716 Ma (Figs. 2A and 3; Zhou et al., 2004; Lan et al., 2014, 2015a). In shelf facies, the Chang’an Formation is absent and only the Gucheng Formation (or its correlatives, the Gucheng Member of the Fulu Formation and the Tiesi’ao Formation; Fig. 3) is preserved as a 3- to 15-m-thick diamictite. The Chang’an and Gucheng glaciations are separated by interglacial deposits assigned to the Liangjiehe Formation (or its correlatives, the Liangjiehe Member of the Fulu Formation and the Xieshuihe Formation; Fig. 3), which consist of sandstone with a thickness ranging from 20 to 300 m. The age of onset of the Gucheng glaciation has been constrained to <691 ± 12 Ma by dating of the upper Xieshuihe Formation in western Hunan Province (Lan et al., 2015b). Deposits on the Yangtze Block belonging to the Nantuo (Marinoan) glaciation are more widespread than those of the Jiangkou (Sturtian) glaciation, with thicknesses ranging from tens to several hundreds of meters (BGMRGZAR, 1985; BGMRGZP, 1987). The age of onset of the Nantuo glaciation has been estimated radiometrically as 654.5 ± 3.8 Ma (Zhang et al., 2008). The Cryogenian succession in northern Tongren City, eastern Guizhou Province, comprises the Liangjiehe, Tiesi’ao, Datangpo, and Nantuo formations in ascending order. The Liangjiehe Formation, which consists of a basal feldspathic-lithic arenite overlain by 50–150 m of quartz arenite, is present only in the central part of the study area (around Datangpo and Taoying villages), disappearing to the north and south (around Lengshuixi and Bapan villages, Fig. 1B). The Tiesi’ao Formation consists of 1 to 15 m of massive diamictite or dolomitic diamictite. The Datangpo Formation is subdivided into three members, the 1st Member consisting of 0.5– 15 m of Mn carbonate and Mn-bearing shale or 2–4 m of dolomite, the 2nd Member consisting of 1–20 m of pyritic black shales (but lacking in some sections with a basal dolomite layer), and the 3rd Member consisting of 100–700 m of gray mudstones and siltstones (Zhou et al., 2016). At most locales, the Datangpo Formation shows no evidence of intraformational hiatuses, suggesting a continuously deposited succession. The large changes in stratal thickness were controlled by the paleobathymetry of the Rift Basin, which comprised a series of NE-SW-trending horsts and grabens, characterized respectively by thinner and thicker deposits (Fig. 2B; BGMRGZP, 1987; Zhou et al., 2016). The Nantuo Formation consists of
100 m of glaciomarine diamictite.

3. Materials and methods 3.1. Study sections Five new locales in northeastern Guizhou were analyzed in our study, including three outcrop sections (JJS-1, BP-1, and BP-2) and

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Fig. 1. (A) Location of study area in China. (B) Geologic map of Neoproterozoic strata of northeastern Guizhou Province and locations of study sections.

two drillsites (ZK01 and ZK4207) (Fig. 1B). The first four locales were located on structural horsts within the Nanhua Basin, with ZK01 being located on the margin of a horst and, thus, possibly rep-

resenting somewhat deeper waters than the other three horst sections (JJS-1, BP-1, and BP-2). The fifth locale (ZK4207) was located in a structural graben (Fig. 2B).

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Fig. 2. (A) Location of study area within the Late Neoproterozoic Nanhua Rift Basin of South China (modified from Wang and Li, 2003). (B) Paleogeographic reconstruction of study area, modified from Zhou et al. (2016) and Yu et al. (2016).

Fig. 3. Schematic Cryogenian stratigraphic framework of the Nanhua Basin, South China. Stratal data, including distribution of Mn-rich deposits, are from BGMRGZAR (1985), BGMRGZP, 1987 and BGMRHNP (1988). (See above-mentioned reference for further information).

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The JJS-1 section is located on a roadside just 1 km south of Jiangjunshan Village (28°080 4700 N, 108°530 5100 E) (Fig. 1B). The section begins in the upper part of the thick sandstones of the Banxi Group, which is unconformably overlain by a 1.5-m-thick dolomitic diamictite of the Tiesi’ao Formation. This layer is conformably overlain by a
4-m-thick layer at the base of the Datangpo Formation consisting of massive white sandy to black micritic dolomite, with the latter containing a
30-cm-thick bed of pisolitic dolomite. The dolomite layer is sharply overlain by sandstone and Mnbearing shale belonging to the 1st Member of the Datangpo Formation, overlain by black shale and yellow siltstone belonging to the 2nd and 3rd members of the same formation (Fig. 4). The ZK01 drillsite is located 4 km southeast of Lengshuixi Village (28°080 2300 N, 108°520 2500 E) (Fig. 1B). The drillcore penetrated to a total depth of 328 m, reaching the Banxi Group. From its base upward, the drillcore contains quartz sandstone of the upper Banxi Group, overlain unconformably by a 1.8-m-thick diamictite of the Tiesi’ao Formation. This layer is unconformably overlain by a 1.0m-thick laminated siltstone and a cap carbonate bed, both belonging to the Datangpo Formation. The cap carbonate can be subdivided into a lower, 0.5-m-thick laminated dolomite and an upper, 1.5-m-thick massive dolomite. The lower interval contains roll-up structures, characterized by syn-sedimentary deformation of laminated dolomite, which are typical of Sturtian cap carbonates (cf. Corsetti and Lorentz, 2006). The cap carbonate is covered by
250 m of black clayey siltstone, which is overlain by Nantuo (Marinoan-equivalent) diamictites (Fig. 4). Sections BP-1 and BP-2 are located on the west side of Bapan Village (27°420 4300 N, 108°560 3500 E) (Fig. 1B). The cap carbonate deposits at these locales show similar characteristics, both being composed of a lower unit of dolomitic diamictite and an upper unit of massive dolomite. The lower and upper units are 1.2 and 1.3 m thick, respectively, at BP-1, and 0.4 and 0.2 m thick, respectively, at BP-2. The cap carbonate at BP-1 is overlain by a 3-m-thick black shale containing a Mn-carbonate lens (Fig. 4). The ZK4207 drillsite is located 10 km southwest of Songtao County (28°090 1500 N, 108°120 1000 E). The total thickness of the Datangpo Formation in this drillcore is 370 m, comprising 13 m of Mn-rich strata (1st Member), 27 m of black shale (2nd Member), and 330 m of siltstone (3rd Member). The 1st Member consists of horizontally laminated, interbedded Mn-carbonate (ore) and Mnshale layers (Fig. 4). See Yu et al. (2016) for a full description of this locale. At ZK01, JJS-1, and ZK4207, the main stratigraphic units were sampled for both petrographic and geochemical analyses. At BP-1 and BP-2, 5 hand samples were taken from each section for petrographic study (BP-1-1 to -5, and BP-2-1 to -5). At ZK01, 31 samples were collected: 8 samples from the laminated siltstone layer (ZK01-11-1 to -16-2), 22 samples from the cap carbonate (ZK0117 to -31-1), and 1 sample from the gray siltstone (ZK01-32). At JJS-1, 16 samples were collected: 7 samples from the lower sandy dolomite (JJS-1-1 to -7), 3 samples from the black dolomite (JJS-18, -9, and -11), and 1 sample from the pisolitic dolomite (JJS-1-10) that collectively comprise the cap carbonate, and 5 additional samples from the upper sandy dolomite (JJS-1-12 to -16). Sample JJS-T was taken from the tuff layer in the overlying Mn-bearing shale for zircon dating analysis. At ZK4207, 14 samples were collected from the Mn-carbonate layer of the 1st Member of the Datangpo Formation for carbon isotopic analysis. For comparison with the present study sections, stratigraphic, geochronological, and carbon isotopic data were compiled from published sources for the Heishuixi, Zhailanggou, and Datangpo sections (Figs. 1B and 2B). The data include zircon TIMS ages from Zhailanggou (Zhou et al., 2004), zircon SIMS ages from Heishuixi (Yin et al., 2006), and carbon isotopic data for Mn-carbonate samples from Datangpo and Zhailanggou (Chen et al., 2008; Zhou and

Du, 2012). The Liangjiehe Formation is present at Datangpo and Zhailanggou but not at Heishuixi. At these three locales, the Tiesi’ao Formation is 10 to 25 m thick and composed of massive diamictite containing poorly sorted mm- to cm-sized clasts. The diamictite is matrix supported, and the clasts are subangular to subrounded and mostly of sedimentary origin (Dobrzinski and Bahlburg, 2007). The unconformably overlying Datangpo Formation consists of 5 to 10 m of interbedded Mn-carbonate and Mnshale (1st Member), 10 to 15 m of black shale (2nd Member), and 200 m of gray clayey siltstone (3rd Member) (Zhou et al., 2004; Yin et al., 2006; Zhou and Du, 2012). 3.2. Analytical methods The isotopic composition of inorganic carbon was determined using a Thermo Finnigan MAT 253 isotope ratio mass spectrometer (IRMS) at the State Key Laboratory of Geological Processes and Mineral Resources (GPMR), China University of GeosciencesWuhan. Isotopic compositions are reported in standard delta notation relative to the Vienna Pee Dee Belemnite (VPDB) standard. Precision is better than 0.2‰, based on multiple analyses of laboratory standards. Samples for organic carbon analysis were treated with 10% HCl to remove carbonate, rinsed with distilled water, and freeze-dried at 40 °C overnight. Total organic carbon (TOC) was determined using an Analytik Jena Multi EA 4000 type C-S analyzer at the State Key Laboratory of Biogeology and Environmental Geology (BGEG), China University of Geosciences-Wuhan. Analytical precision is ±0.2%. The isotopic composition of organic carbon was determined using a Thermo Finnigan MAT 253 IRMS (isotope ratio mass spectrometer) at GPMR. Isotopic compositions are reported in standard delta notation relative to the VPDB (Vienna Pee Dee Belemnite) standard. Precision is better than 0.2‰, based on multiple analyses of laboratory standards. Zircons in sample JJS-T were separated by standard techniques prior to casting in epoxy mounts, polishing and imaging in transmitted light. Carbon-coated mounts were imaged by cathodoluminescence (CL) to observe the internal structure of the grains and to guide in situ laser-ablation analysis (Fig. S1). U–Pb geochronology and trace element concentrations of zircon grains were analyzed by laser ablation inductively-coupled plasma mass spectrometry (LA-ICP-MS) using operating conditions as given in Liu et al. (2008, 2010a,b). Both CL and LA-ICP-MS analyses were carried out at GPMR. Laser sampling was performed using a GeoLas 2005. An Agilent 7500a ICP-MS was used to acquire ion-signal intensities. The laser was operated at a wavelength of 193 nm, a frequency of 5 Hz, an energy density of 10 Jcm2, and with a spot size of 32 lm. Argon was used as the make-up gas and mixed with the carrier gas (He) via a T-connector before entering the ICP-MS. Nitrogen was added to the central gas flow (Ar + He) to decrease the detection limit and improve precision (Hu et al., 2008). Each analysis included a background acquisition interval of 20 s (gas blank) followed by a sample data acquisition interval of 50 s. The Agilent Chemstation was utilized for the acquisition of each individual analysis. Off-line selection and integration of background and analytical signals, time-drift correction, U-Pb dating and zircon trace-element corrections were performed using ICPMS-DataCal (Liu et al., 2010a,b). Zircon analyses made use of three standards: (1) NIST SRM 610 was used to calibrate trace elements and was analyzed at the beginning and end of each analysis session; (2) the zircon GJ-1 was used to monitor the data quality and was analyzed three times after NIST SRM 610 at the beginning and two times before NIST SRM 610 at the end of each analysis session; and (3) the zircon 91500 was used as an external standard for UPb dating and was analyzed twice for every eight sample analyses. Time-dependent drifts of U-Th-Pb isotopic ratios were corrected

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Fig. 4. (A) Lithologic profiles of Cryogenian strata in northeastern Guizhou Province, South China. (B) Detailed lithostratigraphy of the Cryogenian Tiesi’ao Formation and 1st and 2nd members of Datangpo Formation. Stratal and age data are from Zhou et al. (2004), Yin et al. (2006), Zhou and Du (2012), and this study.

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using a linear interpolation with time for every eight analyses of the zircon 91500 standard (Liu et al., 2010a), and the preferred U-Th-Pb isotopic ratios of this standard are from Wiedenbeck et al. (1995). The time-drift correction was explained in detail by Liu et al. (2010a) and is given along with uncertainty calculation methods in the Supplemental Information file. In this study, the zircon 91500 standard yielded a weight-averaged age of 1062.4 ± 5.6 Ma (n = 18, MSWD = 0.01), and the zircon GJ-1 standard yielded a weight-averaged age of 604.7 ± 5.1 Ma (n = 5, MSWD = 0.22) (Fig. S2, Table S1). Ages and uncertainties of 206 Pb/238U ages in our study conform to the accuracy for standards in Liu et al. (2010a,b), but both 91500 and GJ-1 have relatively poor reproducibility in 207Pb/206Pb ages (Table S1). 208Pb/232Th ratios in the sample were calculated based on the 91500 standard using measured values for 91500 assuming the same degree of discordance as for U-Pb ratios. Since the Th content of zircon 91500 (
26 ppm) is much lower than that of zircons we analyzed (100– 400 ppm), the younger Th/Pb ages of most samples are the result of calibration against 91500 (Jackson et al., 2004). Concordia diagrams for U-Pb isotopic ratios and weighted mean calculations were made using Isoplot/Ex ver. 3.75 (Ludwig, 2008), and correlation coefficients between 206Pb/238U and 207Pb/235U were calculated following the protocol of Schmitz and Schoene (2007). 4. Sedimentology of the Sturtian glacial to post-glacial transition 4.1. Petrographic observations 4.1.1. Sturtian glacial deposit The Tiesi’ao Formation consists of diamictite and dolomitic diamictite in the study sections. At ZK01, the 1.8-m-thick diamictite is massive and grain supported, composed mainly of aluminosilicate material, and exhibits decreasing clast size upsection (Fig. 5A and B). In the lower part of the diamictite, the clasts are composed mainly of sandstone fragments from the underlying Banxi Group, have long axes ranging from 2 to 40 mm, and exhibit poor sorting and rounding; the matrix consists of aluminosilicate silt and mud and pyrites nodules. In the upper part of the diamictite, clasts are smaller (2–20 mm). At JJS-1, BP-1, and BP-2, the diamictite layer is dolomitic (Fig. 5C–F) with a matrix-supported fabric consisting of micritic dolomite crystals (Fig. 5G–H). Clasts are mainly composed of poorly sorted, gravel-sized (2–50 mm) lithic fragments and sand-sized quartz grains. In some cases, fractures in the pebbles are filled with dolomitic diamictite matrix (Fig. 5F), which has been interpreted as evidence of frost shattering (Spence et al., 2016). At ZK01, the diamictite is uncomfortably overlain by a 1-m-thick gravel-bearing laminated sandstone (Fig. 6A and B) followed by a sandstone layer (Fig. 6C) in which some thin dolomitic layers (30–100 lm) can be observed (Fig. 6D). 4.1.2. Cap carbonate The Tiesi’ao Formation diamictite is overlain by syn-deglacial to post-glacial sediments of the Datangpo Formation. At ZK01, these sediments consist of thinly laminated (0.5–1.0 mm) dark gray to gray fine-grained dolomite containing microfractures filled with organic matter (Figs. 6D, and 8A). Some of the dolomitic layers (e.g., overlying the sandstone layer in ZK01) exhibit softsediment deformation and local oversteepening (to 80°) (Fig. 6E). These features may represent ‘‘roll-up structures”, which are widely found in cap carbonates (Corsetti and Lorentz, 2006), although their exact origin in the present study units is difficult to determine from drillcore segments. Massive, white to black dolomite beds with a micritic or microspar fabric and cut by dolomite-filled veins were observed in all

four study sections (Figs. 6F, 7A, E, F, H, 8B). Sandy dolomite is present at JJS-1 and BP-1, where sand- and silt-grade quartz grains can be observed within the micritic matrix (Fig. 8C). Pisolitic dolomite was found only at JJS-1, consisting of a single 15-cm-thick grainsupported layer that is distinctly lighter in color than the underlying and overlying massive dolomites (Fig. 9A). The pisoids are spherical to ellipsoidal in shape, range from 1.5 to 2.5 mm in diameter, exhibit well-developed concentric cortices, and either have quartz silt nuclei or lack nuclei altogether. Silty quartz and hematite grains are common in the pisolitic dolomite (Fig. 9B and C). 4.1.3. Strata overlying cap carbonate The strata overlying the cap carbonate deposit of the basal Datangpo Formation are varied in character. At JJS-1, the cap carbonate is overlain by fine-grained (0.1- to 0.25-mm-diam.) sandstone (Figs. 7C and 8D) that grades upward into Mn-bearing shale (Fig. 7D). At BP-1, the cap carbonate is overlain by black shale of the 2nd Member of the Datangpo Formation, whereas at BP-2 and ZK01, the cap carbonates are overlain by clayey siltstone of the 3rd Member. The absence of black shale at BP-2 and ZK01 may have been due to development of local exposure surfaces at these two sites, producing an unconformity between the cap carbonate of the 1st Member and siltstone deposits of the 3rd Member. At some locales, the Datangpo cap carbonates are overlain by Mn-deposits. At JJS-1, a Mn-rich layer is 0.5 m thick, brown to black in color, and massive internally, showing comfortable contacts with the underlying sandstone and overlying siltstone layers. It contains a thin tuff composed of horizontally oriented volcanic glass shards (Figs. 7D and 8F). In BP-1, a Mn-rich lens (30 cm  20 cm) present within black shales of the 2nd Member (Fig. 7G) contains MnO of 2.4–2.7%, Al2O3 of 14–17%, and SiO2 of 60–62%. This feature is closer to a Mn-bearing shale than a Mn-carbonate. From microscopic observations (Fig. 8E), terrigenous detrital minerals such as quartz, feldspar, and clay minerals (mostly illite) of silt or finer size are dominant in the samples. Rhodochrosite (MnCO3) formed mainly around particles of organic matter, with the diameters of spherical rhodochrosite crystals ranging from 2 to 10 lm. Pyrite crystals were also observed in these samples. In contrast, Mn-ore samples in ZK4207 contain more rhodochrosite with higher MnO content (12–30%), lower Al2O3 content (3.6–12.6%) and lower SiO2 content (20.1– 40.5%) (Yu et al., 2016). 4.2. Sedimentological interpretations The Tiesi’ao Formation shows different petrographic characteristics in grabens versus horsts of the Nanhua Basin. In graben sections (e.g., Heishuixi, Zhailanggou and Datangpo), the Tiesi’ao Formation consists of synglacial diamictites exhibiting a matrixsupported fabric and a variable content of mm- to cm-sized clasts (Boulton and Deynoux, 1981; Dobrzinski and Bahlburg, 2007). These deposits represent a distal glaciomarine setting on a continental shelf that was influenced by ice-rafting processes or debris flows of reworked sediments (Boulton and Deynoux, 1981). In horst sections (e.g., ZK01), a grain-supported fabric and bouldersized clasts provide evidence of deposition via subglacial melting of grounded glaciers (Dobrzinski and Bahlburg, 2007). A gravelbearing sandstone layer above the diamictite in ZK01 records turbiditic deposition during icesheet retreat (Boulton and Deynoux, 1981). The dolomitic diamictites in the JJS-1, BP-1 and BP-2 sections are a product of the deglacial transgression. The source of carbonate matrix in Tiesi’ao Formation tillites is uncertain. One possible mechanism is erosional recycling of preglacial carbonate deposits (Fairchild and Spiro, 1990; Fairchild et al., 1993). In eastern Guizhou Province, carbonate and calcareous shale are widely found in the Tonian Hongzixi Formation of the Upper Banxi Group (BGMRGZP, 1987). Erosional recycling of these

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Fig. 5. Drillcore and field photos. (A) Basal Tiesi’ao diamictite at ZK01: note pyrite nodules, 20–60-mm-diam. gravel, and siltstone-mudstone matrix. (B) Upper Tie’siao Formation at ZK01. Dolomitic diamictites in the (C) BP-1, (D) BP-2, and (E) JJS-1 sections. (F) Polished surface of dolomictite sample from JJS-1: gravel clasts consist mainly of sandstone, and the matrix is composed of dolomite. Notice pebble fragments in the red circles; some fractures in the pebbles are filled with dolomitic matrix. (G-H) Poorly rounded and sorted lithic clasts in a matrix of micritic dolomite in the dolomitic diamictite at JJS-1. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

carbonate deposits represents a potential source of the carbonate matrix in the Tiesi’ao Formation. In the Datangpo cap carbonate, the lithofacies association of massive dolomite, laminated dolomite, pisolitic dolomite and sandy dolomite indicate a subtidal depositional environment

(Corkeron, 2007). The pisolitic dolomite is observed only in the JJS-1 section. Pisoids (or ‘‘giant ooids”, e.g., Trower and Grotzinger, 2010) are widely present in Precambrian carbonate deposits, probably due to a combination of low nuclei supply, high cortex growth rates, and strong water agitation (Sumner

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Fig. 6. Drillcore photos of ZK01. (A) Laminated sandstone above diamictite; the contact is shown by a dashed white line. (B) Gravels in the laminated sandstone layer (red arrows). (C) The polished surface of the laminated sandstone layer. (D) Thin section of laminated dolomite, in which thin dolomitic layers can be observed. (E) Laminated dolomite; notice the high dip angle. (F) Massive dolomite; red arrow indicates top. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

and Grotzinger, 1993; Trower and Grotzinger, 2010). The presence of hematite within and between the pisoids and the relatively lighter color of this bed suggest an oxic environment of deposition. At ZK01, laminates in dolomite of the basal cap carbonate overlies laminated sandstone at a high dip angle (Fig. 6E), suggesting that these features may be ‘‘roll-up structures” (Corsetti and Lorentz, 2006). The origin of roll-up structures found in cap carbonates has been variously attributed to subaerial exposure resulting in teepee development (Aitken, 1991), mat formation by chemosynthetic and/or heterotrophic microbes (Kennedy et al., 2001; Corsetti and Kaufman, 2003), and shallow burial (quasisynsedimentary) deformation related to sediment compaction (Corkeron, 2007). The petrographic observations of the present study provide no evidence for subaerial exposure. Another explanation attributes roll-up structures to burial compaction, because the underlying laminated sandstone is mud-rich, the cap carbonate formed a relatively impermeable seal, resulting in the development of an overpressured zone that caused compactional folding of the carbonate layer (Corkeron, 2007). However, we observed no vertical fractures or veins in the laminated sandstone beneath the cap carbonate (Fig. 6C). Microbial activity may be the most likely origin of these laminated, somewhat deformed structures (cf. Kennedy et al., 1998). The observed differences in lithology and sedimentary structures between the four study sections located on horsts of the Nanhua Basin reflect variation in depositional water depths. At ZK01, located on the deep margin of a horst, the diamictite is a glaciomarine deposit, and the overlying syn-deglacial succession contains turbidites below the cap carbonate deposit. At the shallower JJS-1, BP-1, and BP-2 sections, dolomitic diamictite layers are a product of the deglacial transgression and, thus, are probably slightly younger than the diamictite deposit at ZK01. In these sections, all of which are located in the central horst area (Fig. 2B), the diamictite has a dolomitic matrix (Fig. 5C to H), suggesting the influence of increased seawater carbonate alkalinity during the deglaciation.

5. Isotopic results 5.1. Zircon U-Pb dating A total of 56 zircon U-Pb ages were generated from sample JJST, which is from a tuff contained within overlying Mn-shales of the basal Tiesi’ao Formation, about 0.5 m above the cap carbonate deposits. The age of this tuff thus provides a minimum age for the cap carbonate. Of the 56 analyses, 42 of them fall on or close to the concordia trend (discordance 10%) (Fig. 10A), most discordant analyses that have 206Pb/238U ages in the range of 630– 690 Ma have 207Pb/206Pb ages that are displaced to the right of Concordia. Considering the relatively poor precision of the 207 Pb/206Pb on standards (91500 and GJ-1), these discordant ages may also need to take part in the geochronological analyses. Age uncertainties for individual analyses are given as 2r values in the data table and concordia plots. Zircon U-Pb isotopic compositions are presented in the Supplemental Information file. The zircon U-Pb ages in sample JJS-T range from 1779 to 626 Ma but fall naturally into three groups (Fig. 10A–C). The interval 691– 626 Ma contains a cluster of 42 ages yielding a weight-averaged age of 662.7 ± 6.2 Ma (MSWD = 7.8). Another 5 analyses define a second, smaller cluster in the age range of 807–758 Ma. Considering that major phases of widespread magmatism in South China occurred at 820–800 Ma and 780–745 Ma (Li et al., 2003), it is reasonable to associate these zircon ages with those events rather than with the 691–626-Ma age cluster. A third set of 5 analyses is scattered widely over the age range 1112–1779 Ma without any obvious age peak. These zircon ages are likely to represent recycling of older zircons from a magma reservoir, and, thus, they do not have any significance for dating the tuff. One zircon grain (JJS-T-41) gives an anomalous Eocene age (39.6 Ma), which likely to represent a contaminant. The JJS section preserves only the lower part of the Datangpo Fm., and the section is covered by Quaternary sediments, as a consequence of which downward percolating surface water may have introduced this anomalous grain into the study unit.

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Fig. 7. Field photos. (A) Cap carbonate of basal Datangpo Formation at JJS-1; geology hammer in red circle is 30 cm long. (B) Pisolitic dolomite layer in cap carbonate at JJS-1. (C) Sandstone layer covering cap carbonate at JJS-1. (D) Tuff layer within Mn-bearing shale above sandstone layer at JJS-1. (E) Massive dolomite and underlying dolomitic diamictite at BP-1. (F) Massive dolomite at BP-1. (G) Mn-shale lens in black shale above dolomite layer at BP-1. (H) Massive dolomite at BP-2. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

5.2. Carbon isotopes and TOC The carbon and oxygen isotope and TOC data for the present study sections are given in Table 1, and carbon isotope and TOC profiles are shown in Fig. 11. Carbon isotope data for the

Zhailanggou (Chen et al., 2008) and Datangpo (Zhou and Du, 2012) sections are also shown in Fig. 11. At JJS-1, seven samples from the lower part of the cap carbonate deposit (JJS-1-1 to -7) show low inorganic carbon isotope (d13Ccarb) values (2.34 to 0.17‰) (Fig. 11A). Nine samples from the upper

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Fig. 8. Thin-section photos. (A) Laminated dolomite at ZK01. (B) Massive dolomite at JJS-1, showing micritic texture. (C) Sandy dolomite at JJS-1. (D) Sandstone layer covering the cap carbonate at JJS-1. (E) Mn-shale lens at BP-1. (F) Tuff layer at JJS-1. OM = organic matter, Rds = rhodochrosite, Qtz = quartz, Ill = illite.

part of the cap carbonate (JJS-1-8 to -16) record a trend toward higher d13Ccarb (0.05 to +3.23‰). At ZK01, d13Ccarb was generated for 19 dolomite samples (ZK01-21-1 to -30), and organic carbon isotopes and TOC data were generated for all 31 samples (Fig. 11B). d13Ccarb values are relatively 13C-depleted, with most samples between 1 ‰ and +1‰ and three negative excursions corresponding to samples ZK01-22 (4.52‰), ZK01-26 (3.12‰), and ZK01-29-2 (1.43‰). As at JJS-1, the d13Ccarb profile for ZK01 shows an overall positive trend upsection in the dolomite deposit. At JJS-1, TOC values range from 0.05 to 0.70% with a peak corresponding to sample JJS-1-8. d13Corg values range mainly from 28 to 24 ‰ but show positive excursions corresponding to samples JJS-1-2 (25.32‰), JJS-1-7 (22.10‰), and JJS-1-10 ( 24.56‰). D13Ccarb-org values range from 20.33 to 30.56‰, with positive excursions corresponding to samples JJS-1-3 (27.59‰) and JJS-1-8 (29.50‰) and a general trend toward

higher values in the upper part of the section (samples JJS-1-10 to JJS-1-16). At ZK01, TOC contents range from 0.02 to 0.17%, with a peak corresponding to sample ZK01-21-2. The laminated sandstone unit yielded nearly uniform d13Corg values of 30.63 to 29.63‰, whereas the overlying dolomite unit shows positive extrusions corresponding to samples ZK01-21-3 to ZK01-23-2 (from 4 to 2 ‰) and ZK01-28-2 (from 2 to 0‰). D13Ccarb-org values range from 22.03 to 29.83‰, with positive excursions corresponding to samples ZK01-21-1 to ZK01-21-2 (from 28.99 to 29.09‰) and ZK01-25–1. A d13Corg-vs-d13Ccarb crossplot (combining data from the JJS-1 and ZK01 sections) exhibits a slope close to unity 34:0 13 (1:021:16 0:83 ) and an intercept (D Ccarb-org) of 26:718:7 (Fig. S3B). Carbonate isotope and TOC values of the rhodochrosite (Mncarbonate) deposits differ from those of the cap carbonate dolomites. For instance, d13Ccarb is distinctly 13C-depleted (6.90 to

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Fig. 9. (A) Pisolitic layer in the JJS-1 section. (B–C) Spherical to elliptical pisoids with diameters ranging from 1.5 to 3 mm. Hem = hematite, Qtz = quartz.

Fig. 10. (A) Detrital zircon U-Pb concordia diagram for sample JJS-T from tuff layer at JJS-1. (B) Weighted average age for dominant peak in sample JJS-T (n = 42). (C) Probability density distribution diagram for sample JJS-T.

5.79‰ for ZK4207, 9.4 to 5.6‰ for Zhailanggou, and 10.28 to 7.06‰ for Datangpo) relative to dolomite samples from JJS-1 and ZK01 (4.52 to +3.23‰) (Chen et al., 2008; Zhou and Du, 2012). The d13Corg values of the Mn-carbonate deposits at Zhailanggou and ZK4207 range from 33.7 to 29.9‰ with a

mean of 32.87 ‰ and, thus, are more 13C-depleted than the dolomite samples (31.48 to 22.10‰). The TOC content of the Mn-carbonate deposits ranges from 1 to 3% (Chen et al., 2008), which is much higher than for the dolomite samples (0.05– 0.70%).

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Table 1 Carbon isotope and TOC data from section JJS-1, drillcores ZK01 and ZK4207 in Nanhua Basin, South China. Sampling position (m)

Sample

d13Ccarb ‰

d18O ‰

d13Corg ‰

Dd13Ccarb-org ‰

TOC %

Sampling position (m)

Sample

d13Ccarb ‰

d18O‰ ‰

d13Corg‰ ‰

Dd13Ccarb-org ‰

TOC %

312 311.84 311.62 311.52 311.48 311.46 311.44 311.37 311.28 311.22 311.18 311.12 311.06 311 310.05 310.9 310.8 310.73 310.68 310.63 310.58 310.52 310.45 310.43 310.37 310.3 310.25 310.15 310 309.85 309.3

ZK01-11-1 ZK01-11-2 ZK01-12-1 ZK01-12-2 ZK01-13 ZK01-14 ZK01-16-1 ZK01-16-2 ZK01-18 ZK01-19 ZK01-21-1 ZK01-21-2 ZK01-21-3 ZK01-22 ZK01-23-1 ZK01-23-2 ZK01-24-1 ZK01-24-2 ZK01-25-1 ZK01-25-2 ZK01-25-3 ZK01-26 ZK01-27-1 ZK01-27-2 ZK01-28-1 ZK01-28-2 ZK01-29-1 ZK01-29-2 ZK01-30 ZK01-31-1 ZK01-32

/ / / / / / / / / / 1.78 2.39 3.86 4.52 2.84 2.64 2.36 2.1 0.08 0.77 1.32 3.12 0.01 0.49 0.33 0.05 0.78 1.43 0.04 / /

/ / / / / / / / / / 5.67 5.66 8.33 10.72 5.88 4.92 5.7 6.58 4.56 4.28 6.31 8.56 6.11 4.93 6.3 6.57 6.86 9.43 6.54 / /

30.65 30.3 30.03 29.88 29.81 29.8 29.79 29.78 29.78 29.63 30.77 31.48 26.94 30.53 26.7 26.44 28.3 29.54 29.91 28.1 30.17 29.94 25.44 25.81 23.24 21.98 24.02 24.25 24.33 27.35 26.93

/ / / / / / / / / / 28.99 29.09 23.08 26.01 23.86 23.8 25.94 27.44 29.83 28.87 28.85 26.82 25.45 26.3 22.91 22.03 23.24 22.82 24.37 / /

0.12 0.14 0.10 0.10 0.10 0.07 0.09 0.11 0.08 0.07 0.17 0.06 0.06 0.10 0.05 0.04 0.07 0.02 0.11 0.09 0.13 0.09 0.04 0.04 0.02 0.02 0.02 0.02 0.02 0.03 0.05

4 3.8 3.6 3.4 3.2 3 2.8 2.6 2.4 2 1.6 1.2 0.8 0.4 0.2 0

JJS-1-16 JJS-1-15 JJS-1-14 JJS-1-13 JJS-1-12 JJS-1-11 JJS-1-10 JJS-1-9 JJS-1-8 JJS-1-7 JJS-1-6 JJS-1-5 JJS-1-4 JJS-1-3 JJS-1-2 JJS-1-1

3.23 1.44 0.61 0.59 0.22 0.76 0.05 0.05 0.22 1.77 0.17 1.38 0.98 0.57 2.34 2.32

4.19 6.16 5.93 5.35 4.54 5.69 7.33 7.16 4.44 13.69 11.83 9.48 12.77 10.13 13.2 8.12

29.28 28.43 28.43 28.16 27.98 27.84 27.80 27.33 26.12 25.87 25.75 25.35 25.32 24.56 24.15 22.10

30.56 29.24 28.59 29.02 26.34 26.51 24.61 25.30 29.50 20.33 23.98 24.49 26.86 27.59 22.98 26.11

0.05 0.08 0.07 0.13 0.12 0.36 0.07 0.18 0.70 0.05 0.06 0.09 0.09 0.05 0.09 0.13

900.76 899.83 899.12 897.83 897.02 896.22 895.22 893.93 892.71 891.8 891 890.11 889.17 888.32

ZK4207-2 ZK4207-4 ZK4207-6 ZK4207-8 ZK4207-10 ZK4207-12 ZK4207-14 ZK4207-16 ZK4207-18 ZK4207-20 ZK4207-22 ZK4207-24 ZK4207-26 ZK4207-28

6.90 6.61 6.69 6.25 5.99 6.05 6.10 6.74 6.74 6.69 6.65 6.04 5.92 5.79

7.774 9.752 8.537 9.516 9.662 9.922 9.978 10.381 10.192 9.511 9.403 9.057 9.206 11.18

32.84 32.56 32.43 32.08 32.12 31.82 31.84 32.65 32.94 33.09 32.32 31.87 31.81 31.76

25.95 25.95 25.74 25.83 26.13 25.77 25.74 25.91 26.20 26.40 25.67 25.83 25.89 25.98

2.78 3.46 1.98 1.98 2.29 1.64 1.74 2.19 2.22 1.66 2.17 2.72 2.53 1.91

6. Discussion 6.1. Geochronological constraints on cap carbonate age Recent geochronological data constrain the age of termination of the Sturtian glaciation. A SIMS U-Pb zircon age of 667 ± 5 Ma was generated for a reworked ashfall layer in the upper Scout Mountain Member of the Pocatello Formation, southern Idaho, setting a minimum age for the upper diamictite and cap carbonate at that locale (Fanning and Link, 2004). Another SIMS U-Pb zircon age of 667.3 ± 9.9 Ma was yielded by a volcanic tuff within the basal Datangpo Formation at Heishuixi, Guizhou Province, South China (Yin et al., 2006). A TIMS U-Pb zircon age of 662.9 ± 4.3 Ma was obtained from a volcanic tuff within the basal Mn layer of the Datangpo Formation at Zhailanggou, eastern Guizhou Province, South China (Zhou et al., 2004). A re-Os age of 662.4 ± 3.9 Ma was acquired from black shale of the Twitya Formation that immediately overlies the Sturtian cap carbonate of the Rapitan Group, Mackenzie Mountains, Canada (Rooney et al., 2014). A re-Os age of 659.0 ± 4.5 Ma was generated for the basal Taishir Formation, which immediately overlies diamictite of the Maikhan-Uul Formation and its cap carbonate in Mongolia (Rooney et al., 2015). A reOs date of 657.2 ± 5.4 Ma was acquired from the basal black shale of the Aralka Formation, which directly overlies Areyonga Formation glacial deposits in the Amadeus Basin, Australia (Kendall et al., 2006). The LA-ICP-MS U-Pb zircon age of 662.7 ± 6.2 Ma from the present study is in line with the geochronological results above, especially the U-Pb dates, which are on average
5 Myr older than the re-Os dates. Considering that TIMS is a more reliable geochronological method with better accuracy and properly propagated uncertainties, within the uncertainties of the existing geochronological data the termination of the Sturtian glaciation was a globally synchronous event, dating to
660 Ma.

Although we obtained a LA-ICP-MS U-Pb zircon age with a small internal error (1%), larger uncertainties are associated with interlaboratory reproducibility of results. According to Li et al. (2015), the uncertainty of LA-ICP-MS zircon U-Pb dates is ±2–4% (2r) for analyses of the same sample in different laboratories, which is equivalent to ± 13.3–26.6 Myr in the age of the tuff dated in the present study. However, stratigraphic and geochronologic data in the study area provide complementary evidence for the depositional age of the cap carbonate deposits, which are located just above known Sturtian diamictites and below or correlative with Mn-carbonate deposits at the base of the Datangpo Formation. In other sections within the study area, volcanic tuffs located 3 to 5 m above the contact of the Tiesi’ao Formation diamictite and the overlying Mn-deposits of the basal Datangpo Formation have yielded consistent U-Pb ages of 662.9 ± 4.3 Ma (Zhailanggou section, TIMS age) and 667.3 ± 9.9 Ma (Heishuixi section, SIMS age) (Zhou et al., 2004; Yin et al., 2006). These radiometric dates and stratigraphic relationships collectively constrain the depositional age of the basal Datangpo Formation cap carbonate to
670 ± 10 Ma. 6.2. Water-depth dependence of carbon isotope variation Carbonate d13C profiles can provide key information about the nature of processes that operated to generate Neoproterozoic syn-deglacial cap carbonates. The positive d13Ccarb excursion in cap carbonates immediately above Sturtian glacial deposits is a globally synchronous signal (Rose and Maloof, 2010) (Fig. 12). In Mongolia (Siberian Craton), cap carbonate of the Tsagaan Oloom Formation (Tayshir member) records a d13Ccarb shift from 2.1 or 1.5 to +0.5 or +3‰ over a 7- to 20-m interval (Macdonald, 2011; Johnston et al., 2012). In northwestern Canada (Laurentian Craton), cap carbonate at the base of the Twitya For-

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Fig. 11. Carbonate d13C profiles for the 1st Member of the Datangpo Formation. (A) JJS-1, (B) ZK01, (C) ZK 4207, (D) Zhailanggou (data from Chen et al., 2008), and (E) Datangpo (data from Zhou and Du, 2012).

mation records a d13Ccarb shift from 2.31 or 2.5 to 0.96 to 0.5‰ over 33 m (Johnston et al., 2012; Hoffman and Schrag, 2002). In Arctic Alaska (Laurentian), Sturtian cap carbonate in the Katakturuk Dolomite records a d13Ccarb shift from 0.8 or +2‰ to +1.8 to +4‰ in 5 m (Macdonald et al., 2009). In Namibia (Congo Craton), cap carbonate in the Rasthof Formation records a d13Ccarb shift from ca. 5 to +3‰ over 20 m, with a further positive shift in strata overlying the cap carbonate (Halverson et al., 2002). In southern Australia (Australian Craton), cap carbonate below the Tindelpina shale records a d13Ccarb shift from ca. 4– 0‰ (Giddings and Wallace, 2009). In Brazil (Sao Francisco Craton), cap carbonate of the Sete Lagoas Formation records a d13Ccarb shift from 4.5 to +1.3‰ over
20 m (Vieira et al., 2007). Thus, Sturtian cap carbonates globally are characterized by positive d13Ccarb shifts averaging
5‰, from ca. 5 to 2‰ at the base to ca. 0 to +3‰ at the top of the cap carbonate, although with variation of several per mille between sections. The d13Ccarb records of Sturtian cap carbonates at JJS-1, ZK01, ZK4207, Zhailanggou, and Datangpo show both similarities to and differences from these global d13Ccarb records (Fig. 11). All five of the Nanhua Basin locales show upsection shifts toward higher d13Ccarb values, but the magnitude of the shifts ranges from
1 to 5‰. Significantly, there is a pronounced bimodality in the ranges of observed d13Ccarb values: the shallow sections (JJS-1 and ZK01) show a heavier range (i.e., ca. 4 or 2‰–0 or +2 ‰) whereas the deep sections (ZK4207, Zhailanggou, and Datangpo) exhibit a lighter range (from ca. 10 or 7‰ to 6 ‰; Fig. 11). The shallow sections of the Nanhua Basin thus yield d13Ccarb ranges that conform to those of Sturtian cap carbonates globally, whereas the deep sections yield unusually 13Cdepleted C-isotopic compositions. Profiles of d13Corg also show depth-related differences in the Nanhua Basin, with heavier and more variable values in the shallow sections (ca. 31 to 22‰) relative to the deep sections (ca. 33 to 32‰; Fig. 11). On the other hand, average

D13Ccarb-org values are similar for the shallow sections (ca. 26 ± 4‰ for JJS-1 and ZK01) and deep sections (
26 ± 0.5‰ at ZK4207, and
25 ± 1‰ at Zhailanggou), although the former show considerably greater variability than the latter (Fig. 11). These patterns indicate that water depth was an important control on both d13Ccarb and d13Corg records, implying a stratified and poorly mixed ocean during the Sturtian deglaciation (at least in South China). The depth-related gradient in d13C within the Nanhua Basin has several potential origins. First, a large vertical gradient in d13CDIC may have developed owing to a combination of high surfacewater productivity, an active biological pump, and strong watercolumn stratification (Jiang et al., 2007; Li et al., 2017), a mechanism that yields gradients of 7–20‰ in the modern Black Sea and Framvaren Fjord (Song et al., 2013). This mechanism is consistent with the high TOC contents of Mn-carbonate deposits and black shales of the lower Datangpo Formation, which may reflect elevated surface-water productivity and organic carbon sinking fluxes during the Sturtian deglaciation. Related to this mechanism, a coupled oxidation of organic matter and reduction of Mn-oxides (Eq. (1)) would have generated 13C-depleted dissolved inorganic carbon (DIC) (to ca. 12‰; Maynard, 2003; Neumeister et al., 2015). Large amounts of low-d13C Mn-carbonates (rhodochrosite) in the deepwater sections of the Nanhua Basin are consistent with this mechanism (Yu et al., 2016). A second possibility is influx of 13C-depleted dissolved organic carbon (DOC) from the deep ocean into the Nanhua Basin. Decoupling of carbonate and organic carbon d13C records has been cited as evidence of such a global-ocean DOC pool (Swanson-Hysell et al., 2010). However, the carbonate and organic carbon d13C records of the deepwater Mn-carbonate deposits of the basal Datangpo Formation show strong positive covariation (r = +0.91, n = 14, p(a) < 0.01) (Figs. 11, S3B). Although there might be other factors influencing such coupling, the paired carbonate and organic carbon d13C records of Nanhua Basin sections appear to provide no support for this mechanism.

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Fig. 12. Paleogeographic reconstruction of Cryogenian world (Li et al., 2013) and distribution of Sturtian cap carbonate deposits, the d13Ccarb change, thickness and data source for each site are show in the labels.

6.3. Relationship of shallow cap-carbonate dolomites to deep Mncarbonates The newly discovered shallow dolomitic cap carbonates of the Nanhua Basin are considered to be approximately correlative with the previously studied syn-deglacial Mn-carbonates of the deeper (basinal) facies (Yu et al., 2016) based on both field relationships and C-isotope records. Stratigraphically, all of these deposits are found sandwiched between glacial diamictites of the underlying Tiesi’ao Formation and black shale or siltstone of the overlying Datangpo Formation across the study area (Fig. 4). With regard to upsection trends in their C-isotope records, most (4 of 5) shallow and deep locales exhibit similar-magnitude positive shifts in d13Ccarb (3–5‰), suggesting a response to a common forcing. The more limited increase in d13Ccarb upsection at ZK4207 (
1‰) may indicate that this section represents only a fraction of the time of the other four sections, or possibly that seawater d13CDIC evolved to a lesser degree owing to semi-restriction of the deep watermass of this graben site. Despite probable coeval accumulation, the Sturtian capcarbonate dolomites on horsts of the Nanhua Basin and the Mn-carbonates concentrated in adjacent grabens are heterogeneous deposits having different formation mechanisms. Shallow cap carbonates are thought to have formed due to deglacial increases in the carbonate alkalinity of seawater triggered by an enhanced flux of finely weathered carbonate from continents to the ocean (Hoffman and Schrag, 2000) or gas hydrate destabilization (Kennedy et al., 2001, 2008). In contrast, the Mncarbonates in grabens are likely to have formed as a consequence of increasing sediment porewater alkalinity triggered by Mn, Fe, and sulfate reduction during early diagenesis (Yu et al., 2016). This process involved a coupled oxidation of organic matter and reduction of Mn-oxides (Maynard, 2003; Neumeister et al., 2015):

2MnO2 þ CH2 O þ HCO3 $ 2MnCO3 þ H2 O þ OH

ð1Þ

Per this reaction, about 50% of the carbon in Mn-carbonates is derived via oxidation of organic matter in the sediment. Formation of large amounts of low-d13C Mn-carbonates (rhodochrosite) in deepwater sections of the Nanhua Basin is consistent with this mechanism. The post-glacial formation of rhodochrosite (Mn-carbonates) in grabens of the Nanhua Basin was explained via an episodic ventilation model by Yu et al. (2016). In this model, hydrothermally sourced Mn accumulated in the anoxic watermass during the Sturtian glacial interval. After deglaciation, a redox-stratified watermass developed, leading to precipitation of Mn-oxides in grabens through episodic ventilation by density flows. Co-burial of the Mn oxides with organic-rich sediments resulted in concurrent organic matter oxidation and precipitation of secondary Mncarbonates owing to elevated porewater Mn2+ concentrations and alkalinity generated through microbial sulfate reduction and microbially mediated Mn reduction (cf. Polgári et al., 2012, 2016). Much of the microbial decay of organic matter and formation of Mn-carbonates is likely to have occurred close to the sediment-water interface, i.e., at an early diagenetic stage. This hypothesis has the advantage that, if local redox variation was the primary control on manganese accumulation, the diachroneity of Mn-carbonate deposits within the Nanhua Basin can be explained (Fig. 3). 6.4. Models of Neoproterozoic cap carbonate genesis At least four models have been proposed to account for formation of syn-deglacial cap carbonates during the Neoproterozoic: (1) the Snowball Earth hypothesis, (2) the ‘plume-world’ hypothesis, (3) the gas hydrate destabilization hypothesis, and (4) the sediment starvation hypothesis. These hypotheses yield testable predictions that may help to identify which model(s) are most robust. The Snowball Earth hypothesis ascribes cap carbonates to large inputs of carbonate alkalinity to the oceans during deglaciation of a heavily glaciated Earth (Hoffman et al., 1998; Hoffman and Schrag,

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2002). In this model, atmospheric pCO2 rose to high levels during glacial intervals because of the near-shutdown of the global hydrologic and silicate weathering cycles. During deglaciation, a rapid disappearance of sea ice combined with higher temperatures and increased acid rainfall triggered elevated rates of both carbonate and silicate weathering, increasing the alkalinity of seawater. Reduced productivity in the icebound ocean during the glaciation led to the initially low d13Ccarb at the base of the cap carbonate deposits, but increasing productivity during the deglacial stage contributed to rising d13Ccarb (Hoffman and Schrag, 2000; Higgins and Schrag, 2003). Given large additions of alkalinity through enhanced weathering, seawater d13CDIC would have been relatively well buffered and capable of only slow change. The plume-world hypothesis ascribes cap carbonates to rapid precipitation (i.e., over a few thousand years) from low-salinity ocean-surface waters generated through glacial melting (Shields, 2005). In this hypothesis, carbonate alkalinity in the oceansurface layer increased due to weak pH buffering capacity, limited ionic complexation and increased ionic activity, leading to large and rapid increases in carbonate saturation levels. In essence, this mechanism is similar to that of the Snowball Earth hypothesis, although the rate of cap carbonate formation in shallow facies is accelerated greatly. The gas hydrate destabilization hypothesis invokes melting of methane clathrates during deglacial warming, followed by anaerobic oxidation of methane (AOM) in conjunction with microbial sulfate reduction (Kennedy et al., 2001, 2008). This process can cause an increased flux of carbonate alkalinity to the oceans, potentially triggering cap carbonate formation. In this case, the low values at the base of the cap carbonate represent peak inputs of DIC derived from 13C-depleted methane, and the positive shifts upsection represent waning influence of this source of DIC. The extremely negative d13Ccarb values (to 45‰) of cements in sheet cracks, fractures, and cavities in basal cap carbonates and some specific sedimentary structures (e.g., domal or tepee-like and tube-like structures) in Marinoan cap carbonate successions are seen as evidence of methane seep events (Wang et al., 2008a). The sediment starvation hypothesis, which was recently proposed by Spence et al. (2016), considers cap carbonates to be a condensed deposit formed during deglacial transgressions (cf. Nummedal and Swift, 1987; Wignall, 1991). In this scenario, the duration of the deglaciation and accompanying cap carbonate deposition were on a timeframe of
105–106 yr, rather than
104 yr as in other scenarios (Spence et al., 2016). Although there are few chronological constraints on the duration of cap carbonate formation at present, a timeframe of a million years for Neoproterozoic deglaciations is fundamentally improbable given that deglaciation rates are typically rapid owing to forcing by strong positive climate feedbacks (e.g., the ice-albedo feedback; Alley et al., 2003). Furthermore, at longer timescales orbital influences on climate (e.g., the ca. 20-kyr precession or 100- or 413-kyr eccentricity cycles) should have left some imprint on cap carbonate records (e.g., a cyclicity in d13Ccarb; cf. Sageman et al., 2006), of which no observations have been reported to date for Neoproterozoic cap carbonates (including in the present study). The present study sections appear to be records of a single transgression and maximum flooding surface (msf) preserved in the black shale overlying the cap carbonate deposits (Fig. 4). We therefore view skeptically this proposed mechanism of cap carbonate formation. The carbonate and organic d13C data of the study sections may be useful in evaluating the four hypotheses above. The plumeworld hypothesis makes one prediction that is consistent with our observations for Sturtian cap carbonates in South China, i.e., greater d13C variability in shallow sections relative to deep sections, owing to more limited buffering of the former. However, other predictions of this mechanism are not borne out, e.g., more

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rapid deposition of cap carbonates in the ocean-surface layer than at depth. In fact, the shallow sections in our study contain thinner cap carbonates (
1.5–4 m) than the deep sections (
6–12 m), despite nearly equal-magnitude shifts in d13Ccarb; these observations suggest that all Sturtian cap carbonates in the Nanhua Basin formed over approximately the same interval of time, and that the deep deposits accumulated at rates
2–8 faster than the shallow deposits. Spatial variation in d13Ccarb and thickness of the Sturtian cap carbonates in South China may favor the gas hydrate destabilization hypothesis, which would have been characterized by spatially irregular methane clathrate distribution. Assuming methane release primarily to deep watermasses, it may also account for shallow-deep differences in cap-carbonate thickness and d13Ccarb values (see Section 6.3). Section ZK01 shows a rapid d13Ccarb decrease (from 2 to 4‰) over
20 cm at the base of the cap carbonate, which is consistent with another prediction of this model. However, most of our d13Ccarb records exhibit continuously increasing d13Ccarb upsection, and typical methane-seep-related structures, e.g., domal or tepee-like structures and tube-like structures, are not present in our sections. In summary, we find that aspects of both the Snowball Earth and gas hydrate destabilization hypotheses may be able to account for the majority of stratigraphic, sedimentary, and isotopic observations in our study sections. Indeed, these models are not necessarily mutually exclusive, and they could have operated in tandem to produce observed patterns of C-isotope variation globally. 6.5. Model for Sturtian cap carbonate Formation in Nanhua Basin Petrographic and geochemical data from the present study sections are evaluated in the context of existing hypotheses to generate an integrated model of Sturtian cap carbonate formation in the Nanhua Basin. At the end of the Sturtian glaciation, stagnated glacial tills were present on horst surfaces, and the tills from melting terrestrial icesheets formed dolomitic diamictite deposits. A marine transgression occurred as coastal ice melted, producing a meltwater layer that covered horst areas (Fig. 13A and B). Following the end of the deglacial transgression, dolomitic diamictite accumulation on horsts yielded to deposition of dolomite with d13C values typical of cap carbonates globally (2 to 4‰) (Fig. 13C). The transition from the dolomitic diamictite to the dolomitic cap carbonate in the Nanhua Basin provides evidence for increasing seawater alkalinity during the deglaciation. Simultaneously, low-d13C (7 to 10‰) authigenic Mn-carbonate deposits formed in the adjacent grabens, probably through decay of large quantities of organic matter under suboxic conditions (Fig. 13C). The relationships discussed in Sections 6.2 and 6.3 suggest that, during the Sturtian deglaciation, the Nanhua Basin was stratified (although subject to episodic deepwater ventilation), deepwaters were dominantly ferruginous (or ‘‘suboxic”, as shown by intense Mn-oxide reduction), and surface-water productivity and organic carbon sinking fluxes were high (Yu et al., 2016; Li et al., 2012). Deepwater d13CDIC was low probably owing to water-column stratification, an active biological pump, and intense organic carbon remineralization (cf. D’Hondt et al., 1998; Song et al., 2013), although methane inputs to the deep watermass are also possible (Kennedy et al., 2001, 2008). Given the probable role of productivity in establishing a vertical d13CDIC gradient in the Nanhua Basin, it is reasonable to interpret secular variation in d13C in the study sections in terms of productivity variation. Seawater d13CDIC is estimated to decline to ca. 5‰ when marine primary productivity ceases (D’Hondt et al., 1998), and collapse of marine productivity during the Snowball Earth interval is regarded as a proximate cause of low d13Ccarb in Neoproterozoic cap carbonates globally (Hoffman and Schrag, 2000, 2002). The observation that rising

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Fig. 13. Model for Sturtian cap carbonate formation in the Nanhua Basin of South China. (A) During the Sturtian glacial interval, exposed horsts were covered with land-based ice, and the glaciomarine Tiesi’ao Formation was deposited in the grabens. (B) During the deglaciation, land-based ice melted, rising sea levels flooded the horsts, a surface meltwater layer formed, and dolomitic diamictite was deposited. (C) In the deglacial interval, a redox-stratified water column formed in the Nanhua Basin, triggering precipitation of cap carbonate on horsts and deposition of Mn-rich deposits in grabens.

d13Ccarb in Sturtian cap carbonates of the Nanhua Basin stabilized around 0 to +3‰, which are typical values for Phanerozoic marine carbonates, suggests that, by the end of cap carbonate formation, the global carbon cycle had reached an equilibrium at levels of marine productivity and continental weathering similar to younger geologic epochs. In this case, there must have been a large increase in productivity between the base of the cap carbonate (d13Ccarb = 2 to 4‰) and its top (d13Ccarb = 0 or +2‰). If operative at a global scale, this process would have induced changes in seawater d13CDIC in both shallow and deep watermasses. In fact, both shallow and deep facies of the Nanhua Basin exhibit similar mean D13Ccarb-org values (
25–26‰) that show no distinct secular trend (despite greater variability in shallow facies; Fig. 11). These patterns are characteristic of productivity control of global seawater d13CDIC (cf. Kump and Arthur, 1999). During the Sturtian deglaciation, marine productivity would have been driven higher by deglaciation of the global-ocean surface and by release of fine nutrients from continental ice masses (Hoffman and Schrag, 2002; Kunzmann et al., 2013). A major increase in marine productivity would have removed CO2 from seawater, driving the carbonate equilibrium toward CaCO3 satura-

tion and precipitation (Zeebe and Wolf-Gladrow, 2001), thus triggering precipitation of cap carbonates, which would have formed mainly in shallow tropical areas where carbonate saturation levels were highest. Although Neoproterozoic paleogeographic reconstructions are tentative (see Supplementary Information), it appears that most or all cap carbonates formed within
50° of the paleo-equator (Fig. 12 and Fig. S5), which is consistent with global-scale changes in seawater alkalinity leading to supersaturation of seawater with respect to carbonates in the tropics and subtropics (which may have occupied a wider latitudinal range in the warm Neoproterozoic world than today). Strictly speaking, shifts toward higher d13Ccarb are due to increased Corg burial (Kump and Arthur, 1999), but this process often correlates with higher productivity. In fact, the Sturtian deglaciation may have promoted both higher productivity and enhanced Corg burial through (1) warming and stratification of the oceans, and (2) transgressive sea-level rise, which is commonly linked to enhanced Corg burial (Wignall, 1991). Lithofacies relationships in the Nanhua Basin are consistent with elevated productivity in the aftermath of the Sturtian glaciation. Organicrich Mn carbonates and black shales with TOC contents of 1.5–

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5.4% are widely distributed in graben areas of the Nanhua Basin and sometimes overlie cap carbonates in the horst areas (e.g., section BP-1) (Yu et al., 2016). Microbial fossils (Yin, 1990) and biomarkers (Wang et al., 2008b) present in these Mn-carbonates and black shales provide direct evidence of intense microbial activity. These observations emphasize the probable role of elevated marine productivity in controlling changes in seawater alkalinity and formation of Sturtian cap carbonates in the Nanhua Basin of South China. 7. Conclusions Sedimentary observations, zircon U-Pb ages, and C isotope data demonstrate that the dolomite at the base of the Cryogenian Datangpo Formation in the Nanhua Basin of South China represents a cap carbonate deposit associated with the end-Sturtian deglaciation. The massive, laminated, pisolitic, and sandy dolomites in the cap carbonate sequence formed in a shallow subtidal setting. Carbon isotopic profiles from the cap carbonate record a large positive d13C excursion following the Sturtian glaciation, probably caused by increases in marine productivity and organic carbon burial rates during the deglacial interval. Dolomitic cap carbonates of the basal Datangpo Formation are found mainly on horst blocks of the Nanhua Basin, and the correlative strata in the adjacent grabens were Mn-carbonate deposits. Whereas dolomitic cap carbonates are primary marine deposits that may have formed due to the increased alkalinity of deglacial seawater, the Mn-carbonate deposits consist of rhodochrosite precipitated during early diagenesis probably in response to the decay of abundant organic matter under suboxic conditions. Acknowledgments We are grateful to Prof. Yuanbao Wu, Prof. Yongsheng Liu and Dr. Hao Deng in CUG for their helps on LA-ICP-MS method explanation. The editor and two anonymous reviewers are thanked for their constructive suggestions which have greatly improved the quality of this paper. This research is supported by China Geological Survey (CGS) ‘‘Geological and Metallogenic Background in the Southeastern Margin of the Upper Yangtze Block (No. 12120114016701)” Ministry of Land and Resources of the People’s Republic of China ‘‘Deep Prospecting and Metallogenic System in the Southeastern Margin of the Upper Yangtze Block (No. 201411051)” Project. Appendix A. Supplementary data Supplementary data associated with this article can be found, in the online version, at http://dx.doi.org/10.1016/j.precamres.2017. 03.011. References Aitken, J., 1991. Two late Proterozoic glaciations, Mackenzie mountains, northwestern Canada. Geology 19, 445–448. Alley, R.B., Marotzke, J., Nordhaus, W.D., Overpeck, J.T., Peteet, D.M., Pielke, R.A., Pierrehumbert, R.T., Rhines, P.B., Stocker, T.F., Talley, L.D., Wallace, J.M., 2003. Abrupt climate change. Science 299 (5615), 2005–2010. Boulton, G.S., Deynoux, M., 1981. Sedimentation in glacial environments and the identification of tills and tillites in ancient sedimentary sequences. Precambrian Res. 15, 397–422. Brasier, M., McCarron, G., Tucker, R., Leather, J., Allen, P., Shields, G., 2000. New U-Pb zircon dates for the Neoproterozoic Ghubrah glaciation and for the top of the Huqf Supergroup, Oman. Geology 28, 175–178. Bureau of Geology and Mineral Resources of Guangxi Zhuang Autonomous Region (BGMRGZAR), 1985. Regional Geology of Guangxi Zhuang Autonomous Region. Geological Published House, Beijing, China, pp. 23–37 (in Chinese with English abstract).

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